Handbook of Stable Isotope Analytical Techniques

HANDBOOK OF STABLE ISOTOPE ANALYTICAL TECHNIQUES Front Cover: Cover photograph shows a multiple collector inductively...

Author: Pier A. de Groot

123 downloads 2128 Views 104MB Size Report

This content was uploaded by our users and we assume good faith they have the permission to share this book. If you own the copyright to this book and it is wrongfully on our website, we offer a simple DMCA procedure to remove your content from our site. Start by pressing the button below!
Report copyright / DMCA form

DOWNLOAD PDF

HANDBOOK OF STABLE ISOTOPE ANALYTICAL TECHNIQUES

Front Cover: Cover photograph shows a multiple collector inductively coupled plasma mass spectrometer (ICP-MS). Copyright of photo by IRMM, Retieseweg, 2440 Geel, Belgium. Publication kindly permitted by IRMM and acknowledged by the editor.

HANDBOOK OF STABLE ISOTOPE ANALYTICAL TECHNIQUES VOLUME 1

Pier A. de Groot editor Economic Geology Research Institute, School of Geosciences, University of the Witwatersrand, Johannesburg, South Africa and Delta Isotopes Consultancy, Pastoor Moorkensstraat 16, 2400 Mol-Achterbos, Belgium (present address)

2004

ELSEVIER Amsterdam - Boston - Heidelberg - L o n d o n - New York - Oxford Paris - San Diego - San Francisco

- Singapore

- Sydney

- Tokyo

ELSEVIER B.V. Sara Burgerhartstraat 25 P.O. Box 211,1000 AE Amsterdam, The Netherlands

ELSEVIER Inc. 525 B Street, Suite 1900 San Diego, CA 92101-4495 USA

ELSEVIER Ltd The Boulevard, Langford Lane Kidlington, Oxford OX5 1GB UK

ELSEVIER Ltd 84 Theobalds Road London WC1X 8RR UK

9 2004 Elsevier B.V All rights reserved. This work is protected under copyright by Elsevier B.V, and the following terms and conditions apply to its use: Photocopying Single photocopies of single chapters may be made for personal use as allowed by national copyright laws. Permission of the Publisher and payment of a fee is required for all other photocopying, including multiple or systematic copying, copying for advertising or promotional purposes, resale, and all forms of document delivery. Special rates are available for educational institutions that wish to make photocopies for non-profit educational classroom use. Permissions may be sought directly from Elsevier's Rights Department in Oxford, UK: phone (+44) 1865 843830, fax (+44) 1865 853333, e-mail: [email protected]. Requests may also be completed on-line via the Elsevier homepage (http://www.elsevier.com/locate/ permissions). In the USA, users may clear permissions and make payments through the Copyright Clearance Center, Inc., 222 Rosewood Drive, Danvers, MA 01923, USA; phone: (+ 1) (978) 7508400, fax: (+ 1) (978) 7504744, and in the UK through the Copyright Licensing Agency Rapid Clearance Service (CLARCS), 90 Tottenham Court Road, London W1P 0LP, UK; phone: (+44) 20 7631 5555; fax: (+44) 20 7631 5500. Other countries may have a local reprographic rights agency for payments. Derivative Works Tables of contents may be reproduced for internal circulation, but permission of the Publisher is required for external resale or distribution of such material. Permission of the Publisher is required for all other derivative works, including compilations and translations. Electronic Storage or Usage Permission of the Publisher is required to store or use electronically any material contained in this work, including any chapter or part of a chapter. Except as outlined above, no part of this work may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without prior written permission of the Publisher. Address permissions requests to: Elsevier's Rights Department, at the fax and e-mail addresses noted above. Notice No responsibility is assumed by the Publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. Because of rapid advances in the medical sciences, in particular, independent verification of diagnoses and drug dosages should be made. First edition 2004 Reprinted 2005

Library of Congress Cataloging in Publication Data A catalog record is available from the Library of Congress. British Library Cataloguing in Publication Data A catalogue record is available from the British Library. ISBN: 0 444 51114 8

Working together to grow libraries in developing countries www.elsevier.com I www.bookaid.org www.sabre.org : L S E V I ER

~ooK 1...... "..... iAID ...... 1

~abrc" l - ~"O t l l ~ d ~ t

"

101

@ The paper used in this publication meets the requirements of ANSI/NISO Z39.48-1992 (Permanence of Paper). Printed in The Netherlands.

Those incredible deltas

They measured day and night On their super isotope machine And ended that analytical fight With zero errors, as you imagine Then ... they had actually begun Without switching the bloody thing on. C. Brenninkmeijer

This Page Intentionally Left Blank

Dedication to S.M.F. Sheppard

Simon M. F. Sheppard

viii


Dedication The two volumes on Stable Isotope Techniques are dedicated to my colleague and former tutor Simon M.F. Sheppard. He was the person who introduced me, during my study period in the Centre de Recherches P6trographiques et G6ochimiques, at Vandoeuvre-l~s-Nancy, France, to the wonderful world of stable isotopes applied to earth science, especially in the field of geochemistry and mineral deposits. It was, on reflection a critical period in developing my career, motivating me to understand the full ramifications of stable isotope chemistry. Since that time my interest in stable isotope chemistry has constantly progressed. If Simon had not stimulated me to work in this field, I would not have organised this publication presenting the essential guidelines of analytical procedures and techniques for the measurement of stable isotope ratios in samples for a large number of scientific disciplines. Simon Sheppard, born April 16, 1938, at Salisbury, England, was educated at the University of Cambridge, England, where he gained a BA Tripos in Natural Sciences in 1962. His first studies in isotope geochemistry began at McMaster University in Canada, under the guidance of Henri Schwarcz gaining his PhD in 1966 on "Stable isotope (C, O) geochemistry of metamorphic rocks". From 1966 to 1968 Simon was Research Fellow at the California Institute of Technolog~ Pasadena, U.S.A., where he first worked with Samuel Epstein and Hugh Taylor Jr. Here he completed two important publications on porphyry copper mineralization linked to fluid-rock interactions (Sheppard et al., 1969, 1971). Simon was then appointed to an Assitent Professorship from 1968 until 1970, at the University of Texas, Austin, U.S.A. He then returned to the Uinted Kingdom to develop stable isotope research facilities at the Scottish Universities Research and Reactor Centre (SURRC), East Kilbride, Scotland. In 1976 Simon was invited to the Centre de Recherches P6trographiques et G6ochimiques (CRPG), Vandoeuvre-16s-Nancy, in France, by Bernard Poty with the intention of coupling fluid inclusion research with stable isotope systematics. This scientific challenge could not be resisted and so Simon left SURRC to join CRPG sponsored by CNRS, the French national research organization, ultimately becoming the 'Directeur' of CRPG in 1980 for ten years. In 1991 Simon was appointed to the newly founded l~cole Normale Sup6rieure (ENS) in Lyon, France, where he continued his stable isotope research work until his recent retirement in September 2003. During his scientific career, Simon Sheppard has become highly regarded as a stable isotope geochemist in Europe and worldwide. His main area of interest is in the origin and transfer of aqueous, carbonic, sulfurous and silicate fluids in the mantle, crust and hydrosphere. Related to this theme is the use of fluid inclusion analytical methods pioneered by Alain Weisbrod and Bernard Poty at CRPG. Simon always has been fascinated by the application of stable isotope studies to economic geolog~ particularly material transfer by fluids and stable isotope exchange processes (fluid -


ix

mineral interactions). His extensive experience in different scientific areas is also shown by his contribution of two chapters in Volume L Part 1, Review and Discussion on Developments in Stable Isotope Analytical Technologies, where Simon is the co-author of Chapter 2, on "Analysis of Fluids from Clays and Sediments", and author of Chapter 4:6, on "The Experimental Determination of Isotopic Fractionations'. This is the first compilation on experimental methods on this subject to be published. Besides possessing a substantial list of frequently cited publications Simon has been Associate Editor for Precambrian Research (1977- 1989), Geochimica Cosmochimica Acta (1979 - 1985), and is still active as Associate Editor for Lithos (from 1984) and Chemical Geology (from 1991). During my contact with Simon over the years, I learned that Simon is an 'amateur expert' in architecture. I well remember his request to see the 'Rietveld - house' in Utrecht, the Netherlands (now protected by UNESCO as world cutural heritage site). Living in France Simon enjoys fine wine, and has developed a curiosity for goat cheese. Currently he is writing a book on French cheeses which should be published in the near future.

References Sheppard S. M. F., Nielsen R. L. & Taylor H. P. Jr. (1969) Oxygen and hydrogen isotope ratio of clay minerals from porphyry copper deposits. Econ. Geol., 64: 755-777.

Sheppard S. M. F., Nielsen R. L. & Taylor H. P. Jr. (1971) Hydrogen and oxygen isotope ratios in minerals from porphyry copper deposits. Econ. Geol., 66: 515-542.


Introduction

INTRODUCTION

The honour for the creation of this two volume book series on stable isotope analytical techniques must be given to James O'Neil. At the very start of this project I did not realize that he, together with Zachary Sharp, had begun writing a book on a very similar subject. Originally James O'Neil and Zachary Sharp were among the first authors I invited to write a chapter for my book series. It was in this way that I discovered about their own book project, which already had been in progress for some time. It was Zachary Sharp in particular who convinced and motivated me, supported by James O'Neil, to continue my project, after he recognized that the books I had in mind had a different approach and aimed at a far larger range of elements (isotopes) than their own. The prospective publication by James O'Neil and Zachary Sharp will have the form of a textbook and is concentrated more on a purely geochemical field. I am grateful for their "open-minded attitude". At the start, the aim of these books was to concentrate on stable isotope analytical methods of purely geochemical interest, but very soon it became clear it was not easy to draw a line between purely geochemical methods with geological applications and methods used in other fields of science. Considering the interest isotope chemists working in other disciplines could have in a handbook on analytical techniques made me decide to extent the contents of these books, and to include a wide range of other disciplines, where stable isotope analysis is used for different purposes. Disciplines of interest, besides geochemistry, are: anthropology, archaeology, agronom~ atmospheric science, biology, bio(geo)chemistr~ climatology, drug detection methodology, ecology~ environmentology, food science or alimentology (e.g. detection of adulterat-ion), forensic science, hydrolog}~ marine science, medical science, metallurgy~ meteoritic science, metrolog~ nutrition studies, palaeontolog~ petrochemistry, pharmacology, planetary science, and toxicology. The range of elements chosen is purely arbitrary but dependent on the choice of the invited authors. Since new methods were actually developed for Br isotope analysis, by analogy with C1 isotope methods, I decided the element of Br to be the upper limit for these books. This also considerably extended the range of basic techniques. For instance, stable isotope analysis of elements such as Li, B, Mg, Si, K, Ca, Ti, Cr, Fe, Ni, Cu, Zn, Ga, Ge, and to some extent Se, concentrated historically more on solid source mass spectrometry rather then gas source mass spectrometry. Other tools for isotope measurement cross this classical boundary between "solid source mass spectrometry (MS) - gas source MS", such as secondary ionization mass spectrometry (SIMS), inductively coupled plasma mass spectrometry (ICP-MS), laser related mass spectrometry systems, fast atom bombardment mass spectrometry (FAB-MS), or nuclear activation analysis (NAA) techniques. Far less common techniques for stable isotope analysis, such as glow discharge mass spectrometry (GDMS), accelerator mass

xii

Introduction

spectrometry (AMS), and some forms of optical spectrometry, are developed and some already are commonly used, while others are at different stages of development. For example, optical analytical methods were developed for gas samples as a faster but less precise method, parallel with the mass spectrometry technique. Stable isotope analytical methods were developed soon after the discovery of the existence of isotopes. For example, in 1934 there was the discovery of deuterium by Harold Urey for which he received the Nobel prize. Early techniques were based on determination of isotopic ratios by densimetric, gravity, electric resistivity, pycnometry type of methods. The development of a usable mass spectrometer, by Nier, and improved by McKinney and coworkers in the late 1940's - early 1950's, gave an important impuls for the use of stable isotope techniques in scientific studies. These early methods were generally complicated, time consuming procedures, and had relatively low precision and accuracy. First improvements were on precision of methods, and soon attempts were made to simplify preparation procedures. In early methods sample size was comparatively large, in the order of a few to 10's or 100's of mgs. Decrease of sample size was another aim, while improving analytical methods. Accuracy of methods is controlled by certified standard or reference materials (see Volume L Part 2, Chapter 40). Standardization of techniques and correct calibration methods are an important concern and need continuous attention to avoid comparison of isotopic values based on different or badly calibrated scales. Reduction of sample size was made possible first by use of so-called "static mass spectrometers" with a single inlet, thus avoiding the continuous pumping of samples while not being actively measured and basically reducing the needed gas volume by half. This technique is still in use at present, for example in stepped heating procedures (Volume L Part 1, Chapter 13) or in fluid inclusion analysis techniques in geological, geochemical, meteoritic or planetary studies. The development of secondary ionization mass spectrometry (SIMS) and laser techniques introduced the possibility of analysing in situ for specific type of samples. The first generation SIMS were limited in the elements on which stable isotopes could be measured caused by low mass resolution characteristics and precision was comparatively low. High resolution SIMS has recently been developed (e.g SHRIMP-II and Cameca ims-1270), increasing the number of isotopes which can be measured, and with improved precision compares with the first generation machines (see Volume L Part 1, Chapter 30). The introduction of laser technology, after the pioneering work by Ian Franchi, Douglas Rumble or Zachary Sharp for stable isotope measurement decreased the sample size considerably. Techniques for analyzing powdered or grain size samples or analyzing on a microscopic scale in an in situ mode are available with their own specific characteristics and limitations. Single grain or spot analysis inside grains is made possible with the laser and the SIMS techniques. First laser types used were infra-red (IR) Nd:YAG and CO2 lasers. Newer developments are with ultra-violet (UV) lasers such

Introduction

xiii

as excimer lasers, quadrupoled Nd-YAG lasers, Ar-F or Kr-F2-Ne or Xe-C1 gas mixture lasers, and doubled frequency Cu-vapour lasers (Volume L Part 1, Chapter 20 or Volume II, Chapter 6-1.5.1). The development for measuring organic, fluid or solid samples is moving a different wa N into on-line systems (originally a converted elemental analyser (EA) was used) with oxidizing, reducing or pyrolyses reactors, eventually in a combined order depending on the sample material and the gas of interest for isotopic measurement, and with application of a carrier gas (generally He, seldomly H2 or N2 are considered; Ar may introduce problems in the ion source of a MS by sputtering effects) in a socalled continuous flow (CF) system to transport reaction gases through the system. The carrier gas may contain 02 for oxidation purpose in combustion (oxidizing) systems. Addition of gas chromatographs (GC's) for separation or purification of sample materials a n d / o r for purification of effluent gases after reaction in the EA section of these on-line systems became a common feature (see alsoVolume I, Part 1, Chapter 8). The advantage of such techniques is the very small sample size needed, the high number of samples that can be analyzed in short time periods compared with classical methods, the possibility of automation of these systems, reducing labour intensity (and thus costs) for analyzing, and the option to combine the measurement of different effluent gases for different isotopic ratios in consecutive way. Moreover, with application of CF-IRMS techniques there is no need for vacuum conditions, as was the case in precursory techniques. Increasingly, special designed EA's are used in on-line systems for analysis of organic materials and fluids. New developments also include inorganic materials such as" phosphates, sulfates, and nitrates for oxygen, sulfur and/ or nitrogen isotopes. Automation, as mentioned earlier, is another trend in analytical techniques. A large group of materials are suitable for such automated systems (e.g. organic materials, carbonates, water or fluid samples) while other materials are not suitable (e.g. rock or mineral samples for O- and H-isotope analysis, including fluorination systems). Another trend is to combine measurement of several isotopes in a sample in an online system. Organic matter is suitable for such an approach, and combinations of some of H-, 0-, C-, S-, and N-isotopes can be applied. This places some constraints on the MS-side in such systems, where an MS must be able to jump from one m/z ratio to another quickly and to handle the measured peaks for samples and references, including background and other corrections by advanced computer automation. In the fast development of laser technology, tunable diode lasers must be mentioned for possible application on a number of materials by optical spectroscopic methods as a highly probable technique in the future (see also Volume L Part 1, Chapters 33 and 34 for optical systematics). Other analytical techniques and tools have been developed but details of all these developments are not given here in this introduction. As presented in the foregoing

xiv

Introduction

sections, the modern tendency is to analyse on smaller samples (including in situ analysis), preferably for a major part automated, and faster analysis per sample. Analysis of a number of different isotopic ratios on the same sample, generally in a sequential way, are applied either on separated phases in effluent gases produced from samples in reactors, or by sequential analysis of separated sample compounds, or by subsequent handling of the same sample with different treatments (e.g. gas equilibration methods). This handbook consist of two volumes: The first, edited volume contains two parts. Part I includes contributions presenting 'subjective' reviews on analytical techniques for specific stable isotopes or materials, reviews on stable isotope analysis by selected machines, descriptions of specialized and novel methods in stable isotope analytical techniques. Readers are guided to modern analytical techniques and are advised which techniques are the best to use for specific materials or conditions. Part 2 includes matters that are not strictly confined to analytical techniques themselves but related to analysis of stable isotopes, such as" views on the development of mass spectrometers and ion source stability, matters concerning isotopic scales, standards and reference materials, calibration and correction matters, a review on experimental isotopic fractionation determination and directives for setting up a laboratory. Appendices present the internet-based stable isotope discussion list named: 'isogeochem list', the also internet-based stable isotope fractionation calculator, and information on suppliers of stable isotope reference materials The second volume aims to present an encyclopedic overview of stable isotope analytical techniques in an 'objective' way. The chapters in this volume are intended to be complementary to the chapters in the first volume. In the second volume analytical techniques from historical times up to the most recent developments, are presented as a classical order of elements. Short descriptions of methods and diagrams of analytical devices are presented. Many classical techniques, of which several were never used in an operational form or became obsolete or forgotten, are included. Many of these older techniques formed the basis for present-day techniques. They also may improve the understanding of the development of analytical techniques which are used in preference today. Much of the experience from the old technology can be useful in applying to, or in constructing modern analytical systems.

Acknowledgement

I like to express my gratitude to all who were helpful to me during the preparation of these two book volumes. The list is far too long to mention everyone, but I hope that I will be forgiven if their names are not specifically highlighted. All help was important to me, unrelated to the size of this help. Finally, I like to thank all the reviewers, whose comments and suggestions considerably improved the quality of this publication.

xv LIST OF C O N T R I B U T O R S J.K. Aggarwal E.A. Atekwana G. Beaudoin G.E. Bebout M. Berglund S. Borella S.R. Boyd M.E. Bi~ttcher W.A. Brand T.D. Bullen K.L. Casciotti L.-H. Chan

C.C.Y. Chang M. Coleman

L. Dallai P. De Bi6vre J. Diemer T. Ding T. Durakiewicz H.G.M. Eggenkamp G.D. Farquhar

D.F. Ferretti M.L. Fogel K.S. Gan

Institute of Mineralogy, Mtinster University, Corrensstrasse 24, D-48149 M/~nster, Germany Department of Geology, Indiana University Purdue University, 723 W. Michigan Street, SL 122, Indianapolis, IN 46202-5132, USA D6partement de G6ologie et de G6nie G6ologique, Universit6 Laval, Qu6bec, GIK 7P4, Canada Department of Earth & Environmental Sciences, Lehigh University, Bethlehem, Pennsylvania 18015, USA European Commission- Joint Research Centre, Institute for Reference Materials and Measurements (IRMM), Retieseweg, 2440 Geel, Belgium Climate and Environmental Physics, Physics Institute, University of Bern, Sidlerstrasse 5, 3012 Bern, Switzerland CRPG-CNRS, 15 rue Notre Dames des Pauvres, B.P. 20, 54501 Vandoeuvreles-Nancy Cedex, France Max-Planck-Institute for Marine Microbiology, Department of Biogeochemistry, Celsiusstr.1, D-28359 Bremen, Germany Max-Planck-Institute for Biogeochemistry, PO Box 100164, 07701 Jena, Germany Water Resources Division, MS-420, U.S. Geological Survey, 345 Middlefield Road, Menlo Park, CA 94025, USA Water Resources Division, U. S. Geological Survey, 12201 Sunrise Valley Drive, Reston, VA 20192, USA Department of Geology & Geophysics, Louisiana State University, Baton Rouge, LA 70803-4101, USA Water Resources Division, U.S. Geological Survey, 345 Middlefield Rd, MS 434, Menlo Park, CA 94025, USA Postgraduate Research Institute for Sedimentology, University of Reading, UK, current address Center for Life Detection, Jet Propulsion Laboratory, California Institute of Technology, 4800 Oak Grove Drive, MS 183-301, Pasadena, CA 91109-8099, USA CNR-Instituto di Geologia Ambientale e Geoingegneria, Sez. Roma "La Sapienza", P.le Aldo Moro 5, 00185 Rome, Italy Institute for Reference Materials and Reference Measurements, JRC-European Commission, B-2440 Geel, Belgium European Commission, Joint Research Center, Institute for Reference Materials and Measurements, Retieseweg, B-2440 Geel, Belgium Institute of Mineral Resources, Chinese Academy of Geological Sciences, Beijing 100037, P. R. China Los Alamos National Laboratories, Condensed Matter & Thermal Physics Group, Mailstop K764, Los Alamos, NM 87545, USA Department of Geochemistry, Faculty of Earth Sciences, Utrecht University, P.O.Box 80021, 3508 TA, Utrecht, The Netherlands Environmental Biology Group, Research School of Biological Science, Institute of Advanced Studies, Australian National University, GPO Box 475, Canberra, ACT 2601, Australia National Institute of Water and Atmospheric Research, Wellington, New Zealand Carnegie Institution of Washington, Geophysical Laboratory, 5251 Broad Branch Rd., NW, Washington, DC 20015, USA Environmental Biology Group, Research School of Biological Science, Institute of Advanced Studies, Australian National University, GPO Box 475, Canberra, ACT 2601, Australia

xvi S. Ghelli H.A. Gilg I. Gilmour J.-P. Girard

M. GrSning C. Guillou

S. Halas K. Hashizume

V..M. Holland

J. Horita T.R. Ireland T.M. Johnson J.C. Johnston H.R. Karlsson C. Kendall E. Kerstel H. Kipphardt R.V. Krishnamurthy H.R. Krouse K. Leckrone

C. Ldcuyer

M. Leuenberger R. Lucchini I,

Lyon

SPIN, via Tamagno, 3, 42048 Rubiera (RE), Italy Lehrstuhl ffir Ingenieurgeologie, Technische Universit~it M~inchen, Arcisstr. 21, 80290 Mfinchen, Germany Planetary and Space Sciences Research Institute, Open University, Milton Keynes, Buckinghamshire, MK7 6AA, UK BRGM, Department Analysis and Mineral Characterization, BP6009, 45060 Orldans cedex 2, France International Atomic Energy Agency, Agency's Laboratories Seibersdorf, Isotope Hydrology Laboratory, A-1400 Vienna, Austria European Commission, Joint Research Centre, Institute for Health and Consumer Protection, Physical and Chemical Exposure Unit, 1-21020 Ispra (VA), Italy Uniwersytet Marii Curie-Sklodowskiej, Instytut Fizyki, Pracownia Spektrometrii Mas, P1. M. Curie-Sklodowskiej 1, 20-031 Lublin, Poland Centre de Recherches Pdtrographiques et Gdochimiques, 15 Rue NotreDame des Pauvres, B.P. 20, 54501 Vandoeuvre-16s-Nancy Cedex, France, and Department of Earth & Space Sciences, Osaka University, Toyonaka, Osaka 560-0043, Japan European Commission, Joint Research Centre, Institute for Health and Consumer Protection, Physical and Chemical Exposure Unit, 1-21020 Ispra (VA), Italy Chemical Sciences Division, Oak Ridge National Laboratory, P.O.Box 2008, MS 6110, Oak Ridge, TN 37831-6110, USA Research School of Earth Sciences, The Australian National University, Canberra ACT 0200, Australia Geology Department, 245 Natural History Bldg., MC-102, University of Illinois, Urbana-Champaign, Urbana, IL 61820, USA Iterations, P.O. Box 590805, San Francisco, California 94159, USA Department of Geosciences, and Department of Chemistry and Biochemistry, Texas Tech University, Box 1053, Lubbock, TX 79409, USA Water Resources Division, U.S. Geological Survey, 345 Middlefield Rd, MS 434, Menlo Park, CA 94025, USA Center for Isotope Research, Department of Physics, University of Groningen, The Netherlands Bundesanstalt ffir Materialforschung und -pr~fung (BAM), D-12200 Berlin, Germany Department of Geosciences, Western Michigan University, 1187 Rood Hall, Kalamazoo, MI, 49008, USA Department of Geology and Geophysics, University of Calgary, 2500 University Drive NW, Calgary, Alberta, T2N 1N4, Canada Department of Chemistry, Roosevelt University, 430 S. Michigan Ave., Chicago, IL 60605, USA Laboratoire CNRS UMER 5125, "Paldoenvironnements & Paldobiosph6re", Batiment ,~ Geode ,, Campus de la Doua, Universit6 Claude Bernard Lyon 1, 27-43 Boulevard du 11 Novembre 1918, 69622 Villeurbanne cedex, France Climate and Environmental Physics, Physics Institute, University of Bern, Sidlerstrasse 5, 3012 Bern, Switzerland Institut de Mindralogie et Pdtrographie, BFSH-2, CH-1015 Lausanne, Switzerland Department of Earth Sciences, The University of Manchester, Manchester, M13 9PL, UK

xvii B. Marty

B. Mayer W. Meier-Augenstein G. M~not

G. Michalski J. Miller C.R. Qu~tel M. Rehk/imper

F. Reniero

S. Rezzi

M. Ricci S.J. Sadofsky M. Saurer B. Schnetger

M.A. Sephton Z.D. Sharp S.M.E Sheppard R. Siegwolf S.R. Silva D. Smith

P. Spanel

H. Le Q. Stuart-Williams

B.E. Taylor P.D.P. Taylor

Centre de Recherches P~trographiques et G~ochimiques, 15 Rue NotreDame des Pauvres, B.P. 20, 54501 Vandoeuvre-l~s-Nancy Cedex, France, and Ecole Nationale Sup6rieure de G6ologie, Avenue du Doyen Roubault, 54501 Vandoeuvre-l~s-Nancy Cedex, France Department of Geology and Geophysics, University of Calgary, 2500 University Drive NW, Calgary, Alberta, T2N 1N4, Canada Queen's University Belfast, Environmental Engineering Research Centre, David Keir Building, Belfast, BT9 5AG, UK Institute of Geology, University of Bern, Baltzerstrasse 1, 3012 Bern, Switzerland, present address Woods Hole Oceanographic Institution, Department of Geology and Geophyics, Massachusetts, USA University of California, Department of Chemistry and Biochemistry, San Diego, CA 92039, USA NOAA Climate Monitoring and Diagnostics Laboratory, Boulder, CO, USA, and CIRES, University of Colorado, Boulder, CO, USA European Commission, Joint Research Center, Institute for Reference Materials and Measurements, Retieseweg, B-2440 Geel, Belgium Institute of Isotope Geology and Mineral Resources, ETH Ztirich, NO C61, CH-8092 Ztirich, Switzerland, and Institute of Mineralogy, Mtinster University, Corrensstrasse 24, D-48149 Mtinster, Germany European Commission, Joint Research Centre, Institute for Health and Consumer Protection, Physical and Chemical Exposure Unit, 1-21020 Ispra (VA), Italy European Commission, Joint Research Centre, Institute for Health and Consumer Protection, Physical and Chemical Exposure Unit, 1-21020 Ispra (VA), Italy Department of Biological Sciences, University of Idaho, Moscow, ID 83844, USA Department of Earth & Environmental Sciences, Lehigh University, Bethlehem, Pennsylvania 18015, USA Paul Scherrer Institute, CH-5232 Villigen PSI, Switzerland Carl-von-Ossietzky University, Institute for Chemistry and Biology of the Marine Environment (ICBM), P.O. Box 2503, D-26111 Oldenburg, Germany Planetary and Space Sciences Research Institute, Open Universit3r Milton Keynes, Buckinghamshire, MK7 6AA, UK Department of Earth and Planetary Sciences Northrop Hall, Albuquerque, NM, 87131-1116, USA Laboratoire de Science de la Terre and CNRS-UMR 5570, Ecole Normale Sup~rieure de Lyon, 46 All6e d'Italie, 69364 Lyon, France Paul Scherrer Institute, CH-5232 Villigen PSI, Switzerland Water Resources Division, U.S. Geological Survey, 345 Middlefield Rd, MS 434, Menlo Park, CA 94025, USA Centre for Science and Technology in Medicine, School of Postgraduate Medicine, Keele University, Thornburrow Drive, Hartshill, Stoke-on-Trent, ST4 7QB, UK V. Cerm~ik Laboratory, J. Heyrovsky Institute of Physical Chemistry, Academy of Sciences of the Czech Republic, Dolejskova 3, 182 23, Prague 8, Czech Republic Research School of Biological Sciences, Australian National University, GPO Box 475, Canberra City, ACT 2601, Australia Geological Survey of Canada, Ottawa, Ontario K1A OES, Canada Institute for Reference Materials and Reference Measurements, JRC-European Commission, B-2440 Geel, Belgium

xviii M.A. Teece P. Therrien M.H. Thiemens

S. Toyoda

S. Valkiers B.H. Vaughn R.M. Verkouteren

S. Wankel D.M. Wayne J.W.C. White E Wombacher S.C. Wong

N. Yoshida

C.-E You

State University of New York - College of Environmental Science and Forestry, 1 Forestry Drive, Syracuse NY 13210, USA D6partement de G6ologie et de G6nie G6ologique, Universit6 Laval, Qu6bec, GIK 7P4, Canada Department of Chemistry and Biochemistry, University of California, San Diego, La Jolla, California 92093-0356, USA Department of Environmental Chemistry and Engineering, Frontier Collaborative Research Center, and SORST, Japan Science and Technology Corporation (JST), Kawaguchi, Saitama, Japan Institute for Reference Materials and Reference Measurements, JRC-European Commission, B-2440 Geel, Belgium INSTAAR, University of Colorado, Boulder, CO, USA Surface and Microanalysis Science Division, Chemical Science and Technology Laboratory, National Institute of Standards and Technology, Gaithersburg, MD 20899, USA Water Resources Division, U.S. Geological Survey, 345 Middlefield Rd, MS 434, Menlo Park, CA 94025, USA NMT-15, Pit Disassembly and Nuclear Fuels Technologies, MS E 530, Los Alamos National Laboratory, Los Alamos, NM 87545, USA INSTAAR, University of Colorado, Boulder, CO, USA Institute of Mineralogy, M~inster University, Corrensstrasse 24, D-48149 M~inster, Germany Environmental Biology Group, Research School of Biological Science, Institute of Advanced Studies, Australian National University, GPO Box 475, Canberra, ACT 2601, Australia Department of Environmental Chemistry and Engineering, Frontier Collaborative Research Center, and Department of Environmental Science and Technology, Tokyo Institute of Technology, 4259 Nagatsuta, Midori-ku, Yokohama 226-8502, Japan, and SORST, Japan Science and Technology Corporation (JST), Kawaguchi, Saitama, Japan Department of Earth Sciences, National Cheng Kung University, Tainan, Taiwan, ROC

CONTENTS VOLUME I Dedication Introduction List of Contributors PART 1 Chapter I

Chapter 2

Chapter 3 Chapter 4

Chapter 5

Chapter 6

Chapter 7

Chapter 8

Chapter 9

Chapter 10

Chapter 11

Chapter 12

Chapter 13

Chapter 14

Stable Isotope Analysis of Water and Aqueous Solutions by Conventional Dual-Inlet Mass Spectrometry -J. Horita & C. Kendall Conventional and Less Conventional Techniques for Hydrogen and Oxygen Isotope Analysis of Clays, Associated Minerals and Pore Waters in Sediments and Soils - H.A. Gilg, J-P. Girard & S.M.F. Sheppard Techniques for Stable Isotope Analysis of Fluid and Gaseous Inclusions - L. Dallai, R. Lucchini & Z.D. Sharp Selected Ion Flow Tube Mass Spectrometry (SIFT-MS) and Flowing Afterglow Mass Spectrometry (FA-MS) for the Determination of the Deuterium Abundance in Water Vapour - P. Spanel & D. Smith Natural Abundance 2H-NMR Spectroscopy. Application to Food Analysis - S. Rezzi, C. Guillou, E Reniero, V.M. Holland & S. Ghelli Mass Spectrometric Techniques for the Determination of Lithium Isotopic Composition in Geological Material -Chan, L.-H. Thermal Ionization Mass Spectrometry Techniques for Boron Isotopic Analysis: A Review -You, C.-F. GC and IRMS Technology for 13C and 15N Analysis on Organic Compounds and Related Gases - W. Meier-Augenstein Preparation of Ecological and Biochemical Samples for Isotope Analysis - M.A. Teece & M.L. Fogel Extraction of Dissolved Inorganic Carbon (DIC) in Natural Waters for Isotopic Analyses - E.A. Atekwana & R.V. Krishnamurthy Compound Specific Isotope Analysis of the Organic Constituents in the Murchison Meteorite - M.A. Sephton & I. Gilmour A New Method for the Isotopic Examination of Sub-Milligram Carbonate Samples, Using Sulphamic Acid (NH2.SO3H) at Elevated Temperatures - H. Le Q. Stuart-Williams Determination of the Abundance and Stable Isotopic Composition of Trace Quantities of C and N in Geological Samples: The Practice and Principles of Stepped-Heating at High Temperature Resolution - S.R. Boyd Stable isotope measurements of atmospheric CO2 and CH4 - B.H. Vaughn, J. Miller, D.F. Ferretti & J.W.C. White

Page vii - ix xi - xiv xv - xviii

1 - 37

38 - 61 62 - 87

88 - 102

103 - 121

122- 141

142- 152

153- 176

177 - 202

203 - 228

229 - 236

237- 255

256 - 271 272- 304

xx Chapter 15 Preparation and Analysis of Nitrogen-bearing Compounds in Water for Stable Isotope Ratio Measurement - C.C.Y. Chang, S.R. Silva, C. Kendall, G. Michalski, K.L. Casciotti & S. Wankel Chapter 16 615N Analyses of Ammonium-Rich Silicate Minerals by Sealed-Tube Extractions and Dual Inlet, Viscous-Flow Mass Spectrometry - G.E. Bebout & S.J. Sadofsky Chapter 17 Nitrogen Isotopic Analyses at the Sub-Picomole Level Using an Ultra-low Blank Laser Extraction Technique - K. Hashizume & B. Marty Chapter 18 Mass Independently Fractionated Ozone in the Earth's Atmosphere and in the Laboratory -J.C. Johnston & M.H. Thiemens Chapter 19 Site-specific Nitrogen Isotope Analysis in N20 by Mass Spectrometry - S. Toyoda & N. Yoshida Chapter 20 Fluorination Methods in Stable Isotope Analysis -Bruce E. Taylor Chapter 21 Oxygen Isotope Analysis of Plant Water Without Extraction Procedure - K.S. Gan, S.C. Wong & G.D. Farquhar Chapter 22 Oxygen Isotope Analysis of Phosphate - C. L6cuyer Chapter 23 Pyrolysis Techniques for Oxygen Isotope Analysis of Cellulose M. Saurer & R. Siegwolf Chapter 24 Sample Homogeneity and Cellulose Extraction from Plant Tissue for Stable Isotope Analyses S. Borella, G. M6not & M. Leuenberger Chapter 25 Analytical Methods for Silicon Isotope Determinations T. Ding Chapter 26 Procedures for Sulfur Isotope Abundance Studies - B. Mayer & H.R. Krouse Chapter 27 Direct Measurement of the Content and Isotopic Composition of Sulfur in Black Shales by Means of Combustion-Isotope-Ratio-Monitoring Mass Spectrometry (C-irmMS) M.E. BOttcher & B. Schnetger Chapter 28 Summary of Methods for Determining the Stable Isotope Composition of Chlorine and Bromine in Natural Materials - H.G.M. Eggenkamp Chapter 29 Selenium, Iron and Chromium Stable Isotope Ratio Measurements by the Double Isotope Spike TIMS Method T.M. Johnson & T.D. Bullen Chapter 30 SIMS Measurement of Stable Isotopes T.R. Ireland Chapter 31 Stable Isotope Analysis by Multiple Collector ICP-MS - M. Rehk~imper, F. Wombacher & J.K. Aggarwal Chapter 32 Different Isotope Ratio Measurement Applications for Different Types of ICP-MS: Comparative Study of the Performance Capabilities and Limitations C.R. Qu6tel & J. Diemer Chapter 33 Isotope Ratio Analysis Techniques using Photoionization as a Source of Ions - I. Lyon Chapter 34 Isotope Ratio Infrared Spectrometry E. Kerstel -

-

-

-

-

-

-

-

305 - 347

348 - 360

361 - 374

375 - 389 390 - 399 400 - 472 473 - 481 482 - 496 497- 506

507- 522 523 - 537 538- 596

597 - 603

604 - 622

623-651 652-691 692 - 725

726- 745

746- 758 759- 787

xxi Chapter 35 Glow Discharge Mass Spectrometry: F u n d a m e n t a l s and Potential Applications in Stable Isotope Geochemistry - D.M. Wayne Chapter 36 The Use of Molecular Sieves in Stable Isotope Analysis - H.R. Karlsson Chapter 37 Introduction to Isotope Dilution Mass Spectrometry (IDMS) - M. Berglund

PART 2 Chapter 38 Mass Spectrometer H a r d w a r e for Analyzing Stable Isotope Ratios - W.A. Brand Chapter 39 Techniques of Ion Current Stabilization in Isotope Ratio Mass Spectro-

Chapter 40 Chapter 41

Chapter 42 Chapter 43

Chapter 44 Chapter 45 Chapter 46

Chapter 47

metry - S. Halas & T. Durakiewicz International Stable Isotope Reference Materials - M. GrOning The Nature and Role of Primary Certified Isotopic Reference Materials: A Tool to U n d e r p i n Isotopic M e a s u r e m e n t s on a Global Scale - P.D.P. Taylor, P. De Bi6vre & S. Valkiers Traceability in Isotopic M e a s u r e m e n t s - H. K i p p h a r d t Strategies and Practicalities in the Production and Use of Gas Isotope Standard Materials - R.M. Verkouteren Data Corrections for Mass-Spectrometer Analysis of SO2 - M. Coleman O x y g e n Isotope Corrections in Continuous-Flow M e a s u r e m e n t s of SO2 - K. Leckrone & M. Ricci Experimental M e a s u r e m e n t of Isotopic Fractionation Factors and Rates and M e c h a n i s m s of Reaction - S.M.F. Sheppard Laboratory Set-Up for GC-MS and Continuous-Flow IRMS - W. Meier-Augenstein

788 - 804 805 - 819 820 - 834

835 - 856

857 - 873 874- 906

907- 927 928 - 943

944- 956 957- 970 971 - 991

992- 1037 1038 - 1042

A p p e n d i x A Isogeochem list A p p e n d i x B The Web Stable Isotope Fractionation Calculator - G . Beaudoin & P. Therrien A p p e n d i x C Suppliers of Reference Materials

1043 - 1044

References

1053- 1181

Subject Index

1183- 1234

1045 - 1047 1048 - 1052


PART 1 Review and Discussion on Developments in Stable Isotope Analytical Technologies A Guidance to Modern Analytical Methods and Related Matters


Handbook of Stable Isotope Analytical Techniques, Volume 1 P.A. de Groot (Editor) 9 2004 Elsevier B.V. All fights reserved.

CHAPTER 1 Stable Isotope Analysis of Water and Aqueous Solutions by Conventional Dual-Inlet Mass Spectrometry Juske Horital & Carol Kendall2 1 Chemical Sciences Division, Oak Ridge National Laboratory, P.O.Box 2008, MS 6110, Oak Ridge, TN 37831-6110, USA 2 Water Resources Division, U.S. Geological Survey, 345 Middlefield Rd, MS 434, Menlo Park, CA 94025, USA e-mail: [email protected] ; [email protected]

1.1 Introduction The foundation of various analytical methods for the stable isotope composition of water and other aqueous samples (natural abundance, 1H. 2H (D) - 99.985" 0.015 atom%, and 1 6 0 " 170 9180 -- 99.762 90.038 90.200 atom%) was established during the Manhattan Project in the U.S.A., when large amounts of heavy water were produced for nuclear reactors (see Kirshenbaum, 1951, for a detailed account). From early on, there was great interest in the oxygen and hydrogen isotopic compositions of water, because they are the ideal tracers of water sources and reactions. The increased analytical precisions made possible by the subsequent development of modern gas-source isotope-ratio mass spectrometers with dual-inlets and multi-collectors, have caused the proliferation of new analytical methods and applications for the oxygen and hydrogen isotopic compositions of water. These stable isotopes have found wide applications in basic as well as applied sciences (chemistry, geology, hydrology, biology, medical sciences, and food sciences). This is because water is ubiquitous, is an essential and predominant ingredient of living organisms, and is perhaps the most reactive compound in the Earth. In this article, we review recent developments and refinements of analytical methods for preparing waters and other aqueous samples of different origins for the measurement of the oxygen and hydrogen isotopes by conventional dual-inlet, dynamic gas-source isotope-ratio mass spectrometry. Earlier review articles include those by Gonfiantini (1981), Wong & Klein (1986), Platzner (1997), and Coplen (2001). During the past decade, other in-situ (laser-ablation and secondary ion), static, quadrupole, and continuous-flow mass-spectrometric techniques have been developing rapidly. We briefly discussed emerging techniques of continuous-flow mass-spectrometry because the same preparation methods are employed in both dual-inlet and continuous-flow mass-spectrometry.

2

C h a p t e r 1 - J. Horita & C. K e n d a l l

The size, nature (free water or physically/chemically bound water within a matrix), chemical composition (pH, salts, and other dissolved constituents), and isotopic abundance of aqueous samples vary widely depending on their type, origin, and history. Examples are hydrologic (water vapor, precipitation, surface waters, soil waters, groundwater, geothermal water), geologic (fluid inclusions, hydrous minerals, dissolved water in minerals), biological (plasma, urine, saliva, human milk, breath water, plant water), agricultural (juices, wine, milk) and laboratory (synthetic, experimental) fluids. Waters enriched (and sometimes depleted) in deuterium and/or 180 are used for hydrological and biological tracer studies, where isotopic effects and fractionation are not of concern. Some meteorites have extremely wide ranges of oxygen and hydrogen isotopic compositions, reflecting nucleosynthesis and planet-forming processes in the early stages of the universe. Most isotopic studies, however, deal with aqueous samples whose isotopic compositions are within the normal range of terrestrial isotopic compositions (-400%0 < 6D < +50%0 and -50%0 < 6180 < +10%o on the VSMOW-SLAP scale). Investigations are made of natural (biologic, geochemical, and physical) and laboratory processes that cause a variety of kinetic and equilibrium isotope effects (isotope fractionations). For such studies dealing with relatively small isotopic variations, precisions of < 1 - 2%0 (6D, 62H) and < 0.1 - 0.2%0 (6180) are required. In addition to the aspects regarding water/aqueous samples discussed above, the choice of an analytical method depends on many other factors and requirements (expertise/equipment, time/labor, and precision/accuracy). While the conventional analytical techniques developed for natural waters in the 1950's and 1960's are still among the most precise methods, recent developments are shifting toward simplified procedures or automated analyses of small-size samples. In particular, automated preparation systems (available from commercial mass spectrometer companies) connected to the mass spectrometer (i.e., "on-line" systems) not only reduce the time and labor required for a large sample throughput, but also eliminate systematic errors of human origin (Barrie & Prosser, 1996; Brand et al., 1996). Automation, of course, requires substantial capital investments other than the mass spectrometer to keep up with state-of-the-art techniques. With the increasing popularity of stable isotope techniques in many disciplines and the vast amount of isotopic data produced, data-correction procedures, quality control, and inter-laboratory standardization of isotopic data are becoming crucial issues.

1.2 Collection, transportation, extraction, and storage of water samples Great care must be exercised in the collection, transportation, extraction, and storage of water samples prior to their isotopic analysis with techniques described in the following sections. First, it is critical that the sample collected actually be representative of the feature of interest. For example, groundwater wells need to be pumped for a time sufficient to remove the stagnant water in the well before sampling, and samples from deep rivers should be depth-integrated or sampled at weirs or other convergence points. Samples do not need to be filtered or chilled to preserve the isotopic compositions, but samples may need to be poisoned by HgC12 to suppress microbial activity. The most serious problem is evaporation, which can be avoided by quickly transferring the sample into an appropriate container. For large-volume samples of

Stable IsotopeAnalysisof Water and Aqueous Solutions ...

3

hydrological, biological and agricultural fluids, the use of glass or high-density polyethylene bottles with tight caps is recommended. The bottle size should closely match the size of the sample (i.e., small dead volume); small amounts of water in big bottles can readily fractionate by isotope exchange with vapor. From our experience, caps with conical inserts or ones with Teflon liners are the most reliable. Glass bottles should not be filled entirely to the top if there is any chance of freezing or of large pressure changes (such as might be expected in under-pressured airplane luggage compartments) during transportation. One convenient and reliable way to transport large numbers of bottles is to put them back into the original cardboard trays, wrap the trays in bubble paper, and then put the trays in insulated ice chests or coolers (at room temperature), packed securely with plastic peanuts. Long-term storage in leaky containers or low-density polyethylene bottles results in the alteration of isotopic composition due to evaporation and diffusion (Stewart, 1981; Rozanski & Rzepka, 1991). It is reported that the isotopic composition of water stored in some high-density polyethylene bottles with a thin (2mm) wall changed over a period of 2 years, due apparently to diffusion through the wall (Stuiver et al., 1995). For long-term storage, waxing the caps (dip bottle tops in melted paraffin and store upside-down) or torchsealing the waters in glass ampoules should be considered. Samples of waters with known isotopic compositions should be stored along with the samples as a check on possible fractionation. Samples containing large amounts of organic matter (e.g., oil) may need to be stored in cool and dark locations to prevent degradation. Atmospheric water vapor in air can be sampled by opening a valve of a pre-evacuated flask or by circulating air through a flask for a few minutes. Alternatively) water vapor can be collected by slowly pumping air through a glass or metal trap cooled at -60~ or below (Craig & Gordon, 1965; Ehhalt, 1974; White & Gedzelman, 1984; Helliker et al., 2002). A filter may be installed upstream to prevent contamination by rain, snow, or ice. The flow rate, coolant temperature, and other aspects (geometry of a trap, glass bead filler to increase surface area, etc.) must be carefully considered to ensure quantitative trapping of water vapor, depending on the temperature and humidity of air. Various desiccants (silica-gel, molecular sieves, Mg(C104)2, etc.) may be used to collect water vapor without a coolant, but they invariably contribute background water, when heated for the recovery of water. Soil and plant samples should also be kept in the same types of tight containers and bottles with small dead-volumes until extraction and analysis. Water contained in soil and plants can be extracted by various methods (vacuum and azeotropic distillation, squeezing/pressure extraction, and centrifugation) prior to isotopic analysis. Comparison of several studies obtained by these different extraction methods highlight potential problems associated with each method, especially for samples with low water contents (Revesz & Woods, 1990; Ingraham & Shadel, 1992; Walker et al., 1994; Aragu~is-Aragu~is et al., 1995; Landon et al., 1999). Direct CO2 and H2 equilibration methods recently developed for soil and plants samples offer an alternative, promising approach as discussed below (Scrimgeour, 1995; Hsieh et al., 1998; McConville et al., 1999; Koehler et al., 2000).

4


Brines, and agricultural and biological fluids contain high concentrations of dissolved salts and other compounds (sugar, alcohol, protein, etc.). In the decomposition methods for the isotopic analysis of waters discussed below, water must be extracted from these samples prior to isotopic analysis or as part of a single-step extraction/ reduction procedure. It is often very difficult to extract all water from these samples by means of vacuum or azeotropic distillation without potential isotopic exchange and fractionation, even at elevated temperatures. Centrifugation helps remove suspended materials from samples of blood, juice, and other substances. Some organic compounds (light hydrocarbons, oil, and tar) must be removed from natural waters because they can poison metal/chemical reagents used in the decomposition methods and can interfere with mass spectrometric measurements. The separation can be achieved by extracting them with water-insoluble solvents or adsorption on warm paraffin wax. CO2 and H2 equilibration methods are preferred methods because these methods require little pretreatment of aqueous samples. In addition to these analytical problems, the presence of certain kinds of dissolved constituents can change the measured isotopic composition of water, due to "the isotope salt effect", as discussed below. Geologic aqueous samples (fluid inclusions in minerals, water dissolved in minerals and glass, hydrous minerals) require the extraction of water prior to isotopic analysis by means of crushing, thermal decrepitation, and thermal heating (Godfrey, 1962; Roedder et al., 1963; Vennemann & O'Neil, 1993). Because these methods are bulk extraction techniques, the presence of different types of water (inclusion water and dissolved water in minerals) and different generations of fluid inclusions pose serious analytical problems (e.g., Kazahaya & Matsuo, 1985). Complete extraction of water is often a difficult task (e.g., Ihinger et al., 1994). Reactions occurring during extractions (e.g., adsorption of water to mineral surface, thermal reactions among C-O-H gases) could also alter the isotopic composition of inclusion water. The amount of water in these samples is often a limiting factor for isotopic analysis because conventional dual-inlet mass spectrometry requires water samples of a 10 tlmol (0.2 tlL). With decreasing sample size, contamination and memory effects from extraction and preparation systems become increasingly problematic. An extraction step, if necessary, is an important part of overall analytical procedure, and errors associated with these processes have to be critically evaluated. 1.3 Review of analytical methods

1.3.1 Hydrogen Isotopes Hydrogen gas (H2) is the preferred gaseous species for D / H ratio measurements in gas-source isotope ratio mass spectrometry, although other forms of gases (water, methane, acetylene, ethane, propanol) have been used with limited success. Many metals (U, Zn, Zn-CaO, W, Mg, Mn, Cr) and carbon have been tested for quantitative conversion (reduction) of water to H2 gas at elevated temperatures (400 - 1000~ the water-oxygen is converted to metal oxides or CO/CO2 (see reviews by Kirshenbaum, 1951; Wong & Klein, 1986). M + xH20 ~ xH2 + MOx

M: metal or carbon

[1.1]

Stable Isotope Analysis of Water and Aqueous Solutions ...

5

Quantitative conversion of the water to H2 is very important to avoid or minimize any potential isotopic fractionation. High concentrations of dissolved salts and compounds in brines, and in biological and agricultural fluids often interfere with the quantitative conversion reactions as discussed below. Conversion methods can be divided into (1) dynamic methods where multiple aliquots of water are reacted oneby-one with the metal reagent in some kind of reactor, and (2) static batch methods where water samples are reacted separately with a metal reagent in a closed vessel. A conceptually different method of analysis is to equilibrate the hydrogen isotopes between water samples and H2 gas. One main difference between conversion and equilibration methods is that the 6D values produced by conversion methods are composition values whereas the 6D (and 6180 values) produced by equilibration methods are activity values. This important distinction will be discussed further in section 1-4.

1.3.1.1 Dynamic conversion~decomposition methods Conventional U and Zn reactor methods

Among the metals, uranium (U) and zinc (Zn) have been most successfully used for D / H ratio measurements of natural waters in a dynamic conversion system (U: Bigeleisen et al., 1952; Stewart & James, 1981; Sajjad &Tasneem, 1983; Wong et al., 1984; Vaughn et al., 1998; Zn: Graft & Rittenberg, 1952; Friedman, 1953; Horibe & Kobayakawa, 1960; Schiegl & Vogel, 1970; Lyon & Cox, 1980; Dubois, 1985; Morse et al., 1993). These techniques, especially the uranium reduction method, can provide very high precision (lo = 0.2%o, Craig & Gordon, 1965), but require intensive labor and care because of the nature of largely manual procedure. Although dynamic reduction methods with uranium and zinc have been a standard method in the past, their use has significantly declined during the last decade. Recentl~ Cr (Gehre et al., 1996a, b), platinized Mg (Halas & Jasinska, 1996), and Mn (Tanweer & H a n , 1996; Shouakar-Stash et al., 2000) have also been reported as suitable reducing reagents for conversion. The basic procedure for using a uranium furnace is as follows (Bigeleisen et al., 1952)" uranium turnings (depleted in fissile 235U) are broken into small pieces, cleaned with HNO3 and distilled water, and packed tightly into a U-shaped or double-walled quartz tube, with quartz and/or copper wool loosely packed at each end (and perhaps in the middle) to prevent dispersal of the uranium oxide dust that is formed during the reaction (see Bigeleisen et al., 1952; Gonfiantini, 1981). The furnace is heated to 600 - 800~ in a vacuum line and the introduced water is passed once or more through the furnace by moving a liquid nitrogen dewar back and forth between small U-traps on either side of the furnace, and heating the glass line (by heat tape, heat gun, or torch). To ensure that the H2 gas introduced into the mass spectrometer is homogeneous, the gas may be mixed by use of a Toepler pump or a magnetically-operated rotor pump (Schiegl & Vogel, 1970). The U in a vessel should be replaced every few hundred analyses to avoid memory effects caused by the uranium oxides.

6

Chapter 1 - J. Horita & C. Kendall

The resulting H2 gas can be transferred to a reservoir or mass spectrometer inlet by expansion, automated Toepler pump, adsorption to charcoal with liquid nitrogen, or reaction with uranium metal at 80~ to form uranium hydride, which can be decomposed by heating to 700~ (Friedman & Hardcastle, 1970). For analyzing ~2 ~L water samples using the former method, several grams of charcoal (made from coconut shells) is put in glass sample vessels, and the charcoal is completely degassed at 350~ prior to use (J.R. O'Neil, pers. commun., 1999). H2 gas is transferred to the sample vessels with liquid N2, and then admitted into a mass spectrometer without heating the charcoal. There is about a 2%o fractionation in 6D values for H2 gas of normal sample size (a 100 ~mol). However, this does not pose a problem as long as H2 gas prepared from standard and sample waters are processed in the same way. It was observed that for small samples (ca. < 20~mol), 6D values of H2 gas transferred with charcoal became progressively lower (J.R. O'Neil, pers. commun., 1999).

On-line preparation methods Large sample throughput for D / H ratio measurements can be accomplished by (1) off-line, multi-sample preparation systems (section 1-3.1.2) and (2) automated, online preparation systems (i.e., connected to a mass-spectrometer). One early waterreduction unit using uranium metal was installed in the inlet system of a mass spectrometer so that the H2 produced was directly introduced to it (Nief & Botter, 1959). This design was further improved for multi-sample measurements (Thurston, 1970, 1971; Hartley, 1980; Thurston & James, 1984, Gehre et al., 1996b) and for simultaneous measurements of D / H and 1 8 0 / 1 6 0 with twin mass spectrometers (Hagemann & Lohez, 1978; Wong et al., 1984; Barrie & Coward, 1985). This method can be fully automated by the use of an autosampler (Cr: Brand et al., 1996; U: Vaughn et al., 1998). The Aqua SIRA (VG Isogas Ltd, UK) U-reduction system described by Wong et al. (1984) was designed for deuterium-enriched saline fluids derived from clinical and nutritional tracer studies. To reduce memory effects, each sample is injected and analyzed 6 - 20 times, with an average of 7 injections per sample. Samples and standards were injected sequentially (i.e., there is no dual inlet) with precisions of about 0.6 ppm (4%o) for deuterium-enriched samples, which is satisfactory for most biomedical studies. By analyzing a water reference between every 5 samples, precisions as good as 0.2%o for oxygen and hydrogen isotopes are obtainable (Hagemann & Lohez, 1978); however, Wong et al. (1984) report optimum precisions for natural abundance samples of 1.1%o and 0.4%o for hD and ~180, respectively, with their system. Although use of U reduction methods seemed to be on the wane, the successful development of a high-precision, automated, high-throughput U reduction system (Vaughn et al., 1998) may generate renewed interest in the method. Unlike most former methods, only small quantities of uranium are required: 4000 injections (350 400 samples) per 0.5g. During analysis, a 10 - 50 ~L aliquot of water (depending on the size of the sample loop used) is drawn from each septa-topped vial using a Gilson HPLC autosampler (Gilson, Middleton, WI, USA). The autosampler dispenses the water into a 6-port HPLC injection valve. When each sample is injected 4 times, and each vial in a set is reanalyzed twice (called two "tours" of the sample set), a reproduc-


7

ibility of about 0.3%o can be achieved (Vaughn et al., 1998). The authors attribute the high precision of their method to maintaining a highly repeatable sample size (+ 1%) by using an HPLC injection valve (instead of using a syringe), keeping the fractionation during sample evaporation constant, devising methods to reduce and correct for memory effects, and monitoring changes in "machine slope" (%o-scale shrinkage) over time. A comparison of the precision and accuracy of this method with the H2-water equilibration method (described below) for 165 Antarctic water samples analyzed by both methods (Hopple et al., 1998), showed that the 6D values of 94% of the samples analyzed by the two methods differed by less than 3.5%o, and 6% of the values differed by as much as 5%0. The "cutoff" value of 3.5%0 was chosen by the authors because it represents the 3~J standard deviation of most 6D analysis methods. Chromium (Cr) has also been used as a reducing agent (Gehre et al., 1996a, b; Nelson & Dettman, 2001) with good analytical precisions (1%o), and Brand et al. (1996) successfully automated the method using an autosampler. Chromium can be used both for off-line preparation where the sample is collected in glass ampoules, or as an on-line method where 1 ~L water samples are injected directly into the Cr furnace, and the gas flows from there into the expansion bellows of the mass spectrometer system (Gehre et al., 1996b). The Cr reactor for on-line reduction consists of about 50 g of chromium powder (Patinal < 0.3 mm, Merck) in a quartz furnace tube that is heated to 900-1000~ depending on sample type (Gehre et al., 1996b). The method is suitable for a variety of aqueous and organic compounds (gaseous or liquid); however, the temperature of the furnace and reaction time must be adjusted for each type. For example, water samples react very rapidly (80 sec from injection to starting the measurement) whereas methane requires 45 minutes for complete reduction (Gehre et al., 1996b). A disadvantage of the Cr method is that the reactors need to be refilled with Cr after 100 - 200 samples (Gehre et al., 1996b; Brand et al., 1996). Recently, Morrison et al. (2001) tested a Cr reactor method with continuous-flow isotope-ratio mass spectrometry. Water samples as small as 50nL (typically 3gL) was injected with an autosampler into a Cr reactor heated at 1050~ in a quartz tube, and H2 produced was carried with He gas into a mass spectrometer equipped with an electrostatic energy filter. They reported precision of _~0.5%0 with small to negligible memory effects. Recently, Shouakar-Stash et al. (2000) developed a Mn reaction unit installed on an inlet of an isotope-ratio mass spectrometer. Water, brines, natural gas, and organic solvents were reacted with Mn at 900~ for 20 sec (water) to 9 min (chlorinated solvents), and H2 produced was introduced directly to a mass spectrometer. Precision of ~ 1.02.0%0 was obtained. Using 5 gL of water per sample, 200 reductions can be obtained from 50 g of Mn. Advantages of Mn include availability~ low cost, lack of any pretreatment, and lack of memory effect for samples differing by less than about 200%0. Memory effects can be avoided by flushing with 5 gL of sample prior to an analysis. The most recent and very promising technique of water analysis is a high-temperature reduction/pyrolysis method in the presence of carbon, nickel (Ni) and other metals. In the presence of carbon, water reacts to form CO, (CO2), and H2. Ni and other metals appear to catalyze this reaction, and the production of CO2 is limited at

8


high temperatures. Tobias et al. (1995) described a Ni-metal furnace to reduce water to H2 at 850~ for a continuous-flow mass spectrometer; a heated (330~ palladium (Pd) foil "filter" was used to separate the resulting H2 from the carrier gas and other impurities. This Ni reduction at + 1000~ is finding application in off-line preparation with Ni pyrolysis bombs, where H2 diffuses out to a quartz tube around the Ni tube (Gray et al., 1984; Motz et al., 1997). On-line measurements of D / H and/or 1 8 0 / 1 6 0 in waters, organic matter, or inorganic matter in continuous-flow mode, especially when automated by coupling a high temperature elemental analyzer (sometimes called a "pyrolysis" unit) or gas chromatograph (Begley & Scrimgeour, 1996, 1997; Werner et al., 1996; Koziet, 1997; Farquhar et al., 1997; Kelly et al., 1998; Br6as et al., 1998; Loader & Buhay, 1999; Kornexl et al., 1999a, b; Hilkert et al., 1999), are rapidly becoming popular methods in many laboratories for the analysis of small water samples. Recently, Sharp et al. (2001), and Eiler & Kitchen (2001) applied the high-temperature carbon pyrolysis/Hecarrier gas continuous-flow mass spectrometry to the determination of D / H and/or 180 / 160 ratios of nano- to pico-liters of water samples. Coupled with laser ablation or heating, it is becoming reality to analyze both D / H and 180/160 ratios of minute amounts of water in geologic samples (fluid inclusions and hydrous minerals) with unprecedented spatial resolution. There is an active ongoing debate about the best type of pyrolysis reactors and the optimum pyrolysis temperature for different types of samples. This method is described further in Chapter 3.

Memory effects

A major drawback of all dynamic systems for the conversion of water to H2 is memory effects caused by (1) incomplete removal of samples from a preparation system between analyses, (2) memory from a metal reactor itself, and (3) adsorption of water onto the walls of the system, followed by isotopic exchange between the following water sample and this reservoir of adsorbed water. Memory effects of up to 4% difference in 6D values between two consecutive samples are reported, depending on the type and amount of metal and water (Bigeleisen et al., 1952, Graft & Rittenberg, 1952; Hartley, 1980; Lyon & Cox, 1980; Wong et al., 1984). A detailed study of the memory effect was carried out by Morse et al. (1993), using a dynamic zinc reduction method with BDH zinc (BDH Chemicals Ltd., Poole, Dorset, UK)(Figure 1.1). Their study concluded that the furnace blank (0.11 ~mol H2) was insignificant compared to the amount of H2 contributed by adsorption of water from the previous sample (1.2 gmol); the adsorbed water contributed as much as 3 - 10% of the total H2 yield, depending on sample size, and that the amount of contamination was independent of original sample size. If the memory effect is caused by a constant amount of the previous sample (Morse et al., 1993), then the measured 6D of any sample can be satisfactorily corrected by applying a "blank" correction of this amount, having the gD value of the previous sample. The correct 6D value of the sample can be calculated by the mass-balance equation: 6Dsam = [~Dm (Hsam + Hblk)- hDprev Hblk] / Hsam

[1.2]


9

Figure 1.1 - Memory effects of 6D values of SLAP after VSMOW from a Zn reduction unit. The memory effect increases with decreasing sample size, showing the contamination of 1.2 ~mol H2 from the previous samples regardless of sample size. After Morse et al. (1993)..

where (~Dmis the measured value of the sample, ~)Dprevis the measured value of the previous sample, and Hsam and Hblk are the amounts of hydrogen from the sample and blank (memory effect), respectively. Todd (1955) showed that there are two types of water in glass systems, surfacebound adsorbed water and water that diffuses out of the glass at temperatures > 400~ Surface-bound water is strongly bound and does not exchange with hydrogen gas (Graft & Rittenberg, 1952). Heating to 400~ was required to completely degas surface-bonded water, and heating to 100~ only removed about 25%; the amount of water bonded to glass increased with age of the glass and exposure to humid conditions (Todd, 1955). Other causes of memory effects include traces of water left in syringes from previous samples (syringes should be flushed several times with sample to avoid this) or from adsorbed atmospheric moisture, gas trapped in dead volumes of valves, and incomplete conversion in the reduction furnace. Memory effects can be reduced by heating the glass walls, reducing the internal surface area of the system, and by successive injections of the same sample. It appears that the continuous-flow isotope-ratio mass spectrometry with a He-carrier gas is less susceptible to the memory effects. 1.3.1.2 Static batch reduction methods Zn During the last decade, batch reduction methods using Zn in closed glass reservoirs gained popularity, and some methods were automated. Recently, similar methods with platinized Mg (Halas & Jasinska, 1996) and Mn (Tanweer &Han, 1996) were developed and are discussed below. The "grandfather" of all the many versions of zinc batch reduction methods in use today was the method of Coleman et al. (1982) who first identified a type of zinc shot (Analar zinc, BDH Chemicals Ltd., UK) that could

10


be reliably used in a batch method. The basic method is as follows" zinc shot is sieved to retain the-30 to +60 mesh size grains, rinsed in 30% nitric acid, washed in distilled water, dried, and outgassed under vacuum at about 300~ Some users insist that the cleaned zinc be kept under vacuum (maybe even heated) prior to use, and others leave it exposed to the air for years with no problems. About 0.25 g of zinc is put in the bottom of a Pyrex vessel that has a stopcock with a Teflon plug (Figure 1.2); the vessel is connected to a vacuum line, evacuated, and outgassed at about 100~ The closed vessels are then removed from the line, the plugs of the stopcocks are removed, the vessels are filled with dry nitrogen (to avoid contamination with atmospheric moisture), and 1-10 iLL of water is injected onto the zinc using a syringe. The plugs are then inserted quickly, the closed vessels returned to the vacuum line, the water frozen with liquid N2, and the vessels thoroughly evacuated. The closed vessels are then put in a heating block at 450~ (m.p. - 419.5~ for about 30 minutes for quantitative reduction of the water to H2. The vessels are then attached to the inlet system of the mass spectrometer and analyzed. Modifications of the Coleman method abound. Most groups use modifications of the original Coleman vessel type (Stanley et al., 1984; Florkowski, 1985; Wong et al., 1987b; Penman & Wright, 1987; Tanweer et al., 1988), including a very large version designed for a 60-port automated gas handling system (Kendall & Coplen, 1985), and a vessel with a sidearm for soil samples (Turner & Gailitis, 1988). However, other groups report on the successful use of Pyrex or quartz/Vycor tubes (Kendall & Coplen, 1985; Sudzuki, 1987; Vennemann & O'Neil, 1993; Schimmelmann & DeNiro, 1993: Dem6ny, 1995; Yang et Figure 1.2 - Variations of the Coleman reaction vessel for batch Zn reduction method. Wall bubble for samples of biological fluids (urine, plasma, saliva, human milk, etc.)(Wong et al., 1987) and sidearm for porous samples (sand, soils, etc. ) (Turner & Gailitis, 1988). The original vessel by Coleman et al. (1982) had neither bubble nor side arm.


11

al., 1996b; Karhu, 1997). Tanweer et al. (1988) suggested modifying the method by increasing the reaction temperature to 460~ and using 10 times the stoichiometric amount of zinc. Other workers use higher temperatures for glassware outgassing (up to 500~ zinc outgassing (up to 350~ and zinc oxidation (up to 500~ Water introduction methods also differ. Instead of removing the vessels to inject the water, water can also be introduced into the vessels on-line, either by direct injection through a septum, or by freezing water derived from decomposition of minerals or organic matter, or by breaking open microcapillaries containing water. Alternatively) the water can be syringed into hairpin-shaped microcapillaries which are loaded into the vessels, along with glass-enclosed iron bars, at the same time as the zinc; a magnet is used to raise and then drop the bar to break open the capillary just before putting the vessels into the block heater (Kendall & Coplen, 1985). Most workers have either used BDH zinc or a zinc available from Biogeochemical Laboratory at Indiana University (variably referred to as "Bloomington," "Indiana," or "Hayes" zinc)l. A comparison of the analytical precisions for water and human-fluid samples prepared using the "original" BDH zinc and Indiana zinc (Wong et al., 1992a) showed no significant difference; Vennemann & O'Neil (1993) also reported no significant differences in yields or 6D values with the two zinc types. The method can be used for many saline and hypersaline waters, including biomedical fluids (Wong et al., 1987b; Tanweer, 1993a), without prior distillation, but with some limitations (see section 1-4). It was observed that low-pH waters (e.g., acid lake water, water extracted from some hydrous minerals) react incompletely with Zn resulting in poor precisions.

Reactivity of Zn The choice of zinc reagent is critical to the successful use of the method. A number of groups have reported difficulties with a variety of zinc used for batch reduction (Coleman et al., 1982; Stanley et al., 1984; Kendall & Coplen, 1985; Florkowski, 1985; Wong et al., 1992a). In fact, many groups have problems with recent batches of BDH zinc (e.g., Wong et al., 1992a). Tanweer (1993a) reported that a new batch of BDH zinc with 0.3 - 1.5 mm grain size required larger amounts of zinc (+ 2.5 g) for 8 ~L water, compared to that of the original batch (0.25 g), and that reaction temperature must be increased from 460 ~ to 480~ One early comparison of the metal compositions of a number of types of analytical grade zinc found that the only significant difference among the brands was that BDH zinc contained more lead (Kendall & Coplen, 1985). Subsequent analyses of batches of BDH zinc have revealed other possibly important trace contaminants. A recent comparison of the trace contaminants in BDH and SHG zinc (Outokumpu Zinc Co., Kokkola, Finland) showed that BDH zinc contained much higher quantities of sodium (Karhu, 1997). The emission spectrographic analyses performed by Kendall & Coplen (1985) showed no difference in sodium contents among zinc types, perhaps because the detection limit for sodium was so high (0.1%) whereas it was at 0.1 to 100 ppm (depending on element) for the other metals.

Contact Dr. A. Schimmelmann, Dept. of Geol. Sci., Biogeochemical Lab., Indiana University, Bloomington, IN 47405-1405, U.S.A, e-mail: [email protected].

12


Several groups have tried to enhance the reductive properties of zinc by the addition of trace contaminants. Addition of I - 1.5 g lead granules to l g aliquots of otherwise unusable brands of cleaned zinc, or to already-used but recleaned BDH zinc (remelting not necessary), produced moderately acceptable results (Kendall & Coplen, 1985); replacements of the lead with tin or cadmium granules resulted in incomplete reduction. The so-called Indiana zinc has been produced since about 1985 by melting zinc with the addition of undisclosed contaminants, and then reforming the zinc mixture (Hayes & Baker, 1986; Hayes & Johnson, 1988), and is commercially available from Biogeochemical Laboratory of Indiana University (see the above address). This modest commercial effort was initiated in about 1986 to remedy the problem many laboratories were experiencing getting reliable results using different batches of BDH zinc. Karhu (1997) determined that the addition of trace quantities (~ 200 ppm) of sodium can also improve the reductive capability of an otherwise unusable zinc. Zinc-sodium mixtures were made by melting weighed aliquots of SHG zinc and various amounts of sodium (Merck) in a borosilicate tube under a constant flow of Ar (the Ar is necessary to prevent oxidation and eventual ignition of Na during heating); about 200 ppm sodium produced optimum results (reproducibility of + 0.7%o). After cooling, the cylinders were lathed and the thick lathings (1 - 1.5 mm thick) were broken into 15 mg chunks which could be used without further preparation. Samples were then prepared using a sealed-tube method (Kendall & Coplen, 1985; Vennemann & O'Neil, 1993), where about 30 mg of zinc per I mL of water was added to each 6 mm OD quartz tube (Karhu, 1997). Heating the tubes to 250~ prior to opening them into the mass spectrometer was found sufficient in reducing the small systematic bias resulting from adsorption of H2 by zinc (Karhu, 1997), in contrast to the temperature of 480~ recommended by Dem6ny (1995). Many groups now routinely make reliable zinc by adding Na to commercially available zinc (J.R. O'Neil, person, comm., 2000). Remelting zinc to improve the purity or to alter surface features has also been attempted by a few groups. SEM photographs of thoroughly cleaned zinc grains showed that BDH zinc had a much smoother and homogeneous surface (Kendall & Coplen, 1985; Tanweer, 1990) than other types examined. Several types of zinc were melted in long Vycor tubes in an attempt to produce zinc with identical surface configurations (Kendall & Coplen, 1985). When cooled, the impurities were concentrated at one end and removed. The solid, purified, zinc cylinders were lathed to produce fine turnings which were recleaned and tested; about I - 1.5 g of reformed zinc from two other sources behaved just as well as BDH zinc (Kendall & Coplen, 1985). Sublimation of an otherwise unusable zinc on the walls of a vessel or tube prior to reaction with water is another way to produce a usable zinc (Kendall & Coplen, 1985; Sudzuki, 1987). For example, melting of pieces of zinc wire (Nihon Denkyu Kogyo, Japan) under vacuum to decompose any zinc hydroxide, followed by introduction of the water and sublimation of the zinc onto the walls to enhance reaction rates, produces results comparable to the Coleman et al. (1982) method (Sudzuki, 1987). Noto & Kusakabe (1995) reported that Zn powder (no grain size information available) can also be used successfully after washing with nitric acid and degassing in vacuum at 250~ Hence, there is ample evidence that many other types of zinc are potentially


13

usable but the reforming or sublimation of the zinc is time consuming. These experiments suggest that successful water reduction and zinc oxidation is highly dependent on surficial casting-produced features and not on just chemical composition.

Hydrogen reservoir in Zn There is some evidence for a source of exchangeable H in zinc. In one of the few published studies that made a direct comparison of 6D values produced by the conventional U furnace method and the zinc batch method (Kendall & Coplen, 1985), it was noted that one batch of BDH zinc produced 6D values for VSMOW and SLAP that were within 1 - 2%0 of the values obtained by uranium reduction, whereas three other BDH batches showed considerable scatter and produced 6D values that were 2 10%o lower than those determined with U reduction. Similar differences among different batches of Zn have been observed in many other laboratories. Schimmelmann & DeNiro (1993) also reported that gD values of water prepared with the Indiana zinc in Pyrex tubes at 500~ were up to 14%o lower than those of U-reduction method. The source of exchangeable H in the zinc could apparently be removed by melting the zinc under vacuum (Kendall & Coplen, 1985). Florkowski (1985) showed that there was a strong "amount effect " with some types of zinc (including one batch of BDH zinc), causing large (> 20%0) fractionations for 0.8 g of zinc; some zinc types and batches also showed significant differences among various fractions < 1.5 mm. Subsequent work by others at the IAEA lab (Tanweer et al., 1988) has confirmed the existence of small differences in 6D values for waters prepared with various weights of BDH zinc, but suggested that the large depletions in D seen in earlier studies might be an artifact of incomplete cleaning/degassing of the zinc. Furthermore, they concluded that the amount of BDH zinc (10 times the stoichiometric amount) and a conversion temperature of 460~ not the grain size, provides the optimum conditions for precise and accurate results (Tanweer et al., 1988). Lower stoichiometric ratios (0.1 to 0.2 mg zinc for 8 - 12 ~L water) produced low 6D values because of incomplete reaction, and high values also produced low gD values (Tanweer et al., 1988). The theory that the low 6D values associated with excess zinc was caused by absorption of hydrogen in the zinc (Florkowski, 1985) is supported by the results of some subsequent experiments (Tanweer et al., 1988; Dem6ny~ 1995). Schimmelmann & DeNiro (1993) demonstrated that "Bloomington zinc" leached with nitric acid released 44 gmol H2/g with 6D values of -163 to -229%0 upon heating to 450~ even after drying in vacuum at 100~ It has been suggested that one of the main reasons why many laboratories experienced poor reproducibility with BDH zinc is because of failure to follow the cleaning procedures outlined by Coleman et al. (1982) and explained in more detail in Tanweer et al. (1988), resulting in incomplete removal of adsorbed water and oxides on zinc surfaces (Tanweer, 1990). Ultrasonic cleaning before the acid wash has been suggested as a means to remove the fine particles (Turner & Gailitis, 1988) that can be a minor factor in producing lower 6D values than obtained using coarser fractions (Tanweer et al., 1988). Although they only report data from a single batch of BDH zinc, Tanweer et al. (1988) suggest that proper cleaning of the zinc eliminates the variations in performance that had been observed with different batches of BDH zinc.

14


Hydrogen reservoir in the glasses The presence of a source of H on and/or within Pyrex glass has been demonstrated by several experiments. Pyrex vessels, preheated to 100 ~ - 200~ under vacuum, and then filled with H2 to a pressure of 33 KPa (25~ with a 6D of-122%o required 8 days of continuous heating at 430~ to reach a steady-state composition of about-175%o; pretreated vessels filled with H2 with a 6D of-690%o showed an increase to -665%o in 24 h, but had not reached a steady-state composition in 8 days (Kendall & Coplen, 1985). These results strongly suggest a source of hydrogen within the Pyrex with a 6D value intermediate between the two isotopic compositions of H2. These degrees of isotopic exchange were found regardless of whether the vessels had previously been baked for several hours with hydrogen of-122 or -689%o, suggesting that the reservoir was large and the exchange rate was slow. Turner & Gailitis (1988) also measured the extent of gaseous exchange with the Pyrex vessels by degassing vessels at 100~ adding H2 with a known 6D at 19 KPa (25~ and baking the vessels at 450~ for 2 hours. The heating resulted in a shift of 0.4 - 2%0, with no significant difference in results for vessels that had previously been exposed to H2 with 6D values showing a range of about 250%0. Use of Vycor (or quartz) vessels or sealed-tubes was found to cause significantly less isotopic exchange than Pyrex vessels or tubes by several thorough studies (Kendall & Coplen, 1985; Vennemann & O'Neil, 1993; Karhu, 1997), supporting the existence of a sizable reservoir of exchangeable hydrogen in borosilicate glass. However, it should be noted that because in many of the experiments described in the paragraphs above the glass was not heated to 400~ (the temperature apparently required to completely degas surface-bonded water, according to Todd, 1955) between experiments, it is possible that some of the results attributed to exchangeable hydrogen in the glass might be due to surface-bonded water. The degree of exchange is correlated with the surface to volume ratio and time; H2 in 6 or 9 mm sealed-tubes is more exchanged than H2 in large vessels at a constant pressure (Kendall & Coplen, 1985). Vycor and quartz behaved identically and showed much less exchange with the H2 of samples than Pyrex (Kendall & Coplen, 1985; Sudzuki, 1987; Vennemann & O'Neil, 1993; Karhu, 1997). Hence, quartz or Vycor are an obvious, but expensive, choice for a relatively trouble-free sealed-tube technique.

Considerations for routine-analysis The effect of hydrogen isotope exchange with Zn reagents and/or glass during the reaction and storage is to decrease the difference between the 6D of waters (i.e., "shrink the %o-scale")(Kendall & Coplen, 1985; Turner & Gailitis, 1988). Storage of reacted (and unreacted) samples in Vycor tubes for 18 months did not affect the 6D of the gas (the reacted samples were NOT reheated before analysis), making Vycor tubes ideal for researchers who stockpile samples (Kendall & Coplen, 1985). This experiment is in sharp contrast to the findings of Dem6ny (1995) who reported that heating quartz or Pyrex tubes to 480~ immediately prior to introduction of the H2 into the mass spectrometer was necessary to eliminate a systematic bias in the 6D that he attributed to adsorption of H2 into the zinc (Hayes zinc), and the somewhat similar findings of Karhu (1997) who reported that heating of the tubes to 250~ was needed to remove a systematic bias with the sodium-spiked SHG zinc. These results suggest


15

that the degree of apparent absorption into zinc may depend on zinc type. BDH zinc may be less subject to absorption; Kendall & Coplen (1985) and Vennemann & O'Neil (1993) reported no significant amount effect with BDH zinc. Nevertheless, if the same sample volumes and water/zinc ratios are maintained, a linear normalization correction will remove the bias caused by the use of borosilicates and any possible zinc absorption effects. In fact, many users of zinc batch methods may be completely unaware that their version of the method is routinely producing results that may be biased by as much as 10 - 20%o. They might only discover this fact if they compare their zinc-batch results with results obtained using other reduction methods in their own laboratory (which is, unfortunately, rarely done), send their breakseals to another laboratory for analysis, or wonder about apparent No-scale shrinkage. Since the degree of bias (fractionation) depends on the details of the preparation system used, when such gas samples prepared in one laboratory are analyzed in a different laboratory where the mass spectrometers are calibrated by using gas samples prepared using a slightly different version of the zinc method, the values must be normalized using water reference standards prepared along with the water samples (in the first laboratory). One of the keys to successful use of the zinc method is to maintain consistent analytical conditions. It is a very unforgiving method where seemingly small methodological changes can produce large problems with reproducibility. Users should follow procedures of a published method exactly~ or else make whatever modifications seem necessary~ test the modified method thoroughly, and once the method is working fine, change nothing. Even under ideal conditions in a laboratory where there are experienced workers with the zinc method, the fussy method occasionally produces ~D values that are 5 - 50%0 fractionated. Hence, it is recommended that whenever possible, samples be prepared in duplicate and/or any meteoric-derived samples that deviate significantly from the GMWL (Global Meteoric Water Line) be considered for reanalysis. The standard zinc batch method has been modified for the analysis of pore water in soils by addition of a sidearm positioned below the stockcock (Figure 1.2), which can be loaded with soil samples contained in small reusuable glass tubes (Turner & Gailitis, 1988). The side arm remains above the top of the heating block, and the soil water evaporates to react with the zinc. For samples up to 150 gL and water contents < 20%, the modified technique requires a reaction time of 2.5 h. Insufficiently long reaction times, combined with water amounts > 30 gL, result in significant decreases in the 6D values due to incomplete reduction of water. There is a consistent bias in the 6D values of water extracted from kaolinite, apparently from an additional source of exchangeable H from the clays, that can be corrected for by use of standards of known composition (Turner & Gailitis, 1988). Other metals Problems with the reliability of batches of BDH zinc have prompted searches for other reagents that might be suitable for use in a batch reaction mode. Chromium can

16


be used for the successful analysis of small (1 gL) water samples, either in an off-line batch mode or in an automated dynamic mode (Gehre et al., 1996b)(section 1-3.1.1). The use of magnesium (Mg) powder activated by a thin film of platinum (Pt) for analyzing water samples was first reported by Deqiu & Zhengxin (1985); this method was described in more detail by Halas & Jasinska (1996). The basics of the method are as follows" granular Mg (0.5 - 2 mm) is sieved, vacuum roasted, coated with dissolved platinum by a rather complicated procedure, the treated granules are dried and outgassed, and then about 4 gL of water is reacted with 120 mg of Mg-Pt reagent at 400~ for I h. The Pt coating step involves the use of several hazardous chemicals, including an acetone-ether solution that has to be distilled prior to use, and aqua regia. Previous attempts to use magnesium for either batch or dynamic reduction had been stymied by the production of Mg(OH)2 on the surface; the Mg(OH)2 can be decomposed at about 550~ but this would preclude the use of Pyrex reaction vessels (Halas & Jasinska, 1996). Coating the Mg with platinum black (Mg-Pt) may result in total decomposition of water to H2 at 400~ because the Mg-Pt is so reactive that the formation of hydroxides does not occur (Halas & Jasinska, 1996). The preliminary experimental data reported by Halas & Jasinska (1996) show some promise, but more work is needed to demonstrate optimum conditions for long-term precision and accuracy. They report that the method is rather sensitive to the amount of reagent used (perhaps because of diffusion of hydrogen into the metal), and that they hope that it might be suitable for analyzing untreated brines because of the high reactivity of the reagent. Manganese (Mn) shows much more promise in that it requires almost no pre-treatment and a preliminary report shows analytical precision of 0.4-0.8%o (Tanweer & Han, 1996). Although two types of Mn were tested, the 0.1 mm size Mn available from Fluka Chemicals Ltd. (Buchs, Switzerland) was shown to be a better choice than the finer-grained Mn from Merck Chemicals (Darmstadt, Germany). Optimum conditions using Coleman-type vessels (Coleman et al., 1982) appear to be 8 gL of water, 0.8 g of Fluka Mn, and a reaction time of 40 minutes at a temperature of 520~ (Tanweer & Han, 1996). Major advantages of the method are that the Mn requires no pretreatment, does not seem to show an amount effect (for 0.1 mm Fluka zinc). Furthermore, it is possible that other brands of Mn could be used, if they were sieved and rinsed to remove the fine powders that apparently are responsible for the amount effects. Shouakar-Stash et al. (2000) developed a Mn reactor unit installed on an inlet of an isotope-ratio mass spectrometer (see section 1-3.1.1). Recently, Ward et al. (2000) showed that lithium aluminum hydride can be used to reduce water for D / H ratio measurements by the reaction: LiA1H4 + 4H20 --* LiAI(OH)4 + 4H2

[1.3]

9 - 11 /~L of water was reacted with 0.04 to 0.05g of LiA1H4 in a vacutainer, and the reaction occurred instantly apparently at room temperature. Because LiA1H4 contributes to hydrogen in H2 produced, D / H measurements must be standardized with VSMOW-SLAP or other standards prepared by the same procedure. They reported precision of + 4 - 8%0 for waters with natural D / H ratios.


17

Van Kreel et al. (2000) and Previs et al. (2000) report D / H ratio measurements of biological fluids enriched in D20 after converting water to acetylene with calcium carbide CaC2 + H20 --~ CaO + C2H2

[1.4]

The reaction was complete within a few second at room temperature, carried out in a closed container (e.g., Vacutainer). The mass ratio of m / e - 27 (C2HD) to 26 (C2H2) was determined with a mass spectrometer. The current precision (few tens of %0), however, limits its application.

1.3.1.3 H2-water equilibration H2-water equilibration methods were originally developed during the 1930's1940's for determining deuterium contents of heavy water by equilibrating it with H2 gas in the presence of Pt catalysts (Kirshenbaum, 1951). A major problem was a sharp decline in the catalytic activity of Pt catalysts used due to coverage of the surface by water. During the 1970's, new types of Pt catalysts were developed, which retain a hydrophobic surface because of the matrix materials used (Teflon, styrene divinyl benzene, etc.). These Pt catalysts significantly enhanced D / H exchange between liquid water and H2 (Rolston et al., 1976). Ohsumi & Fujuno (1986), and Horita (1988) reported precise (lo = 0.8%o) D / H ratio measurements of natural waters, including s. Subsequently, Horita et al. (1989) demonstrated that the H2-water equilibration method with the hydrophobic catalysts can be readily automated using a gas equilibrator designed for the automated CO2-water equilibration method (Figure 1.3). Several investigators tested and improved this method for natural waters (Coplen et al., 1991; Ohsumi, 1991; Brand et al., 1996; Thielecke et al., 1998; Bourg et al., 2001) and for biological and agricultural fluids (Scrimgeour et al., 1993; Coplen & Harper, 1994; Br6as et al., 1996; Thielecke & Noack, 1997). The major advantages of the H2-water equilibration method are: (1) no pretreatment of water samples is required regardless of the type of aqueous solution (e.g., fresh water, brine, biological and agricultural fluids), except for removal of dissolved H2S; (2) very precise data (lo < 0.5%o, in external precision) can readily be obtained; and, (3) the entire procedure can be readily automated with an on-line gas equilibrator, thus reducing required labor and time, and reducing human errors. Both D / H and 180/16 0 ratios of water can be determined successively with the same gas equilibrator. Gas equilibrators and multi-preparation systems with septum-sealed glass vials (e.g., Vacutainers) for gas equilibration methods are currently available from commercial mass spectrometer firms (Analytical Precision, Finnigan MAT, Micromass, PDZ Europa). The gas equilibrator is specially designed for fully automated high-precision gas equilibration methods, and the multi-preparation system is versatile for many other sample preparations. Recently, the H2-water equilibration method has also been adapted to a continuous-flow mass spectrometer with a unique geometry (Prosser & Scrimgeour, 1995). The addition of a small energy filter on conventional continuous-flow mass spectrometers to better separate the He and H2 peaks has allowed the modification of the equilibration systems to operate under continuous-

18


Figure 1.3 - Schematic of an automated, on-line gas equilibrator for H2- and CO2-water equilibration methods. Air in the equilibration vessels is p u m p e d out through a capillary tube to minimize a loss of water. Temperature control of + 0.05 ~ and + 0.1~ or better is required for H2- and CO2-water equilibration methods, respectively.

flow mode using He as a carrier gas. Precisions of 1%o for hydrogen and 0.1%o for oxygen or better are presently being reported by the commercial vendors.

Exchange reactions, catalysts and normal procedures Hydrogen isotope exchange between gaseous H2 and liquid water proceeds via two consecutive reactions; HD(g) +H20(v) ~=~H2(g) + HDO(v) and

[1.5]

19

Stable Isotope Analysis of Water and Aqueous Solutions ... H D O ( v ) + H20(1) e,, H 2 0 ( v ) + HDO(1)

[1.6]

The overall reaction is HD(g) + H20(1 ) k H2(g ) + HDO(1)

[1.7]

where k is the rate constant. Pt-catalysts catalyze the isotope exchange reaction [1-5], which is otherwise extremely sluggish. The fractionation factor, a - (D/H)water/(D/ H)H2, of the overall reaction [1-7], is very large (3.81 at 25~ Rolston et al., 1976), resulting in extremely low D / H ratios (ca. 0.004%) of H2 gas equilibrated with natural waters (0.015%). The D / H fractionation factor is also very sensitive to temperature, and precise temperature control (___0.05~ is required. The original hydrophobic Pt catalyst used by Horita (1988) and Horita et al. (1989) is Hokko Beads (125 - 250 mm diameter styrene divinyl benzene porous resin beads coated with 3 wt% Pt), manufactured by Shoko Ltd. (Minato-ku, Japan). Another type (1 mm, I wt% Pt) is also available. Finnigan MAT (Bremen, Germany) provides Pt catalysts manufactured by MS-Analysentechnik (Berlin, Germany). The Atomic Energy of Canada Limited (AECL) of Chalk River National Laboratory (Ontario, Canada) also produces similar hydrophobic Pt catalysts. The hydrophobic Pt catalysts are generally carefree, and can be used many times after washing with distilled water and drying in air. It is reported that heating the I mm Hokko Beads at 60~ resulted in the loss of catalytic activity, whereas heating at 60 ~ - 100~ has no effect on the powdered Hokko Beads and other hydrophobic catalysts (ISOGEOCHEM1 discussion). Dissolved HzS is known to poison the catalysts; hence, water samples suspected to contain H2S should be treated by reacting the water with Cu metal or AgNO3 overnight (Coplen et al., 1991). Although the human nose is very sensitive to low concentrations of H2S, prolonged exposure is harmful; hence, water samples should not be routinely tested by sniffing. Heating the beads that were exposed to H2S at 230~ for 2 hr in a H2 stream can restore the activity (Ohba & Hirabayashi, 1996). Among biological and agricultural fluids, alcohol in wine is reported to interfere with mass spectrometric measurements (Br6as et al., 1996). However, it is not clear whether alcohol poisoned the catalysts or interfered with mass spectrometric measurements. The treatment of urine samples with activated charcoal is also recommended. The catalysts are electrostatic, and can be better handled by gluing to a stick (Coplen et al., 1991) or containing them in a Teflon bag (Ohba & Hirabayashi, 1996). The suppliers of the hydrophobic Pt catalysts are: 9 Shoko Co. Ltd. (attn: Wataru Maruyama)" 3-8-3 Nishi-Shimbashi, Minato-ku, Tokyo 105, Japan. Tel. +81 3-3459-5106; Fax (81) 3-3459-5081

Information on the ISOGEOcHEM list can be found in Appendix A of Part 2 in this Volume of the Handbook.

20

Chapter 1 - J. Horita & C. Kendall U.S. agent of Shoko" ICON Services Inc., Stable isotopes and labeled compounds (attn: John Kilby), Ox Bow Lane, Summit, NJ 07901, U.S.A. Tel. +1 (800) 322-4266; Fax +1 (908) 273-0449; www.iconisot.ios.com

9 MS Analysentechnik: c/o Bogen-Elektronik, Potsdamer Str 12-13, 14163 Berlin, Germany. Fax +49 30 8226072 9 AECL: Marketing and Sales/Nuclear Products and Services (attn: Chris Knight), Chalk River National Laboratory; Chalk River, Ontario, Canada K0J 1J0. Tel +1 (613) 584-8811 (ext 6029); Fax +1 (613) 584-1438. For normal operation, 1 to 5 mL of water or aqueous sample is loaded into a reaction vessel, together with a few to tens of mg of the Pt catalyst: the amount of the catalyst required probably depends on the type and manufacture. Large amounts of water are preferred because then (1) loss of water during the evacuation of air in a reaction vessel through a capillary tube is negligible when a gas equilibrator is used, (2) the evaporation of water within a reaction vessel (typically 10 - 20 mL) is negligible, and (3) there is little change in the D / H ratio of water samples during isotopic exchange between H2 and water. However, water samples as small as 0.1 - 0.25 mL have been successfully analyzed (Coplen et al., 1991; Thielecke et al., 1998). The vessels are then evacuated either after freezing the water or through a capillary tube without freezing. H2 gas is admitted into the vessels to a given pressure, usually one atmosphere. It is recommended that the same amounts of water and H2 gas be used for the entire set of standard and sample waters, so that isotopic changes caused by the above three processes should be constant and accounted for by standardization (normalization). Valves to the reaction vessels are closed and the vessels are immersed into a water or air bath. The bath temperature should be near room temperature (generally at 18 40~ and controlled to within + 0.05~ because the water-H2 D / H fractionation factor is very sensitive to temperature (6.3%o per ~ The recommended maximum + 0.05~ deviation in temperature corresponds to ~ 0.3%o in 6D (H2) values. The use of hollow plastic balls on the surface of the water bath can drastically reduce temperature fluctuations. A bath temperature below room temperature may be preferred in order to avoid the condensation of water inside the vessels. However, the above three factors (water loss, evaporation, and isotopic exchange) should be evaluated carefully for any experimental procedure, and suitable corrections made to the data, if necessary. Shaking the vessels is not advisable, because unlike CO2-water equilibration discussed below, it does not facilitate the isotope exchange reaction, and because it could spread the catalyst powders, plugging the capillary tube. After isotopic equilibration, H2 gas is expanded into a sample-side reservoir of a mass spectrometer through a Utrap cooled with dry ice or liquid N2 that removes water moisture. When the sample of water is small (< 1 mL), water loss during the evacuation of the vessels could alter the isotopic composition of water. This effect must be critically evaluated depending on the exact procedure used (i.e., vessel volume, evacuation time, geometry of capillary tube, etc.). To minimize problems with small-size samples, the analyst should: (1) use small-volume reaction vessels to minimize the loss of water


21

to the vapor phase, and (2) use the same amount of water for standard and sample waters, so that the effect of water loss will largely cancel out. In either case, the accuracy of ~SD values may not be as high as that of large-volume samples. On the other hand, the effect of isotopic exchange on 6D values of waters can be readily corrected by the following equation, similar to that used in the CO2-water equilibration method (Craig, 1957). [1.81

5corr -- (l+p / ot)6meas - (p / Ot)Sinit

where i~corr, 6meas, and Dinit a re 6D values of corrected, measured, and initial H2 gases against a working standard H2. The p is the molecular H2/water ratio, and c~= (D/ H)water / (D / H)H2 at equilibrium (3.81 at 25~ The overall D / H isotope exchange rate between H2 and water follows a first-order reaction, and the following equation is applied; f n o n - e q u i l - ( ~t - ~ ) / ( 6 0

- 6~) - exp(-kt)

ln(fnon-equil) - ln(6t - ~ ) / (~0 - ~ )

- -kt

[1.9a] [1.9b]

where fnon-equil is the fraction of non-equilibrated H2 gas (1.0 at t = 0, and 0 at t - ~), and 60, 6t, and ~ are the 6 values of H2 at the initial time, time t, and at the equilibrium time, respectively. In the presence of the hydrophobic Pt catalysts, the half-reaction time ( t l / 2 - 0.693.k-1, the time needed to reach a fnon-equil value of 0.5) is only 2.5 min (Horita et al., 1989). This means that after 25 min. (i.e., ten times the tl/2), the extent of isotopic equilibrium attained between H2 gas and liquid water is 99.9% (i.e., fnon-equil - 0.1%). Putting the beads in a Teflon bag slows down the isotope exchange rate (tl/2 - 40 min, Ohba & Hirabayashi, 1996). In contrast, isotopic equilibration takes 3 days using non-hydrophobic Pt-A1203 catalysts (Scrimgeour et al. 1993). The activities of different hydrophohic Pt catalysts have not yet been compared.

Special considerations There are several aspects to be considered for D / H ratio measurements with the H2-water equilibration method. The H2 gas used as a working standard in dual-inlet mass spectrometers should have a 6D value close to that of H2 equilibrated with the samples and standards. For example, the 6D value of H2 gas equilibrated with VSMOW and SLAP at 25~ is-740 and -850%o on the VSMOW-SLAP scale, respectively. Commercial H2 gas cylinders have a wide range of 6D values (-100 to -800%o), depending on manufacturing processes: electrolysis of water generally produces high purity gases with 6D values in the range of-100 to -200%o, while H2 from petrochemical processes (e.g., thermal cracking) has more negative 6D values (< -600%o). Because the D / H ratios of the H2 gas equilibrated with natural waters are very low (D/H = 0.004% at 25~ the relative contribution of H3 + (formed in the source) becomes a very large part of the total M / e - 3 (HD + and H3 +) signal in the mass spec-

22


trometer. The production of H3 + is proportional to the square of H2 pressure in an ionization chamber of a mass spectrometer, and is in the range of 5 to 10ppm per 10-9 A of M / e - 2 with many commercial gas-source isotope-ratio mass spectrometers. Thus, when D / H ratio measurements were conducted at 5 x 10-9 A of M / e = 2, there is 25 to 50 ppm contribution of H3 + to the approximate 80 ppm of HD +. Thus, attaining a small and, more importantly, very stable H3 + contribution during the course of a mass-spectrometry session is crucial to precise measurements of D / H ratios with the H2-water equilibration method. Possible gradual drifts of the H3 + contribution or other ion source conditions with time can be easily monitored and corrected by measuring the D / H ratios of H2 gas equilibrated with a standard water several times during a session. With precise temperature control (< + 0.05~ of the water/air bath and a linear time-correction of small drifts in the H3 + contribution, an external precision of + 0.5 - 1.0%o (lo) or better can be readily achieved (Brand et al., 1996). This precision is nearly as good as that of conventional uranium-reduction methods, and is better than that of a batch Zn-reduction method in many laboratories. A set of two or more laboratory standard waters, which are calibrated on the VSMOW-SLAP scale (as discussed in the section on calibration) should always be analyzed together with the samples. Since the H2-water equilibration method does not require any pretreatment of samples (except for removal of H2S) prior to analysis, this technique is a preferred method for various aqueous samples with high concentrations of dissolved salts and compounds. Scrimgeour (1995) used the H2-water equilibration method to determine D / H ratio of water in soils and plants by directly equilibrating with H2 gas without any extraction of water. A soil sample yielded results consistent with that of an azeotropic method, but twig samples from shrubs and trees did not equilibrate completely with H2. Incomplete isotopic exchange is due partly to the low catalytic activity of the PtA1203 used, and this technique needs further investigation. Koehler et al. (2000) determined 6D values of pore water in clay-rich core samples directly using the H2-water equilibration method. With the Hokko beads, isotopic equilibrium between H2 and pore water was attained within 4 h, and a high precision (< + 1%o) was obtained. The analytical methods for D / H ratios are summarized in Table 1.1.

1.3.2 Oxygen isotopes For oxygen isotope analysis of water and other aqueous samples with dual-inlet, gas-source isotope-ratio mass spectrometry, C02 is a preferred gas species. Oxygen gas is preferred for the measurement of all three isotopes (160, 170, and 180). Water can be directly injected to a mass spectrometer, but this causes a large memory effect and, in the case of saline waters, rapid corrosion of the filament (Wong et al., 1984). CO is used as an analyte in continuous-flow isotope-ratio mass spectrometry using methods where oxygen in water or organic materials is reduced ("pyrolyzed") with C to form CO.

Table 1.1 - S u m m a r y of analytical methods for D / H ratios of water and aqueous samples, which are currently in use or u n d e r development.

Method

Reagent

Reduction

U

Reduction

Zn

Reduction

Zn

Reduction

Cr

Reduction

Mg/Pt

Reduction

Mn

Reduction

C

Reduction

LiA1H4

Reduction

CaC2

Electrolysis Equilibration H2 / Pt

Reaction

T (~

Reaction time (min)

Sample size (~L)

Precision (%o, lc~)

U + 2H20--* UO2 + 2H2 in reactor Zn + H 2 0 ~ ZnO + H2 in reactor ( w / q t z sand) Zn + H 2 0 ~ ZnO + H2 in closed tubes

400-800

15-30

0.5-10

0.2-1.0

380-450

15-30

0.5-10

1.0-2.0

450-550

10-60

1-10

0.5-2.0

2Cr + 3 H 2 0 ~ Cr203 + 3H2 in reactor

850-950

1-2

1-4

0.5-2.0

Mg + H 2 0 ~ MgO + H2 in closed tubes M n + H20--* MnO + H2 in reactor / closed tubes C + H 2 0 ~ C O / C O 2 + H2 in Ni bomb or pyrolysis furnace LiA1H4 + 4H20 --* LiAI(OH)4 + 4H2 in closed tubes CaC2+H20 --* CaO+C2H2 in closed tubes H 2 0 --* H2 + 1/202 H20(liq) ~=~ H2

Comment

9 o r~

> Use of Toepler / charcoal / expansion 0.5- 1% m e m o r y effect Use of Toepler / charcoal / expansion 1 - 2% m e m o r y effect Zn selection, degassing, Zn / water ratio critical Dissolved salts / organic interfere Commercial a u t o m a t e d unit on-lined to MS 1% m e m o r y effect Dissolved salts / organic slightly interfere Need to prepare platinized Mg Mg / water ratio i m p o r t a n t Grain size critical (100-200 ~m optimum)

400

60

4-10

2 (?)

520-900

0.5-40

5-8

0.7-1.5

950

20

5

2.0

ambient

O

O ~. 9

Commercial a u t o m a t e d unit on-lined to MS

Complete electrolysis Commercial a u t o m a t e d unit on-lined to MS No sample pretreatment (except for H2S) t,J

24


1.3.2.1. Conversion methods

Many chemicals have been used to convert water for oxygen isotope analysis: guanidine hydrochloride (Boyer et al., 1961; Dugan et al., 1985; Dugan & Borthwick, 1986; Wong et al., 1987a), BrF5/C1F3/CoF3 (O'Neil & Epstein, 1966a; Bottinga & Craig, 1969; Blattner, 1973; Gulens et al., 1985; Dugan & Borthwick, 1986; Suvorova & Dubinina, 1996; Baker et al., 2002), and carbon with metal catalysts (Majzoub, 1966; Bariac et al., 1982; Brenninkmeijer & Mook, 1981; Ferhi et al., 1983; Gray et al., 1984; Edwards et al., 1994). BrF5 and C1F3 decompose water to 02, which is usually converted to CO2 with graphite. Other chemicals (Br2, HgC12-Hg(CN)2, and Na2S208) were also tested in the past. Recently, Meijer & Li (1998) improved an electrolysis method for precise 6170 (+ 0.07%o) and ~180 (+ 0.10%o) analysis of water. All these decomposition methods are designed to convert small amounts (< 10 ~L) of water quantitatively to 02, CO and CO2. Among these decomposition methods, the guanidine hydrochloride method is the only one routinely used in some isotope laboratories for the oxygen isotope analysis of water, due to the relative ease of its procedures and the high precision (+ 0.10%o) attainable. The carbon pyrolysis technique, which converts water exclusively to CO at high temperatures (1000 - 1300~ has recently been adapted for isotopic analysis of water, other fluids, organic matter, and minerals with a continuous-flow mass spectrometer (Brand et al., 1994; Werner et al., 1996; Begley & Scrimgeour, 1996, 1997; Koziet, 1997; Farquhar et al., 1997; Br6as et al., 1998; Loader & Buhay, 1999; Houerou et al., 1999; Kornexl et al., 1999a,b; Sharp et al., 2001); this new method is described elsewhere in this book (Volume II, Part 3, Chapter 6-2.3.6). In the guanidine hydrochloride method, water is directly converted to CO2 (Dugan et al., 1985; Wong et al., 1987a). About 100 mg of very pure guanidine hydrochloride is placed in a Pyrex tube, which is then evacuated and heated to melt the guanidine; the tube is then cooled, about 10 gL of water is frozen into the tube, any contaminants are pumped away, and the tubes are sealed with a torch. The sealed tube is then heated in a furnace at 260~ for 8 - 16 h (or longer), for the water and guanidine to react together to produce ammonia and CO2. As the tubes later cool below 70~ the gases further react to form solid ammonium carbamate (NH4NH2CO2). The cooled, pre-scored, reaction tube is then either placed inside an evacuated, ball-jointed, tube-cracking assembly (Dugan et al. 1985) or is mounted in an apparatus containing a flexible tube-cracker (Wong et al., 1987a). Both of these units contain a separate reservoir containing about 0.5 mL of 100% H3PO4. After the tube is cracked open, the unit is removed from the vacuum line and the lower part of the unit placed in an oven at 80~ for about an hour for the ammonium carbamate to react with the phosphoric acid, and decompose to produce CO2, ammonium chloride, and ammonium phosphate. The CO2 is then purified on a vacuum line and analyzed. The method can be fussy and special precautions to observe include purifying the guanidine by melting it twice, and never letting any of the other reagents directly contact the H3PO4. Erratic results, including anomalously high ~13C values of the CO2, can be caused by the accidental production of C12 gas. With BrF5 or C1F3, small amounts (5 - 10 ~L) of water are converted to 02 in the same manner as samples of silicates and oxides (O'Neil & Epstein, 1966a; Bottinga &


25

Craig, 1969; Blattner, 1973; Gulens et al., 1985; Dugan& Borthwick, 1986; Suvorova & Dubinina, 1996). Water samples have to be transferred quantitatively to an Ni bomb and the contamination of atmospheric vapor has to be avoided. Three to 10 times the stoichiometric amounts of the reagents are used, and the reaction can proceed readily at 80 - 350~ within a few minutes. The 02 is usually converted to CO2 with heated graphite. Alternatively, 02 can be directly introduced into an isotope-ratio mass spectrometer as an analyte, and the both 170/160 and 180/160 ratios of water can be determined. This method is useful for the study of non mass-dependent isotopic fractionation and extraterrestrial waters.

1.3.2.2 Direct introduction of water Majzoub & Nief (1968) first described an isotope ratio mass spectrometer for direct 180/160 measurement of water where the H2180 + and H2160 + ion currents were monitored. Although high precision results (+ 0.08%o) were achieved, the entire inlet and source units had to be heated to minimize memory effects. Hagemann & Lohez (1978) built a twin mass spectrometer system for simultaneous oxygen and hydrogen isotope measurement of water, using the direct injection method for 180/160 and online uranium reduction method for D / H (discussed above). A commercial version of the twin mass spectrometer (Aqua SIRA, VG, UK) was built and tested (Wong et al., 1984; Barrie & Coward, 1985). However, the direct introduction method has never become a routine technique. 1.3.2.3 C02-water equilibration The CO2-water equilibration method has been most widely used for the oxygen isotope analysis of large-volume water samples (hydrologic, agricultural, and biological fluids). This method, based on the classic work by Harold Urey and his colleagues (Cohn & Urey, 1938, Mills & Urey, 1939; 1940), was first applied to natural waters in the early 1950's (Baertschi, 1953; Dansgaard, 1953; Epstein & Mayeda, 1953). This method is often referred to as the "Epstein-Mayeda" method. The CO2-water equilibration method is a simple and very precise method for obtaining 6180 values of water. A few mL of water are placed in a vessel equipped with a stopcock. The air is removed by successive freezing, evacuating, and then thawing the water sample. A few 100 ~moles of CO2 are then frozen into the vessel, the stopcock closed, and the vessel placed in a water or air bath where the temperature (usually 25~ is controlled to about + 0.1~ The vessels are periodically shaken to speed up the equilibration rate. After several hours or days, the vessels are returned to a vacuum line, the CO2 is separated from the water by freezing the water first in liquid nitrogen and then a dry ice-alcohol slush, and then analyzed. Roether (1970) eliminated the time-consuming freezing and thawing steps that were required to remove air from the vessels without fractionating the water, by evacuating air from the reaction vessel through a capillary tube. This innovation minimized oxygen isotope fractionation of the water to about 0.1%o. With this technique, Fairbanks (1982) achieved a high precision (_+0.03%0) for seawater samples. The Roether technique laid the foundation for the automated oxygen and hydrogen isotope analysis of water samples by on-line gas equilibration (Buzek, 1983; Chiba et al., 1985;

26


Brenninkmeijer & Morrison, 1987; Wong et al., 1987b; Horita et al., 1989). The CO2water equilibration technique was also modified for rapid manual preparation of multiple water samples, using inexpensive labwares such as plastic syringes and preevacuated glass vials (Yoshida & Mizutani, 1986; Graber & Aharon, 1991; Socki et al., 1992).

Exchange reactions The overall oxygen isotope exchange between gaseous CO2 and water proceeds via two reactions. ke

Gas exchange:

CO2(g ) ~ CO2(aq ) kh

[1.10] kl

m

Hydration / dissociation: CO2(aq ) + H20(1 ) ~ H 2 C O 3 r H + + H C O 3

[1.11]

The overall reaction is, ktot C1602(g) + H2180(1) r C18016 O(g) + H2160(1)

[1.12]

The hydration reaction (kh) is orders of magnitude slower than the dissociation reaction (M), and the enzyme carbonic anhydrase is known to catalyze the hydration reaction. When a reaction vessel is not shaken, the gas exchange (ke) is rate-limiting in the overall reaction (Roether, 1970). .~ However, with increasing fie~ quency of shaking (i.e., an ~ increase in the effective sur~ face area and mixing of the ~ solution), the overall exchange ~ rate increases rapidly, and the ~ hydration reaction becomes ~ rate-limiting (Figure 1.4:). ~• Roether (1970) described the ~ overall rate constant (ktot) for ~ each case; o ~

Figure 1.4 - Effect of shaking on overall rate constant (Mot, min-1) of 180/160 exchange in CO2-water equilibration method. With increasing shaking cycles, the reaction changes from gas exchange-limited to hydration-limited processes. After Roether (1970).

O

27

Stable Isotope Analysis of Water and Aqueous Solutions ... ktot - ke'CO2(diss) / CO2(g)

(gas exchange limited)

[1.13]

ktot - 1 / 3"kh'CO2(diss) / [CO2(g) +CO2(diss)]

(hydration limited)

[1.14]

where CO2(diss) is total dissolved CO2 (CO2(aq) + H2CO3 + HCO3-). The characteristic time (ktot -1) decreased an order of magnitude with increasing shaking frequency, from that of the gas exchange reaction to that of hydration (18 min.) at 18~ (Roether, 1970) (Figure 1.4). The ktot also depends on the amount of water. Under normal conditions, vigorous shaking decreases significantly the half-life time of the isotope exchange (tl/2 - 0.693"ktota1-1) from 100 -400 min. to 10 - 20 min. Thus, the oxygen isotope equilibration between gaseous CO2 and water can be complete after 2 - 3 hr (99.9% equilibrium after ten times h/2) rather than I - 2 days, when not shaken vigorously (Figure 1.5). For detailed discussion on the kinetics of oxygen isotope exchange between gaseous CO2 and water, see Roether (1970) and Taylor (1973). The required equilibration time must be established in each laboratory for the conditions and types of samples anticipated, rather than relying on published times. The oxygen isotope exchange between CO2 and alkaline waters (pH = 9), where carbonate ion (CO32-) iS the dominant inorganic carbon species, is very slow (Mills & Urey, 1940). A large fraction of CO2 gas introduced into the vessel is absorbed into the water. For these reasons, the pH of alkaline waters needs to be adjusted to below 6 - 7 with acid (anhydrous H3PO4 or other acids). High salt contents also slow down the isotope exchange rate, and vigorous shaking of the vessel becomes necessary (see the section below).

Figure 1.5 - Examples of 180/160 exchange in CO2-water equilibration method. Shaking decreases a half-life time (tl/2=ln2.ktot -1) by a several factor. Salinity and pH affect significantly overall isotope exchange time. The exact half-life time also depends on many factors (geometry of a vessel, shaking amplitude and frequency, amount of water and CO2, and chemistry of water). Data of pure water from Gonfiantini (1981). A Dead Sea brine is of Na-Ca-K-C1 type with 320 g/L salinity (NaC1 - 1.95 molal, KC1 - 0.15 molal, MgC12- 2.00 molal, and CaC12- 0.49 molal).

28


The equilibrium oxygen isotope fractionation factor ~ = (180 / 160)CO2(g) / (180 / 160)H20(1) is 1.0412 at 25~ (Hut, 1987), and its temperature dependence is 0.2%0/~ Thus, temperature control of + 0.1~ is adequate for precise measurement. The oxygen isotope composition of the dissolved CO2 species (CO2(aq), H2CO3, HCO3-) is different from that of gaseous CO2, and degassing of dissolved CO2 should be minimized during the extraction from the vessel. Depending on the ratio of CO2 to water and their initial isotopic compositions, the isotopic composition of the water generally changes during the isotopic exchange with CO2. This effect should be corrected, if necessary, as with the H2-water equilibration method (equation [1.8])(Craig, 1957). The isotopic composition of water samples should be measured along with laboratory standard waters, which are in turn calibrated on the VSMOW-SLAP scale (see discussion below).

Small samples For small amounts of water (< 0.1 mL), water loss during evacuation through a capillary tube of an automated gas equilibrator a n d / o r evaporation of water inside a reaction vessel (typical volume 10 - 20 mL) becomes significant. Isotopic exchange with CO2 gas also causes a significant change in the isotopic composition of water. To circumvent these problems, Kishima & Sakai (1980) developed a "micro-scale" CO2water equilibration method (a.k.a the MCE method) for small amounts (0.5 - 10 ~L) of water using small-volume (0.5 mL) reaction vessels and applying the corrections required to account for isotopic exchange and unavoidable evaporation. Rozanski et al. (1987) tested this method. Ohba (1987) and Socki et al. (1999) further improved this technique using a sealed Pyrex glass tube as a reaction vessel. The correction equation for evaporation and isotope exchange is; ~corr -- C~. p(~meas - ~init) + 103"(1+10-3"6meas)[(1 + (a'-l)~,] - 103

[1.15]

where p is the ratio of oxygen atoms in CO2 to that in H20 in the vessel, and ~, is the ratio of vapor to liquid water in the vessel. The a - (180/160)CO2(g)/(180/160)H20(1) = 1.0412 (Hut, 1987) and a' - (180/160)H20(v) / (18O/160)H20(1 ) - 0.99074 at 25~ (Horita & Wesolowski, 1994). The 5cor, ~meas, and ~init are 6180 values of corrected, measured, and initial CO2 gases, respectively, relative to working standard CO2. When the evaporation of water is negligible (~, ~ 0), this equation is the same as equation [1.8]. The tl/2 ranges from 2 to 6 hr for 0.5 to 10 ~L of water (Kishima & Sakai, 1980), which is comparable to those of the original CO2-water equilibration method. Contamination of water from moisture and other sources has to be avoided while loading the sample into a reaction vessel. For accurate and precise data, the transfer and separation of water and CO2 have to be complete. Furthermore, the amount of water and CO2 introduced to a reaction vessel must be measured precisely to make the needed corrections. With these precautions and the above corrections, the 5180 values of water samples as small as 0.5 ~L can be determined precisely (lo - 0.05 0.1%o). Using CO2 with a 6180 close to that expected after equilibration can significantly shorten reaction times and improve analytical precisions. In the absence of liquid water, the isotope exchange between CO2 and water vapor is prohibitively slow,


29

and the reaction has to be catalyzed by Pt at high temperatures (Dostrovsky & Klein, 1952). One important advantage of the micro-scale CO2-water equilibration method is that the 6D value of the same water sample can be determined using a reduction (Zn, U, etc.) method after the removal of the CO2 gas. In fact, if water samples prepared for the micro-scale CO2-water equilibration method cannot be weighed (or otherwise measured) to sufficient precision, the yield of H2 gas must be measured to correct the data. Most of the decomposition methods discussed above for the ~5180 analysis of small amounts of water cannot be used for 6D analysis of the same water sample. An exception is the high-temperature (1450~ carbon pyrolysis method, by which 6D and 6180 analysis can be done on a single sample (Motz et al., 1997). Water is converted to H2 and CO within an elemental analyzer, which are then separated by GC and introduced into a mass spectrometer with He carrier gas. A peak-jump between H2 and CO enables a sequential measurement of 6D and 6180 values. Splitting small water samples into two or more aliquots for separate analyses is not recommended because of probable fractionation associated with the splitting. Scrimgeour (1995), Hsieh et al. (1998), McConville et al. (1999), and Koehler et al. (2000) applied CO2 equilibration methods to water in soil and plants. Oxygen isotope equilibrium between CO2 and soil water was attained within I - 2 days. The precision ranged from + 0.12%o (McConville et al., 1999) to + 0.3 - 0.4%o (Hsieh et al., 1998; Koehler et al. (2000). The CO2-water equilibration method was also successfully adopted for continuous-flow mass spectrometry for as small as 40 ~L of water, using septum-sealed glass vials (Analytical Precision, Finnigan MAT, Micromass, PDZ Europe; Fessenden et al., 2002). Leuenberger & Huber (2002) developed a novel on-line continuous-flow mass spectrometric method for 180/160 analysis of ice cores. CO2 was injected into a stream of liquid water generated by layer-by-layer melting of an ice core. After dissolved CO2 reaches an (near) isotopic equilibrium with water at 50~ dissolved CO2 was separated through a gas-permeable membrane and carried with a He to a mass spectrometer. Leuenberger & Huber (2002) attained a precision of better than 0.1%o with a resolution of I - 3 cm of an ice core. The analytical methods for 180/160 ratios are summarized in Table 1.2.

1.4 Effects of dissolved compounds High concentrations of dissolved compounds in water (brine, biological and agricultural fluids) can pose serious analytical problems. In the decomposition methods for 6D and 6180 analyses of water and other aqueous samples discussed above, water must first be extracted from the samples. If the water is injected directly into a reaction unit, serious contamination and memory effects would result. The most common method for the extraction of water is vacuum distillation. However, it is not always easy to extract water quantitatively from high-salinity brines, juices, and blood even with prolonged heating at elevated temperatures. Incomplete extraction of water from

Table 1.2 - S u m m a r y of analytical m e t h o d s for 1 8 0 / 1 6 0 ratios of w a t e r a n d a q u e o u s samples, w h i c h are currently in use or u n d e r d e v e l o p m e n t . Method

Reagent

Reaction

Equilibration

CO2

H20(liq) 100

0.05- 0.1

C o m m e r c i a l a u t o m a t e d unit on-lined to MS S h a k i n g facilitates the reaction Slow reactions for brines

Equilibration

CO2

H20(liq) ~=~ CO2

2 0 - 30

3 - 10 hr

0.5-10

0.05-0.1

For small s a m p l e s Corrections for v a p o r i z a t i o n a n d CO2 e x c h a n g e

260 80

8 - 10hr I hr

10

0.1 - 0.15

5 - 10

0.1 - 0.2

Conversion

guanidine

Conversion

BrF5(C1F3)

Conversion

Electrolysis

graphite / diamond

g u a n i d i n e + 2 H 2 0 --* N H 4 N H 2 C O 2 + NH4C1 N H 4 N H 2 C O 2 + NH4C1 + H3PO4 ~ CO2 + NH4C1 + (NH4)3PO4 H 2 0 + BrF5 --* 1 / 2 0 2 + BrF5 + 2HF 0 2 + C ~ CO2

T (~

80 - 350

Few to 30

Sample size (~L)

Precision (%o, l o )

Comment

R e a g e n t s explosive r

C + H 2 0 --~ C O / C O 2 + H2 950 - 1200 (in Ni b o m b or pyrolysis tube) 350 2 C O ~ CO2 + C H 2 0 ~ H2 + 1 / 2 0 2

Reaction time (min)

?

20 15

2 - 10

40

1000

0.2 - 0.3

0.1

C o m m e r c i a l a u t o m a t e d unit on-lined to MS

r~

Partial (lg), chemical analysis by X-ray fluorescence, atomic

Conventional and Less Conventional Techniques for Hydrogenand OxygenIsotope ...

47

absorption spectrometry and/or inductive coupled plasma techniques can be performed. The latter techniques usually can detect smaller amounts of contamination than XRD or TEM. DTA-TGA analysis and FTIR spectrometry may also be very useful to identify and quantify minerals present in mixtures of clays, in particular clay minerals closely related chemically or polytypes (i.e., kaolinite, dickite, nacrite). Yields of organic carbon removal can be analyzed using conventional techniques, e.g., coulometric titration. Purity checks are most critical as very small amounts of contamination may result in significant shifts of measured isotopic compositions from those of the clays and erroneous interpretations (Girard & Fouillac, 1995).

Effects of separation treatments on the isotopic composition of clays Most of the treatments presented above have been verified not to affect significantly (i.e. beyond analytical uncertainty) the isotopic composition of clays. Many of these tests were done in a variety of studies, over the years in a patchy fashion, and using different experimental conditions and different clay reference materials depending on availability for each particular study. Many were conducted at a time when refined mineralogical techniques for the characterization of clay minerals were not routinely used. A comprehensive validation of each of the different chemical treatment is yet to be carried out on well characterized international clay standards with the ambition to investigate specimen of variable chemistry and/or crystallinity from all groups of clay minerals. The current state of knowledge with regards to the absence or existence of significant isotopic effect for most common clay minerals is summarized in Table 2.1. References of relevant studies in which demonstrative validation tests were reported are indicated or shown in the section Appendix. The reader is referred to these publications for the details of the experimental conditions. Investigators with an interest in modifying experimental conditions of any particular chemical treatment or in applying a chemical treatment to clay samples significantly different (chemically- or crystallinity-wise) from those used in the reference study are advised to conduct their own validation tests.

2.3 Isotopic analysis of clay minerals 2.3.1 Hydrogen isotopes Clay minerals have various reservoirs of hydrogen. These include hydrogen in hydroxyl groups, adsorbed water and in some cases interlayer water. Adsorbed and interlayer water of clays exchange isotopically rapidly with atmospheric vapor at ambient temperatures (Moum & Rosenqvist, 1958; Savin & Epstein, 1970). Consequently, they must be removed from OH hydrogen prior to analysis. For most clay minerals, pre-heating in vacuum at temperatures of about 200~ for two or more hours removes most of these waters without isotopic exchange with the structural OH hydrogen (Savin & Epstein, 1970). However, evidence for significant hydrogen isotope exchange between interlayer water and OH-hydrogen during degassing of halloysite were detected by Lawrence (1970). Additionally, recent experimental studies (Longstaffe & Mizota, 1999; Hsieh & Yapp, 1999) document rapid hydrogen isotope exchange between water and halloysite, especially the hydrated 10/~ form, even at

Table 2.1 - Effect of selective dissolution and other treatments on the isotopic composition of clays and associated minerals Mineral

Treatment

Effect on 6180

Effect on 6D

Recommended o u t g a s s i n g T~

U s e f u l references

Illite

Ac-Ac, H202, DCB, P1-Ash

none

none

200-250~

Illite/smectite Smectite

1N-HC1 Ac-Ac, H202, DCB Ac-Ac, H202, DCB, P1-Ash

none none none*

? none none-strong

200-250~ 180-250~

none none* none none none none small

none none none-slight small ? none ?

200-250~ 200-250~

Halloysite

Ac-Ac, H202, DCB Ac-Ac, H202, P1-Ash DCB Am-Ox 1N-HC1, 6N-HC1 Na-polytungstate densimetry H202, DCB

100~

Eslinger (1971), Yeh (1974, 1980) Hogg et al. (1993), Taieb (1990) Eslinger (1971) Eslinger (1971), Yeh (1974, 1980) Yeh (1980), Taieb (1990), Stern et al. (1997), Hyeong & Capuano (2000) Eslinger (1971), Yeh (1980) Yeh (1980), Ta~'eb (1990) Bird et al. (1992) Giral (1994), Appendix Giral (1994), Appendix Eslinger (1971), Bird et al. (1992), Giral (1994) Appendix, Girard et al. (2000) Hsieh (1997), Hsieh & Yapp (1999)

Goethite Gibbsite

5M N a O H DCB

none none

strong none

100-120~ 110-130~

Yapp (1991), Chazot (1997) Bird et al. (1989)

Chlorite

Kaolinite

Ac-Ac = buffered acetic acid 1N; H202 = h y d r o g e n peroxide 30%; DCB = Na-dithionite-citrate-bicarbonate; Am-Ox = a m m o n i u m oxalate; P1-Ash = plasma asher; * 9presumably because of the absence of effect on 6D value.

4~ oo

r

to !

4~ r

t


49

room temperature. Thus, D-H ratios of halloysite obtained by conventional techniques are probably only of limited value. Minor D-H exchange problems during degassing have also been reported by Lawrence (1970) and Fagan & Longstaffe (1997) for some smectites. Marumo et al. (1995) showed that the conventional degassing temperature of 200~ may not be sufficient to remove completely interlayer water from some smectites. Especially, poorly crystallized smectites seem to retain considerable amounts of interlayer water at temperatures at which dehydroxylation starts. Residual interlayer water after incomplete degassing is strongly enriched in deuterium with 6DSMOW values up to + 29%o (Marumo et al., 1995). Thus, significant contamination from residual D-enriched interlayer water has to be expected in the case of incomplete interlayer water removal. Thermogravimetric or differential thermogravimetric analyses in vacuum, stepwise heating experiments and precise calculation of hydrogen yields are very useful to constrain appropriate temperature cuts for separating interlayer and hydroxyl water of such samples. Dehydroxylation of clay minerals under vacuum generally starts at temperatures above 350 to 400~ (Brindley & Lemaitre, 1987). The previously degassed samples are further heated without exposure to the atmosphere either with a resistance furnace or an induction oven to temperatures above 900~ The expelled water is collected with a liquid nitrogen trap. Clay minerals with high Fe2+ concentrations, e.g. some chlorites and smectites, can produce significant amounts of free hydrogen during dehydroxylation which has either to be oxidized to water using for example a CuO furnace or collected with a Toepler pump. The dehydroxylation water is subsequently reduced to hydrogen using, for example, hot uranium, zinc or other metals (e.g., Bigeleisen et al., 1952; Vennemann & O'Neil, 1993; see Volume IL Part 3, Chapter 1-2.3/12 of this book) and its isotope composition is measured. Analytical reproducibility is typically better than +2 %o for kaolinites, but often worse than that for smectites and some illites.

2.3.2 Oxygen isotopes The conventional method to extract oxygen from clays is the fluorine oxidation, or fluorination, technique developed by Taylor & Epstein (1962) and Clayton & Mayeda (1963). Typically ~ 10 mg of clay is reacted with excess F2 or BrF5 in a nickel tube at 500-600~ overnight. C1F3 has also been used as a fluorinating agent with some success (Borthwick & Harmon, 1982). The liberated oxygen is cleaned of extraneous reaction gases by standard cryogenic techniques and converted to CO2 over hot graphite prior to analysis on the mass spectrometer. Because clays are not refractory the fluorination reaction under such conditions is complete. Analytical uncertainty, based on reproducibility of measured 6~80 values, typically is +0.2-0.3 %o. As for hydrogen analysis, one practical difficulty of oxygen analysis of clays is to remove adsorbed water, and interlayer water when dealing with expandable clays, prior to fluorination reaction. Most of the non-structural water can be removed by placing the samples in a dried atmosphere (drybox with P205) for 12 hours to 24 hours (Savin, 1967, Savin & Epstein, 1970). This step can be extended to a few days for minerals such as smectites, smectite-bearing mixed-layers, vermiculite and halloysite

50

Chapter 2 - H.A. Gilg, J.-P. Girard & S.M.F. Sheppard

in order to reduce the amount of residual interlayer water as much as possible. Additional removal of non-structural water is done by heating clay samples in the nickel tubes under vacuum to a temperature of 150-250~ for I to 2 hours (Savin, 1967; Savin & Epstein, 1970). Recommended outgassing temperatures for common clays and associated minerals are given in Table 2.1. Alternatively, a prefluorination treatment of a few minutes at room temperature with a small amount of reagent may be applied (Clayton & Mayeda, 1963). Prefluorination at higher temperatures and/or for longer periods of time is not advisable because it might cause partial reaction of the clay and loss of structural oxygen (Hamza & Epstein, 1980; Hogg et al., 1993; Girard & Savin, 1996). On the other hand, if the prefluorination treatment is too short, water adsorbed on clays and on the internal wall of the Ni tubes may not be completely removed. Best suited conditions (time, amount of reagent, temperature) should be determined by the investigator. The possibility of using the laser fluorination techniques recently developed to extract oxygen from small amounts (< lmg) of silicates and oxides (Sharp, 1990) has not yet been extensively investigated on clays. The laser technique consists of loading many samples in a single reaction chamber, and reacting each of them in a sequence by heating with a laser in the presence of reagent (F2 or BrF5). Because clays are susceptible to passive fluorination (partial reaction at room temperature) the laser technique does not appear well suited. Cross-contamination of oxygen from the different samples present in the chamber may occur. We know of a few attempts to laser fluorinate clays with variable degrees of success. Unpublished preliminary experiments by one of the authors (JPG) using the 1 - 10 ~m fraction of Macon kaolinite (Girard & Savin, 1996) gave encouraging results. Average 6180 of thirteen laser determinations performed in three different loads was within 0.1%o of the conventional 6180 value and the associated standard deviation was 0.4%o. Even better results with an analytical precision similar to the conventional fluorination technique (0.2 to 0.3 %o) were reported from non-swelling clay minerals, such as kaolinite, dickite or illite analyzed at the Scottish Universities Research and Reactor Centre, East Kilbride (Maliva et al., 1999; A.E. Fallick, 2000, pers. comm.). Addition of an airlock module to a standard laser fluorination system allows samples to be transferred and reacted in the reaction chamber individually (Spicuzza et al., 1998b). This approach yielded good result for whole rock powders containing clays and holds promise for laser analysis of clay minerals (J.W. Valley, 1998, pers. comm.). The potentials of in-situ oxygen isotope measurements of clay minerals using secondary ion mass spectrometry have been explored by Williams & Hervig (1997) and Williams et al. (2001a). Their preliminary study on two standard clay minerals, montmorillonite (SWy-1) and kaolinite (KGa-1), however, revealed large analytical errors (~1 to 3 %o, lo) and mineral-dependend instrumental mass fractionations.

2.3.3 Intracrystalline fractionation of oxygen isotopes Two types of structural oxygen can be distinguished in clay minerals: hydroxyl oxygen (OH oxygen) and non-hydroxyl oxygen (non-OH oxygen). Hydroxyl oxygen mainly occurs in A1-OH, Fe-OH and Mg-OH bonds located in the octahedral sheets of


51

clays and in the brucite-like layer of chlorites. It represents 45% of the stoichiometric oxygen in kaolinite and chlorite, and 17% in illite. The non-OH oxygen occurs in M-OM bonds, where M is a cation other than H (commonly Si, A1, Mg, Fe...), and is found in the tetrahedral sheets and in the bonds bridging tetrahedral and octahedral sheets. Because the chemical bonds in which these two types of oxygen are involved are different, their isotopic composition (180/16 0 ratios) must be different (Taylor & Epstein, 1962). The fractionation between OH and non-OH oxygen is referred to as intracrystalline oxygen isotope fractionation and should be expected to be temperature dependent. In principle it ought to be possible to determine the temperature and the 6180 value of the water from which a clay mineral formed, directly from the 6180 values of OH and non-OH oxygen. Intracrystalline fractionation of oxygen isotopes in clays therefore constitutes a potential single mineral geothermometer. Two analytical approaches, partial fluorination and thermal dehydroxylation, have been used in an attempt to measure the intracrystalline fractionation of natural clays (Savin, 1967; Hamza & Epstein, 1980; Bechtel & Hoernes, 1990; Girard & Savin, 1996). In the partial fluorination approach, the clay sample is only partly reacted by use of an insufficient amount of reagent, an insufficient reaction temperature or an insufficient reaction time. This approach is based on the assumption that OH groups react faster than the remainder of the structure, and that OH oxygen is liberated prior to non-OH oxygen (Savin, 1967). In the thermal dehydroxylation approach, the clay is dehydroxylated by heating in vacuum and the liberated water is collected (by freezing) and analyzed isotopically (Bechtel & Hoernes, 1990; Girard & Savin, 1996). To be successful, the dehydroxylation process must occur with no kinetic fractionation during the liberation of water (only 50% of the OH oxygen is driven off as water) and no exchange between OH and non-OH sites during diffusion of the water through the particles. The results of analytical studies (Savin, 1967; Hamza & Epstein, 1980; Bechtel & Hoernes, 1990; Girard & Savin, 1996) agree with theoretical calculations (Zheng, 1993) from a qualitative standpoint. All indicate that OH oxygen is significantly depleted (up to a few tens of per mil) in 180 relative to non-OH oxygen in a same clay. They also suggest that intracrystalline fractionation decreases with increasing temperature of formation (Hamza & Epstein, 1980; Bechtel & Hoernes, 1990). However, the accuracy of intracrystalline fractionation measurements conducted with each of the two approaches remains questionable. Intracrystalline fractionations measured for similar low temperature (supergene) kaolinites in four different studies range from 17 to 27 %o, and are much smaller than values of ca. 40 %0 predicted by theoretical calculations (Zheng, 1993). In an extensive and detailed study of kaolinite and dickite, Girard & Savin (1996) showed that both partial fluorination and thermal dehydroxylation were associated with potential flaws previously overlooked and related to intrinsic properties of the clay samples, including particle size distribution and crystallinity. Partial fluorination experiments conducted under variable conditions (temperature, reagent amount, reaction duration) suggested that liberation of OH and non-OH oxygen was simultaneous and lead to unreliable determinations of the intracrystalline fractionation. Thermal dehydroxylation experiments yielded consistent results and was con-

52


cluded to be a more promising approach than partial fluorination, provided the less than I gm size fraction of the clay was used. A similar comparative study conducted on serpentine by Plas & Frfih-Green (1997) lead to the same conclusions. Although application of the approach to natural systems (Bechtel & Hoernes, 1993; Plas & Frfih-Green, 1997; Harris et al., 1999) and experimental studies (Fortier et al., 1994) seems to give reasonable results, additional developmental work is necessary before the method can be routinely used and applied as a single-mineral geothermometer.

2.3.4 Controlled isotope exchange technique The controlled isotope exchange technique (CIE) was developed by Labeyrie & Juillet (1982) for oxygen isotope analysis of biogenic opal (diatom silica) and was later applied to a Mg-rich smectite (stevensite) by Escande et al. (1984) . The method is based on the concept that such minerals contain two reservoirs of oxygen. One that isotopically exchanges readily with the ambient water (or water vapor) even during preparation and is related to interlayer water and silanol groups (Escande et al. 1984). The second reservoir does not show such an easy isotope exchange with the environment and its isotope composition reflects the conditions of mineral formation. In the CIE method, the weakly bound oxygen is exchanged with water vapors of known isotopic composition under controlled conditions (200~ and 24 h for stevensite) using a special setup (Figure 2.2). Following the controlled isotope exchange experiment, the sample is rapidly calcinated at 1000~ (and thus dehydroxylated) to prevent re-equilibration with water vapor during loading in a conventional silicate oxygen isotope extraction line. The precision thus obtained is +0.3 %o. From the experiments using isotopically different water vapors, it is possible to calculate the fractionation factor between easy exchangeable oxygen and water vapor at 200~ and additionally the


53

percentage of exchangeable oxygen in the clay. The latter varied between 40% and 2.5% for stevensites synthesized at temperatures between 25 and 400~ (Escande, 1983). Because the oxygen isotope values of the dehydroxylated clay mineral is measured, the results of this technique cannot be easily compared to samples analyzed by the conventional degassing method (section 2.3.2). In addition, sample dependent isotope exchange between OH oxygen and non-OH oxygen during calcination cannot be excluded (Girard & Savin, 1996). A similar controlled isotope exchange approach was developed for D / H measurements of goethites (Yapp & Poths, 1995) and halloysites (Hsieh & Yapp, 1999). The CIE method is very time consuming and cannot be regarded today as a standard technique for clay minerals, but it has potentials for clay minerals with considerable amounts of easy exchangeable oxygen or hydrogen, such as halloysites or some smectites.

2.4 Isotope analysis of pore water in clay-rich rocks The isotope analysis of pore waters in clay-rich rocks is very important in understanding the movements and chemical evolution of interstitial waters in low-permeability rocks or soils, but also their effects on the alteration of rocks. Liquid water occurs in very different "pools" or "compartments" within such rocks: a) free, mobile, interstitial or pore water, b) surface-adsorbed water and c) structural water in hydrous minerals and organics. Some pools of loosely bound water in clay-rich rocks are physically not separable, such as surface-adsorbed water and structural interlayer water in swelling clays or channel water in zeolites. These are here referred to as "sorbed" water. It has been demonstrated that sorbed water fractionates hydrogen and oxygen isotopes with pore waters in a rock (e.g. France-Lanord & Sheppard, 1992; AraguasAraguas et al., 1995). A technique of estimating the isotope composition of "sorbed" water in clay-rich rocks by mass balance calculations is given in France-Lanord & Sheppard (1992) and will not be discussed further here. Various techniques have been developed to analyze the isotopic composition of pore waters from low-permeability rocks (Anonymous, 2000). Each one has its potentials, but also its limitations. The techniques and some of their problems are discussed in the following paragraphs.

2.4.1 Physical techniques of water extraction High-pressure squeezing techniques to extract pore waters from clay-rich sediments are documented for example in Manheim & Sayles (1974), Entwistle & Reeder (1993), Cuevas et al. (1997), B6ttcher et al. (1997a) and here (Figure 2.3). Maximum stresses obtained using a hydraulic press are in the order of 100 to 200 MPa. The efficiencies of pore-water extraction increase with the initial moisture content (measured as the weight percentage of water loss during oven-drying at 105~ and vary between 0 and 70%. The minimum moisture content necessary to extract any pore water from an unconsolidated sediment using high-pressure squeezers is about 7 to 15% (Entwistle & Reeder, 1993; B6ttcher et al. 1997a), but may be higher for consolidated shales. Ultra-centrifugation (e.g. Litaor, 1988; Walker et al. 1994) is an alternative to squeezing for relatively water-rich and unconsolidated samples. Although these physical extraction techniques may not allow a complete extraction of the pore waters, they are generally less affected by fractionation processes compared to distillation processes. Ultrafiltration effects by a clay membrane (Coplen & Hanshaw,

54


1973), however, may result in some minor isotope fractionation (< 3%o for 6D and < %0 for 6180) during extraction from clay-rich samples. 2.4.2 Distillation methods In the vacuum distillation method (e.g., Jusserand, 1980), the sample is heated to temperatures ranging from 35 to 200~ under vacuum and the released water collected in a vessel cooled with liquid nitrogen. A variant of this technique has been used by Moreau-Le Golvan et al. (1997) for indurated shales with low water content (< 5 wt. %). The samples are crushed to variable grain sizes, mostly less than 5 mm. Contact time with the atmosphere was varied (15-60 min.). The chosen vacuum distillation temperature was 60~ and the duration 14 hours. A significant effect of grain size and atmospheric exposure time was detected mainly for 6180, but not for 6D values. The mean absolute standard deviation obtained for 6D values of 44 pore water extracts from a shale was + 1.3 %0. In the azeotropic distillation method (e.g., Revesz & Wood, 1990; Leaney et al., 1993) an immiscible hydrocarbon (e.g. toluene, hexane or kerosene) is added to the soil sample forming an azeotrope mixture. Heating of these mixtures leads to boiling at temperatures of less than 100~ The resulting vapor mixture is condensed and the extracted soil water forms an immiscible liquid in the condensation tube. Both methods require complete extraction of the water as the distillation process involves major isotopic fractionation if the amount of water remaining in the samples is greater than 2% of the original present (Araguas-Araguas et al., 1995). Both temperature and extraction yield affect the isotope composition of the distillate. Various comparative tests of the different distillation techniques have been published (e.g., Ingraham & Shadel, 1992; Walker et al., 1994; Araguas-Araguas Figure 2.3 - Hydraulic press for the extraction of pore waters from clay-rich rocks.

Conventional and Less ConventionalTechniques for Hydrogenand OxygenIsotope ...

55

et al., 1995). The results show that with increasing clay mineral content, decreasing grain size and decreasing water content the inter- and intra-laboratory discrepancies of isotope measurements increase. High-temperature distillation yields larger deviations than low-temperature distillation techniques. The main problem with these techniques is that in addition to mobile or free pore waters also some of the adsorbed and interlayer water with its distinct isotopic composition (France-Lanord & Sheppard, 1992) is extracted. 2.4.3 Direct equilibration method The direct equilibration method for the analysis of 6180 values of soil waters is described in detail by Hsieh et al. (1998). It is a variant of the classic CO2-water equilibration technique (e.g. Epstein & Mayeda, 1953). A soil sample of 6 to 10 g, which is sterilized by exposure to gamma rays, is placed in a glass vessel on a vacuum line. After freezing the soil sample with ethanol-dry ice slurry, the vessel is opened to vacuum. Several thawing-pumping-freezing cycles are needed to release non-condensable soil gases. A defined volume of tank CO2 is added to the frozen soil sample and is allowed to equilibrate for 12 to 48 hours with the sample at a thermostated water bath temperature of 25~ An aliquot of the CO2 is taken with a gas-tight syringe through a rubber septum and injected into the mass spectrometer. The calculation of the soil water 6180 value using mass balance consideration and appropriate fractionation factors are documented in the Appendix A of Hsieh et al. (1998). The reproducibility of the method is between 0.3 and 0.4%o. McConville et al. (1999) and Koehler et al. (2000) developed automated techniques for measuring both 6180 and 6D values of porewaters by direct CO2 and H2 equilibration. The D / H composition of pore waters from clays-rich samples is determined by equilibration of the sample with H2 during 4 hours using a Pt catalyst at 25~ The 6D results are with 1%o of those derived from waters collected from piezometers (Koehler et al., 2000). 2.4.4 Radial diffusion cell method This relatively new method was developed by Van der Kamp et al. (1996) to determine the isotopic composition, chemistry and effective porosities for groundwaters in aquitards. The method is based on diffusive exchange between interstitial water in a cylindrical samples and water in a reservoir placed along the axis of the sample. The sealed diffusion cell has an outer diameter of 7.2 cm and a length of ca. 12 cm. The drilled out central reservoir (diameter of 2.5 cm, about 25 mL) is stabilized with a porous or perforated plastic liner and filled with isotopically distinct distilled water (test fluid). Equilibrium between test fluid and pore water is reached after about 60 days in the investigated samples. The equilibrated reservoir water is removed with a syringe and measured for its isotope composition. The central reservoir is refilled again with the initial test fluid. The equilibration, extraction and refilling is repeated several (about 4 to 5) times. Mass balance calculations and plots of net mass added versus isotope composition of the repeated experiments allow the determination of the initial isotope composition of the pore water. Experiments at room temperature (25~ and sampling temperatures (4~ showed no discernible effects. The radial diffusion method was tested by Van der Kamp et al. (1996) for D / H of pore water in a clay-rich glacial till. The cell results are within 2 to 3 %o for 6D of accumulated pore

56


waters collected in situ from piezometers.

2.5 Isolation and isotope analysis of minerals intimately associated with clays 2.5.1 Silica minerals Quartz, cristobalite, tridymite or opal can be significant constituents of finegrained sediments, such as clays, soils, or shales, and argillically altered hydrothermal systems. The isotopic compositions of these silica minerals can be highly variable and yield information about their genesis (e.g., Sheppard et al., 1971; Blatt, 1987; Mizota 1996), but are also important for isotope mass balance calculations of non-separable mixtures containing silica minerals. The isolation of quartz generally follows the flow sheet presented in Sridhar et al. (1975). After a treatment of the sample with 6 N HC1 at 100~ (removal of carbonates, hydrous iron oxides, some amorphous material and clay minerals) and gravimetric size fractionation (settling, centrifugation), the sample is fused with powdered sodium pyrosulfate. The melt is dissolved with 3 N HC1 (removal of phyllosilicates) and washed. Residual amorphous silica and feldspars are dissolved by shaking in H2SiF6 that has been cleaned from HF by reaction with sized 1-100 ~m commercial ground quartz (Jackson et al., 1976). After dissolution of fluorates using saturated H3BO3 and the purity checks by XRD and SEM, the quartz isolates are weighted for quantitative determination of the quartz content. This isolation procedure does not affect the 5180 of the quartz (Syers et al., 1968; Sridhar et al., 1975). Cristobalite or tridymite is separated from quartz in size fractions > 2/~m before pyrosulfate fusion using a fluid with a density of ca. 2.4 g /cm3. The lighter cristobalitebearing fraction is given an acid-base-fluosilicic (ABFS) treatment (Henderson et al., 1972): following several treatments with 6N HC1 at 80~ and later with 0.5 N NaOH at 100~ residual minerals like feldspars are removed with 30% H2SiF6 at 18~ These treatments can also be used for isolation of amorphous biogenic silica, such as diatom or phytolith opal, or diagenetic opal-cristobalite (Henderson et al., 1972). These waterand OH-bearing forms of amorphous or semi-amorphous silica, however, contain significant amounts of easily exchangeable oxygen. Therefore, they need a special analytical procedure, such as rapid high-temperature dehydration (e.g., Labeyrie, 1974; Wang & Yeh, 1985; Brandriss et al., 1998), controlled isotope exchange (e.g., Labeyrie & Juillet, 1982; Juillet-Leclerc & Labeyrie, 1987) or step-wise fluorination (Matheney & Knauth, 1989).

2.5.2 Hydroxides and oxyhydroxides of iron and aluminum The purification of iron (III) oxides and oxyhydroxides, specially goethite or hematite, from clay-rich rocks for oxygen isotope analysis can be achieved by selective dissolution of silicates using hot 5 M NaOH (Yapp, 1991). Chazot (1997) conducted tests on the effect of the 5M NaOH treatment on 5180 values of one natural and one synthetic goethite. Tests were performed on the goethites alone as well as on mixtures of goethite and kaolinite, and using a solution markedly depleted in 180 (Figure 2.4). The results indicate that the 5M NaOH treatment has no effect on oxygen isotopes when done at 80~ for 4.5 hours. In contrast, goethite 5180 value may be significantly modified, presumably due to isotopic exchange with the NaOH solution, when treatment is done at 100~

Conventional and Less ConventionalTechniquesfor Hydrogenand OxygenIsotope ...

57

(~ 18OSMOW ( ~ o ) -16

~ ~ I ~ l

-14

"7-'~

-12 I

-10

-8

I

~

NaOH solution estimate

untreated goethite

H

-6 I

-4 I-

H

goethite + 5M NaOH at 80~ for 4.5 hours goethite + kaolinite + 5M NaOH at 80~ for 4.5 hours

!

goethite + 5M N aOH at 100~ for 4.5 hours synthetic goethite

natural goethite

Figure 2.4 - Measured 6180 values of untreated and treated goethites and goethite-kaolinite mixtures using the 5M NaOH method modified after Yapp (1991). Data from Chazot (1997). Heavy-liquid and high-field magnetic separation are also useful for the enrichment of these minerals. However, a complete isolation may not be possible in all cases, especially for aluminum oxides and oxyhydroxides (gibbsite, boehmite etc.) and require material balance calculations using chemical analyses and other selective dissolution methods (e.g., Bird et al., 1989, 1992; Yapp, 1991, 1998). Special techniques are necessary for hydrogen isotope analysis of goethite including removal of organics with H202 and degassing in vacuum up to 3 hours at 100~ (Yapp & Pedley, 1985) and controlled isotope exchange of "high temperature" nonstoichiometric hydrogen (Yapp & Poths, 1995). For oxygen isotope analysis, the goethite is prefluorinated with BrF5 at room temperature to remove nonstoichiometric water (Yapp, 1987). Goethite may also contain small amounts of a Fe(CO3)OH component. Its carbon isotopic composition can be determined by stepwise decarbonation(-dehydration) in vacuum at ca. 230~ (Yapp & Poths, 1991; 1993). 2.5.3 Titanium oxides

Anatase and rutile can be isolated from kaolinite-rich rocks by boiling in 6N HC1 and consequently dissolving the remaining impurities using H2TiF6 at 45~ (Sayin & Jackson, 1975; Bird et al., 1992, 1993). Residual quartz or mica can be removed from the titanium oxides using a magnetic separator, as minor Fe is substituted in TiO2.

58

Chapter 2 - H.A. Gilg, J.-P. Girard & S.M.F.

Sheppard

2.5.4 Phosphates Mizota et al. (1992) determined the oxygen isotope composition of both Ca- and Al-forms of soil phosphates by selective extraction. Aluminum-bound phosphate is extracted with 0.5 M NH4F solution at pH 8.2 and room temperature for 24 hours. A subsequent extraction with 0.5 M HC1 dissolves completely the Ca-forms of phosphate, mainly apatite (Williams et al., 1967). The two extracts were purified, the phosphate precipitated as BiPO4.0.5H20 and finally fluorinated following the procedures of Tudge (1960). 2.5.5 Carbonates, organic matter and isotopic analysis of whole rock shales The H- and O-isotope compositions of whole rock shales cannot usually be directly analyzed by conventional whole rock techniques unless the ubiquitously associated carbonates and organic matter are only present as trace constituents (< few tenths of wt. %). During H-isotope analysis using induction furnace heating of the crucible, hydrocarbons can distill in the reaction tube, leading to incomplete hydrogen yields. Although heating the reaction tube with an external electric furnace may eliminate this problem, the maximum working temperature of the furnace may be insufficient to liberate all the hydrogen from the minerals, again leading to incomplete yields. During the extraction of oxygen gas using a fluorinating agent, carbonate carbon can react with oxygen and fluorine to give COF2 and COBr2 compounds. The yields of molecular oxygen gas are thus too low. For these reasons a multi-technique approach is necessary. Carbonates The C- and O-isotope compositions of associated carbonates can be determined using the classic H3PO4 acid attack on the whole rock powder. The resulting CO2 gas, however, may contain contaminants arising from reactions of the acid with organic matter and sulfur minerals. Because of mass interference in the mass spectrometer, the measured 645 value will not necessary be equal to the ~13C value. These contaminants can be removed by reacting the sample gas with silver phosphate (Smith & Croxford, 1975; Charef & Sheppard, 1984). Organics and whole rock analysis The total organic matter of a rock cannot usually be separated in a form suitable for isotopic analysis. The techniques developed to separate kerogen from a rock by destroying the minerals are so drastic (e.g. HC1 and HF 6N attack; Durand & Nicaise, 1980) that some fraction of the organic matter is probably also destroyed. An alternative approach is to determine the H- and C-isotope composition of the total organic matter by difference: analysis of the whole rock sample after removal of carbonates (clays plus organics) plus analysis of the whole rock sample after removal of both carbonates and organics by, for example, plasma ashing. The whole rock minus carbonates sample can be analyzed in the following way. Two reaction tubes in silica glass are interconnected (Figure 2.5). In one of the tubes the "whole rock" sample is mixed with copper oxide (CuO) and in the other only copper oxide is present. In the tube with the sample, a layer of copper oxide is placed above the sample. Before introducing the copper oxide to the tubes it was degassed at 800~ to reduce the carbon

Conventional and Less Conventional Techniques for Hydrogen and Oxygen Isotope ...

59

blanks. The tube with the sample is degassed at 120~ to remove adsorbed water. The extraction is started by heating the tube with copper oxide by itself to 850-900~ to generate a partial pressure of oxygen in the two reaction tubes. Then the sample tube is heated progressively to 1000~ so that all the volatile constituents are oxidized and gases such as CH4, H2S and other hydrocarbons cannot form. The sample tube is kept at 1000~ for about one hour. The gases are then passed through two liquid nitrogen Figure 2.5 - Apparatus for the extraction of total hydrogen and carcooled traps to com- bon from organic-rich whole rocks by oxidation, using copper oxide pletely separate the (CuO). condensable gases (H20, CO2, SO2) from the excess oxygen and other non-condensable gases. Finally the water and carbon dioxide are separated from each other and purified, and then analyzed isotopically in the usual way. The above technique where the sample is always in a significant oxygen atmosphere during the heating stage gives more satisfactory results than the sealed tube method. In the sealed tube method the sample and copper oxide mixture are sealed in a silica tube and then heated in a muffle furnace at 1000~ After cooling down the tube is opened under vacuum with a tube cracker and the water and carbon dioxide are separated and purified in the usual way. Hydrogen yields, usually presented as H20 +, are similar to or often lower than the two reaction tube method. This may be due to back reactions and hydrated mineral formation in the sealed tube during cooling. The H- and C-isotope compositions of the organic matter are calculated from the yield and isotopic data derived from the analyses of the whole rock sample and whole rock minus organic matter sample. Combined with the C- and O- isotope analyses of the carbonates and mass balance considerations, the H-, C- and O-isotope compositions of both the whole rock and principal constituents can be derived. This multi-

60


analysis approach may not be essential if the organic matter content of the whole rock is less than about 0.6 wt. % and carbonate is only a trace mineral. 2.6 Conclusions

Over the past three decades, significant progress has been made in the development of physical and chemical extraction and purification techniques of clays for subsequent stable isotope analysis. The flow sheet (Figure 2.1) combines these techniques into a standard procedure. Most of these treatments do not to alter the isotopic composition of the clay minerals in a significant way (see Table 2.1). However, for certain clay minerals and some treatment procedures, rigorous tests are missing. There are also difficulties in separating mixtures of some clay minerals, such as kaolinite and illite, or of different generations of the same mineral in a sample. Specific techniques for hydrogen and oxygen isotope analysis of clay minerals have been established and yield satisfactory results for most minerals with a precision of ca. + 0.2 %o for 6180 and + 2 to 3 %o for 6D. However, the precision of isotope analysis is often lower for clay minerals containing interlayer water, such as halloysites or some smectites, because either they show significant isotope exchange between OH groups and interlayer water during degassing or have important amounts of easy exchangeable oxygen and/or hydrogen. For these minerals, alternative techniques, for example a modification of the controlled isotope exchange technique of Labeyrie & Juillet (1982), should be developed and improved. The isotopic analysis of pore waters in clay-rich rocks is still in its infancy. We suggest that if the pore water content is sufficiently high, physical extraction techniques, such as high-pressure squeezing, are preferable to distillation techniques, as the latter techniques are not able to separate of free pore water and the isotopically distinct adsorbed and interlayer water of clay minerals. Promising alternatives are the newly developed direct equilibration and the radial diffusion method.

61

Conventional and Less Conventional Techniques for Hydrogen and Oxygen Isotope ...

Appendix Effect of Ammonium Oxalate (AM-OX) and Na-dithionite-citrate-bicarbonate (DCB) treatments for removal of iron oxide and hydroxide on the isotopic composition of kaolinite. The kaolinite used for the test is the Macon reference kaolinite described in Girard & Savin (1996). Treatment

untreated untreated DCB DCB DCB AM-OX AM-OX

N u m b e r of treatments

Temperature (~

Duration (h)

0 2 yield (%)

6180 (%0)

H 2 0 yield (%)

0 0 1 1 1 3 10

25 55 70 -

50 1 3

97 97 104 100 99 103 96

21.0 21.2 21.2 20.9 21.1 21.4 21.6

15.2 14.7 14.3 14.7 14.7 12.2 14.7

6D (%o) -54 -54 -56 -57 -57 -60 -60

Effect of Na-polytungstate densimetry (NaPT) on the isotopic composition of kaolinite Treatment

untreated NaPT NaPT

Temperature (~ Room Room

Duration (days)

0 2 yield (%)

6180 (%o)

H 2 0 yield (%)

6D (%o)

5 5

99 98 104

21.7 21.6 21.6

15.0 14.4 12.4

-59 -54 -60

Handbook of Stable Isotope Analytical Techniques, Volume 1 P.A. de Groot (Editor) 9 2004 Elsevier B.V. All rights reserved.

CHAPTER 3 Techniques for Stable Isotope Analysis of Fluid and Gaseous Inclusions Luigi Dallail, Raffaele Lucchini2 & Zachary D. Sharp3 1 CNR-Instituto di Geologia Ambientale e Geoingegneria, Sez. Roma "La Sapienza", P.le Aldo Moro 5,

00185 Rome, Italy 2 Institut de Min6ralogie et P6trographie, BFSH-2, CH-1015 Lausanne, Switzerland 3 Department of Earth and Planetary Sciences Northrop Hall, Albuquerque, NM, 87131-1116, USA e-mail: [email protected]

3.1 Introduction

Combined stable isotope and fluid inclusions studies of ore deposits and vein systems have greatly increased our understanding of the fluid phase involved in mineral deposition (e.g. Kerrich et al., 1978; Selby & Nesbitt, 1996). The isotope composition of a fluid phase can be calculated on the basis of mineral-fluid equilibrium fractionation at a given temperature, from the 6-values of minerals that formed in equilibrium with the fluid itself. This procedure is not valid when the mineral experienced open system behaviour and original isotopic equilibrium is no longer preserved (e.g. Frezzotti et al., 2000). A more direct approach is to measure the isotopic composition of the fluid remaining as fluid inclusions in a mineral" this fluid generally represents a sample of the fluid trapped within a mineral or a sediment at the time of their formation. The significance of the isotope data of the inclusion fluid is based on the assumption that its isotope composition has been preserved over geological time; that is neglegible isotopic exchange between inclusion fluid and host minerals has occurred. The assumption of preserved fluid composition has been tested in several studies. The stable isotope composition of waters extracted from different minerals within a given lithology are the same (e.g. Vityk et al., 1993; Genty et al., 2002; Naden et al., 2003), and when accurately measured, the isotopic composition of inclusion fluids reflect calculated equilibrium composition of the host mineral phase. For these reasons in the last fifteen years the number of the stable isotope laboratories performing stable isotope analyses of fluid inclusions has noticeably increased and the amount of new data is rapidly growing (see Figures 3.1-3). The aim of this paper is to discuss the various methods of fluid inclusion extraction, their drawbacks and problems, and the effort that have been made to improve their accuracy.

Techniques for Stable Isotope Analysis of Fluid and Gaseous Inclusions

3.2 Methods of extraction Stable isotope analysis of fluid inclusions involves two separate steps: a) quantitative extraction of fluid phase from the host mineral; b) mass spectrometric analyses of the extracted phase. The least ambiguous analyses can be performed on macroscopic fluid inclusions. Piperov & Penchev (1982) described a procedure of withdrawing the inclusion water by puncturing the inclusion itself: they introduced a capillary into the inclusion and drew the water out. Similarly Lazar & Holland (1988) and Genty et al., Figure 3.1a,b,c - Compilation of measured 6D values of fluid inclusions extracted from a) plutonic and b) & c) metamorphic rocks and veins. The reference number on the X-axes corresponds to the reference number in the Appendix. Temperatures refer to homogenization temperatures. The 6D values of the fluid inclusions from magmatic rocks (a) fall in the compositional field of the magmatic waters (- 40 < 6D < - 80). The isotopic composition of fluids extracted from inclusions in contact metamorphic rocks (b) also fall in the magmatic range. The 6D values of inclusions with higher TH (closer to the pluton) are those typically found in primary fluid inclusions of plutonic rocks. The hydrogen isotope composition of metamorphic rocks (b) shows the open-system vs. closed-system behaviour of the investigated rocks. In regional metamorphic rocks, hydrogen isotope data are more scattered in samples with low homogenization temperatures, refleting meteoric or seawater sources. At higher Th values, 6D values converge on those of magmatic origin.

# reference

63

64 3.2a,b,c - Compilation of measured 613C values of fluid inclusions extracted from magmatic (a) and metamorphic (b, c) rocks and veins. The reference numbers on the X-axis correspond to the reference numbers in the Appendix. Fluid inclusions of magmatic rocks (a) have measured ~13C values in the range of the magmatic CO2 assumed on the basis of carbonatite 613C values (- 3 t o - 8%o). The very low values were interpreted in terms of mixing in the source regions of juvenile carbon and CO2 derived from subducted sediments (Mattey et al., 1984). The 613C values of the CO2 produced by metamorphic reactions (b) vary over a broad range. Generally 613C values < - 8%o are related to oxidation a n d / o r partial exchange reactions with organic matter. Fluid inclusions from carbonate-bearing rocks (marbles, calciteschists) have the highest ~13C values (from- 4.9%o to + 5%o). The 613C values of fluid inclusions from schistose and gneissic rocks are scattered, likely reflecting mineralogical heterogeneities and different carbon sources. The ~)13C values of inclusions extracted from meta-igneous rocks are slightly depleted compared to the magmatic range, and may indicate the involvement of external CO2 or partial exchange with organic matter during metamorphism. The isotopic compositions reported for contact metamorphic rocks (c) are influenced both by the mineralogical and the external CO2 imput as mentioned above. Figure

Chapter 3 - L. Dallai, R. Lucchini & Z.D. Sharp


65

Figure 3.3- Correlation of the measured 6180 vs. 6D values of aquaeous fluid inclusions ex-tracted from halites and speleothems. Also shown is the global meteoric water line (GMWL). Data plotting to the right of the GMWL are related to evaporation. The reference data are reported in the Appendix: Horita & Matsuo, 1986 (A); Ohba & Matsuo, 1988 (El) Horita, 1990 (11); Yang et al., 1995 (~); Yang et al., 1996b ( , ) ; Dennis et al., 2001 (A); Naden et al., 2003 (O); Vityk et al., 1993 (O).

(2002) were able to drill a hole into the inclusions walls (halite crystals and spelothem calcite, respectively), and extract the inclusion fluid with a micropipette. In these cases (Madan galenas, Red Sea halites, and Villars c speleothems, respectively), the inclusions were millimetric in size, a situation which is rarely encountered. More common fluid inclusions are at the micron scale, so that an entire population of inclusions must be released to generate enough material for analysis. The methods used for bulk fluid inclusion extractions are thermal (vacuum heating to decrepitation, pyrolysis) a n d / o r mechanical (crushing or grinding in vacuum). Both methods are destructive. As of now, only a few attempts have been made on non-destructive analytical methods, such as Raman spectroscopy (Rosasco et al., 1975; Marshall et al., 1994), but instrumental detection limits and analytical reproducibility lead to results that are not satisfactory for geological investigation. The crushing method is based on the principle of breaking the minerals hosting the inclusions in order to break and open the inclusions themselves (Roedder et al., 1963). Volatiles trapped within the minerals are released in a sealed, evacuated vessel (e.g. stainless-steel tube, steel or pyrex ball-mill) and then cryogenically transferred to a vacuum line for isotope extraction. Thermal decrepitation method involves heating the samples in a vacuum line to sufficently high temperatures for the inclusions to crack open due to high internal pressure (Roedder et al., 1963). Specific advantages and drawbacks of the mechanical

66


and the thermal methods of extraction make these methods complementary. Once the inclusion fluids have been extracted, the analytical methods generally used for the stable isotope analysis of gaseous and aqueous phases are applied to convert H-C-N-O-S compounds to gases suitable for analysis in the mass spectreometer, namely CO2, H2, N2, and SO2. Oxygen isotope analysis of fluid inclusions water have been performed according to different methods: micro C O 2 - water equilibration (Kishima & Sakai, 1980); fluorination (O'Neil & Epstein, 1966b); guanidine hydrochloride (Dugan et al., 1985); continuous-flow water-reduction (Sharp et al., 2001). Hydrogen isotope analyses have been made either by Zn reduction (Coleman et al. 1982), uranium reduction (Bigeleisen et al., 1952; Friedman & Smith, 1958), and most recently, water reduction in continuous flow systems (Prosser & Scrimgeour, 1995; Burgoyne & Hayes, 1998; Sharp et al., 2001). The details concerning the methods mentioned above are reported in specific chapters of these books and will not be treated hereafter. 3.3 Thermal decrepitation The procedure of heating the host mineral and decrepitating inclusions has been successfully employed to extract inclusion fluids from numerous minerals, including quartz, fluorite, barite, sphene, galena, halite and calcite. With regards to the latter two phases further considerations will be added in the drawbacks paragraphs. Thermal decrepitation is a simple and inexpensive method: samples are placed in a quartz tube and evacuated. The tube is then heated using a thermocouple-monitored furnace. The degree of heating necessary to cause decrepitation is determined on the basis of the pressure changes in the vacuum line. Generally, heating continues until a stable pressure is reached. The schematic of a simple and efficent extraction line is shown in Figure 3.4 (after Lucchini, 1997, unpublished). The quartz tube is heated to high temperature prior to sample loading to remove any water that could desorb from the walls. Samples are loaded and thoroughly degassed under vacuum and low temperature heating (100-150 ~ C). Further heating is applied to extract the fluid inclusions. In some cases, decrepitation allows for selective extractions of distinct fluid inclusion populations (Hattori & Sakai, 1979) by heating the sample to different temperatures (stepwise heating, see also Chapter 13). The extraction of fluids belonging to different stages of the rock evolution is possible by step-heating when the temperatures of decrepitation for each generation are distinct. For instance if late stage fluid inclusions trails are observed in a sample, decrepitation temperature may be as low as 100-120~ The gas released by this generation can be either measured or discarded. The furnace temperature can be raised progressively to reach the estimated temperature of decrepitation. However, decrepitation experiments at high temperatures (T > 750 ~ C) have shown that additional release of water or other volatiles may occur once the decrepitation temperature range is significantly overstepped. Piperov & Penchev (1973) carefully monitored the reactions taking place after they decrepitated H20 - CO2 inclusions in an allanite crystal at temperatures between 350 ~ to 600~ Almost no additional gas was released until 900~ where a sharp increase of H20 and CO2 production was observed. They interpreted the newly


Figure 3.4 - Schematic of the extraction lines for fluid inclusions. Upper figure illustrates the extraction line for analysis of hydrogen isotope ratios from water or methane, and carbon isotopes from CO2 or methane. Lower figure is for determining the 6180 values of water. H20 is equilibrated with CO2 derived from the reservoir. The configuration of a single line can be changed to accommodate either type of analysis.

67

68


released gas in terms of decrepitation of sub-micron inclusions and release of hydrogen gas dissolved in the crystal structure. If the sample was then reheated, a considerable amount of hydrogen was released at temperature above 500~ with a sharp hydrogen release at 800 to 900~ Variable amounts of the other gas species, such as CH4 and CO were also present; thus gas species other than the expected C O 2 and H20 were derived from chemical reaction during heating. The amount of CO was greater at higher temperature of decrepitation. Water can be trapped both as fluid inclusions and structurally bond water (bubble water) during mineral growth, the isotopic composition of water released from high temperature decrepitation resulting in mixed 6D values. Simon, (2001) grinded and thermally decrepitated at temperature of 1200~ aliquots of hydrothermal quartz with different grain size, in order to evaluate the contribution of bubble water to the 6D values of the released water. Water yields obtained by thermal decrepitation decrease as a function of grain size because water from large fluid inclusions can be lost while grinding the sample. Because the structurally bonded water is D-depleted, the final hydrogen isotope composition is influenced by the volumetric ratio between fluid inclusion water and bubble water. The gD values of water extracted by thermal decrepiatation decrease with decreasing grain size fraction because the bubble water fraction becomes dominant in fine grain material. Most accurate 6D estimates of the fluid inclusion water are obtained when fluid inclusion/bubble water ratio is high (minerals formed at medium to low temperature), and when coarse-grained material is decrepitated. Release of CO2-rich fluid inclusion from quartz at temperatures of 1100~ has been successfully performed (see Nesbitt & Muehlenbachs, 1995); at such high temperature CuO was used to oxidize possible carbon species; this technique is described in the oxidative pyrolysis paragraph (see below). Decrepitation accomplished by melting of the host mineral fails to exctract the inclusion fluids: a quartz xenolith bearing high density pure CO2 inclusions within a basalt was heated and melted at temperatures in excess of 1200~ (Dallai et al, unpublished). No inclusion decrepitation and no CO2 was released was observed until T > 1460~ was reached. At this temperature the quartz (and the quartz-glass tube) melted, and several newly formed bubbles were observed into the molten sample (Figure 3.5). These bubbles were interpreted in terms of decrepitated gas unable to escape from the viscous mineral but no CO2 was detected by Raman spectroscopy of these newly formed bubbles, because the CO2 density was too low. Only an H20 vapour phase was produced during heating above 600~ Because inclusions contained only high-density CO2, H20 desorption from mineral and/or sample holder, or dissolved water in basalt were inferred as possible sources. Very high temperature vacuum volatilization has been proven reliable to extract fluid inclusions from different mineral phases such as halite (Knauth & Beeunas, 1986), plagioclase (+ quartz + apatite), and clinopyroxene (+ olivine; Kelley & FrtihGreen, 2001). Care shall be taken to quarry all the impurities (i.e. organic matter,


69

Figure 3.5 - Photograph of the melted quartz sample (now glass) with newly formed vescicles (scale 400:1).

hydrous phases). This latter procedure is often difficult and time consuming. Vacuum crushing method (see below) is generally advised for fluid inclusion extraction from halites. Recently, Sharp et al., (2001) have obtained accurate and reproducible data by vaporizing extremely small halite samples in a helium stream, and processing the evolved water in a CuO furnace to covert any H2 to H20. The water is then cryofocused and released into a high T reduction furnace that allowed for quantitative conversion of H20 to H2 and CO gas. Both 6D and 6180 values could be determined from the same sample by this technique.

Significant drawbacks of decrepitation are: 1) the method is generally not applicable to fluid inclusions hosted by hydrous minerals (Kazahaya & Matsuo, 1985); 2) the possibility exists for chemical reactions to occur between released gases during heating repetitions; 3) low density inclusions that homogenize to a vapor phase will not decrepitate because of the flat isochores; 4) problems are encountered when extracting calcite and/or dolomite-hosted inclusions (calcite starts decomposing at around 450~ CO2 gas may be produced from carbonate breakdown and add to CO2 from fluid inclusions. L6cuyer & O'Neil, (1994) observed small amounts of CO2 being produced during decrepitation of H20-bearing fluid inclusions in carbonate shells. Above 500~ the amount of CO2 evolved was extensive and the amount of CO2 was correlated with the temperature of decrepitation. However, no correlation between the 6180 values of the extracted H20 and the amount of CO2 (or the temperature) was found. In general it is advised not to heat carbonate-hosted inclusions above 150~ (Dennis et al, 2001). With regards to D / H determination in calcite-hosted inclusion water (and calcitebound water), it has been noted that high-T decrepitation leads to ~SD values that are significantly lower than the actual inclusion values (Yonge, 1982). Matthews et al., (2000) reported an isototopic fractionation of- 30.0 _+1.4 %0 between the 6D values of

70


inclusions in cave deposits and the values of the waters from which they grew. A slightly lower value (- 22.1 + 3.9 %0), was reported by Yonge, (1982); this difference may be related to the different temperature of heating (Yonge used 700~ whereas Matthews et al. used 900~ and/or to the different behaviour of the analyzed cave deposits (Matthews et al., 2000).

Reproducibility for

hydrogen isotope measurements of fluid inclusions Hattori & Sakai, (1979) and Jenkin et al., (1992) reported a precision better than 2%0 (2 o). For carbon isotope composition of fluid inclusions Kreulen (1980) reported a precision between 0.5 and 1.0 %0 (2 ~J). The reproducibility of some of the analyses perfomed on the extraction line of Figure 3.4 is reported in Table 3.1.

3.4 Oxidative Pyrolysis The stepped heating procedure at very high temperatures can be applied in order to disrupt the crystal structure of the mineral and extract the gas dissolved in the crystal. The oxidative pyrolysis method is mainly applied to the recovery of the reduced carbon species in minerals formed at high-temperature (e.g. olivine, pyroxene), typically in volcanic glasses. The h-values of the gas species extracted with this method represent the isotopic compositions of gas mixtures, namely gas released from the fluid inclusions and gas phase dissolved into the crystal. Therefore measured carbon isotope compositions are correlated to the abundance of these two components, and possible effects of isotopic fractionation need to be evaluated (Nadeau et al., 1990). The method involves loading the sample in a pre-baked quartz glass combustion vessel, which is then evacuated. The sample is heated to 400 ~ - 500~ in order to remove any surfacial organic contamination (Mattey et al., 1984). The temperature is raised further (800 ~ to 1200~ in the presence of a low partial pressure of 02, (10 - 20 mbars) evolved from heating CuO. Oxygen converts the solid and reduced carbon to CO2. Separation of CO2 from other gas species (SO2, H 2 0 ) is achieved by standard cryogenic methods (Nadeau et al., 1990; Kelley & Friih-Green, 2001). Mattey et al., (1984; 1989) reported that all the carbon is released from basalt glasses after 1 hour pyrolysis at 1200~ repeated extractions at 1300~ of pyrolyzed samples showed almost no further carbon release. Watanabe et al., (1984) reacted the samples for 24 hours in the presence of 0 2 to extract CO2-rich inclusions from olivine crystals (oxygen was used in order to convert any graphitic carbon left in the crystal after inclusion decrepitation). The partial pressure of oxygen in the reaction furnace is used to convert the reduced forms of released carbon to CO2 for isotopic analyses (Watanabe et al., 1983; Mattey et al., 1984:; 1989; Nadeau et al., 1990). The reported reproducibility of the oxidation pyrolysis method is 1.2%o for the ~13C of fluid inclusions in olivine crystals analyses (Watanabe et al., 1984). Mattey et al. (1984:; 1989) performed their isotopic analyses on a static mass spectrometer in order to measure very small amount of gas (Carr et a1.1986). The overall precision they obtained on the measured 613C of carbon dissolved in basalt glasses was + 5.0%0.

Sample Wt (g)

Mineral Tdecr.('C ) Gas phase extracted from fluid inclusions c02 H20 N.C. mbar

A5.1 A5.1 A5.1 83.2 B3.2 B3.3 B3.3 MR19 MR19 MR17 MR17 MR25 MR25 r141 1-141 r104 r104

%, vol mol. (MS) C02 H20 NC

?, vol mol. (other methods) c02

H20

NC

21.3

77.4

0.2

51.O

40

9

49

48

4

55

41

4

~ ~ ~ C Volt P DMAT251 B ~ DSM O Wmbar

2.6 2.5 2.7 4.1 2.7 3.6 4.1 1.1 1.3 1.0 1.1 0.8 1.0

Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz

540 550 600 600 600 600 600 520 520 520 520 520 520

20.2 19.2 19.9 220 120 98.4 120 8.9 9.5 8.2 11.0 16.0 17.0

-3.0 -3.2 -3.3 5.4 5.2 4.3 4.6 2.4 2.5 -3.1 -3.0 0.0 0.1

6.0 4.8 8.0 6.4

Qtz Qtz Qtz Qtz

550 550 550 550

0.2 3.3 7.1 4.8

-5.4 -5.6 -3.4 -3.7

0.972 1009

-35.2 -36.6

80 50

20.7 19.5

76.6 78.8

2.7 1.7

6741 3313 5312 7056 0.099 0.102 0.110 0.176 0.105 0.125 Volt DeltaE 2620 1548 840 593

-73.6 -75.8 -19.8 -22.8 2.7 -2.6 -29.3 -32.4 -96.8 -99.7

380 400 450 440 100 150 50 70 110 100

17.9 16.8 27.8 29.7 47.2 46.2 45.0 40.9 60.1 58.4

80.8 82.2 71.1 68.4 43.7 41.3 50.3 54.6 32.8 35.7

1.2 1.0 1.0 1.9 9.1 12.5 4.7 4.5 7.1 5.9

31.5 28.8 37.0 33.4

0.01 0.01 0.27 0.13

4.3 4.9 14.0 15.8

95.6 94.9 80.9 79.9

0.1 0.1 5.0 4.2


Table 3.1 - Example of data reproducibility using the extraction line as in Figure 34.2 for fluid inclusions extracted from quartz veins. VMAT251 and VDeltaE are referred to the hydrogen gas voltage (mV) in the Faraday cup 2 with bellow loo?, open.

71

72


3.5 Crushing or grinding in vacuum

Several detailed papers exist which describe the method of crushing minerals in vacuum and inclusion fluids release for stable isotope analyses (i.e. Andrawes & Gibson, 1979; Kita, 1981; Kazahaya & Matsuo, 1985; Horita & Matsuo, 1986; Norman & Sawkins, 1987; Ohba & Matsuo, 1988; Dennis et al., 1998). Mineral crushing was initially performed in aluminum tubing. The samples were loaded in the tube; the tube was then evacuated and squeezed until the content had been finely ground and the volatiles released (Roedder et al, 1963). By using this method low gas yields were generally obtained, and the tube could not be re-used. Ball-mill apparata and/or stainless steel crushing cells have been successively adopted. The ball-mill devices made by Kita (1981) and Ohba & Matsuo, (1988), are shown in Figure 3.6a,b, respectively. They consist of pyrex cylinders filled with aluminum, pyrex or stainless steel ball which moves up and down by hand shaking, magnetic breaker or any motorized device. More recently a crushing cell has been developed (Dennis et al., 2001)" such a device consists of a stainless steel flanged tube with a hardened base, containing a piston that is electromagnetically raised and lowered (Figure 3.6c). Halite and calcite minerals have been successfully analyzed using the ball mill and the crushing cell, whereas the crushing efficency on fluorite, quartz or harder mineral phases is generally unsatisfactory. Using the apparatus of Kita, (1981) the uncrushed quartz fractions after 1 hour milling was 60%. The fraction left ucrushed by Ohba & Matsuo's ball-mill, (1988) was 44%. Further attempts are needed to constrain the efficency of the crushing cell. The main drawbacks of crushing are selective crushing, gas generation during crushing, and adsorption. Crushing very small inclusions is problematic and selective release of volatiles from bigger inclusions may occur (L4cuyer & O'Neil, 1994). A correlation between the size of inclusions and the 8180 of the extracted waters has been observed, the smaller inclusions having lower oxygen isotope values due to exchange with the host mineral and/or diffusional 180 loss from inclusion to the host quartz (Ohba et al., 1995). Therefore measured isotopic compositions of inclusions from incompletely crushed samples may result in slightly enriched 8180 values. The major drawback of crushing is adsorption, which takes places mainly when H20-vapour is released from fluid inclusions and interacts with the newly formed surfaces of the powdered mineral. These surfaces can be chemically active; oxygen and hydrogen isotope fractionation occurs due to the H20 vapor adsorption, particularly on quartz and fluorite (Barker & Torkelson, 1974). It is worthwhile to note that no adsorption has been observed in absence of crushed sample, thus "adsorption occurs only when new surfaces are being created" by fracturing the mineral (Ware & Pirooz, 1967). The results obtained by grinding the sample in a "pepper-mill"-like device showed that "the smaller the final grain size, the more enriched in deuterium the extracted water", and that "the larger the final grain size, the smaller the water yield" (Simon, 2001). Crushing quartz grains above 0.6 mm in size resulted in constant water yields, and in 8D values similar to the ones obtained by thermal decrepitation. The positive correlation between the 8D values of water extracted and the amount of surfaces formed by grinding was interpreted in terms of adsorbed water being D-depleted (Simon, 2001). Similar conclusions were inferred from 8180 values of desorbed water (post-crushing heating) lower than 8180 values of water collected upon crushing (Matsuo, 1991). Interestingly, the latter experiment yielded contrasting results for hydrogen isotope


Figure 3.6a, b, c - Sectional views of the Kita's (1981), Ohba & Matsuo's (1988) and Dennis et al., (2001) crushing devices.

73

74


fractionation (6D values of desorbed water higher than 6D of collected water; Matsuo, 1991). In order to reduce adsorption, the crushing device should be on-line with a vacuum extraction line (adding a cold finger at liquid nitrogen temperature to the crushing apparatus to collect the released gas results in significantly higher gas yields (Ohba & Matsuo, 1988)), and heated to moderate temperatures (150 ~ - 200~ during the whole procedure (Dennis et al., 2001; Simon, 2001). The adsorbed gas can be released by heating the powdered crystal after completing the crushing procedure; as adsorption is temperature-dependent and very high temperatures (up to 600~ are necessary to completely desorb water from silica surfaces (Hockey & Pethica, 1961; Moore & Rose, 1973). However, there is no consensus about the possibility to obtain a quantitative recovery of the adsorbed water on crushed material was from heating the powdered material; Kazahaya & Matsuo, (1985), Horita & Matsuo, (1986), and Matsuo, (1991) reported successful experiments from heating halite samples to 180 ~ to 220~ Conversely, water desorption from crushed quartz was considered unsatisfactory for accurate isotope analyses by Ohba & Matsuo, (1988). These authors addressed the problem of measuring the original oxygen isotope composition of inclusion waters, due to the incorporation of oxygen atoms into the newly formed quartz surfaces during adsorption. Mathematical corrections to the measured 6D and ~180 values were adopted, assuming isotopic fractionation of the adsorbed water could be modelled as a Rayleigh fractionation process. Even so, results were less than ideal. Continous flow techniques can be employed to transport the released gas in an inert gas stream to the cryogenic trap (e.g. Li & Shi, 1983), or directly into a gas chromatograph (Andrawes & Gibson, 1979) or a mass spectrometer (e.g. Sharp et al., 2001). Andrawes & Gibson (1979) reported that quantitative water extraction was possible by using a helium flow-assisted crushing device. However chromatographic analysis of other gases simultaneously with water was unsuccessful and water needed to be measured separately. Still, different gases are adsorbed to different extent. Experiments by Barker & Torkelson, (1974) and Matsuo, (1991) indicate that CO2, N2 and CH4 adsorption on powdered grains is negligible. Crushing has been proven accurate and reproducible: reported precision for ~)13C analysis of CO2 from fluid inclusions hosted in quartz is + 0.5%o or better (Kreulen, 1980; Matsuo, 1991). Measured 6180 and 8D values of water rich inclusions hosted in halites and carbonates have reproducibilities within + 1.6%o and + 1.7%o, respectively (Matsuo, 1991) and + 0.3, + 2.0%o (Yang et al., 1996b). Horita & Matsuo, (1986) reported 6180 and 6D values of water hosted in halites differing by - 0.2 %o and - 3.0 %o from the values of the host brine. Duplicate 6D analysis of water inclusions in speleothems agree to + 3.0%o (Harmon et al., 1979; Dennis et al., 2001).

34.6 Cryogenic separation and molar volumes calculation C O 2 , H20, CH4 and N2 gas species released during decrepitation and/or crushing need to be separated prior to analysis. Different methods are used for the various gas species. Most involve a simple vacuum extraction line illustrated in Figure 3.4: water and CO2 are released and frozen into the trap closest to the original sample (Trap A)

75


using liquid nitrogen. They can be cryogenically separated using liquid nitrogen and ethanol-dry ice slush, respectively. Methane does not freeze at liquid nitrogen temperature and is free to expand throughout the line. CH4 is oxidized at T - 800~ in the cuprous oxide furnace. Cu20 is used instead of CuO because it has much lower 02 vapor pressure at 800~ During oxidation, methane is converted to CO2 and H20 and frozen into the trap adjacent to the furnace (Trap B). A small piece of platinum in the furnace catalyzes the reaction, increasing the oxidation rates. Again CO2 can be separated from the water cryogenically and frozen into a sample container for 813C and 8180 analyses (the 8180 value of the oxidized methane has no significance). The water

Samples containing hydrocarbons CO2, H20, N2, hydrocarbons Thermal Decripitation ~ Production in some cases of CO, 02, H2S, SO2 / Crushing / Cracking of Hydrocarbons

I Condensible

Separation with liquid Nitrogen

[ Non-Condensible ] N2, CH4, CO ] Oxidation with CuO at 800~ CH4+202~CO2+2H20 2CO+O2-+2CO2

CO2 H20

Separation with liquid Nitrogen Condensible Non-Condensible

Separation with dry ice + ethanol

CO2

I

I

H20

I

I

~ Introduction of CO2

I

I

N2

I

I

Adsorb on zeolite trap at LN2 temp.

Equilibration at 25~ 24hrs CO2, H20

I CO2, H20 I


Non-Condensible

I

I C~


CO2 ~r

I

I

I

8180(H20)

I

H20

H20 Reaction of H20+Zn 500~ 30min

I Reaction of H20+Zn 500~ 30min

813C(CO2) 8180(CO2)

Condensible

8H(H20)

815N

813C(CH4 - CO)

Figure 3.7- Schematic diagram for procedure of gas treatment and separation.

8H(CH4)

76


F i g u r e 3.8a, b - Schematic of the laser-based fluid inclusion extraction line. The fluid inclusion is opened during laser 'drilling'. Helium gas passes through the heated sample chamber, incoroporating water released during exposure of the fluid inclusion. Water and CO2 are collected on the liquid nitrogen trap in a 6-way valve assembly. After 3 minutes, all water is transferred. The 6-way valve is switched and the trap is heated, releasing the water as a coherent 'pulse'. Water passes through a 1450~ microfurnace filled with glassy carbon causing complete reduction of water to H2 and CO, through a gas chromatographic column (5A molecular sieve) to separate the reaction products, and finally into the mass spectrometer for isotopic analysis. (b) Schematic of reaction chamber. Helium enters and exits the chamber through two holes in the bottom. Laser radiation is admitted through a UV-grade silica window. The seal is maintained with a Viton O-ring, affixed firmly on the top of the sample chamber.

can be frozen into a tube with zinc for reduction to ZnO and H2 for hydrogen analysis (Coleman et al., 1982). Alternatively, very small amounts of water can be equilibrated with CO2 for oxygen isotope determinations (Kishima & Sakai, 1980; Yu, 1991), equilibrated with H2 gas for hydrogen isotope determinations (Horita, 1988; Horita et al., 1989) or exchanged with guanidine for oxygen isotope determinations (Yang et al., 1996b). The general techniques described above allow for 6D and ~13C values from methane, 613C and 6180 values from CO2 and 6D and 6180 values from water to be deter-


77

mined. Remaining non condensable N2 gas can be adsorbed onto 5/k molecular sieve cooled to liquid N2 temperatures for N2 analysis. A schematic step-diagram for gas separation is reported in Figure 3.7. Yields can be determined manometrically.

34.7 Further Developments The choice of the method of extraction depends on the mineral hosting the inclusions and the fluid characteristics. Crushing and decrepitation are complementary methods (e.g. Hattori & Sakai, 1979; Alderton & Harmon, 1991). However these two conventional methods bear obvious limitations, such as the generally inability to separate different generations of inclusions. Analysis of individual inclusions has only recently been made successfully using a laser GC system (see below). Several attempts have been made using Raman microprobe, based on the relationship between the different isotopic peaks for CO2 and the isotopic ratios of the carbon species (Rosasco et al., 1975; Dhamelincourt et al., 1979; Marshall et al., 1994). It is our experience that the measured isotopic ratios for a single fluid inclusion using Raman spectroscopy are quite precise, but inaccurate due to varying sample geometry, depth to inclusion and thickness. Thus, the results of this method are still unsatisfactory. Recently, Sharp et al., (2001) developed a method for microliter water analysis based on carbon reduction, merging the advantages that gas chromatography and continuous flow technology provide for mass spectrometric analysis techniques. Water samples are entrained in a helium stream and carried through a furnace filled with glassy carbon and heated to 1450~ The high temperature heating causes reduction of the water to H2 and CO, which are then separated in a gas chromatographic column and directly measured for D / H and 180/160 ratios in a mass spectrometer. Heating fluid inclusion-bearing samples in the He-stream is easily incorporated into the system. The system has recently been modified to incorporate a UV laser to open individual large (~ 100 mm diameter) inclusions (Figure 34.8a,b). In this way different generations of inclusions can be analyzed.

Aknowledgements

The authors are indebt to Paul Dennis, who provided insights to the crushing procedures. Reviews by J. Horita, A. Matthews, J. O'Neil, W. Yang, and an anonymous referee improved the original version of the manuscript. Financial support by the Swiss National Science Foundation (21-049302.96) is acknowleged.

Appendix C o m p i l a t i o n of p u b l i s h e d d a t a o n fluid inclusions 6D, 613C a n d 6180 analyses of minerals f r o m rocks in different geological settings. # (1) Location

Method(2) Sample

Contact metamorphism

Mineral

H o s t rock

Met. g r a d e

T~

6DSMOW%o max min

150 150 200-440 250 250 250 250 350 350

-105 -83 -43 -78 -69 -51 -47 -61 -59

-41

Ap Dol Qtz Cal Qtz Sulfate Py Qtz Qtz

Limestones Calc-silicate Flysch Limestones Limestones Limestones Hornfels Limestones Calc-silicate

Qtz-Cal(def.)

Bulk

Metasediment

Green.

90-150

-158

-61

D

Qtz-Cal(undef.)

Bulk

Metasediment

Green.

90-150

-90

-30

Lake Louise, Canada

D

Qtz-Ank-CC

Bulk

Metasediment

Prehnite

90-185

-150

-120

4

Lake Louise, Canada

D

Qtz-Ank-Py

Bulk

Metasediment

Prehnite

90-185

-91

-72

5

Rocky Mountains, $52 ~ Canada

C

Qtz

Limestones

Amph.

100-186

-150

-100

6 7 8

South Cornawall, UK Connemara, Ireland Purcell Mountains, canada

D D D

Low grade

110-150 118-250 149-333

-80 -32 -170

-49 -27 -61

C C C C C C C C C

1 2 3 4 5 6 7 8 9

Mines Gasp6, Quebec, Canada Weolag Tu-Mo deposit, Korea South Cornawall, UK Mines GaspG Quebec, Canada Mines Gasp6, Quebec, Canada Mines Gasp6, Quebec, Canada Weolag Tu-Mo deposit, Korea Mines Gasp6, Quebec, Canada Weolag Tu-Mo deposit, Korea

1

Jasper, Canada

D

2

Jasper, Canada

3

Skarn Qtz-vein (Tu-Mo) Qtz-vein(W-Sn) Skarn Skarn Skarn Qtz-vein (Tu-Mo) Skarn Qtz-vein (Tu-Mo)

Regional metamorphism

Dae Hwa W-Mo Mine, South Korea 10 Dae Hwa W-Mo Mine, South Korea 11 South Cornawall, UK

9

Qtz-vein

Flysch Qtz Qtz-vein(Pb-Zn) Qtz, Cal, Brt Metagabbro Qtz-vein Metasediment Qtz Qtz-vein

Amph.

-23

-43

Reference

Shelton, 1983 So et al., 1983 Wilkinson et al., 1995 Shelton, 1983 Shelton, 1983 Shelton, 1983 So et al., 1983 Shelton, 1983 So et al., 1983 Nesbitt & Muehlenbachs, 1997 Nesbitt & Muehlenbachs, 1997 Nesbitt & Muehlenbachs, 1997 Nesbitt & Muehlenbachs, 1997 Nesbitt & Muehlenbachs, 1995 Wilkinson et al., 1995 O'Reilly et al., 1997 Nesbitt & Muehlenbachs, 1997 Shelton et al., 1987

n.s. Qtz-vein(W-Mo)

Cal

Orthogneiss

150-225

-101

n.s. Qtz-vein(W-Mo)

Dol

Orthogneiss

150-225

-75

-71

Shelton et al., 1987

Qtz-vein(Sn-Cu)

Qtz

Flysch

150-250

-25

-25

Wilkinson & al 1995

D

Low grade

Appendix continued >

(3

!

c~ c~ ~,,,i ~ ~,,,i o

N

> Appendix continued

C~

12 Rocky Mountains Trench, N52 ~ Canada 13 Rocky Mountains, N52 ~ Canada 14 Proterozoic West R. M. T., $52 ~ Canada 15 Proterozoic West R. M. T., N52 ~ Canada 16 Mount Robson, Canada

C

Qtz-vein

Qtz

Metasediment

Green-Amph

150-250

-150

-120

C

Qtz-vein

Qtz

Sandstone

Green-Amph

150-300

-130

-40

C

Qtz-vein

Qtz

Metasediment

Amph.

150-300

-60

-150

C

Qtz-vein

Qtz

Metasediment

Green-Amph

150-320

-150

-90

D

Qtz-vein

Qtz

Metasediment

Prehnite

160-260

-155

-62

17 Cariboo Mountains, Canada

D

Qtz-vein

Qtz

Metasediment

Green.

165-330

-142

-109

18 Malton Gneiss, N52 ~ Canada

C

Qtz-vein

Qtz

Gneiss

Green-Amph

170-210

-80

-40

19 Slide Mountain Terrane, N52 ~ Canada 20 West Onimeca Ext. Compl, $52 ~ Canada 21 Kootenay Arc, $52 ~ Canada

C

Qtz-vein

Qtz

Metabasalts

Green-Amph 200-280

-150

-120

C

Qtz-vein

Qtz

Amph.

200-280

-90

-150

C

Qtz-vein

Qtz

Amph.

200-300

-100

-130 -87

Massive sulfide Massive sulfide Qtz-vein Qtz-vein Qtz-vein

Wooklawn, Canada Wooklawn, Canada Soutern Alps, New Zeland Soutern Alps, New Zeland Rocky Mountains Trench, $52 ~ Canada 27 Barkerville Terrane, N52 ~ Canada 28 Central Alps, Switzerland 29 Tete jaune cache, Canada

C C D D C

30 Dae H w a W-Mo Mine, South Korea 31 Onimeca Ext. Complex, $52 ~ Canada

n.s. Qtz-vein(W-Mo)

22 23 24 25 26

Cal Qtz Cal Qtz Qtz

Metavolcanic Metavolcanic Qtz-Feld schists Qtz-Feld schists Metasediment

Green. Green. Amph. Amph. Amph.

200-320 200-320 200-320 200-320 205-295

-127 -89 -68 -59 -140

-53 -44 -150

C

Qtz-vein

Qtz

Metasediment

Green-Amph 220-280

-150

-120

D D

Qtz-vein Qtz-vein

Qtz Qtz

Syenite Metasediment

Green. Green.

240-280 250-270

-76 -74

-33

Chal

Granitic Gneiss

250-300

-75

250-320

-100

C

Qtz-vein

Qtz

Amph.

-140

Nesbitt & Muehlenbachs, 1995 Nesbitt & Muehlenbachs, 1995 Nesbitt & Muehlenbachs, 1995 Nesbitt & Muehlenbachs, 1995 Nesbitt & Muehlenbachs, 1997 Nesbitt & Muehlenbachs, 1997 Nesbitt & Muehlenbachs, 1995 Nesbitt & Muehlenbachs, 1995 Nesbitt & Muehlenbachs, 1995 Nesbitt & Muehlenbachs, 1995 Lusk & Krouse, 1997 Lusk & Krouse, 1997 Jenkin et al., 1994a Jenkin et al., 1994a Nesbitt & Muehlenbachs, 1995 Nesbitt & Muehlenbachs, 1995 Mullis et al., 1994 Nesbitt & Muehlenbachs, 1997 Shelton et al., 1987

9 9

9

r o C~

~,,,i ~

9

Nesbitt & Muehlenbachs, 1995 Appendix continued > ~D

> Appendix continued 32 Onimeca Ext. Complex, $52 ~ Canada 33 Canadian Cordillera, Canada 34 Canadian Cordillera, Canada 35 Canadian Cordillera, Canada 36 Dae Hwa W-Mo Mine, South Korea 37 Okanagan Valley, British Columbia 38 Brunswick, Canada 39 Brunswick, Canada 40 South Cornawall, UK 41 Naxos, Greece 42 Central Alps, Switzerland 43 Western Carpathians 44 Western Carpathians 45 Western Carpathians 46 Central Alps, Switzerland 47 Central Alps, Switzerland 48 Wilson Terrane, NVL, Antarctica 49 Dae Hwa W-Mo Mine, South Korea 50 Dae Hwa W-Mo Mine, South Korea 51 Central Alps, Switzerland 52 Naxos, Greece 53 Central Alps, Switzerland 54 Central Alps, Switzerland 55 Central Alps, Switzerland 56 Central Alps, Switzerland 57 Soutern Alps, New Zeland 58 Western Carpathians 59 Naxos, Greece 60 Naxos, Greece 61 Naxos, Greece

O C D D D n.s.

Qtz-vein

Gneisses

Amph.

Qtz Qtz Qtz Sche

Oceanic Terranes "Suspect Terranes Island Arc Terranes Granitic Gneiss

Qtz

Graphitic schists

C Massive sulfide C Massive sulfide D Qtz-seg C Qtz-seg D Qtz-seg D Qtz-seg D Shear zone D Qtz-seg D Qtz-seg D Qtz-seg D Qtz-vein n.s. Qtz-vein(W-Mo)

Cal Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Wo

Metavolcanic Metavolcanic Flysch Schists(chl) Marble Granitoides Granitoides Metasediment Phillite Schists Pelite Granitic Gneiss

n.s. Qtz-vein(W-Mo)

Qtz

Granitic Gneiss

D C D D D D D D C C C

Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz Qtz

Orthogneiss Marble Syenite Gneiss Phillite Schists Qz-feld schists Granitoides Bio schist Marble Marble

D

Qtz-carb-vein(Au) Qtz-carb-vein(Au) Qtz-carb-vein(Au) Qtz-vein(W-Mo)

Qtz

Qtz-vein(Au)

Qtz-vein Qtz-seg Qtz-vein Qtz-vein Qtz-vein Qtz-vein Qtz-vein Qtz-vein Qtz-seg Qtz-seg Qtz-seg

268-286 Green. Green. Low grade Amph. Green. Green. Green. Green. Amph. Amph. Green.

Green. Amph. Green. Amph. Amph. Amph. Amph. Green. Amph. Amph. Amph.

250-330

-120

-150

250-350 "250-350 250-350 260-300

-158 -141 -123 -74

-125 -130 -93 -60

-148

-121

270-370 270-370 290-330 350 300-400 300-400 300-400 300-400 320-340 320-340 350-420 350

-169 -169 -18 -70 -59 -57 -39 -35 -90 -42 -84 -65

-88 -125 -10

375-390

-78

-73

Shelton et al., 1987

380-430 385 400-420 400-450 400-450 400-450 400-500 450-500 530 565 620

-65 -88 -61 -68 -54 -51 -42 -64 -59 -66 -69

-39

Mullis et al., 1994 Rye et al., 1976 Mullis et al., 1994 Mullis et al., 1994 Mullis et al., 1994 Mullis et al., 1994 Jenkin et al., 1994a Hurai et al., 1997 Rye et al., 1976 Rye et al., 1976 Rye et al., 1976

Nesbitt & Muehlenbachs, 1995 Nesbitt et al., 1989 Nesbitt et al., 1989 Nesbitt et al., 1989 Shelton et al., 1987 Zhang et al., 1989

-34 -27 -76

-98

-47 -41 -38

Lusk & Krouse, 1997 Lusk & Krouse, 1997 Wilkinson et al., 1995 Rye et al., 1976 Mullis et al., 1994 Hurai et al., 1997 Hurai et al., 1997 Hurai et al., 1997 Mullis et al., 1994 Mullis et al., 1994 Frezzotti et al., 2000 Shelton et al., 1987 r

!

,o

c~ ~..L. ~..L.

N Appendix continued >

> Appendix continued

=r

62 Naxos, Greece C 63 Bushveld complex, South Africa C 64 North Thompson fault, Canada D

Qtz-seg Quarzite Qtz-vein

Qtz Qtz Qtz

Gneiss Metasediment Metasediment

Amph. Green. Amph.

680

-78 -34 -158

-146

Rye et al., 1976 Schiffries & Rye, 1990 Nesbitt & Muehlenbachs, 1997

r

Retrograde alteration South Cornawall, UK Connemara, Ireland Connemara, Ireland Connemara, Ireland

r~

D D D D

Qtz-vein Rock Rock Rock

Qtz Qtz Qtz Cal

Flysch Metagabbro Granite Metagabbro

Low grade Green. Green. Green.

265-315 230-400 230-310 275

-28 -27 -26 -22

-13 -29 -18

Wilkinson et al., 1995 Jenkin et al., 1992 Jenkin et al., 1992 Jenkin et al., 1992

-55 -44 -43 -42 -41 -27 -23 -9 -34 -58 -81 -73 -163 -76 -65 -142 -62 -39 -38 -71 -65 -79 -69 -78 -62

-10 -3 -15

Knauth & Beeunas, 1986 Knauth & Beeunas, 1986 Knauth & Beeunas, 1986 Knauth & Beeunas, 1986 Yang et al., 1996b Knauth & Beeunas, 1986 Knauth & Beeunas, 1986 Knauth & Beeunas, 1986 Shemesh et al., 1992b Shemesh et al., 1992b Yang et al., 1995 Yang et al., 1995 Yang et al., 1995 Yang et al., 1995 Yang et al., 1995 Madu et al., 1990 Schwarcz et al., 1976 Schwarcz et al., 1976 Schwarcz et al., 1976 Schwarcz et al., 1976 Schwarcz et al., 1976 Harmon et al, 1979 Harmon et al, 1979 Harmon et al, 1979 Harmon et al, 1979

o o

Sedimentary rocks Palo Duro Basin, Texas USA Palo Duro Basin, Texas USA Palo Duro Basin, Texas USA Death Valley, USA Qaidam Basin, China Verdi Valley, Arizona USA Lyons, Kansas USA Salado Fm, New Mexico USA Mount Hemon, Israel Mount Hemon, Israel McArthur Pass, Canada Vermillion Pass, Canada Nahanni Butte, Canada Berry, Canada Kotaneelee, Canada Snowbird, British Colunbia M a m m o t h Cave NP, Kentuky Crystal Cave, Bermuda San Luis Potosi, Mexico N o r m a n Bone Cave, W Virginia Grapevine Cave, W Virginia Coldwater cave, Iowa N o r m a n Bone Cave, W Virginia Tumbling Creek Cave, Missouri M a m m o t h Cave NP, Kentuky

D Cpl C D D D D D D D C C C C C D C C C C C C C C C

Cal-vein Cal-vein Dol cement Dol cement Dol cement Dol cement Dol cement Qtz-vein(Au, Sb) Cal speleothem Cal speleothem Cal speleothem Cal speleothem Cal speleothem Cal speleothem Cal speleothem Cal speleothem Cal speleothem

Halite Halite Halite Halite Halite Halite Halite Halite Cal Cal Dol Dol Qtz, Dol Dol Dol Qtz Cal Cal Cal Cal Cal Cal Cal Cal Cal

Evaporite Evaporite Evaporite Evaporite Evaporite Evaporite Evaporite Evaporite Carbonates Carbonates Dolomite Dolomite Dolomite Dolomite Dolomite Listwanites

35 64 142 145 163 166-175 170 240

-45

-3 -12 -40 -74 -72 -42 -139 -39 -8 -15 -52 -51 -67 -52 -34 -39


o ~,,io

C3 r~

9

o

oo

> Appendix continued Mt. Seldom, Israel England Yunan Province, China Yunan Province, China Yunan Province, China Hubei Province, China Hubei Province, China Qinghai Province, China Qinghai Province, China Searles Lake, California Great Lakes North Atlantic Pacific California St Marteen, Caribbean Sea Florida Keys Soreq cave, Israel Soreq cave, Israel Soreq cave, Israel Soreq cave, Israel Villars cave, France Clamouse cave, France

to

C Cal speleothem C C M D C M C M C Biogenic Crb. D Biogenic Crb. D Biogenic Crb. D Biogenic Crb. D Biogenic Crb. D Cal speleothem D Cal speleothem D Cal speleothem D Cal speleothem D Cpl Cal speleothem Cpl Cal speleothem

Plutonic rocks 1 2 3

Connemara, Ireland D Connemara, Ireland D San Diego & Riverside C Counties, US 4 San Diego & Riverside C Counties, US 5 San Diego & Riverside C Counties, US 6 Connemara, Ireland D 7 SW granites, England D/C 8 Bushveld complex, South Africa C 9 SW-Indian Ridge D 10 SW-Indian Ridge D

Halite Cal Halite Halite Halite Halite Halite Halite Halite Halite Cal Cal Cal Cal Cal Cal Cal Cal Cal Cal Cal

-53 -51 -68 -91 -86 -38 -82 -29 -45 -80 -80 -44 -50 -35 -75 -24 -28 -20 -38 -41 -35

Evaporite Evaporite Evaporite Evaporite Evaporite Evaporite Evaporite Evaporite

-34 -39

-78

74 -60 -26 -50 -20 -10 -21 -23 -16 -34 -39 -33

Ohba & Matsuo, 1988 Dennis et al, 2001 Horita & Matsuo, 1986 Horita & Matsuo, 1986 Horita & Matsuo, 1986 Horita & Matsuo, 1986 Horita & Matsuo, 1986 Horita & Matsuo, 1986 Horita & Matsuo, 1986 Horita, 1990 Lecuyer & O'Neil, 1994 Lecuyer & O'Neil, 1994 Lecuyer & O'Neil, 1994 Lecuyer & O'Neil, 1994 Lecuyer & O'Neil, 1994 Matthews et al., 2002 Matthews et al., 2002 Matthews et al., 2002 Matthews et al., 2002 Genty et al., 2002 Genty et al., 2002 r

~r

Granite Granite Gabbro

125-205 270-340 520-565

-45 -24 -58

-17 -17 -51

O'Reilly et al., 1997 O'Reilly et al., 1997 Taylor et al., 1979

Aplite-peg

Gabbro

565-700

-78

-57

Taylor et al., 1979

Graphic-peg

Gabbro

565-700

-52

-50

Taylor et al., 1979

600

-24 -55 -47 -52.0 -50.0

-14 -48

O'Reilly et al., 1997 Alderton & Harmon, 1991 Schiffries & Rye, 1990 Kelley & Friih-Green, 2001 Kelley & Friih-Green, 2001

Qtz-vein Qtz-vein Pocket peg

Qtz-vein(Mo) Rock Pegmatoid Rock Rock

Qtz, F1, Brt Qtz, Cal

Qtz Qtz Qtz Plg Plg

Granite Granite Gabbro O1-Gabbro Gabbro

243-252 259-292

-27.0 -35.0


!

~,,i~ ~,,i~

N

> Appendix continued 11 SW-Indian Ridge

D

Rock

Plg/rock

12 SW-Indian Ridge 13 SW-Indian Ridge

D D

Rock Rock

Plg/rock Pig

~r

O1-Gabbro Norite Pegm-Gabbro Felsic rock

234-268 -54.0

-33.0

Kelley & Fr~ih-Green, 2001

333-404 -71.0 190-318 -44.0

-7.0 -28.0

Kelley & Fr~ih-Green, 2001 Kelley & Fr~ih-Green, 2001

Hydrothermal alteration SW granites, England

D/C

SW granites, England

D/C

Amba-Dongar, India

D

SW granites, England Tui mine, New Zeland Tui mine, New Zeland SW granites, England

D/C D D D/C

Bushveld complex, South Africa Bushveld complex, South Africa Tui mine, New Zeland Tui mine, New Zeland Babine Lake, British Columbia Babine Lake, British Columbia Babine Lake, British Columbia Bushveld complex, South Africa Beregovo deposit, Ukraine

C C D C n.s. D D C C

Milos Island, Greece

C

# (1) Location

Qtz-vein (Pb-Zn-Ti) Qtz-vein (Pb-Zn-Ti) Fl-vein

Granite

110-135

-45

Qtz

Granite

110-135

-36

2

Alderton & Harmon, 1991

115-150

-55

-46

230-400 250 250 250-300

-39 -24 -17 -41

-16

Palmer & William-Jones, 1996 Alderton & Harmon, 1991 Robinson, 1974 Robinson, 1974 Alderton & Harmon, 1991

300-600 300-600 350 350 400

-63 -38

-101 -100 -94 -29 -52

Schiffries & Rye, 1990 Schiffries & Rye, 1990 Robinson, 1974 Robinson, 1974 Zaluski et al., 1994 Sheets et al.. 1996 Sheets et al.. 1996 Schiffries & Rye, 1990 Vityk et al., 1993

-0.3

Naden et al., 2003

F1

Qtz-vein (Sn-W) Qtz-vein (Au) Qtz-vein (Au) Qtz-vein (Cu-Pb-Zn) Hydro-vein Qtz-vein Qtz-vein (Au) Qtz-vein (Au) Qtz-vein(CU) Qtz/Cc-vein Qtz/Cc-vein Qtz-plug Mineralized Qtz-vein Vein

Method(2) Sample

F1

Qtz Qtz Sph Qtz

Carbonatite Granite Andesite Andesite Granite

Qtz Qtz Gal Qtz Qtz Qtz Qtz Qtz Qtz

Gabbro Green. Green. Gabbro Andesite Andesite Rhyodacite Porphyry Cu-dep. Porphyry Cu-dep. Gabbro Rhyolite tufts

200-250

-71 -53 -70 -35 -151 -153 -150 -48 -94

Qtz

Tuffs-Ignimbrites

150-284

-7.8

Mineral

Host rock

Met. g r a d e

T~

Alderton & Harmon, 1991

-9

~)13CPDB%o min max

Or~

9 9

~,,d.

9

oo

o o~ C~

o

Reference

Contact metamorphism 1 2 3 4

Dome de l'Agout, France Dome de l'Agout, France Weolag Tu-Mo deposit, Korea Dome de l'Agout, France

C C C C

Rock Rock Qtz-vein (Tu-Mo) Qtz-and-seg

Bulk Bulk Qtz Bulk

Calcsilicate Calcsilicate Calcsilicate Schists(bio)

350 Horn.

-5.3 -4.9 -3.4 -14.7

1.2 0.8

Kreulen & Schuiling, 1982 Kreulen & Schuiling, 1982 So et al., 1983 Kreulen & Schuiling, 1982 Appendix continued >

OO

> Appendix continued 5 6 7 8 9

South Cornawall, UK Soutern India Dome de l'Agout, France Dome de l'Agout, France Weolag Tu-Mo deposit, Korea

4~ Qtz-vein(W-Sn) Qtz Rock Qtz-Grt Qtz-seg Qtz Qtz-seg Qtz Qtz-vein (Tu-Mo) Py

D C C C D C D D D D D D D P D C D D C P D C D C C D C C D

Qtz CalcSchistss Qtz-vein Qtz Marble Qtz-seg Qtz Marble Qtz-seg Qtz Marble Qtz-seg Qtz Marble Qtz-vein Qtz Dolomites Rock Qtz Flysch Qtz-vein(Pb-Zn) Qtz Flysch Qtz-vein(Sn-Cu) Qtz Flysch Qtz-vein Qtz Anphibolites Qtz-vein Qtz, Cal, Ba Metagabbro Qtz-vein Qtz Granitoides Shear zone Qtz Granitoides Qtz-vein Qtz Pegmatite Rock Qtz Orthogneiss Qtz-vein Qtz Orthogneiss Qtz-vein Qtz Orthogneiss Qtz-vein Qtz Pelite Qtz-vein Qtz Gneiss Qtz-seg Qtz Gneiss Rock Qtz Paragneiss Qtz-vein Qtz Paragneiss Qtz-vein Qtz Metasediment Qtz-vein Qtz Schists(Bt) Qtz-seg Qtz Schists(Chl) Qtz-seg Qtz Schists Qtz-vein Qtz Schists Qtz-seg Qtz Schists(graphitic) Qtz-seg Qtz Schists(graphitic) Qtz-vein(Au)

Regional metamorphism 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29

Tauern Window, Austria Naxos, Greece Naxos, Greece Naxos, Greece Wilson Terrane, NVL, Antarctica Naxos, Greece South Cornawall, UK South Cornawall, UK South Cornawall, UK Tauern Window, Austria Connemara, Ireland Western Carpathians Western Carpathians southern india Tauern Window, Austria Central Alps, Switzerland Central Alps, Switzerland Wilson Terrane, NVL, Antarctica Naxos, Greece southern india Central Alps, Switzerland Central Alps, Switzerland Western Carpathians Naxos, Greece Dome de l'Agout, France Central Alps, Switzerland Naxos, Greece Naxos, Greece Okanagan Valley, British Columbia

Flysch Gneiss Gneiss(mu) Gneiss(sill) Hornfels

D P C C C

200-440 Gran. 250

Amph. Amph. Amph. Amph. Amph. Amph. Low grade Low grade Low grade Amph. Green. Green 9 Gran. Amph. Green. Green. Green. Amph. Gran. Green. Green. Green. Amph. Green. Green. Amph. Amph.

400-600 385 565 620 110-150 150-250 290-330 400-600 118-250 300-400 450-500 400-600

680

300-400 530

380-690 380-690 268-286

-11.9 -13.3 -8.5 -10.3 -15.2

-9.0 -6.3 -12.0 -9.4

Wilkinson et al., 1995 Jackson et al., 1988 Kreulen & Schuiling, 1982 Kreulen & Schuiling, 1982 So et al., 1983

-3.8 -2.9 1.7 3.6 -0.1 -7.2 -13.9 -3.9 -14.1 -7.0 -16.8 -10.5 -9.2 -9.5 -5.3 -10.5 -7.0 -3.4 -3.5 -15.8 -6.5 -4.7 -11.8 -8.0 -16.6 -9.4 -5.0 -13.0 -8.9

-1.5

Hoefs & Morteani, 1979 Rye et al., 1976 Rye et al., 1976 Rye et al., 1976 Frezzotti et al., 2000 Kreulen, 1980 Wilkinson et al., 1995 Wilkinson et al., 1995 Wilkinson et al., 1995 Hoefs & Morteani, 1979 O'Reilly et al., 1997 Hurai et al., 1997 Hurai et al., 1997 Harris et al., 1993 Hoefs & Morteani, 1979 Hoefs & Stalder, 1977 Hoefs & Stalder, 1977 Frezzotti et al., 2000 Rye et al., 1976 Harris et al., 1993 Hoefs & Stalder, 1977 Hoefs & Stalder, 1977 Hurai et al., 1997 Rye et al., 1976 Kreulen & Schuiling, 1982 Hoefs & Stalder, 1977 Kreulen, 1980 Kreulen, 1980 Zhang et al., 1989

2.5 5.0

-7.7 -1.8 -3.4 -4.0 -3.1 -6.2 -4.4 -9.0 -2.0 -2.0 -1.0 -5.0 -8.2


r

!

C~

N

> A p p e n d i x continued 30 Soutern Alps, N e w Zeland

t~ t~

D

Qtz-vein

Qtz

Schists(Qtz-Feld)

Amph.

200-320

-11.7

-8.2

Jenkin et al., 1994a t~

Volcanic rocks 1 2 3 4 5 7 6 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27

Loihi S e a m o u n t Scotia sea Mariana t r o u g h Loihi S e a m o u n t Indian ridge Mid Atlantic Indian ridge Mid Atlantic Indian ridge Juan de Fuca Explorer S e a m o u n t h Hawai Juan de Fuca Ridge Sea Rise east pacific rise Kilauea East rift C a y m a n Rise Famous zone Mid Alantic Ridge Galapagos Galapagos Ridge Loihi S e a m o u n t C a y m a n Rise Famous zone Mid Alantic Ridge Kilauea East rift Famous zone Mid Alantic Ridge Rita zone East pacfic rise Loihi S e a m o u n t N o r t h w e s t e r margin, N o r t h America 28 Marianna arc

P P P P P P P P C P P P P C P P C P P P P P P P C P P

Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock Rock

Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Glass Bulk

P

Rock

Glass

C

Pocket peg

P

Rock

Alkali basalt BABB BABB Basanite MORB MORB MORB MORB MORB MORB OIB OIB Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Tholeiite Trans basalt Xenolites

-8.3 -12.0 -11.8 -5.8 -11.7 -5.3 -10.5 -7.8 -5.8 -7.4 -5.4 -6.7 -24.7 -9.1 -8.8 -7.1 -7.0 -6.4 -6.4 -6.3 -6.3 -6.2 -6.1 -6.0 -5.6 -5.9 -9.9

-6.5 -9.0

-7.3 -6.7 -5.3

-3.5 -22.3 -4.2 -4.9 -6.0 -5.5 -5.6 -5.8 -5.1

-5.8 -0.7 -4.0

Exley et al., 1986 Mattey et al., 1984 Mattey et al., 1984 Exley et al., 1986 Mattey et al., 1989 Mattey et al., 1984 Mattey et al., 1989 Exley et al., 1986 Mattey et al., 1989 Mattey et al., 1984 Mattey et al., 1984 Mattey et al., 1984 Sakai et al., 1984 Pineau & Javoy, 1983 Des Marais & Moore, Des Marais & Moore, Pineau & Javoy, 1983 Des Marais & Moore, Sakai et al., 1984 Exley et al., 1986 Sakai et al., 1984 Sakai et al., 1984 Sakai et al., 1984 Des Marais & Moore, Pineau & Javoy, 1983 Exley et al., 1986 N a d e a u et al., 1990

~r r~

o 9 r~

9 w,,~.

1984 1984

o 9

1984 C~

o

1984

-29.7

-24.7

Mattey et al., 1984

-15.5

-3.0

Taylor et al., 1979

-17.1

-8.3

Miller & Pillinger, 1997

Plutonic rocks San Diego & Riverside Counties, US Cornubian Batholith, England

Gabbro Qtz

Granite

520-565

A p p e n d i x continued >

OO

> Appendix continued Carrock Fell, England Xihuashan Cornubian Batholith, England Dome de l'Agout, France Dome de l'Agout, France Connemara, Ireland Southern india SW-Indian Ridge SW-Indian Ridge SW-Indian Ridge

P P C C C D D D D D

Rock Rock Rock Qtz-seg Peg Qtz-vein Rock Rock Rock Rock

Qtz Qtz Qtz Qtz Bulk Qtz, F1, Ba Qtz Plg/ rock Plg/ rock Pig / rock

SW-Indian Ridge SW-Indian Ridge Soutern India

D D P

Rock Rock Rock

Plg/ rock Plg/ rock Qtz-Grt

D D

Qtz pebbles Qtz pebbles

Sedimentary rocks Stanleigh Mine, Canada Venterrdorp Contact Reel South Africa Snowbird, British Colunbia, Canada Providencia, Mexico Providencia, Mexico

D

Qtz-vein(Au, Sb)

C C

Granite Granite Granite Granite Granite Granite Charnockite O1-Gabbro Gabbro O1-Gabbro Norite Pegm-Gabbro Felsic Vein Charnockite

125-205

Conglomerates Conglomerates

-16.0 -9.7 -8.3 -11.8 -2.7 -18.7 -12.4 -11.7 -10.2 -10.4

-16.0 -3.0 -7.4 -8.9 -9.5 -8.1 -1.9 -2.4 -4.6

Miller & Pillinger, 1997 Miller & Pillinger, 1997 Miller & Pillinger, 1997 Kreulen & Schuiling, 1982 Kreulen & Schuiling, 1982 O'Reilly et al., 1997 Harris et al., 1993 Kelley & Friih-Green, 2001 Kelley & Fr/.ih-Green, 2001 Kelley & Fr/ih-Green, 2001

-23.9 -24.9 -11.9

-1.7 -8.4 -6.3

Kelley & Fr/ih-Green, 2001 Kelley & Fr/ih-Green, 2001 Jackson et al., 1988

-8.0

1.5

1.1 -8.6

Qtz

Listwanites

240

-10.0

Sph Cal

Limestones Limestones

290-365 330-350

-7.0 -11.0

T~

~}180%o

Vennemann et al. 1992 Vennemann et al., 1992 Madu et al., 1990 Rye & O'Neil, 1968 Rye & O'Neil, 1968

r =r r~

# (1) Location

McArthur Pass, Canada Vermillion Pass, Canada Nahanni Butte, Canada Berry, Canada Kotaneelee, Canada Dabsun Lake, China Mt. Seldom, Israel Inghilterra ? Yunan Province Yunan Province

Method(2) Sample

C C C C C M C C C M

Dol Dol Dol Dol Dol

cement cement cement cement cement

Cal speleothem

Mineral

H o s t rock

Met. g r a d e

min Dol Dol Qtz, Dol Dol Dol Halite Halite Cal Halite Halite

Dolomite Dolomite Dolomite Dolomite Dolomite Evaporite Evaporite Evaporite Evaporite Evaporite

142 145 163 166-175 170

-13.7 -11.4 -18.3 -5.7 -3.7 0.93 -1.9 -8.44 -2.2 -21.9

Reference !

max

-11.1 -11.2 -1.7 1.26 2.6 -5.2

Yang et al., 1995 Yang et al., 1995 Yang et al., 1995 Yang et al., 1995 Yang et al., 1995 Yang et al., 1996b Ohba & Matsuo, 1988 Dennis et al., 2001 Horita & Matsuo, 1986 Horita & Matsuo, 1986 Appendix continued >

=r

~,,ao

N

> Appendix

continued

Hubei Province Hubei Province Qinghai Province Qinghai Province Searles Lake, California Providencia, Mexico Providencia, Mexico Providencia, Mexico Kaneuchi deposit, Japan Beregovo deposit, Ukraine

C M C M C C C C C/D C

Milos Island, Greece

C

Qtz-vein Mineralized Qtz-vein Vein

Halite Halite Halite Halite Halite Sph Qtz Cal Qtz Qtz Qtz

Evaporite Evaporite Evaporite Evaporite Evaporite Limestones Limestones Limestones Rhyolite tufts

290-365 330-315 330-350 150-350 200-250

3.3 -14.2 6 -16.8 -9.8 5.8 -4.5 -0.2 -1.2 -7.1

14.1 6.2 -3.7 0.8 -0.3 -0.9

Horita & Matsuo, 1986 Horita & Matsuo, 1986 Horita & Matsuo, 1996 Horita & Matsuo, 1996 Horita, 1990 Rye & O'Neil, 1968 Rye & O'Neil, 1968 Rye & O'Neil, 1968 Ohba et al., 1995 Vityk et al., 1993

Tuffs-Ignimbrites

150-284

-7.8

-0.3

Naden et al., 2003

1) Reference n u m b e r in Figures 3.1abc and3.2abc; 2) Extraction method, C: crushing; D" decrepitation; D / C: both crushing and decrepitation; P: pyrolysis; M: melting; CpI: capillary; n.s." not specified. 3) TH~ h o m o g e n i z a t i o n temperature for the fluid inclusions or calculated temperature of trapping. Mineral abbreviations according to Kretz, (1983).

o o r~

>

9

r~ o

9


CHAPTER 4 Selected Ion Flow Tube Mass Spectrometry (SIFT-MS) and Flowing Afterglow Mass Spectrometry (FA-MS) for the Determination of the Deuterium Abundance in Water Vapour Patrik Spanell & David Smith2 I V. Cerm~ik Laboratory, J. Heyrovsky Institute of Physical Chemistry, Academy of Sciences of the Czech Republic, Dolejskova 3, 182 23, Prague 8, Czech Republic 2 Centre for Science and Technology in Medicine, School of Postgraduate Medicine, Keele University, Thornburrow Drive, Hartshill, Stoke-on-Trent, ST4 7QB, U.K. e-mail: I [email protected]; 2 [email protected]

4.1 Introduction In this chapter, we describe novel methods by which on-line, real-time water vapour deuterium abundance can be measured in single breath exhalations and above aqueous liquids, including urine and blood. The need for these measurements became apparent as our trace gas analytical techniques began to be used for the determination of trace gas metabolites in breath for clinical diagnosis and therapeutic monitoring (Smith & Spanel, 1996a, b; Spanel et al., 1998). In particular, it became clear that a rapid, non-invasive method to determine total body water in patients suffering from end-stage renal failure would have great value if it could be used directly in the clinical environment. Deuterium isotope dilution within the body coupled with a new analytical approach offers a route to this objective. Deuterium abundance is conventionally determined in liquid water, urine, saliva and blood samples by first equilibrating a sample of these media with gaseous hydrogen in the presence of a catalyst, thus producing a H2/HD mixture above the liquid. Then conventional mass spectrometry, sometimes coupled with gas chromatography (GCMS) is used to quantify H2 + and HD + ions derived from this mixture (Begley & Scrimgeour, 1997, Part 1, Chapter 1). However, this approach requires long-term sample preparation and relatively laborious analysis. Thus, several days pass from sample acquisition to analysis. The value of a method that would provide an immediate result is obvious. In response to this need, we have developed the selected ion flow tube mass spectrometric method (SIFT-MS) and the flowing afterglow mass spectrometric method (FA-MS) for the on-line, real time determination of the deuterium abundance in water vapour. Both methods involve the flow of thermalised H30+(H20)3 ions along a fast flow tube. These ions react in multiple collisions with molecules of water vapour

Selected Ion Flow Tube Mass Spectrometry(SIFT-MS)and FlowingAfterglow ...

89

introduced into the flow tube and their isotopic composition reaches equilibrium and is analysed by a mass spectrometer located downstream. A typical mass spectrum obtained when mass selected hydronium ions, H3160 + (m/z = 19), are used as precursor ions for the analysis of humid air and breath by the SIFT-MS method (Smith & Spanel, 1996a, b) is shown in Figure 4-1a. Additional ions appear at m / z of 20 and 21 in the product ion mass spectrum (see the spectra in Figure 4-1). These product ions are the isotopomers of H30 + containing respectively D, 170, and 180, which are formed in the helium carrier gas by isotope exchange reactions between the injected, thermalised H3160 + ions and the isotopomers of water, i.e. H2160, HD160, H2170 and H2180. Sequences of ion-molecule reactions also occur that form the hydrated hydronium (water cluster) ions H30+~ H30+.(H20)2 and H30 +~(H20)3 at m / z of 37, 55, and 73 respectively, which appear on the mass spectra together with their corresponding D, 170 and 180 isotopomers at m / z 38 and 39, 56 and 57 and 74 and 75 (see Figure 4-1a). Clearlyi there is information in these mass spectra on the isotopic composition of the water present in the carrier gas. Hence, the deuterium content of a water sample introduced into the helium carrier gas can be determined from such spectra. To properly understand that which follows, we need to distinguish between the isotopic composition of the following three "phases"" the liquid water sample (designated by the subscript liq), the water vapour transferred from an aqueous sample headspace into the helium carrier gas (designated by the subscript yap) and the H30+(H20)o,1,2,3 ions that comprise the ion swarm created in the carrier gas (designated by the subscript ion).

4.2 Background science 4.2.1 Partition of HDO between liquid water and its vapour In water containing a low abundance of deuterium almost all the deuterium is contained in HDO molecules. Therefore, in order to determine the deuterium isotope abundance ratio in a liquid water sample, Rlliq = D / ( H + D), by analysing its vapour, the partition of HDO between the liquid and vapour phases needs to be addressed. A difference arises because HDO has a lower saturated vapour pressure than H20 at sub-boiling temperatures. Thus, the deuterium abundance in the headspace vapour, Rlvap, is lower than that in the liquid, Rlliq. The ratio of these parameters, i.e. the temperature dependent dimensionless partition coefficient K 1 - Rlvap/Rlliq, can be calculated from the data and the equations given by Van Hook (1972) and Jancso & Van Hook (1974) using Raoult's law (Atkins, 1990). In Figure 4-2, K1 is plotted as a function of temperature. Similarly, the values of the isotope abundance ratio of 170 in the water vapour, R2vap - 170/(160 + 170 + 180) and that of 180, R3vap - 180/(160 + 170 + 180) are proportional to their corresponding abundance in the liquid R2liq and R3liq. The partition coefficients K 2 - R2vap/R2liq and K3 - R3vap/R3liq derived from the data given by Jancso & Van Hook (1974) are also given in Figure 4-2. Note that K2 is very close to unity and that there is a small but discernible effect for 180 that can be accounted for by K3.

90

Chapter 4- P. Spanel & D. Smith

Figure 4.1 - Spectra obtained by SIFT-MS from water vapour plotted on a semi-logarithmic scale as counts per second (c/s) versus mass to charge ratio (m/z). a) A spectral scan over the m / z range 10 to 100 for tap water when H30 + is injected ( m / z = 19). Note the production of the three hydrates H30+(H20)1,2,3 at m / z values 37, 55 and 73. Note also the appearance of their D, 170, and 180 isotopomers, b) A spectral scan from m / z 72 to 76 for tap water. Note the clear separation of the mass peaks. Mean count rates were obtained for each m / z value by recording and averaging the count rates at four mass settings around each peak (see text), c) A spectral scan from m / z 72 to 76 for a 1% mixture of D20 in tap water. Note the much larger count rate at m / z =74 compared to that in b) for tap water alone

Selected Ion Flow Tube Mass Spectrometry (SIFT-MS) and Flowing Afterglow ...

91

Figure 4.2 -Dimensionless water vapour/liquid water partition coefficients K1, K2 and K3 plotted as a function of temperature for the isotopomers HDO, H2170 and H2180 respectively, calculated with respect to H20 from the data and equations given in Van Hook, 1972, and Jancso & Van Hook, 1974, using Raoult's law (Atkins, 1990). The vertical lines at 30~ and 37~ (body temperature) indicate typical temperatures used for measurements.

4.2.2 Ion molecule reactions; equilibrium between ions and water vapour molecules in the gas phase To determine the deuterium abundance in water vapour using our novel methods, it is essential to understand the ion chemistry that generates the isotopomers of the ions. The isotopic composition of the ion swarm is primarily determined by the kinetics of the ion-molecule reactions. The initial reactions that occur (Spanel & Smith, 2000) are: H30 + + HDO ~,,,i.

9

b,,,i 9

r9 9

r~

C3 r~

L,l

160

Chapter 8 - W. Meier-Augenstein

for sample preparation or sample clean up of complex mixtures for isotope ratio analysis. Again, we are dealing with a form of two-phase partitioning where slight differences in solute/stationary phase interaction leads to mass discrimination. Very recently, Filer presented a comprehensive overview of isotopic fractionation during chromatography (Filer, 1999).

8.2.2. Sample derivatization It is quite remarkable that publications of studies dedicated to the influence of sample derivatization on high-precision CSIA by GC/C-IRMS are few and far between (Meier-Augenstein, 1997a; Meier-Augenstein, 1999). Observations to this effect are usually passed on by word of mouth, be it at conferences and, more recently, as postings to discussion groups on the internet. Since one of this book's intentions is to serve as a reference book, I would like to take this opportunity to bring some of these none the less valuable observations into the public domain. When derivatizating a sample prior to CSIA, one should bear in mind the various ways this can influence the analysis and its results. (I) Isotopic fractionation due to kinetic isotope effects during derivatization. (II) Change of isotopic composition in the target compound(s) due to derivatization agent(s), e.g. tracer dilution. (III) Gas chromatographic consequences. (IV) Potential adulteration of isotope ratio measurement caused by non-quantitative sample conversion.

Isotopicfractionation

During derivatization reactions, kinetic isotope effects come only to bear if one of the reagents reacts non-quantitatively. Any such effect associated with derivatization reactions and theoretical considerations for calculating true ~513C-values from derivatized compounds have been discussed in a recent article (Rieley, 1994). However, in most cases the possibility of kinetic isotope effects during derivatization can be safely ignored. The most widely used derivatization reactions can be divided in three classes, silylation, esterification and acetylation. Silylation, which replaces the hydrogen in a hydroxyl or an amino group with an alkylated silyl group, will not cause a carbon isotopic fractionation since no carbon-containing bond is involved. Since silylation is carried out using a huge excess of the silylation agent it is not likely to cause isotopic fractionation of oxygen or nitrogen either. In esterification and trans-esterification reactions, a carbon centre contained in the final product is involved and, hence, a carbon kinetic isotope effect is possible. These reactions usually employ an excess of derivatization agent, e.g. an alcohol, with a catalyst. Since the carbon of interest is in the compound and these reactions are rapid and usually quantitative, no carbon kinetic isotope effect will be expressed. In acetylation reactions, however, the situation is reversed. The carbon involved in the rate-determining step is in the derivatization agent rather than in the compound of interest. Since the derivatization agent is used in excess and, hence, will not react quantitatively, a carbon kinetic isotope effect at the acetyl carboxy-carbon is possible although the compound of interest may react quantitatively.

GC and IRMS Technology for

13Cand 15N Analysis on Organic Compounds and Related Gases

161

Change of isotopic composition Changes of the 13C/12C isotope ratio due to derivatization are usually a minor consideration unless the true isotopic composition of the target compound(s) must be known. Khalfallah et al. (1993) reported a method to correct for 13C-tracer dilution through carbon added during derivatization of plasma lactate. Demmelmair and Schmidt (1993) described a carbon balance equation to calculate ~3C-values of free amino acids from ~13C-values of their N-acetyl, O-propylate (NAP) derivatives at natural abundance level. However, in most biochemical applications such as in-vivo or in-vitro tracer studies the interest is focused on changes in isotopic enrichment rather than the absolute values. This is easily achieved since measurements are compared against baseline or background samples that have undergone the same sample preparation procedure.

Effects on gas chromatographic separation Next, the effects of derivatization on sample integrity and the quality of GC separation have to be considered. For a variety of reasons, silylation agents have enjoyed increasing popularity over the last decade. Derivatization introducing trimethylsilyl (TMS) or tert-butyl-dimethylsilyl groups (tBDMS) is usually quantitative, reaction conditions are relatively mild, the reactions as such are easily performed and samples can be injected directly after the derivatization has been successfully completed. However, there are other reasons as to why silylation agents should be used with caution when it comes to derivatizing samples for GC/C-IRMS. Derivatization of amino acids and a-ketocarboxylic acids with TMS leads to formation of several derivatives for one compound (i.e. mono-, di-, tri-TMS and cis/trans isomers). Sample stability is another concern. Recent observations with tBDMS-derivatives of amino acids obtained from protein hydrolysates showed that these derivatives can deteriorate in a matter of days despite being stored in dark glass vials at subzero temperature (Hofmann et al., 1995). Derivatization by silylation might also hamper GC separation as the extremely apolar nature of trimethylsilyl (TMS) and tertbutyldimethylsilyl (tBDMS) derivatives can obscure compound characteristics that could otherwise be gas chromatographically exploited. For example, despite the entirely different chemical nature of the pure amino acids, the tBDMS-derivatives of L-aspartic acid and L-4-hydroxyproline exhibit similar gas chromatographic properties. A temperature gradient of 5 ~ starting at 80~ after a 5-minute isothermal step leads to a 100% peak overlap of the two compounds. Only by employing three different temperature gradients separated by two isothermal steps, can a baseline separation be achieved. Furthermore, derivatization with TMS or tBDMS limits one's choice of GC stationary phases to apolar phases such SE30 and SE52/SE54 (for stationary phases and commercial equivalents, please, refer to Table 8.2). Lastly, sample derivatization can interfere with sample conversion into CO2 and N2 thus potentially compromising both accuracy and precision of isotope ratio measurements. For instance, an excessive carbon load introduced by derivatization might result in incomplete combustion thus compromising accurate isotopic analysis. The tBDMS-derivative of L-leucine contains 18 carbons (6 from leucine plus 6 per tBDMS

162


group) and one nitrogen atom per molecule. That means in order to generate I moleequivalent of N2, 36 mole-equivalents of CO2 will be generated. Segschneider et al. (1995) reported that accurate 615N from tBDMS were obtained only when more than 7 nmol of N2 per individual amino acid were introduced into the ion source of the IRMS. For leucine this means that in this case 252 nmol of CO2 were generated; this amount is roughly equivalent to a volume of 5.6 mL. One practical implication of this is that on average the CO2 trap will clog up after three runs only. This excessive carbon load can also lead to a temporary overload of the oxidation catalyst, thus resulting in the additional formation of carbon monoxide. Carbon monoxide cannot be trapped cryogenically and, hence, will enter the ion source together with N2 thus causing isobaric interference with the measurement of N2 isotope ratios (Rennie et al., 1996). Potential interference of silyl groups with the oxidation catalyst should also be taken into consideration. As yet, a potentially negative influence of SiO2, formed during combustion of silylated compounds, on the performance of the oxidation catalyst has not been ruled out conclusively. The formation of siliceous deposits on the oxidation catalyst is one possibility as this would not only reduce its reactive surface area but would also impede gas flow through the combustion reactor. For the same reason, i.e. non-quantitative sample conversion due to interference with the oxygen donor / oxygen sink system, the use of trifluoroacetates (TFA) or heptafluorobutyrates (HFB) should be avoided (Hofmann et al., 1995). Fluorine forms extremely stable fluorides with Cu (CuF2) and Ni (NiF2) thus irreversibly reducing combustion efficiency of the CuO/NiO system. In addition, fluorine irreversibly poisons the platinum that most on-line combustion-IRMS systems use in their combustion reactors. If CuO/Pt is used as combustion catalyst, this effect will show up immediately. Co-injecting equimolar amounts of N-TFA, O-propyl L-leucine and methyldecanoate (both compounds contain 11 carbons) on to a GC/C-IRMS system fitted with a fully oxidized CuO/Pt combustion chamber yielded a peak area for the amino acid that was only half the size of that obtained from the fatty acid (MeierAugenstein, 1997a). In other words, for the TFA-derivatized amino acid only 50% of the expected CO2 yield was produced. Combustion systems based on C u O / N i O / P t seem not immediately effected. This is probably due to the large oxygen sink capacity of the NiO but these systems will deteriorate with time (depending on sample throughput but usually after 100 - 150 sample injections), too.

Overview of derivatization methods Naturally, there are circumstances in which the risks mentioned above can be deemed acceptable. Yet, this decision lies solely with the analyst. Of course, one way of avoiding the problems with derivatization is not to derivatize the sample at all. For free fatty acids, amines and alkanols, this approach involves the use of moderately polar to polar stationary phases and high-temperature GC (Table 8.2). However, not every polar compound is amenable to these techniques, e.g. amino acids, and hightemperature capability of polar stationary phases is limited even when oxygen free helium is used as carrier gas. In the following, amongst others, derivatization

GC and IRMS Technology for 13C and 15N Analysis on Organic Compounds and Related Gases

163

Table 8.2- Stationary phases, their commercial equivalents and typical applications. Stationary Phasea)

Commercial Equivalents

Applications

SE 30 apolar; 100% polymethylsiloxane, 0% functionalities

DB-1, CP-Sil 5 CB, BP1, Rtx-1, HP Ultra-l,

Aliphatic, olefinic and aromatic hydrocarbons; long chain fatty acids and alcohols; waxes; steroids; TMS andtBDMS derivatized compounds.

SE 52 (SE 54)b) apolar; 5% phenyl

DB-5, CP-Sil 8 CB, BP5, Rtx-5, HP Ultra-2

Aliphatic, olefinic and aromatic hydrocarbons; long chain fatty acids and alcohols; waxes; steroids; TMS andtBDMS derivatized compounds; NPP derivatized amino acids.

OV-1701 moderately polar; 7% cyanopropyl, 7% phenyl

DB-1701, CP-Sil 19 CB, BP10, Rtx-1701, HP-1701

Fatty acid methyl esters (FAME); amino acids as NAP derivatives, TFA methylates / propylates, methylor ethyl-chloroformates; alcohols.

OV-17 moderately polar; 50% phenyl

DB-17, CP-Si124 CB, Rtx-50. HP-17

Methyl- and ethylcloroformates of amino acids; ethylchloroformates of organic acids.

OV-225 moderately polar; 25% cyanopropyl, 25% phenyl

DB-225, CP-Si143 CB, BP225, HP-225, Rtx-225

FAME, phenols, ethyl-chloroformates of organic acids; NAP derivatives of amino acids.

Carbowax polar; PEGc)

DB-Wax, CP-Si152 CB, BP20, HP-Wax, HP-20M, Stabilwax

Aldehydes; ketones; esters; alcohols; amines, flavours; FAME; amines.

OV-351 polar; PEG nitroterephthalic acid ester

DB-FFAP, CP-Wax 58 CB, BP21, HP-FFAP

Free fatty acids; flavours; FAME.

Silar 7CP highly polar; 75% cyanopropyl, 25% phenyl

BPX70

FAME; NAP derivatives of amino acids; TFA methylates/ propylates of amino acids.

Silar 10C highly polar; 100% cyanopropyl

CP-Sil 88 CB

FAME.

a) Stationary phases are listed in order of increasing polarity from top to bottom. b) Strictly speaking, SE 54 is not the same as SE 52. However, since there are no commercially available columns coated with 5% phenyl, 1% vinyl polymethylsiloxane, it is usually listed together with SE 52 equivalents (5% phenyl polymethylsiloxane). c) PEG = polyethylene glycol

164


methods for amino acids are mentioned, the products of which will not interfere with one-line sample combustion.

Trimethylsilylates and tert-butyldimethylsilylates (TMS and tBDMS) Silylation agents still enjoy considerable popularity for derivatization of polar functionalities such as hydroxyl- (OH), amino- (NH2) and thiol- (SH) groups. Derivatization introducing TMS or tBDMS groups is usually quantitative, reaction conditions are relatively mild and the reagents can be used off the shelf. In cases where the ready-to-use silylation reagents do not achieve complete derivatization, e.g. OHgroups in sterols, the addition of a reaction promoting base such as pyridine or triethylamine helps to achieve a quantitative reaction. However, as pointed out earlier, silylation should be regarded as a last resort when it comes to sample preparation for GC / C-IRMS. Apart from the potential formation of several derivatives for one compound (i.e. mono-, di-, tri-TMS and cis/trans isomers), derivatisation with TMS or tBDMS adds a large number of additional carbons to the molecule of interest thus, in the case of C isotope ratio measurement, decreasing its 613C-value by 20%0 and more.

Methylation

Methylation is the standard derivatization method for carboxylic acids in general and for fatty acids in particular. With the commercial availability of BF3/methanol solution, there is no need to use the cumbersome HC1- and HBr-methanol solutions. BF3 is a Lewis acid that catalyses the esterification with methanol just as well as HC1 or HBr. Its advantage lies in its high volatility and the virtual absence of corrosive properties.

Ethyl-chloroformates (ECF):

Actuall~ the reaction of amino acids with ethyl chloroformate (ECF) yields their N-carboxy-ethoxy, O-ethylates (Husek, 1995). This is a very fast, easy to use reaction and almost as convenient to perform as silylation reactions. This method can also be used for other organic acids such as hydroxy-, amino- and keto-carboxylic acids (Husek & Matucha, 1997; Husek, 1998). However, in dealing with amino acids this method tends to yield non-quantitative derivatization for Glu, Gln, Tyr, and Hyp. Arg and Asp cannot be recovered at all.

N-pivaloyl, O-isopropylates (NPP): This method, developed by Metges et al., is very reliable and produces good derivatization yields even for problem amino acids such as Glu, Asp, Lys, His and Tyr (Metges et al., 1996; Metges & Petzke, 1997).

N-acetyl, O-proplylates (NAP):

The classic method, however, there are two aspects of the original protocol (Wolfe, 1992) that might cause some inconvenience. First, quantitative removal of HBr is difficult and time consuming. Secondly, esterification using alcohol and inorganic acid requires the presence of a water scavenger and the original protocol suggests

Table 8.3- Overview of widely used derivatization methods together with typical conditions for their GC separation and GC / C-IRMS caveats. C o m p o u n d Class

Derivatization

Typical GC Conditionsa)

Caveats

Long Chain Alcohols

TMS

SE 30:50 ~ (2 min) to 200 ~ at 40~ 200 ~ to 320 ~ at 3 ~/ min

none

r (3

=r 9 o 0-Q

Fatty Acids

Methylation using Methanol/BF3

SE 52 / SE 54:120 ~ (4 min) to 280 ~ at 4 ~

none

H y d r o x y / A m i n o / KetoCarboxylic Acids

TMS or tBDMS

SE 52 / SE 54:80 ~ (5min) to 150 C at 2~ 150 ~ C to 220 o C at 3.5~ 20~ to 300 ~ C

TMS: multiple derivatives for same c o m p o u n d tBDMS: excessive carbon load

Risk of non-quantitative derivatization

r

t,o

>

H y d r o x y / Amino / KetoCarboxylic Acids

Ethylchloroformates

OV-17:60 ~ (3 min) to 300 ~ at 6 ~/ min

H y d r o x y / A m i n o / KetoCarboxylic Acids

Acetyl, methylates

OV-1701:70 ~ (5 min) to 220 ~ at 5 ~/ min

none

9

Amino Acids

TMS or tBDMS

SE 52 / SE 54:50 ~ (5 min) to 150 ~ at 6~ 150 ~ to 300 ~ at 1 2 ~

TMS: multiple derivatives for same c o m p o u n d tBDMS: excessive carbon load

r

o

~..L.

Amino Acids

TFA, methylates TFA, iso-propylates

OV-1701:70 ~ (5 min) to 180 ~ at 3.5~ 180 ~ to 230 ~ at 5~ 10~ to 270 ~

Poisoning of combustion catalyst

Amino Acids

Ethylchloroformates

OV-1701" 60 ~ (3 min) to 100 ~ at 5 ~ 100 ~ to 300 ~ at 10 ~/ min

Risk of non-quantitative derivatization

Table 8.3 continued >

o o

> Table 8.3 continued Amino Acids

N-Pivaloyl, iso-propylates

SE 52 / SE 54:70 ~ (1 min) to 220 ~ at 3 ~ 220 ~ to 300 ~ at 10 ~/ m i n

none

Amino Acids

N-Acetyl, propylates

OV-1701:70 ~ (5 min) to 200 ~ at 4~ 200 ~ to 300 ~ at 6 ~/ min

none

a) The information given in these columns should be read as generalized guidelines providing a starting point for the reader to resolve individual analytical tasks.

r

!

r

~,,~o


167

2,2-dimethoxy-propane. Unfortunately, this compound leads to side reactions, which turn the initially clear solution crimson (this even happens when heating a blank solution of the reagents without amino acids being present). The products of these side reactions and, hence, the colour are almost impossible to remove. For GC/C-IRMS analysis of amino acids as their NAP derivatives, the following protocol is suggested which avoids the pitfalls of the classic method. The first step is a trans-esterification reaction that eliminates the need for a water scavenger as no water is being formed during the condensation step. The amino acid(s) [up to a total amount of 20 ~mol] are dissolved in 400 /~L of propylacetate and 200 ~L of propanol/BF3 (14%). The mixture is heated in a crimped autosampler vial at 110~ for 30 min. From the cooled down solution, solvent and excess reagent are removed in a gentle stream of N2 at 60~ The residue is taken up in 500 ~L of propanol/BF3 and blow dried again. Repeat the procedure with 300 ~L of propanol. This reaction ideally results in a solid, crystalline residue of almost colourless appearance. The amino acid propylates are dissolved in a mixture of 200 ~L of acetonitrile and 100 ~L of 1,4 dioxane. To this solution, 150 ~L of triethylamine (mix well) and 90 ~L of dry acetic anhydride (either freshly distilled or stored over molecular sieve 3 ~) are added. The resulting mixture is heat in a crimped autosampler vial at 60~ for 15 minutes (the reaction time can be increased if necessary, but the temperature must not exceed 60~ After cooling down, 300 ~L of CHC13 are added and the mixture is neutralized with 600 ~L of 0.001 M NaHCO3 in two steps (2 times 300 ~L). The aqueous phase is discarded and the chloroform phase is dried over molecular sieves (3 ~, beads, pearl form). Compound classes of biochemical interest, various derivatization methods and corresponding typical GC conditions are summarized in Table 8.3.

8.3 GC/C-IRMS For isotope ratio measurement the analyte must either be a simple gas such as CO2 or N2, or must be converted into a simple gas, isotopically representative of the original sample, before entering the ion source of an IRMS. It is also important to remember that IRMS, in fact, determines the difference in isotope ratio with great precision and accuracy rather than the absolute isotope ratio. IRMS measurements yield the information of isotopic abundance of the analyte gas relative to the measured isotope ratio of a standard or reference gas. This is done to compensate for mass discriminating effects that may fluctuate with time and from instrument to instrument. From the above it is patently obvious that a GC cannot be directly coupled to an IRMS. The need for sample conversion into simple gases has prompted the design of a combustion interface where the GC effluent is fed into a combustion reactor. This reactor, either a quartz glass or ceramic tube, is typically filled with CuO/Pt or CuO/ NiO/Pt and maintained at a temperature of approximately 820~ or 960~ respectively (Merritt & Hayes, 1994; Merritt et al., 1995a). The influence of combustion tube packing on analytical performance of GC/C-IRMS has been reported by Eakin et al. (1992) and, more recently, by Ellis & Fincannon (1998). To remove water vapour generated during combustion, a water trap is required. Most instrument manufacturers

168


employ a Nation [Perma Pure Inc.] tube for this purpose. Nation is a fluorinated polymer that acts as a semi-permeable membrane through which H20 passes freely while all the other combustion products are retained in the carrier gas stream. Quantitative water removal prior to admitting the combustion gases into the ion source is essential because any water residue would lead to protonation of CO2 to produce HCO2 +, which interferes with analysis of 13CO2(isobaric interference). Very recently, Leckrone & Hayes (1998) published a detailed study of this effect. In contrast to dual-inlet IRMS machines, GC/C-IRMS systems produce almost Gaussian shaped signals. In addition, due to the "chromatographic isotope effect" (Matucha et al., 1991; Matucha, 1995; Filer, 1999) the m / z 45 signal (13CO2) precedes the m / z 44 signal (12CO2) by 150 ms on average (Rautenschlein et al., 1990; Brand, 1996), an effect not observed in ordinary Continuous Flow-IRMS (CF-IRMS) systems. This time displacement depends on the nature of the compound and on chromatographic parameters such as polarity of the stationary phase, column temperature, and carrier gas flow (Meier-Augenstein et al., 1996). Therefore, loss of peak data due to unsuitably set time windows for peak detection and, hence, partial peak integration will severely compromise the quality of the isotope ratio measurement by GC/CIRMS. The same is true for traces of peak data from another sample compound due to close proximity resulting in peak overlap with the sample peak to be analysed. Due to the fact that isotope ratios cannot be determined accurately from the partial examination of a GC peak, high resolution capillary gas chromatography (HRcGC) resulting in true baseline separation for adjacent peaks is of paramount importance for high-precision CSIA. It should be noted that the chromatographic isotope effect is not caused by a vapour pressure effect. It is rather the result of different solute/stationary phase interactions that are dominated by Van der Waals dispersion forces leading to an earlier elution of the heavier isotopomer (Matucha et al., 1991). This difference in chromatographic solute/stationary phase interaction is caused by lower molar volumes of the labelled, and thus heavier, compounds. The reason for the decrease in molar volume is the increased bond strength and thus shortened bond length between 13C-H and, to a lesser degree, 12C-13C,and 12C-H and 12C-12C,respectively. A problem not easily overcome is the partial peak overlap of compounds showing similar characteristics in both chemical and chromatographic respect. There is many an application where closely related compounds simply cannot be resolved resulting in peaks overlap. To make matters worse, CO2 and N2 disperse more freely within the carrier gas stream than their parent organic compounds which can result in overlapping CO2 peaks even for baseline resolved GC peaks. To extract valuable information obscured by such peak overlaps, Goodman & Brenna (1994, 1995) suggested software algorithms for improved data processing. These algorithms were based on curve fitting rather than the summation (the industrial standard) using combinations of exponentially modified Gaussian (E) and Haarhoff/Van-der-Linde (H) functions and were tested on up to 70% valley peak overlap. When the adjacent peaks were of equal abundance (leading peak" trailing peak 1"1) combinations of HE and HH appeared to


169

provide the best recovery of isotope ratios. In the case of unequal abundance in favour of the leading peak (10:1), the HH combination gave the best accuracy. When the abundance was reversed (1"10), the EH combination provided the best accuracy but only for peak overlap up to 40% valley. Despite these encouraging results, curve-fitting algorithms for restoring lost accuracy have not been incorporated into any commercial IRMS data reduction software by the manufacturers. However, any progress in this direction aided by full evaluation of new algorithms for routine use (under 'real life conditions') requires user access to such algorithms and hence their incorporation into IRMS manufacturers' software.

8.3.1 High resolution capillary gas chromatography (HRcGC) As pointed out earlier, baseline separated gas chromatographic peaks are the basis for high-precision CSIA. To achieve this goal, in the first instance, basic gas chromatographic rules must be observed: (1) The polarity of the stationary phase should meet the polarity of the sample constituents (Table 8.2). (2) Column head pressure and, hence, carrier gas velocity should be set to suit column diameter (Table 8.4). (3) Temperature gradients should be chosen that exploit the maximum of the column length (i.e. the longer the column, the slower the temperature rise per minute; Table 8.4). (4) Samples should be injected in splitless mode. However, to avoid peak broadening and peak tailing, the split should be opened 10 seconds after injection. (5) Use of a retention gap (also known as pre-column or guard-column) will prolong the lifetime of the GC column and improve chromatographic performance. The retention gap (a deactivated piece of fused silica capillary tubing, typically 1 to 3 m long) helps focusing the injected sample thus leading to sharper peaks. Table 8.4 - Recommended values for carrier gas velocity and temperature gradient according to column length when using helium as carrier gasa). Column length [m] 10 15 20 25 30 40 50

Elution of methane [s]b) 35 53 70 88 105 140 175

Temperature gradient [~ 2.5 1.65 1.25 1.05 0.84 0.63 0.5

a) Based on working directions given by K. Grob (Grob, 1986). b) Set GC oven temperature to 30 ~ Set split ratio to about 1:30, inject a few ~L of natural gas (or lighter gas) and measure elution time of the first peak (FID signal). Adjust column head pressure to match recommended elution time.

170


Further to these principles, HRcGC techniques such as multi-dimensional capillary GC (MDcGC), enantio-selective GC, porous layer open tubular (PLOT) column GC for analysis of VOCs and high-temperature capillary GC (HTcGC) are powerful tools for high-precision CSIA when used in combination with GC/C-IRMS. Nitz and co-workers were the first to report the advantages of using MDcGC in GC/C-IRMS (Nitz et al., 1992). MDcGC is now, often in combination with enantioselective GC, almost exclusively used in authenticity control of flavours and fragrances by CSIA (Casabianca et al., 1995; Juchelka et al., 1998). Similarly, HTcGC is strongly associated with CSIA of steroids and long-chain fatty acids, e.g. (Woodbury et al., 1998a). It is often inferred that high temperature capillary GC is limited to the use of crossbonded apolar stationary phase such as SE 30 and SE 52 and therefore only of limited applicability. This assumption is mainly based on the maximum allowable operating temperature (MAOT) given by GC column manufacturers. Of course, strictly speaking high temperature capable GC columns should be purpose-made. However, most stationary phases except for polyethylene glycol (PEG) based ones, can be safely used up to 360~ (apolar phases up to 450~ if certain precautions are observed. (I) To avoid damage to the stationary phase at high temperatures, the carrier gas must be free of oxygen and moisture traces. Considering the cost for replacing a GC column, the investment in a high capacity gas purifier will soon have paid off. (II) It is a remarkably little known fact that polar stationary phases are light sensitive (Grob, 1986). Even exposure to indirect daylight or light from fluorescent tubes will lead to column deterioration. The first sign of this happening is usually a dramatic increase in column bleed. (III) Conditioning the column properly prior to high temperature usage will minimize normal column bleed (e.g. at 4 ~ to 300~ maintain at 300~ for 5 hours; repeat but program to 360~ and maintain at 360~ for 2 hours). (IV) If a polar phase is to be used for HTcGC, chose a polysiloxane based phase with a high cyanopropyl content. In contrast to common belief, cyanopropyl substituted polysiloxane phases are more stable and inert than phenyl substituted phases. Kurt Grob suggested that this is caused by high surface tension of phenyl substituted phases which in turn causes problems with film stabilization during column coating (Grob, 1986). Regrettably, the achievements of HRcGC in terms of optimum peak-shape and baseline separation are likely to be impaired during combustion and the subsequent passage through the interface. Changes in tubing diameter and frequent use of unions to connect the various parts of tubing lead to a loss in peak definition (peak broadening; peak distortion) and even to partial peak overlap. All of the aforementioned have a detrimental effect on accuracy and precision of isotope ratio measurement (MeierAugenstein et al., 1996). Very recently, Keith Goodman reported a single-capillary interface design (SCID) which he developed to overcome these problems. As the name suggests, a single capillary was used to connect the GC column to the open-split in front of the IRMS (Goodman, 1998). This capillary was threaded through a furnace and accommodated 2 CuO wires positioned thus as to coincide with the furnace dimensions. So far, this design has been tested for 13C isotopic abundance analysis of n-alkanes.

GC and IRMS Technology for

13Cand 15N Analysis on Organic Compounds and Related Gases

171

8.3.2 Isotopic fractionation and isotopic calibration during GC/C-IRMS analysis Owing to its unique design, it is not possible in GC/C-IRMS to calibrate target compounds against a standard of known isotopic composition, introducing the standard in exactly the same way as the analyte. There are only three feasible means of introducing a standard; (a) addition of reference compounds to the sample, (b) introduction of reference gas pulses to the carrier gas stream or (c) introduction of reference gas pulses directly into the ion source. The results of an extensive study into methods of isotopic calibration by Merritt et al. (1994) emphasised these demands. Comparing the use of internal reference compounds with the introduction of reference gas pulses directly in the ion source of the IRMS, Merritt et al. (1994) found an offset of > 2%0 between the two methods in the case of incomplete combustion and other systematic errors affecting only the analytes. These systematic errors affected both the analytes and the co-injected reference compounds but were not reflected by the external reference gas pulses. Other groups reported similar observations (Caimi et al., 1994; Caimi & Brenna, 1996). In the absence of such systematic errors, Merritt et al. (1994) found that both methods of isotopic calibration gave consistent results as long as multiple reference peaks were used to permit drift correction. Only one reference peak for isotopic calibration, albeit from an internal reference compound, is not enough to compensate for the influence of GC parameters such as analyte/stationary phase interaction or column temperature on measured isotope ratios (Meier-Augenstein et al., 1996). There are several stages during GC/C-IRMS analysis where mass discrimination and, hence, isotopic fractionation can occur. Closer inspection identifies seven potential sources. (1) Isotopic fractionation during sample injection (which can be overcome by oncolumn or time programmed splitless injection). (2) Chromatographic isotope effect. (3) Chromatographic peak distortion (leading and trailing peak tail). (4) Combustion process. (5) Peak distortion of N2/CO2 gas peak during passage of the combustion interface. (6) Changing flow conditions at the open split prior to the IRMS. (7) The IRMS itself. Obviously, the external reference gas pulses only compensate for item (7), whereas internal reference-compounds reflect all of the aforementioned. Recently, a method for isotopic calibration was reported that, provided a combustible gas was used, could reflect the systematic errors caused by items (4) to (7). This method combines the convenience and practicability of external reference gas calibration with the advantage of reflecting the majority of physical influences to which analytes are subjected in a GC/ C-IRMS system (Meier-Augenstein, 1997b).

172


8.4 Examples of high-precision CSIA using GC/C-IRMS High-precision CSIA of 13C isotopic abundance at both natural abundance and low enrichment levels yields measurements of 613C-values with a precision of 0.3%0 or better. Thanks to this high precision, even minute changes in the 13C isotopic composition can be reliably detected. Not surprisingly therefore, GC/C-IRMS has become the method of choice to determine the origin of a given organic compound by measuring its characteristic isotope 'finger print'. One area where HRcGC techniques in conjunction with GC/C-IRMS are unsurpassed is authenticity control of flavours, fragrances and essential oils. The potential of GC/C-IRMS for this kind of application has been very successfully exploited and developed further. The powerful combination of GC/C-IRMS with enantioselective capillary GC and MDcGC for authenticity control of flavours, wines and spirits have been reviewed recently by A. Mosandl (Mosandl, 1995; Mosandl et al., 1995; Mosandl, 1997). Another aspect of authenticity control is food quality control to prove and prevent fraudulent adulteration of a high quality product such as maize germ oil or virgin olive oil with oils of minor quality (Woodbury et al., 1995; Woodbury et al., 1998b; Woodbury et al., 1998a; Angerosa et al., 1997) High-precision CSIA of fatty acids in general is one application of GC/C-IRMS at both natural abundance and enrichment level that has yielded a wealth of information (Meier-Augenstein, 2002). Ecological Sciences have greatly benefited from this technique using CSIA of fatty acids to study terrestrial and marine ecosystems, food webs, even migration patterns of birds. Similarly, Bio-medical Sciences have exploited this technique using 13C-enriched fatty acids to study precursor/product relationships of poly-unsaturated fatty acids in-vivo. Furthermore, GC/C-IRMS was successfully used to trace the origin of narcotic drugs. Variation in 13C isotopic abundance of heroin was found to be dependent on its geographical site of production (Desage et al., 1991). Variations in both ~13C and 615N values provided valuable information to discriminate the different origin of several batches of confiscated 3,4-(methyldioxy)-methylamphetamin (MDMA, Ecstasy) tablets (Mas et al., 1995). Coca leaves from South America were found to vary in their ~13C (-32.4 %o to-25.3 %o) and 615N (0.1-13.0%o) values (Ehleringer at al., 2000). Humidity levels and the length of the rainy season and differences in soils were thought to affect the fixation processes and cause the observed subtle variations in 13C and 15N contents, respectively. In conjunction with the variations of trace alkaloids (truxilline and trimethoxycocaine) contents found in cocaine, researchers were able to correctly identify 96 % of 200 cocaine samples originated from the regions studied. High-precision CSIA might also provide the means for an unambiguous doping test in athletes. Independent work by two groups showed a drop in 613C values of endogenous 5a- and 5b-androstanediol from -26.52%o to -30.21%o, in some cases even


173

down to -32.44%o, after synthetic testosterone was administered (Becchi et al., 1994; Aguilera et al., 1996; Shackleton et al., 1997). This drop in 613C value lasted for up to 10 days before g13C values of androstanediol returned to their former baseline value. Stable isotope analysis of natural gases, long chain alkanes and other fossil biomarkers by GC/C-IRMS is used to determine origin and maturity of natural gas and crude oil sources (Bakel et al., 1994; Bjoroy et al., 1994; Hughes et al., 1995; Berner & Faber, 1996), to identify sources of oil spills and oil pollution (Uzaki et al., 1993; Mansuy et al., 1997) as well as ocean transported bitumen (Dowling et al., 1995) and to characterise refractory wastes at heavy-oil contaminated sites (Whittaker et al., 1995; Whittaker et al., 1996). Applying high temperature capillary GC/C-IRMS, Evershed and co-workers gleaned new insights in paleo-dietary habits and prehistoric lifestyle from organic residues preserved in archaeological artefacts (Evershed et al., 1994; Evershed et al., 1995; Evershed et al., 1997; Raven et al., 1997). In a multidisciplinary approach, variations in 615N values from soil amino acids were used to determine differences in land use during Bronze Age, medieval and early modern times. One of the observations made were the consistently low levels of 15N abundance (615N < 0%o) in the amino acids Thr and Phe from soils of unmanured sites of cereal production. All the other amino acids showed positive 615N values (Simpson et al., 1997). Interestingly enough, a similar pattern in 615N values of free plasma amino acids was found in fasting human subjects (Metges & Petzke, 1997). Of course, this is only a limited selection out of the wide spectrum of GC/C-IRMS applications. However, over the last few years several review articles have collated publications in the field of GC/C-IRMS, to which the interested reader is referred (Preston, 1992; Whittaker et al., 1995; Brenna, 1994; Brand, 1996; Brenna et al., 1997; Meier-Augenstein, 1999).

8.5 Hyphenated GC/C-IRMS Techniques In recent years, the research efforts of different groups working in the field of GCIRMS have focused on extending the scope of on-line CSIA towards the measurement of organic 180/16 0 and organic 2H/1H isotope ratios. Research has also been aiming at high-precision measurement of two different elemental isotope ratios such as 2H/ 180, 13C/180 and 13C/15N, from the same compound source in one analytical run. In addition, hybrid systems have been developed that enable CSIA while at the same time recording a conventional mass spectrum of the target compound to aid its identification (Meier-Augenstein et al., 1994a; Meier-Augenstein, 1995; Hall et al., 1999). Placing a PORAPLOT Q capillary column between the combustion reactor and the IRMS, thus achieving baseline separation of N2 from CO2 by 100 s, 15N/14N and 13C/ 12C isotope ratios from alanine, leucine and phenylalanine could be measured in one single gas chromatographic analysis (Rennie et al., 1996). The excellent separation of N2 from CO2 provided ample time to switch IRMS ion source parameters from N2- to

174


Figure 8.2- Dual isotope measurement of 615N and 613C from alanine (as N-acetyl, O-propyl derivative) during the same analysis. The arrow indicates when the ion source parameters were switched from nitrogen mode (m/z 28) to carbon dioxide mode (m/z 44). A 100 s baseline to baseline separation of N2 from CO2 was achieved by passing the combustion products past-reduction through a PORAPLOT Q column, held at 35 ~ The results obtained from dual-isotope analyses (615N:-7.78 + 0.10%o vs. air; 613C:-40.22 + 0.14%o vs. PDB) were in good agreement with those obtained from separate analyses (615N: -7.86 + 0.38%o vs. air; 613C:-40.20 + 0.21%o vs. PDB).

CO2-mode (Figure 8.2). In 1994, using a GC-based IRMS system, Brand et al. showed that CSIA of 1 8 0 / 1 6 0 ratios was possible by converting oxygen-containing organic compounds on-line to CO by means of a pyrolytic reaction (Brand et al., 1994). The on-line coupling of GC and IRMS via a pyrolysis interface (GC/Py-IRMS) was used for the simultaneous determination of ~513C- and 6180-values for vanilla from different origins (Hener et al., 1998). Farquhar et al. converted bulk plant matter into N2 and CO by an automated on-line pyrolysis-based reaction using nickelized carbon at about 1100~ and separating N2 from CO post-pyrolysis in a GC fitted with a 5 ~ molecular sieve PLOT column (Farquhar et al., 1997). Independently, Begley & Scrimgeour (1997) reported on high-precision 62H and 6180 measurement for water and VOCs by using 20% nickelized carbon to generate both H2 and CO at temperatures of between 1050~ and 1100~ Standard deviation was + 2%o and + 0.3%o for 62H and 6180, respectivel~ with samples ranging from urine, water and VOCs. Their pyrolysis system was based on earlier work that was aimed at simultaneous 62H and 6180 determination from small water and urine samples (0.5 mL) (Begley & Scrimgeour, 1996). Common to both studies was the use of a

GC and IRMSTechnologyfor 13Cand 15NAnalysison OrganicCompoundsand RelatedGases

175

novel IRMS with the high dispersion necessary for separation of the 2HIH+ and 4He+ ion beams (Prosser & Scrimgeour, 1995). Very recently, this highomaSs dispersion IRMS was coupled to a GC via a pyrolysis interface including a 5 A molecular sieve PLOT column to achieve CSIA for 2H of fatty acids. In a technical brochure, 62H-values for 16:0 and 18"1 fatty acids (as methyl esters) from tuna oil were reported as148.5 +_4:.1%o and-155.3 _+1.0%o (vs. VSMOV) respectively (Scrimgeour et al., 1999). This GC/Py-IRMS system has also been applied to the measurement of lipid synthesis in humans, using deuterium oxide (D20) incorporation into fatty acids. Administering safe deuterium enrichments (< 0.5 atom%), the labelled fatty acids did not to contain more than one deuterium per molecule of fatty acid, and showed similar chromatographic behaviour to natural abundance samples. Following overnight incorporation of D20, plateau palmitate enrichments were measured by GC/Py-IRMS with a relative standard deviation of 0.5% (Scrimgeour et al., 1999). A different approach to CSIA for H/D ratios of organic compounds such a ethyl benzene and cyclohexanone was published by Tobias & Brenna (1996). Initially using a two stage reactor interface (CuO at 850~ followed by Ni held at 950~ (Tobias & Brenna, 1996), they found that better precision for 62H was achieved by employing an empty alumina tube held at about 1150~ (Tobias & Brenna, 1997). Since their IRMS was not capable of fully resolving sample derived 2HIH from excess 4He carrier gas, they used a heated Pd filter in conjunction with a make-up pressure unit to prevent He from entering the IRMS. This set-up selectively admitted only hydrogen through the Pd foil membrane into the ion source. The measurement of intra-molecular variations in isotopic abundance due to kinetic isotope effects during biosynthesis is another recent development to extend the scope of GC-IRMS. The group around H.L. Schmidt employed 13C isotope pattern analysis for distinction of natural compounds from corresponding synthetic products (Weber et al., 1997; Weilacher et al., 1996). In 1997, an on-line pyrolysis system for position specific isotope analysis (PSIA) of selected compounds from a complex mixture was described in detail for the first time (Corso & Brenna, 1997). They coupled a GC (GC-1) for sample separation prior to pyrolysis to the GC (GC-2) separating pyrolytic products of the selected sample compound. Furthermore, they installed a valve into GC-2 to permit separated pyrolysis fragments to be admitted to an organic MS for structure analysis of these fragments. 8.6 Conclusion

High-precision CSIA by GC/C-IRMS of gases and organic compounds at natural abundance and low enrichment levels is a powerful tool that provides quantitative information such as bio-availability, assimilation, turnover, incorporation and metabolism in biological and ecological systems. If great care is taken to avoid the pitfalls associated with sampling, sample preparation and sample separation, GC/C-IRMS will yield insights into complex systems offering answers to biochemical, physiological, and environmental questions that cannot be obtained with any other analytical instrumentation.

176

Acknowledgements

Chapter 8 - W. M e i e r - A u g e n s t e i n

The author gratefully acknowledges financial support through the Dietmar-HoppFoundation for Medical Research and Studies (Walldorf, Germany). Many thanks are due to Dr Helen F Kemp for critical reading of the manuscript.


CHAPTER 9 Preparation of Ecological and Biochemical Samples for Isotope Analysis Mark A. Teecel, 2. & Marilyn L. Fogel2** State University of New York - College of Environmental Science and Forestry, 1 Forestry Drive, Syracuse NY 13210, USA 2 Carnegie Institution of Washington, Geophysical Laboratory, 5251 Broad Branch Rd., NW, Washington, DC 20015, USA e-mail: * [email protected], ** [email protected] 1

9.1 Introduction Stable isotope analyses have proven to be a critical source of information for delineating processes in ecological and ecosystem studies (Rundel et al., 1989; Lajtha & Michener, 1994), which canencompass the study of habitats and the interactions of organisms with their environment. The main power of isotopic analyses resides in the ability to study both specific processes and also to trace sources of materials and flows of energy through complex ecological webs. Biological processes, such as photosynthesis and metabolism, fractionate materials and change isotope compositions. For example, plants fractionate carbon during CO2 uptake and therefore carbon isotope studies are a useful tool to study the photosynthetic process. Conversely, at the whole organism level, the fractionation associated with animal feeding is much smaller. As biological or inorganic matter that transfers between components in a particular ecosystem may have unique isotopic compositions (e.g., Fry & Scherr, 1984; Michener & Schell, 1994; Dittel et al., 1997; Chamberlain et al., 1997; Kelly, 1999), isotopes may be useful as indicators of food web structure and trophic level status. For example, the living components of all terrestrial communities, including humans, as well as many marine communities rely on the primary production of organic matter by photosynthesis to provide energy for consumers. Organic matter and energy are then passed along food chains, with energy being dissipated at each trophic level. Stable isotopes can be one of the most effective tools for tying together these ecological processes that affect, in turn, affect the biosphere, the atmosphere, and the geosphere at both local and global scales (e.g. Gearing, 1991; Quay et al., 1989; Paerl & Fogel, 1993). This chapter is intended as an overview of the application of stable isotope techniques to addressing ecological questions. We offer suggestions for ecologists who might be considering including isotope analyses to their work, as well as for more experienced isotope geochemists who are interested in expanding their research into a new field. The chapter expands on previous work including the several books that

178

Chapter 9 - M.A. Teece & M.L. Fogel

have been published on isotope application to ecological studies (Rundel et al., 1989; Lajtha & Michener, 1994; Griffiths, 1998), and specific texts on methodology for stable isotope use in environmental studies (Coleman & Fry, 1991; Knowles & Blackburn, 1993; Boutton & Yamasaki, 1996). Although techniques in measuring stable isotopes have changed to more automated devices, the information included in specific chapters of these books (i.e., Ehleringer, 1991; Boutton, 1991) is still appropriate. The focus of this chapter is on the proper sampling protocols for accurately determining the natural abundance and distribution of stable isotopes, particularly carbon and nitrogen, in ecological systems and does not address methods that include the addition of isotopically labeled compounds at enriched levels. Furthermore, methods for the analysis of H isotopes in different sample matrices are presented elsewhere in this book and BOX

9.1 - C a s e s t u d i e s o f a n a q u a t i c e c o s y s t e m

A simplified representation of the complex interactions and pathways of energy transfer in an aquatic ecosystem. Examples of external energy and substrate sources are also represented.

F i g u r e 9.1 -

C a s e s t u d y 9.1 -

Role of terrestrial carbon in aquatic food webs

Terrestrial carbon enters aquatic ecosystems via land runoff and may be a significant source of carbon to the consumers of the food web. The size and importance of this external source of carbon to the aquatic food web can be assessed by measuring the stable carbon isotope ratios of different elements of the food web. The isotopic composition of phytoplankton will reflect an aquatic signal (e.g. lacustrine or oceanic signal), whereas isotope analysis of dissolved and particulate organic matter will indicate whether such materials are terrestrial in origin. Measuring the isotopic composition of larval fish, whose diets include phytoplankton and particulate organic matter, will indicate the relative importance of terrestrial carbon to the aquatic food web. Furthermore, analysis of planktivorous and canivorous fish and generalist top feeders (e.g. trout, bass, and birds such as gulls) will confirm the role of terrestrial carbon in the aquatic ecosystem. Such approaches were employed in studies of the aquatic food web in Loch Ness in Scotland (Jones et al., 1998), and in an Australian tropical stream (Bunn et al., 1997).

Preparation of Ecological and Biochemical Samples for Isotope Analysis

Case study 9.2 -

179

Nutrient loading in aquatic ecosystems

Management decisions on issues of water pollution frequently revolve around high concentrations of nutrients, their sources and how they affect water quality. A research study can consider a study of phytoplankton blooms in a lake, and ask the question, "What is the source of nutrients supporting a phytoplankton bloom in a lake?" A study of this sort requires measuring the concentration and distribution of nutrients such as phosphate, nitrate and ammonium, in conjunction with measuring levels of bacterial activity and phytoplankton productivity. Nitrogen isotope measurement of phytoplankton and different sources of nutrients will aid in determining the major source of nutrients which supports the phytoplankton bloom. Dissolved inorganic nitrogen has at least four major sources: terrestrial runoff, regeneration from benthic processes, atmospheric deposition, and regeneration in the water column by bacteria. To best characterize each source, their concentrations and 615N values should be monitored periodically (weekly to monthly) over the course of the year as temporal variations occur. In conjunction with these measurements, bacterial activity should be measured to determine the relative importance of these processes over the year. In addition, it is important to determine the dominant species of phytoplankton which comprises the bloom, and also the change in algal population over the year. Isotopic compositions of the different algal species collected will provide insights into the temporal utilization of different nutrient sources with respect to time. The overall result of this type of combined study will indicate the relative importance of particular nutrient sources to phytoplankton productivity in the lake. Examples of these types of studies include those by McClelland et al. (1997) and Cifuentes et al. (1988, 1989).

the reader is referred to the chapters by Horita & Kendall (Chapter 1, this volume) and Volume IL Part 3, Chapter 1-2/3 on hydrogen isotopes. Our chapter is presented in three sections: Section I deals with the development of experimental approaches and the potential uses of stable isotopes, while Section II describes different methods of storage and preparation of samples for isotopic analysis. In Section III we outline chemical methods for bulk and compound specific isotope analyses of biological materials pertinent to ecological studies.

9.2 Section I: Stable isotope biogeochemistry in ecological research Before undertaking a study in this field, a clear understanding of the ecological questions that are of relevance is of the utmost importance. In any particular ecosystem, the number of plant, animal, water, soil, and air samples available for analysis is almost infinite, and although methods for analyzing bulk isotopic ratios are becoming more and more automated (Barrie & Prosser, 1996; Fry et al., 1992; Wong et al., 1992b), resources can easily be wasted when trying to investigate a large-scale ecosystem. Conversely, nothing is more vexing to the scientist, and to the community, as isotopic results that are incomplete in the sense that sources and sinks, and interconnections between matter pools have not been adequately characterized either chemically, biologically, or isotopically. Once hypotheses have been established, the site needs to be assessed biologically, chemically, and physically. Literature surveys are a good source of biological inventories, but the most useful information usually arises from the studies and observations of ecologists who have often worked for many years at a particular site. For example, the U. S. National Science Foundation funds a series of Long Term Ecological Research Sites (LTER) and in recent years, some of these sites have been the focus of

180


isotopic investigations (e.g. Fry, 1991). As in any new scientific endeavor, the first iteration of a large-scale isotopic project will very likely involve some unforeseen problems associated with the analytical design. Many factors need to be considered in an attempt to reduce these potential problems, thus we suggest an approach that can be tailored to the specific questions being addressed. Illustrative examples of how this general approach might be applied to investigations of aquatic (Figure 9.1) or terrestrial (Figure 9.2) ecosystems are also described in Box 9.1 and Box 9.2 respectively. In general, the first step to approaching the isotopic study of an ecosystem requires the consideration of issues that are specific to natural abundance stable isotope analysis of ecosystems and their components, and takes place before actual sampling of B O X 9.2 - C a s e s t u d i e s o f a t e r r e s t r i a l e c o s y s t e m

Schematic of the major components of a typical terrestrial ecosystem. Representative examples of different trophic levels and particular pathways of energy and matter transfer are shown. A well defined stable isotope study can aid in understanding the complex interactions between these components. F i g u r e 9.2 -

C a s e s t u d y 9.3 - The impact of atmospheric carbon dioxide on a terrestrial ecosystem A question relevant to the present day surficial environment is whether a signal of increasing CO2 concentrations in the atmosphere can be found in terrestrial plants, animals, or soils? This question provides an opportunity to study changes in CO2 concentrations over several different timescales by designing a protocol which involves the analysis of carbon from a variety of "pools". For example, the 613C value of plant tissue is related to both the isotopic composition and concentration of atmospheric CO2. This isotopic ratio of plants is set during a single growing season, whereas the 613C value of tree rings provide excellent yearly signals. At longer timescales, the isotopic composition of leaves is passed up the food chain to herbivores, such as deer. Collagen in their bones, or apatite in their teeth take several years to form and will therefore contain an integrative isotopic signal. As a result, a longer-term record of 613C values more in line with the tree ring study can be determined. Lastly, the carbon isotope composition of soil carbonates which are formed over thousands of years, provide a record of long-term changes in climate and CO2 concentration.


181

m a t e r i a l begins. It is i m p o r t a n t to c o n s i d e r the different factors t h a t affect the stable isotope c o m p o s i t i o n of material. Biological a n d b i o c h e m i c a l p r o c e s s e s fractionate isot o p e s a n d therefore h a v e a s u b s t a n t i a l effect on the isotope c o m p o s i t i o n of o r g a n i s m s . O n the other h a n d , the f r a c t i o n a t i o n associated w i t h the c o n s u m p t i o n of organic m a t ter b y h i g h e r trophic level o r g a n i s m s is small, facilitating the use of i s o t o p e s in source identification. Therefore, the isotopic c o m p o s i t i o n of a p a r t i c u l a r m a t e r i a l in an ecos y s t e m will d e p e n d on b o t h its source isotopic c o m p o s i t i o n a n d p r o c e s s - r e l a t e d frat i o n a t i o n s associated w i t h its f o r m a t i o n .

C a s e s t u d y 9.4 - Assessment of the effect of fertilization nitrogen on agricultural crops The primary question is the following: Are nutrients completely available for crop growth? There is also a secondary concern whether fertilizer pollutes ground water and affects other plant and aquatic ecosystems. For this study, characterizing the isotopic and chemical concentrations of separated nitrogen-containing nutrients (i.e. nitrate and ammonium) is critical. A sampling regime should be implemented which includes the collection of plants, crops, soil, and groundwater samples. In conjunction with concentration data, nitrogen isotopes can be used to trace the transport of nutrients and fertilizer nitrogen species into plants and groundwater. Initially, baseline values of concentration and isotopic composition of the nitrogenous nutrients in potential sources, such as groundwater, should be sampled. After application of the fertilizer, soil and groundwater nutrient concentrations should be monitored over time, preferably over the entire course of the growing season, and also subsequent seasons. If there is evidence of elevated nutrient concentrations, then more complicated analysis of the 615N of dissolved nitrogen species should be added to the study (see Box 8.3). Assessing plant uptake of nutrients involves sampling and analyzing the bulk nitrogen isotope composition of plants, both crop and indigenous species, in the agricultural fields on a monthly or weekly basis. In addition, different parts of the plants should be sampled in order to determine the degree of isotopic heterogeneity in a large plant sample. Lastly, at least 5-10 samples of the crop plants from various locations in the field should be collected in the event that the fertilizer application was not uniform. C a s e s t u d y 9.5 - Determination of dietary sources Stable isotopes are increasingly being used to investigate the feeding ecology and behavior of animals and birds. The question may be "What is the relative contribution of freshwater fish to the diet of a particular bird," or "Does the diet of an animal change over the seasons". In addressing these questions, it is imperative that the potential diet sources under investigation exhibit distinct isotopic compositions (i.e. d values of sources differ by more than 2%o). For example, marine species of fish (613C = -17%o + 4) have isotope compositions which are substantially different to freshwater species (613C = -24%o + 4). In such diet studies, potential dietary sources are collected and analyzed and compared with particular tissue samples from the consumer. Similar sample types should be measured (e.g. feathers, muscle tissue, teeth), as the isotopic composition of individual tissue types can vary significantly within a single organism (Gearing, 1991; Tieszen et al., 1983). The carbon isotope composition of an organism should reflect its diet and the nitrogen isotope composition may indicate the trophic level of the organism in the ecosystem. It is important to understand the limitations of these methods, as there can be considerable sources of isotopic variability in the diet and the particular tissue analyzed. Gearing (1991) provides a general review of the study of trophic relationships with stable isotopes and addresses many of the issues of isotopic heterogeneity in sample types. There is a wealth of studies employing stable isotopes in elucidating feeding strategies of a large variety of organisms including birds (Bearhop et al., 1999), fish (Doucett et al., 1999), whales (Hobson and Schell, 1998), crabs (Fantle et al., 1999) and primates (Schoeninger et al., 1998).

182

XQ

0


Differentiation of sources using isotopic ratios: Because stable isotopes are most often used as tracers of different sources in an ecosystem (Case study 9.1), the first consideration should be to establish whether sources truly have unique or even distinct isotopic ratios (e.g. Farquhar et al., 1989a; Quay et al., 1992). As a rule of thumb, if the 6 values of sources differ by less than 1%o, a major isotopic study designed to differentiate their movement in an ecosystem should probably be reconsidered (e.g. Benner et al., 1987). While current instrumentation and methods are capable of distinguishing much smaller differences in isotopic composition (0.2%o for 613C or 0.3%o for 615N), in complex systems involving multiple interactions, the reality of distinguishing sources of materials with such small differences is a concern. Overall, spatial and temporal heterogeneity in 6 for a single material, and isotopic variation between specific compound classes in organisms, makes distinguishing and tracing particular sources of organic matter virtually impossible at the level of less than 1%o. Therefore, it is important to study ecosystems where the ratio of signal (distinct isotope ratios) to noise (e.g. isotopic heterogeneity) is favorable. Isotopic heterogeneity in individual samples" The most important variable that can blur isotopic signals is heterogeneity within a single sample. Individual samples may be isotopically heterogeneous on spatial scales from meters to nanometers, and therefore sampling of material requires careful consideration of these factors. For example, within a single animal, individual tissues (e.g. muscle, blood, hair) have different isotopic compositions (Tieszen et al., 1983), and in a study on isotopic heterogeneity of leaf tissue, we observed that the ~13C value of plant mesophyll is consistent over a leaf, however the veins and petioles show significant deviation. In addition, individual materials are isotopically heterogeneous on a molecular scale with different biochemical compound classes (e.g. proteins, lipids and carbohydrates) having a range of isotopic compositions. This molecular heterogeneity results from different process-related fractionations that occur during the biosynthesis of specific compound classes. Since the development of continuous flow gas chromatography-combustion-isotope ratio mass spectrometry (GC-C-IRMS) for the isotopic analysis of individual compounds (Hayes et al., 1990), the literature is filled with publications defining the breadth and scope of isotopic ratios on the compound specific level. Not only are different biochemical fractions quite distinct in their isotopic compositions (range from 3-5%o), but the individual compounds that comprise the pools also have almost an order of magnitude greater variation (20-30%o) (e.g., Blair et al., 1985; Rieley et al., 1991; Abrajano et al., 1994; Fantle et al, 1999). These isotopic variations between individual compounds generally result from different isotopic fractionations during their biosynthesis. Furthermore, the heterogeneity in such biochemical isotopic fractionations can be overprinted by modification of physiological processes, such as the variations in the ~513Cvalue of phytoplankton as a function of dissolved CO2 concentrations (Bidigare et al., 1997; 1999). These physiological variations can result in shifts of C or N isotope ratios of the bulk material by at least 5%0 (e.g., Ambrose, 1991; Bird et al., 1995; Buchman et al., 1996; Cifuentes et al., 1988). Therefore, unless a very specific biochemical pool if sampled from a physiologically invariant source, then the likelihood of source vari-


183

ations is very real.

3. Spatial variation: Sample heterogeneity governed by physiological or biochemical changes, can be exploited for delineating ecosystem dynamics. For example, the stable carbon isotope compositions of vegetation and air in forest ecosystems are influenced by height. With increasing height, the canopy air becomes progressively enriched in 13C, as the influence of soil-respired CO2 and tropospheric CO2 changes. Therefore, as the carbon source for photosynthetic fixation (i.e. canopy CO2) changes, the isotope values of leaves of individual trees vary as a function of height (Buchmann et al., 1997). Similarly, the isotopic composition of leaves may change with overall elevation (Sparks & Ehleringer, 1997). 4. Temporal variation" The stable isotope composition of components of an ecosystem can vary on a multitude of different timescales from seconds to thousands of years. The temporal variations of an ecosystem and how it responds on a diel, seasonal, or longer time spans rely specifically on isotopic fluctuations (Case study 9.2). In such studies, researchers hypothesize or predict that isotopic variations will be linked to specific biological or chemical processes (Cifuentes et al., 1989; Fogel et al., 1999). Such studies have also led to discoveries of the importance of particular pools of matter or organisms that were previously considered unimportant in the functioning of an ecosystem (e.g. Stapp et al., 1999). In addition, the isotopic composition of individual materials may reflect a specific time period and this parameter is an important one to consider (e.g. Ben-David et al., 1997; Johnson et al., 1998). For example, carbon isotope ratios in leaf tissue reflect a single growing season (e.g. Brooks et al., 1997), whereas the carbonate in apatite of teeth of animal feeding on leaves will have an isotopic signal reflecting several years of an animal's life (0-5 years) (Lee-Thorp & van der Merwe, 1991). Trees with significantly sized trunks can live for several centuries; thus long-term trends are recorded in the 613C values in growth rings (Stuiver & Braziunas, 1987; Bert et al., 1997). The 613C values of soil carbonates, which form over thousands of years, reflect long-term climate change that occurs on continental scales (e.g., Cerling, 1992). If one study included all of these carbon pools, the leaves, teeth, wood, and soil carbonates, multiple ecological processes occurring over very different time signals could be discovered (Case study 9.3). Considering all the issues discussed above and incorporating the experiences of several researchers, we summarize below an approach to a general ecological investigation that involves stable isotope techniques:

Step 1" After defining the research questions or hypotheses, decide whether there is a good possibility that stable isotopes will be able to provide answers to the questions. Is there a good "isotope signal to noise ratio"? Which isotopes will be the most useful? Are there methods already established to analyze the stable isotopes in the ecological materials that are part of the study design?

184


Step 2" Design a survey of the ecosystem to measure baseline isotopic values in the pools that are of interest. Is the study site accessible year round or seasonally? Does the ecosystem vary seasonally, and most importantly, does the scientific question involve studying the ecosystem over a period of time? Is special sampling equipment required? Are permits needed for sampling or importation(e.g. CITES) across international borders? Step 3" Begin the planning stage for collecting, processing, and analyzing samples. Assemble the equipment needed for the collection keeping in mind the storage needed until samples are analyzed (Section II). A list of all reagents, bottles, and chemicals that are needed in the field can be extremely helpful. Ancillary analyses, for example nutrient concentrations, are important considerations that are often overlooked. Step 4" Preliminary sampling trip and isotopic analyses: Take the first set of samples, if the opportunity exists to travel easily to the ecosystem to be studied. Collection of 2-3 times more samples than you intend to analyze will allow you to choose from this array of samples at the time of isotope analysis. If the sampling site is remote, check the availability of equipment at the site, and confirm that all necessary items of equipment and chemicals are on hand. Attempt the first set of isotopic analyses. The first set of analyses should determine whether the analytical procedures are robust, if sources have distinct isotopic compositions and whether the stable isotope techniques used will address the research objectives. Step 5: Rethinking the experiment and the sampling protocols" Often, the first field collection is an adventure that is problematic and incomplete. Analyze the initial data and reformulate the questions and experimental design to reflect the reality of the ecosystem. Consider adding additional parameters, sample types, or time points, at this stage, that now are more important, and eliminate those that will not contribute to the final results. In many instances, the material that is collected consists of a complex mixture (e.g. soil, sediment, filtered material) and therefore a more refined approach to isotope analysis may be required. In such cases, measurement of the isotopic composition of individual organic compounds can provide unique information for studying complex interactions where physical separation and even biological separation are difficult (for example the transfer of dissolved organic carbon from phytoplankton to bacterioplankton). The methods and techniques for such analyses are discussed in section 9.3.3. 9.3 Section II: Methods of sample collection and storage Different types of samples require different methods of collection and storage, and the choice of methods will be influenced by the location of the ecosystem, the availability of cold storage, and the kinds of analyses to be performed. Whatever methods are chosen, it is imperative that they do not affect the isotopic integrity of the material being sampled. Isotopic ratios of organic material are susceptible to alteration if degradative processes take place during storage. Samples that are poorly preserved are at


185

risk of possible bacterial decay, which may result in mineralization or solubilization of organic carbon. Alternatively, contamination can alter isotopic ratios, especially of samples destined for compound specific isotope analysis. Contamination while sampling is minimized through the use of gloves and minimal handling of materials. Geochemical materials with ancient organic matter are sampled and stored at room temperature (e.g. light hydrocarbons in petroleum samples), whereas other materials require more refined techniques and, in some cases, specialized collection equipment is required.

9.3.1 Biological materials Biological sampling is based on capturing and preserving the isotopic and biochemical composition of an organism at a particular point in time. In order to accomplish this, living processes must be effectively inhibited, because at death, organisms release enzymes (e.g. proteases, nucleases, and lipases) that specifically target all of the biochemical classes of compounds. Cold storage is the best method for slowing down living reactions, while also inhibiting degradation pathways. If liquid nitrogen is available on the site, quick-freezing and storage in an enclosed vessel will preserve all of the high molecular weight biochemical information, as well as any isotopic information. Dry ice or freezing in a conventional-20~ freezer are also acceptable methods of cold storage, and are often available on oceanographic research vessels or nearby laboratories. Prior to freezing biological material it is important to rinse any adhering sediment or salt off of the organism with distilled or deionized water. The use of fixatives (e.g formalin or ethanol) to preserve samples is not recommended, as we have found that tissue preserved by the addition of these reagents has an altered carbon isotope ratio, most likely because of binding between the solvent and the tissue. In remote areas, there are several ways to overcome the inconvenient access to cold storage. First, long-life liquid nitrogen dewars can be purchased that maintain liquid nitrogen temperatures of-196~ for at least 30 days. The dewars are initially expensive, but are reusable and should provide several years worth of service. For this method to be useful, there must be a source of liquid nitrogen available. Most universities have reasonably priced liquid nitrogen reservoirs, and dewars such as these require filling only one time. Transporting liquid nitrogen to the field requires careful planning and prior to using such methods the reader is encouraged to understand the potential hazards associated with working with cryogenic liquids. Safety information is available from gas supply companies and also the National Research Council (1981). An alternative to cold storage is drying in the field. Plants can be effectively dried in plant presses on herbarium paper. We have tested this method with samples taken in the Australian desert, which is characterized by low humidity and high temperatures. Drying occurred in 2-3 days, with no change in plant appearance. Care should be taken to use clean blotter paper rather than newspaper, as the ink from the newsprint could contaminate the plant sample. Last, some plant presses are connected to a source of heat to promote drying; a simple light bulb should be chosen over heated charcoal.

186


An isotopic comparison of samples preserved by drying, with those preserved using liquid nitrogen revealed no difference in 615N o r ~13C. Keep in mind that one aspect of drying that may cause complications is the use of heat to accelerate drying, which could act to increase the activity of degradative enzymes. Another complication often encountered in tropical regions is high humidity, which will increase drying times. A third and simple technique to dry and preserve samples in the field is to use desiccants (e.g. silica gel). The sample can be wrapped in foil and placed in a sealed jar or sample bag that contains a desiccant. With frequent replenishment of the desiccant, this method can remove a significant amount of water. Storage of animal tissue is a separate problem, because drying tissue can attract insects that might consume it. As a result ample consideration should be given to the type of tissue to be studied. For example, collagenous or keratinous material from bone, horn, nails, or feathers can be easily rinsed with distilled water and dried without any chemical or isotopic alteration. Teeth are also very resistant tissues that are especially desired because they contain well preserved organic and inorganic isotopic reservoirs. However, if tissue samples are needed, they should be frozen immediately, preferably in liquid nitrogen or put on ice as soon as possible.

9.3.2 Aquatic particulate material One of the most frequently used methods for collecting aquatic material for isotopic analysis is filtration. There are, however, many things that should be considered prior to venturing into the field. First, how much material needs to be collected for the desired analysis? Second, will the filter interfere with the analysis, or contaminate it, in any way? Third, what is the particle size of interest? For on-line isotopic methods for determining bulk C, N, H, and O isotope ratios of organic material, less than a milligram of material is needed, even including replicate analyses. In contrast, for compound specific work, especially if a target molecule (e.g. cholesterol) is sought, 5-10 mg of organic matter is a reasonable amount to collect to allow for replicate analyses. For example, if the environment that is to be studied is rich in biological material, typically a liter of water is filtered for bulk analysis. For sampling larger amounts of material or in regions where there is a very low amount of particulate biological material, large filters are required and often each scientist constructs filtration devices that fit individual needs. Large volume filtration apparatus have been designed to operate in situ, to be connected with complex pumping system, or to be contained in pressurized vessels (Wakeham & Volkman, 1991). The size of the filter can be determined by the user, and for most materials sheets of glass fiber or polycarbonate can be cut to fit the design of the filter holder. At this stage the type of filter material is an important consideration. For a filter to be compatible with an analytical technique it should not contaminate the sample. Most researchers use glass fiber filters, which can be cleaned by heating (450~ for 4 hours), and conventional filter holders (47 mm in diameter), which can be readily obtained from scientific supply companies. Filters containing particles (e.g. GF/F filters) are wrapped in clean aluminum foil, stored frozen, until they are dried or pro-


187

cessed in the laboratory prior to isotopic analysis. If the method requires a specific amount of weighed material, the filter should be chosen that will allow the sample to be recovered quantitatively and with ease (e.g., a Nuclepore filter is the perfect choice, because material can be readily rinsed off the surface of the filter). The particle size is the last consideration for filtration. By the simplicity of their design, filters are only crude ways of separating material by size. The most common filter is the GF/F filter (glass fiber) with a pore size of 1 mm, however, as filtration proceeds, material is trapped in the fibers of glass, so that the effective or nominal pore size is reduced to 0.7 mm. A GF/F filter can effectively trap most phytoplankton, many bacteria, and all zooplankton. Typically, a much larger size filter or mesh is used to separate zooplankton from phytoplankton and bacteria. Mesh sizes > 100 mm will separate the majority of zooplankton, and filter sizes > 20 mm will remove a majority of the microzooplankton. The sample trapped between I mm and 20 mm is the material typically captured on a GF/F filter. It will usually include phytoplankton of all types, detrital material, some attached bacteria, and microheterotrophs including flagellates and rotifers. Unfortunately, many studies in aquatic ecology are aimed at defining trophic relationships among these different groups of organisms, and thus simple filtration fails to provide adequate separation. In such cases, a compound specific approach may be useful, especially if the organisms under investigation produce source-specific biological marker compounds, as outlined in section 9.4.3. However, if the concentration of these source-specific compounds is low, then substantially larger volumes of material would be required, collection of which may not be logistically feasible in the field.

9.3.3 Collection of water samples Water samples collected for the isotopic analysis of ammonium or nitrate should be frozen immediately. Typically one liter of water is collected, and if the sampling location is very turbid or biologically-active (e.g. a river or lake), it is best to filter the sample through a pre-combusted GF/F filter. Seawater samples from the open ocean that contain very little particulate matter can be collected and frozen directly. In addition, a subsample (approximately 25 mls) should be frozen or analyzed for nutrient concentrations immediately (Solarzano, 1969; Strickland & Parsons, 1972). For isotopic analysis, the frozen sample will be returned to a laboratory for subsequent analysis, however long-term storage should be avoided, because concentrations of ammonium, for example, are often lower after storage. If a freezer is not available in the field, water samples should be kept cold on ice. If no method of refrigeration is available, adding I ml of 2 N HC1 to I liter of water will generally preserve the sample by halting bacterial growth and ensuring that the pH of the solution favors the ionic state of ammonium.

9.3.4 Storage of samples

In the laboratory, samples are typically dried and stored in a freezer prior to isotopic analyses. The preferred method of drying is freeze-drying, but alternatives such as oven drying may be employed. Freeze-drying employs a vacuum to remove water from tissues that are kept frozen throughout the procedure, thereby minimizing the

188


possible degradative processes caused by heat drying. However, freeze-dryers typically are not resistant to acid, so that samples can not be pre-treated with acid (e.g. samples containing carbonate carbon) and subsequently dried. For such acid-treated samples, a similar type of drying, which was developed for biological tissues and is termed centrifugal vacuum drying, can be employed. In these instruments, samples are placed into a centrifuge, and as vacuum drying proceeds, the sample is concentrated in the bottoms of the centrifuge tubes. These drying units are generally built with chemically resistant housing and traps so that samples treated with acid or solvents can be dried readily. The most common alternative to freeze-drying is oven drying. Most ecological laboratories house large drying ovens maintained at 80~ with or without flowing dry gas. Drying biological tissue, that is clearly dead and no longer metabolically active, can be accomplished at temperatures between 50 ~ and 80 ~ At lower temperatures, flowing an inert gas such as nitrogen, argon, or helium over the sample both minimizes oxidation, and also decreases drying time, as water is swept efficiently from the oven. Vacuum ovens offer another alternative means of drying, and often desiccants (e.g. P205 or Silica Gel) are added to increase the efficiency of water removal.

9.3.5 Sample preparation for isotope analysis

The initial preparation of samples for isotope analysis is dependent on the type of analysis to be performed. If the isotope composition of the whole sample is sought, then after drying the material, little further preparation is required. Samples that may contain carbonate carbon, however, should be treated with dilute HC1 (0.1N) and thoroughly rinsed with distilled water prior to drying. Carbonate carbon is typically more enriched in 13C than organic carbon and may confuse analyses of organic material if not removed prior to isotope analysis. The carbonate present in dissolved materials or on filters can be removed by incubating the sample for 24 hours in a glass dessicator over fuming HC1. Carbonate can be removed from sediments, which can contain up to 90 wt% CaCO3, by acidifying samples with either concentrated HC1 in silver sample cups (Nieuwenhuize et al., 1994) or with a modified HC1 vapor phase method (Yamamuro & Kayanne, 1995). Complex sample matrices (e.g. soils, leaf litter, stomach contents) should be ground to a fine powder using a mortar and pestle, as the necessity for sample homogeneity cannot be overemphasized. Material frozen with liquid nitrogen in a mortar, or crucible, can be easily ground to a fine powder using a glass pestle. Alternatively, commercially available grinders can be used which pound samples to fine powders under cryogenic conditions (e.g., Spex Mill, Wig-L-Bug). Such equipment is very useful when working with materials such as wood, bone, soils and sediments. If a compound specific approach is to be undertaken then substantial preparation, including specific chemical and biochemical techniques may be performed prior to isotope analysis. These techniques are discussed with relation to specific analyses later in this chapter.


189

9.4 SECTION III: Stable isotope analysis of ecological samples The present chapter concentrates on techniques to measure the isotopic composition of important nutrients, bulk organic matter and the three major classes of organic compounds (lipids, amino acids and carbohydrates). The reader is also directed to Chapters 8 and 23 for additional methods on such subjects. In addition, it may be important to measure the isotope composition of inorganic material in the ecological setting under investigation. The specific techniques required to measure these parameters are discussed in other chapters on isotopes in hydrology (Chapter 1), dissolved inorganic carbon (Chapter 10) and atmospheric gases (Chapter 14).

9.4.1 Isotopic analysis of ammonium Measurement of the two major N-containing nutrients, ammonium and nitrate, is important for nitrogen isotope studies. The analysis of dissolved nitrate is described in a Chapter 15 (applications in Liu et al., 1996; Brandes & Devol, 1997; Chang et al., 1999). There are several methods for ammonium analysis, and in all of them, NH4 + is separated from solutions or samples by either gas phase transfer (i.e., distillation or diffusion) or adsorption and separation by a substrate (e.g., molecular sieve Zeolite BOX 9.3 - Nitrogen isotope analysis of ammonia

Two distinct methods of measuring 15N-NH4 + are available. One method is a modification of the ammonia diffusion method and is described in detail in Holmes et al. (1998). The second method utilizes steam distillation and zeolite adsorption to measure the nitrogen isotopic composition of ammonia (Velinsky et al., 1989). The steam distillation unit requires a hardy refrigeration unit, a source of ammonia-free water for steam generation, a condenser, and a receiving flask. All of the glassware used in this method should be cleaned and rinsed with 1% HC1 and then rinsed thoroughly with distilled water. After the glassware for distillation and receiving is assembled, and prior to preparing the first sample, 200 mls of pure ethanol is distilled into the receiving vessel. This rinse is followed by a subsequent wash with 200 mls of distilled water. Both solutions are discarded after distillation. This washing and rinsing procedure is repeated between each sample. The recovery flask is then prepared as follows: a pre-combusted Pyrex Pasteur pipette is attached to the output tube from the condenser unit. The flask contains 25 mls of dilute HC1 (0.00I N), and care is made to have the tip of the pipette underneath the surface of the acid solution. A filtered water sample (250 mls) is added to the distillation flask with 1 ml of 10N NaOH. The amount of NaOH added will depend on the salinity and the buffering capacity of the environmental water. Once the NaOH is added to the solution, the NH4 + is converted to NH3, so distillation should begin immediately. It is essential that >99% of the ammonia is distilled, and during all phases of this procedure, individual steps should be checked for yields by conventional ammonium analysis (Strickland and Parsons, 1972). It is important to ensure that the pH of the acid trap remains below 6. After distillation is completed, the receiving flask is removed from the unit and the ammonium is adsorbed onto a zeolite molecular sieve. A bed of zeolite is made on a GF/F filter (cleaned by preheated to 550~ for 2 hours) and the ammonium solution passed through the filter. Gravity filtration takes up to 60 minutes but results in approximately 100% recovery. The filtered zeolite is placed in a drying oven at 50~ for no less than 48 hours. If a dry gas can be passed through the oven, water vapor is more efficiently removed. After drying, the zeolite can be scraped off the filters and analyzed by combustion in either a sealed tube or by EA-IRMS. The precision and accuracy of this method is about +0.5%o, with the error probably originating from slight isotopic fractionation of the NH4 + onto the zeolite. For EA analysis, the zeolite should be dried in the presence of either He or Ar, because N2 from air can be trapped within the pores of the zeolite, which interferes with the analysis.

190


W-85) (Box 9.3). All of the procedures that have been developed are relatively simple, however, attention must be paid to almost all of the details, because large isotopic fractionations occur in the transfer of ammonium [NH4 +] or ammonia [NH3] between different solutions. For example, there is a 30%o isotope effect between dissolved ammonium and gaseous ammonia. Therefore, to minimize this potential fractionation during distillation or diffusion it is essential that all (>99%) of the a m m o n i u m / a m m o nia in solution is transferred to the receiving solution. Smaller isotope effects are associated with adsorption. In many of the employed methods distillation and adsorption are often coupled together (e.g. Velinsky et al., 1989), and sloppy techniques with either process can result in decreased analytical precision and accuracy. Moreover, with adsorption methods substantial losses in recovery are possible. Usually, the amount of the sample that is necessary for isotopic analysis requires the methods to be as efficient as possible, with 20 mg of N being the lower limit for routine analysis. Surface water NH4 + concentrations in most aquatic ecosystems are below 5-10 mM, such that 200 to 350 mls of water need to be processed for acceptable precision. Therefore, it is essential to prevent contamination of NH4 +, which can arise from many sources, as ammonia is a gas. In particular, investigators should be keenly aware of any enriched tracer work that is being done in the same laboratory or ship, and if so, separate areas and glassware are warranted. The other sources of possible contamination are from cleaning solutions, smoke, and distilled water. The isotope composition of low levels of ammonium in marine waters can be measured using an adaptation of the ammonia diffusion method. Large volumes of water (up to 4L) are used and during the process, ammonium is converted to ammonia, which is trapped on glass fiber filters (Holmes et al., 1998). To determine the appropriate volume of water required for analysis, it is necessary to know how much ammonium is in the sample (Solarzano, 1969; Strickland & Parsons, 1972; Holmes et al., 1999). This method may be initiated in the field or on-board ship, although diffusions generally last two weeks. Two distinct advantages of this method are that multiple samples can be run concurrently and that the labor involved per sample is minimal.

9.4.2 Isotope analysis of bulk organic matter The isotope composition of bulk tissue samples can be measured using one of two methods. In the first, the organic material is placed in a quartz tube with reduced copper and copper oxide (CuO), sealed under vacuum and combusted in a muffle furnace at 900~ The resulting gases are purified on a vacuum line and are then analyzed by conventional IRMS (for a thorough review see Boutton, 1991). The second more common approach uses a continuous flow system, which comprises an elemental analyzer (EA) directly attached to an IRMS (Fry et al., 1992). Samples are weighed into tin or aluminum boats and loaded into a multiple sample carousel on the EA. The whole sample including the metal boat is combusted and the resulting gases (CO2, N2 and H20) are separated on a gas chromatography column. These gases are then introduced into the IRMS, and both 613C and 615N values can be

Preparation of Ecologicaland BiochemicalSamplesfor IsotopeAnalysis

191

obtained from the same sample. Typical sample sizes range from 100 - 200 mg for animal tissue to 1- 2 mg for sediments. It is good laboratory practice to analyze a standard after every 5 samples to check precision, and ideally the standard should be similar in elemental composition to the samples of interest. Examples of standards include mesquite leaves for the analysis of plant material and collagen for muscle tissue analysis. The EA-IRMS system has many advantages over the traditional sealed tube method, including a rapid throughput of samples (up to 80 samples per day) and sample sizes of less than 100 mg (e.g. organic rich material such as muscle tissue). It should be noted that results from EA-IRMS analysis differ from those of off-line techniques, as a result of an isotopic fractionation which occurs during the introduction of the gases into the mass spectrometer. Therefore, a correction should be applied to the EA-IRMS results which should be calculated on a daily basis through the analysis of known, preferably NBS, standards. In some instances, such as the analysis of muscle, the isotope composition of fatfree tissue may be required to simplify trophic web analysis (e.g. Hesslein et al., 1993). Removal of fats or lipids from whole tissue will affect the 613C values as lipids are typically depleted in 13C relative to bulk material. The techniques employed to remove lipids range from vigorous Soxhlet extraction (e.g. Focken & Becker, 1998) to mild solvent extraction (Hesslein et al., 1993). If the goal of the study is to measure fatfree tissue, then the extraction method employed should aim to remove fats and little else. Generally, this can be easily achieved by a mild extraction technique, such as immersion of the sample in hexane, with gentle agitation and decanting of the solvent. More vigorous techniques may remove other organic compounds, such as the more polar compounds (e.g. cholesterol), which will in turn affect 613C values of the "fat-free" tissue. In addition, vigorous extraction of blood samples may affect 615N values, as compounds such as uric acid and urea may be preferentially extracted (Bearhop et al., 2000). Stable isotope analysis of bulk organic matter has been widely used in ecological research (reviews include Griffiths, 1998; Lajtha & Michener, 1994; Ehleringer et al., 1993; Rundel et al., 1989; Peterson & Fry, 1987). For example, the isotopic composition of plant organic matter can be related to changes in physiology, associated with differences in water use efficiency (Bert et al., 1997; Johnson & Tieszen 1994; Read et al., 1992; Farquhar et al., 1989a). Environmental parameters such as light, water, salinity and air pollution (Buchmann et al., 1996; D u e t al., 1998; Farquhar et al., 1989a) mitigate ~)13C values as well. Stable isotopes of animal tissue are also used to investigate trophic structure, migration, and metabolism (Hobson & Clark, 1992; Koch et al., 1995; Ben-David et al., 1997; Anderson & Polis, 1998; Witt et al., 1998; Schmutz & Hobson, 1998; Wainwright et al., 1998; Jones et al., 1998; Schoeninger et al., 1998; Alisauskas et al., 1998; Marra et al., 1998; Hansson et al., 1997; Nelson et al., 1998; Ostrom et al., 1997).

9.4.3 Compound specific isotope analysis of major biochemicals The isotopic composition of individual organic compounds in ecological and biological samples can be extremely useful in deconvoluting processes and interactions

192


between organisms in complex systems. In studies where bulk isotope analysis may provide an overview of the whole ecosystem, targeting specific molecules will provide information on the role of individual organisms. This approach relies on different organisms synthesizing different molecules. Through numerous chemical and biochemical procedures, these molecules can be separated and their isotopic compositions analyzed individually. The approach is particularly powerful in studies where physical separation of material is difficult such as in the determination of sources to water column particulates in lakes. Filtering particulate matter from the water column provides an overview of the multiple sources to the lake and the isotopic composition of the complete filter reflects an average signal of input. The isotope composition of specific molecules will provide more detailed information pertaining to the importance of different sources. The 613C of individual n-alkanes indicate different types of higher plant sources (Spooner et al., 1994; Rieley et al., 1991), whereas the isotopic composition of a particular suite of sterols may provide information on phytoplankton inputs (Canuel et al., 1997). A more recent application of compound specific isotope analysis is aimed at elucidating particular metabolic pathways in organisms. The isotope composition of individual compounds is dependent on the enzymatic reactions which occur during their biosynthesis, and therefore retains a signal of their pathway of synthesis. Several studies have exploited these techniques to elucidate the pathways of lipid synthesis in organisms such as bacteria (Teece et al., 1999; Summons et al., 1998; van der Meer et al., 1998) and pigs (Stott et al., 1997a). The compound specific approach requires a significantly greater effort to prepare a sample for analysis, and in some cases this additional work may not be worthwhile. For example, the preparation of a single sample for the analysis of the 613C values of individual amino acids by gas chromatography combustion isotope ratio mass spectrometry (GC-C-IRMS) may take up to three days in the laboratory. Typically, the isotope composition of samples are measured in triplicate with each analytical run being up to 60 minutes in length. As a result of these protracted periods of preparation and analysis, it is difficult to acquire large data sets. Secondly, GC-C-IRMS systems are not abundant and are not simple off-line attachments which are free of problems. As with any analytical instrument, these systems require continual monitoring and maintenance, however the results obtained from such analyses are often worth the hard work involved. Therefore, in determining whether such a compound specific isotope approach is warranted, several questions should be addressed: 9Can bulk isotope analysis provide the answer? ~ Is the determination of sources or biochemical pathways important? 9 Are there multiple sources (pathways) that cannot be resolved by bulk techniques? 9Do the source organisms (pathways) synthesize different molecules? 9Are these compounds specific to the source organism (pathway)?

Preparation of Ecological and BiochemicalSamplesfor Isotope Analysis

193

Once the specific compounds of interest have been determined, the question of analysis techniques need to be addressed. The primary technique to analyze the isotope composition of individual compounds is gas chromatography combustion isotope ratio mass spectrometry (GC-C-IRMS). The sample is injected into a gas chromatograph and the compounds of interest are separated on a well-chosen chromatography column. The column eluent passes through a micro-volume furnace where the material is combusted and the resulting CO2 (or N2) is introduced into the mass spectrometer (Hayes et al., 1990; Eakin et al., 1992). The limiting factor in compound specific analysis using GC-C-IRMS is that only volatile compounds can be analyzed as the inlet system is a gas chromatograph. There has been limited success in designing systems where the inlet system is a liquid chromatograph, though commercial experimental systems have been developed and marketed. In many cases, chemical derivatization of a compound is necessary to synthesize a component that is amenable to gas chromatographic analysis, and therefore GC-C-IRMS analysis. Volatile compounds, such as certain classes of lipids, can be directly analyzed, while amino acids and carbohydrates require extensive chemical derivatization. The analytical procedures to separate particular classes of compounds are numerous, accordingly we will present an overview of suggested approaches. It should be noted that in order to obtain reproducible isotope values, individual compounds must be sufficiently separated on the gas chromatography (GC) column prior to mass spectrometric analysis. Isotopic precision depends on adequate, preferably baseline, separation of components. Therefore, it is important to verify both the methods of GC separation as well as IRMS techniques for the isotopic analysis of specific compounds. It is essential to check the analytical procedures used by the repeated analysis of compounds of identical chemical composition. Checking the methodology in this manner should confirm that the specific compounds of interest are quantitatively extracted from the sample, and also indicate whether any isotope fractionation occurs during analysis. Furthermore, it is good practice to add one or more "internal" standard compounds of known chemical and isotopic composition to the sample prior to extraction/analysis. These internal standards should have similar chemical properties to the compounds of interest, however should not be present in the natural sample or coelute with other compounds. The isotopic composition of these compounds can be measured in the final analyses as a check to indicate the reproducibility of the methods used. For example, nC36 alkane is typically employed as an internal standard in hydrocarbon analyses, and nC19:0 fatty acid used in fatty acid analyses. The approach to measuring the isotope composition of individual compounds requires numerous steps and these should be addressed prior to sample collection (Figure 9.3).

Step 1" Collection and storage of samples for analysis should minimize contamination. GC-C-IRMS is a very sensitive technique, with nanogram quantities of material detected (cf. milligram amounts for bulk isotope analysis). Human hands


194

Step I

Sample Collection Freeze-dry I

Dry Sample

]

Bulk Isotopic Analysis

Step 2

Extraction

Complex Mixture of Compounds

Step 3

Chemical Separation

I I

I

CompoundCroup

I I

I

CompoundCroup

Derivatization Scheme

Derivatization Scheme

I Step 4

GC-amenable Compounds [ [ GC-amenable Compounds [ Step 5

Identification

Identification Step 6

I Analysis by GC-C-IRMS

I Analysis by GC-C-IRMS

Step 7

9 . 3 - A generalized analytical scheme for the stable isotope analysis of individual organic compounds in biological samples. The numbered steps refer to analytical procedures that are specific to the compounds of interest, and are discussed in more detail in the text. Figure

Preparationof Ecologicaland BiochemicalSamplesfor IsotopeAnalysis

195

contain many of the lipids and proteins that will be subsequently analyzed, so handling should be minimized and gloves used. The use of chemical preservatives, such as ethanol or formaldehyde, is discouraged and samples should ideally be stored in clean glassware or sterile plasticware. For most analyses, material should be freeze-dried as discussed above.

Step 2" Extraction of the compounds of interest requires the correct extraction method. Hydrophobic lipids and fats are extracted using a series of organic solvents, whereas carbohydrates and amino acids are hydrophilic and sulfuric and hydrochloric acids are employed in their extraction. The chemicals, materials and conditions differ for all compounds and selecting the correct method is paramount in the successful extraction of material. Prior to extraction, (internal) standard compounds should be added to check the methodology. Step 3" The mixture of extracted compounds may be complex and require additional separation techniques to isolate the compounds of interest in greater purity. Several methods of separation can be used including column chromatography, thin layer chromatography, ion exchange chromatography, and electrophoresis. The particular techniques required differ with each compound class and extensive literature can be found relating to these analytical challenges. Step 4: The mixture of compounds may be chemically derivatized to produce compounds, which can be analyzed on a gas chromatograph. Volatile lipid compounds (e.g. hydrocarbons) can be analyzed directly, such as those present in epicuticular waxes of higher plants (Rieley et al. 1991; Collister et al. 1994; Lockheart et al. 1997). However, fatty acids and alcohols require derivatization prior to analysis, in order to attain sufficient chromatographic separation required for reproducible isotopic analysis. Numerous techniques for derivatization are published and depend on the chemical functionality of the particular compounds: Fatty acids are methylated whereas alcohols are typically analyzed as the trimethylsilyl ethers. Step 5: To acquire reproducible isotope numbers, the components of complex mixtures must be adequately separated on the gas chromatography (GC) column to reduce the influence of contaminating or co-eluting peaks during analysis. Separation of compounds using GC requires the correct choice of chromatographic column. There are a multitude of chromatographic columns available from several vendors, varying in polarity and size, and the choice of column is dependent on the chemical properties of the components being analyzed. Nonpolar columns, such as DB-1, HP-1, Ultra-l, are used for hydrocarbon analyses and a column containing a 5% phenyl substituted methylpolysiloxane phase (HP-5, CP-Sil 8C) is useful for analysis of derivatized alcohols and fatty acids. Step 6" The identity of the specific compounds isolated should be confirmed by GCMS prior to isotope analysis. This step will ascertain whether the separation on the chosen GC column is adequate and will also indicate whether any contam-

196


ination was introduced during the analytical scheme.

Step 7: Stable isotope analysis of the individual compounds by GC-C-IRMS. The analyses should be performed in triplicate to determine variations on instrumental analysis. Analysis of standards of similar chemical structure will aid in determining the reproducibility of results. Step 8: If a compound was derivatized prior to analysis, then the isotopic composition of the original compound must be calculated. Derivatization usually involves the addition of a carbon-containing component in order to produce the volatile compound. Calculation of the isotopic composition of the original compound requires consideration of the contribution of the added carbon to be assessed if measuring ~)13C values. Therefore, it is essential to analyze compounds of known chemical and isotopic composition to calculate the effect of the derivatizing agent. For example, fatty acids are methylated prior to analysis and the isotopic contribution of the added methyl carbon is assessed by isotopic analysis of underivatized and subsequently derivatized fatty acid standards (see Abrajano et al., 1994:). Simple mass balance calculations allow the contribution of the 613C-derivatizing agent carbon to be assessed, which can then be applied to the unknown compounds to calculate the original isotope compositions. These steps are general guidelines and what follows are more specific examples of approaches to analyzing specific molecules. The approaches below are suggestions and each method can be altered or customized to the particular question being addressed.

9.4.3.1. Isotope composition of individual lipids

Lipids are a class of hydrophobic compounds that include fatty acids, phospholipids, triacylglycerides, alcohols, hydrocarbons and sterols (e.g. cholesterol). Lipids perform numerous functions including acting as a source of energy (e.g. triacylglycerides) to a protection mechanism against predators (e.g. plant wax esters). The carbon isotopic composition of individual lipids reflects both the isotopic composition of the carbon source utilized by the organism and isotopic fractionations accompanying biosynthesis. Isotopic fractionations which result from the enzyme-mediated reactions during biosynthesis of lipids are ultimately dependent on environmental conditions. Therefore, the isotope composition of individual lipids can provide valuable information on both the carbon source and potential changes in the organisms environment. Such an approach has been used in several studies to determine changes in past vegetation patterns by isotopic analysis of higher plant-derived hydrocarbons (France-Lanord & Derry, 1994; Bird et al., 1995; Lichtfouse et al., 1995; 1997) and lignin-phenols (Goni & Eglinton, 1996) in lake and marine sediments. Extraction of lipid compounds is a relatively simple procedure, however some caution should be applied to the methods employed. Ideally, the samples should be dry, having been freeze-dried as described above. There are two main methods of lipid extraction, both of which use similar organic solvents. The choice of organic sol-


197

vents in the extraction of lipids is important and for most procedures mixtures of dichloromethane and methanol are applicable [dichloromethane should be used rather than chloroform - dichloromethane has similar chemical properties to chloroform but is substantially less harmful]. Glass distilled or Optima grade solvents should be used whenever possible. All glassware should be thoroughly cleaned and rinsed with solvents prior to use to minimize contamination. One method for lipid extraction uses a Soxhlet system, in which clean boiling solvents (mixtures of dichloromethane and methanol) are refluxed through the sample for periods ranging from 30 minutes to 24 hours. This method is best employed when extracting lipids from a complex matrix such as sediments, soils and some woods, where the compounds of interest may be physically trapped or weakly bound in the matrix. For biological material, particularly living tissue (e.g. plants, muscle, and blood), a simpler and less harsh extraction scheme should be used. This second extraction method (Box 9.4) uses a mixture of dichloromethane and methanol at room temperature, with repeated mixing using an ultrasonic tank or vortex mixer (modified from Bligh & Dyer, 1959). Once the lipids have been extracted from the sample, the task of separation and derivatization of the particular compounds of interest needs to be addressed. The total organic extract of a sediment may contain more than 80 different compounds, so that one or more clean-up steps may be required before analysis by GC-C-IRMS. The goal of these procedures is to obtain a sample fraction, which contains the molecules of interest in a mixture that can be adequately separated on the gas chromatograph. In the case of hydrocarbon analysis, a single purification step allows adequate separation of more than 20 compounds (e.g. Freeman et al., 1990; Bjoroy et al., 1991) whereas low concentrations of cholesterol may require multiple clean up steps to isolate the compound in sufficient purity (e.g. Canuel et al., 1997). Similarly, analysis of ligninphenols, a class of compounds that are uniquely synthesized by vascular plants, requires a complex procedure of chemical oxidation and derivatization procedures prior to analysis by GC-C-IRMS (Goni & Eglinton, 1996).

BOX 9.4 - Lipid extraction of plant tissue

Freeze-dried leaves are ground to a fine powder using a pestle and mortar. A mixture of dichloromethane:methanol (1:1) (Optima grade) is added to the plant samples (15-20mg) in a precombusted glass test tube fitted with a Teflon-lined cap. The mixture is thoroughly mixed for up to 10 minutes, using a combination of ultra-sonication and vortex mixing. If required, the sample can be centrifuged (1,000g; 5 mins) to pellet the plant matter, and the solvent removed with a clean glass pipette. This procedure should be repeated 5 times to ensure complete extraction with 5ml of clean solvent mixture being used for each cycle. The extract is transferred to a round-bottomed flask and the solvent removed using a rotary evaporator. The total organic extract (TOE) is redissolved in the minimum dichloromethane required to quantitatively transfer it to a clean precombusted vial fitted with a Teflon-lined cap. The solvent is removed under a stream of nitrogen at room temperature to produce the dried TOE, which is stored in a freezer prior to further treatment.

198


BOX 9.5 - Isotope composition of individual fatty acids from muscle tissue

After extraction of muscle tissue using dichloromethane and methanol, the total organic extract (TOE) is subjected to mild alkaline hydrolysis to release the esterified fatty acids. A 6% KOH solution in methanol (5 ml) is added to the TOE, the tube tightly sealed and allowed to react overnight at 50~ After cooling, 3 ml of double-distilled deionized water is added and the neutral lipids extracted three times (each 5 ml) into a mixture of hexane:diethyl ether (9:1). Neutral lipids are transferred, and stored dry, in glass vials with Teflon-lined caps. The remaining aqueous methanol phase is acidified to pH 2 with 6N HC1 (Pierce Chemicals, constant boiling HC1). The fatty acids present in this fraction are extracted three times (each 5 ml) into a mixture of hexane:diethyl ether (9:1) and transferred to a clean precombusted glass test tube. Solvent is then removed under a stream of nitrogen at room temperature. Fatty acids are subsequently converted to their corresponding fatty acid methyl esters (FAMES). 200ml of a solution of 14% BF3 in methanol (Pierce Chemicals) is added to the fatty acid fraction, the tube tightly sealed and allowed to react for 15 minutes at 60~ After cooling, 3ml of doubledistilled deionized water and 3ml of hexane is added to the mixture. FAMES are extracted into the hexane fraction (3 x 3 ml), transferred to a glass vial and the solvent removed under a stream of nitrogen. Samples should typically be analyzed within two days of preparation using a nonpolar GC column (Ultra-i). Measurement of the ~)13C values of individual fatty acids (see Box 9.5) requires a single clean-up step and a subsequent derivatization step. After organic extraction of the sample, the fatty acid fraction is isolated using ion exchange chromatography and converted to the corresponding fatty acid methyl esters (FAMES). The 613C values of fatty acids have been reported in studies of plants (Collister et al., 1994; Rieley et al., 1991; Vogler & Hayes, 1980), changes in diets (Rhee et al. 1997; Trust-Hammer et al., 1998; Gilmour et al., 1995a), and in ecological studies of mussels and shrimp in hydrothermal vents (Pond et al. 1998; Rieley et al., 1999). The techniques to analyze particular lipid c o m p o u n d classes are reported in the literature (e.g. Blau & Halket, 1993) and require various chemical separation and derivatization techniques. Isotopic analysis of hydrocarbons has been used in studies ranging from palaeoenvironmental reconstruction (Rieley et al., 1991; Freeman et al., 1994; Bird et al., 1995) and elucidation of input sources of vascular plants to coastal sediments (Canuel et al., 1997), to the feeding habits of insectivorous bats (Des Marais et al., 1980). The isotopic analysis of PCBs (polychlorinated biphenyls) and PAHs (polyaromatic hyrdrocarbons) may also be employed to determine sources of toxic chemicals in the environment (O'Malley et al., 1994; Ballentine et al., 1996; Jarman et al., 1998). Alcohols can be analyzed as trimethylsilyl (TMS) derivatives (Jones et al. 1991) and such techniques have been applied to elucidation of dietary sources through the m e a s u r e m e n t of 613C of cholesterol in pigs (Stott et al. 1997a) and fossil whalebones (Stott et al. 1997b).

9.4.3.2. Isotope composition of individual amino acids The amino acids are the building blocks of proteins, and the isotopic composition of these c o m p o u n d s reflect their pathways of biosynthesis. Amino acids are ubiquitous c o m p o u n d s that give very little clue as to their origin or source by their relative abundance alone. The isotopic compositions of these molecules, however, are diverse and can be quite powerful in delineating pathways of biosynthesis or processes of


199

diagenesis (Macko et al. 1987; Engel et al., 1990; Engel & Macko, 1997; Fogel et al., 1997; Fogel & Tuross, 1999; Fantle et al., 1999). Predictable and fundamental isotopic patterns have been studied in a diverse group of organisms from differing ecological and geological contexts. Typically in plants, glycine and aspartic acid are the most enriched in 13C with valine being the most depleted (Fogel et al., 1997; Fogel & Tuross, 1999). The distribution of isotope signatures of amino acids in higher trophic level organisms is dependant on, and indicative of, the food source. The 813C and 815N values of individual amino acids in such organisms may be predicted on the basis of their pathway of biosynthesis (non-essential amino acids) or their direct incorporation from the diet (essential amino acids) (Fogel et al., 1997). Isotopic compositions of amino acids have been analyzed by off-line combustion after the compounds had been separated by high performance liquid chromatography (Abelson & Hoering, 1961; Macko et al., 1987; Hare et al., 1991; Balzar et al., 1997). As stated, liquid chromatography has yet to be routinely interfaced with an isotope ratio mass spectrometer, therefore a majority of the analyses are now performed on derivatized amino acids separated by gas chromatography. Tissues are hydrolyzed with HC1, dried down, and then derivatized by a number of different methods (see Chapter 8). The method detailed in Box 9.6 is straight-forward, robust, and a majority of the major biological amino acids can be separated for isotopic analyses. In addition, the method is suitable for chiral detection of D- and L- amino acids (Silfer et al., 1991; Metges et al., 1996). A disadvantage of this method is the addition of fluorine molecules that provide volatility can contaminate oxidation furnaces and reduce the working lifetime of these furnaces. The isotope composition of amino acids is determined by isotope fractionations associated with the specific enzyme-mediated reactions which occur during their synthesis. Glycine is commonly the amino acid with the most 13C-enriched isotopic comBOX 9.6- Amino acids in bone collagen

Samples of isolated collagen (Ambrose, 1990; Koch et al., 1994) are weighed (1-3 mg), and loaded into pre-combusted hydrolysis tubes, and hydrolyzed under an atmosphere of N2 in lml of 6N constant-boiling HC1 (Pierce Chemical) (20 hours at 110~ The resulting hydrolysates are dried under a stream of N2 at 100~ and the released amino acids esterified by I ml of anhydrous acidified iso-propanol (1 hour at 110~ and subsequently acylated by 0.5 ml trifluoroacetic anhydride (TFAA) in 0.5 ml dichloromethane (10 min. at 110~ Samples are analyzed on a GC-CIRMS system using a non-polar GC column (e.g. Ultra-I). A suite of amino acid standards are derivatized and analyzed with each batch of samples. For carbon isotope analyses, about 2 mg of hydrolyzed protein is injected onto the column. After various splits in the injector (1:10) and the open split (1:2), about 10 ng of carbon per peak enters the ion source of the mass spectrometer. For nitrogen isotope analysis, about 3 mg of sample is injected in the splitless mode, such ca. 100 ng of amino acid is converted to N2. Each sample should be analyzed in triplicate with corresponding standard amino acids. Reproducibility for carbon isotopes ranges from _+0.2%o to _+1.0%, depending on peak size, with a typical error of _+0.4 %o. For nitrogen isotopes, because the derivative contains no added nitrogen, values are determined directly. Standard deviations on the major peaks (e.g. proline, aspartate, glutamate, lysine, and arginine) are about _+0.5%. Well-resolved but very small peaks (e.g. valine and leucine) have higher errors _+1-2%.

200


position. It has a simple and central role in metabolism. In terrestrial plants, glycine is formed from serine during photorespiratory processes. Isotopically heavier carbon is shunted into this pathway with 13C depleted carbon going into subsequent amino acids synthesized in the Krebs cycle. Glycine is a key intermediate in the formation of porphyrins and purines, thus an understanding of its isotopic composition may be critical in understanding the isotopic composition of molecules, such as chlorophyll. Many amino acids are essential for the growth of animals and the isotopic signatures imparted to these particular essential amino acids can be retained as proteins from primary producers are digested. For example, in omnivorous animals, meat or other animal protein is the primary contributor of lysine to the diet. Therefore, the isotopic composition of lysine should indicate the source of dietary protein, whereas the b u l k ~)13C should reflect the primary carbohydrate or energy source (Fogel et al., 1997). Furthermore, isoleucine can be 6-10%o more positive in 613C than leucine and valine in bacterial and algal proteins. The isotopic difference between leucine and isoleucine transfers up the food chain into fish, birds, and mammals and therefore the 613C of these components can be used to trace trophic interactions (Fogel et al., 1997). T h e ~)13C of

alanine may be an indicator of the energy status of an organism. Alanine is one of the most variable amino acids in terms of isotopic composition, and derives its carbon skeleton directly from that of pyruvate. Pyruvate, and alanine, are central metabolites between the TCA cycle and carbohydrate metabolic pathways (e.g., glycolysis in case of heterotrophs or the Calvin cycle of photosynthesis). Lastly, ranges in gl3C of amino acids are also useful indicators of source and processes. For example, in higher plants, a range of 25%0 in 613C is common, and in phytoplankton, the spread in 613C varies from 15 to 20%o, whereas in sharp contrast, cultured bacterial amino acids have a very narrow range of 6-8%o (Macko et al., 1987; Fantle et al., 1999; Fogel et al., 1997; Teece & Fogel, unpublished results). The different ranges of 613C values is governed by isotope fractionations occurring during amino acid biosynthesis and such distributions may be used in future to determine sources of these compounds to natural ecosystems. 9.4.3.3.

Isotope composition of carbohydrates

The largest reservoir of carbon in the biosphere resides in the carbohydrates, and can comprise up to 40% of the dry weight of bacteria and up to 70% of that of vascular plants. Carbohydrates serve as metabolic storage products, as carbon and energy sources in non-photosynthetic metabolism (e.g. starch, inulin, glycogen, mannitol), and are associated with cell walls and membranes that provide protection, stability and strength (e.g. cellulose, hemicellulose, pectins, chitin, agar, peptidoglycans, lipopolysaccharides and glycolipids). Carbohydrates represent the major form of photosynthetically fixed carbon and are critical energy components of food webs. In the oceans, most of the organic material produced by phytoplankton is consumed by herbivorous zooplankton and protozoans in the photic zone, which are in turn consumed by higher trophic level organisms. Carbohydrates, in particular glucose, represent the major organic carbon


201

BOX 9.7- Isotope composition of individual monosaccharides in zooplankton Samples of isolated zooplankton (20rag) are added to 72% H2SO4 (lml) at room temperature for 90 minutes. After dilution to 1.2M H2SO4, the sample is further hydrolyzed for 3 hours at 100~ An internal standard (myo-inositol) is added to the hydrolysate, and both are subsequently neutralized to pH 6.5 with BaCO3. The resulting BaSO4 precipitate is washed and removed by centrifugation, and the neutralized hydrolysate reduced to 2ml by rotary evaporation. The released monosaccharides are reduced to alditols by reaction with 0.5ml of a freshly prepared solution of NaBH4 (500rag in 10ml H20) for at least 12 hours at room temperature. The residual NaBH4 is decomposed by the addition of glacial acetic acid, and after effervescence ceases (pH 5.5), a second internal standard (erythritol) is added, and the solution evaporated to dryness under reduced pressure. Boric acid is removed by repeated additions of methanol, followed by evaporation to dryness. Acetylation is performed in sealed tubes in a solution of pyridine:acetic anhydride (1:1) for 2 hours at 100~ After the addition of water (2ml), the alditol acetates are extracted into dichloromethane (3 x 2ml) and evaporated to dryness under a stream of nitrogen. The derivatives are analyzed in triplicate by GC-C-IRMS using a polar column (e.g. BPX-70) and compared to a suite of standard compounds treated in an identical manner.

component transferred between trophic levels. The isotopic composition of glucose provides an important indicator of the processes and pathways of energy utilization in ecosystems (e.g. Moers et al., 1993). Measurement of the isotope composition of carbohydrates is an analytical challenge, and therefore there are few published reports of analyses. Macko et al. (1990; 1991) measured the 613C of individual monosaccharides by off-line combustion of components which had been previously isolated by liquid chromatography. Carbohydrates are highly hydrophilic and the challenge to produce derivatives, which are amenable to GC analysis, requires several chemical reactions and derivatization steps (Moers et al., 1993). An approach to the analysis of individual monosaccharides is presented in Box 8.7 (modified from Cowie & Hedges, 1984), however as many steps are involved it is paramount to test these procedures with standard compounds to determine whether the reaction scheme results in isotope fractionation during derivative preparation. In the majority of biological material, carbohydrates are enriched in 13C relative to amino acids and lipids, and also bulk tissue. Individual monosaccharides can be up to 10%o enriched in 13C relative to whole tissue, such as in cyanobacterial mats (Moers et al, 1993). In herbaceous and woody pant tissues, cellulose and hemicellulose are typically enriched in 13C by 1-2%o relative to whole plant material (Benner et al, 1987). During the diagenesis of organic matter, the isotopic composition of carbohydrates provides a means to trace sources of these compounds, either resulting from new bacterial production or degradation (Macko et al., 1991). In the future, the isotopic composition of individual monosaccharides may provide additional vital information in tracing sources of organic matter and the transfer of energy in ecosystem studies. The methods for the isotopic analysis of these compounds has been established, and now provides an additional tool for investigating ecological questions on a molecular scale.

'

202

Chapter9 - M.A.Teece & M.L.Fogel

9.5 Conclusions

Because of new technological advancements in the elemental analyzers linked to isotope ratio mass spectrometers (e.g., Kornexl et al., 1999a), stable isotope analysis is quickly becoming a standard measurement for interpretation of various biological and environmental parameters. We easily imagine that in the next ten years these instruments will take their place along side of nutrient autoanalyzers, fluorescence instruments, and UV/V is spectrophotometers. The challenge then arises for ecologists to use stable isotope tools as wisely and as carefully as those who developed the early methods, perfected the techniques, and pioneered their use. Graduate students in ecology are currently learning the "ins and outs" of stable isotope fractionation and systematics, as part of their core education. The compound specific techniques are deeply rooted in chemistry and will probably remain in the domain of geochemists and biochemists. Chemistry departments, on the other hand, are finding that creating links to environmental and biological colleagues through interdisciplinary collaborations open up new ways of examining complex ecosystems. Students in ecology will naturally gravitate to laboratories where these more complex isotopic methods are being practiced and applied to biological problems. The ability to couple the power and speed of elemental analyzer online methods to the specificity of individual compound work will be key for the full integration of stable isotopes as essential parameters for understanding and delineating ecosystems and ecosystem processes.

Acknowledgments The authors would like to thank Henry Fricke, Susan Ziegler, Timothy Filley, Joachim Bebie, Matthew Wooller, and Richard Ash for helpful advice and discussions during the preparation of this manuscript. Helpful comments were provided by the reviewers Brian Fry and Page Chamberlain.


C H A P T E R 10

Extraction of Dissolved Inorganic Carbon (DIC) in Natural Waters for Isotopic Analyses E. A. Atekwanal & R. V. Krishnamurthy2 Department of Geology, Indiana University Purdue University, 723 W. Michigan Street, SL 122, Indianapolis, IN 46202-5132, USA 2 Department of Geosciences, Western Michigan University, 1187 Rood Hall, Kalamazoo, MI, 49008, USA email: I [email protected], 2 [email protected] 1

10.1 Introduction 10.1.1 Dissolved inorganic carbon (DIC) Dissolved inorganic carbon (DIC) consists of CO2(aq), H2CO3, HCO3-, and CO32-. Its concentration in natural waters is governed by the interaction of various carbon species and can be represented by the following reactions (Garrels & Christ, 1965). CO2(g) + H 2 0 *--, H 2 C O 3 H2CO3 ~ H + + HCO3H + + C a C O 3 ,--, H C O 3 - + C a 2+

[10.1a] [10.1b] [10.1c]

The source of CO2(g) initiating the above reaction [10.1a] in natural waters is either atmospheric or biogenic. The biogenic component is predominantly from root respiration or oxidation of organic matter. The carbonate component [10.1c] in ground and surface waters is derived from the dissolution of soil and/or bedrock carbonates. In studies of DIC in natural waters, determining both the concentration and carbon isotopic ratio of DIC (~113CDIC) are useful. The concentrations of DIC and ~)13CDIC in natural waters are controlled by carbon flux, as well as, biogeochemical cycling in the system. Any process that contributes to or removes carbon from natural waters influences the DIC pool. Changes in DIC concentration can occur due to dissolution/precipitation reactions within the aqueous carbonate system, biogenic uptake and release of CO2 and from mixing of carbon from different sources. Changes in the ~)13CDIC would result from fractionation accompanying carbon transformations during the above processes.

10.1.2 Utility of DIC measurements

In hydrologic systems, the concentrations of various species of DIC are indicative of the reaction pathways of waters. Using thermodynamic principles, the concentra-

204

Chapter 10- E.A. Atekwana & R.V. Krishnamurthy

tion and ~)13CDIC can be used to define the DIC and hydrologic evolution of waters (e.g. Deines et al., 1974; Wigley et al., 1978). Examples of a few studies in which DIC has also been utilized include, tracing groundwater recharge (e.g. Landmeyer & Stone, 1995; Cane & Clark, 1998), studies of carbon evolution of groundwater (e.g. Reardon & Fritz, 1978; Chapelle & Knobel, 1985; Rose & Long, 1989; White & Chuma, 1987; McMahon & Chapelle, 1991; Cai et al., 2003), carbon cycling in streams (e.g. Pawellek & Veizer, 1994; Flintrop et al., 1996; Yang et al., 1996a; Atekwana & Krishnamurthy, 1998; Helle et al., 2002; Finlay, 2003), carbon cycling in lakes (e.g., Rau, 1978; Quay et al., 1986; Herczeg, 1987; Weiler & Nriagu, 1987; Wachniew & Rozanski, 1997; Nakamura et al. 1998; Vreca, 2003), determining the extent of surface water-groundwater interaction (e.g. Taylor & Fox, 1996) and assessing microbial influence of groundwater isotopic composition in natural and polluted groundwater (e.g. Chapelle et al., 1988; Chapelle & McMahon, 1991; Nascimento et al., 1997; Hunkeler et al., 1999; Fang et al., 2000; Pombo et al., 2002).

10.1.3 Principles of DIC extraction DIC in natural waters is extracted by converting carbon components into CO2 gas. Quantitative measurements of the evolved CO2 permit determination of DIC concentration; the CO2 is introduced into an isotope ratio mass spectrometer (IRMS) for 613C determination. Isotope ratios are reported in the 6-notation where: ~) (%0)- ((Rsample / Rstandard)-1) x 10 3

[10.2]

R is 13C/12C. Values are reported relative to VPDB (Vienna Pee Dee Belemnite). In studies where the carbon-14 activity of DIC is of interest, the extracted CO2 is prepared for C-14 analysis and measured using counting techniques or Accelerator Mass Spectrometry (AMS). 10.2 DIC extraction

The information presented in this chapter is essentially a guide for selecting an appropriate technique for DIC extraction for isotopic analysis. The focus is on extraction techniques routinely used for DIC isotopic analyses. Mass spectrometric analyses are not discussed because of differences in mass spectrometric types, individual machine configuration and analytical setup. The techniques presented assume that mass spectrometric CO2 analysis is off line, although the procedures described can be adapted to online systems with slight procedural modifications. For micro-extraction of DIC and ~)13CDIC determination, the techniques of Graber & Aharon (1991) and Salata et al. (2000) can be consulted. For extraction of individual carbon components for isotopic analysis from water, the technique of Games & Hayes (1976; 1977) can be used. St-Jean (2003) describes a method for automated 613C analysis of DIC on continuous-flow isotope ratio mass spectrometer (CF-IRMS). The techniques described here do not consider water sampling methods or protocol. It is assumed that sampling of groundwater, surface water (grab or depth specific samples) or pore water in sediments or rock will be done in a manner that preserves sample integrity, such as preventing CO2 degassing from the sample or atmospheric

205

Extraction of Dissolved Inorganic Carbon (DIC) in Natural Waters for Isotopic Analyses

CO2 contamination. Techniques currently used for DIC extraction for isotopic analy-

ses can be categorized as variants of the Gas Evolution, Precipitation, Precipitation and Gas Evolution, and Vapor Phase Equilibration techniques (Figure 10.1). 10.3 Gas evolution technique In the gas evolution technique, DIC in a water sample is converted to CO2 under vacuum conditions by acidification (Mook, 1968; Tan et al., 1973; Games & Hayes, 1976; 1977; Reardon et al., 1979; Hassan, 1982; Graber & Aharon, 1991; Holt et al., 1995; Atekwana & Krishnamurthy; 1998; Cane & Clark, 1998). The yield of CO2 extracted is converted to DIC concentration and the CO2 is introduced into a mass spectrometer for 613CDIC determinations. The main advantage of the gas evolution technique is the simultaneous determination of DIC concentration and the 813CDIC of the same water sample. Although the gas evolution technique is simple, it can be plagued with problems that potentially cause erroneous concentration and 613CDIC determinations. Hassan (1982) has shown that the gas evolution technique seems to suffer from CO2 gain/loss from water in sample containers during shipping and laboratory storage prior to DIC extraction, and incomplete transfer of evolved gas phase in sample containers to the vacuum line

~

$

D][C

1"Coi c

8~sCo,c

$

l~ol c

s~n]~

,.~C03

(

l'~oiC

I-~ a c l ~ e

DIG

HeMsp~e

H~O ~po~

81RaOlC

Figure 10.1 - Flow diagram showing common techniques and major steps involved in DIC extraction and possible measurements associated with CO2 extracted by each method.

206

Chapter 10- E.A. Atekwana & R.V. Krishnamurthy

during extraction. Additionally, water sample storage problems can be compounded by bacterial activity (including photosynthesis and respiration) which could alter DIC concentrations and ~513CDICin samples not stored in opaque glass bottles, cooled to 4~ and/or preserved by the addition of a bactericide such as HgC12 (e.g. Mook, 1968; Yang et al., 1996a; Cane & Clark, 1998). One technique that overcomes these problems, is simple, and allows for accurate determination of both DIC concentrations and 613CD~C (Atekwana & Krishnamurthy, 1998) is detailed below.

10.3.1 The Vacutainer Gas Evolution Technique (Atekwana & Krishnamurthy, 1998) This procedure uses vacutainer septum tubes (blood serum vials) to collect, react, and extract DIC from natural water samples. We call it the vacutainer technique to distinguish it from other gas evolution techniques. The vacutainer tubes are preloaded with 85% phosphoric acid and magnetic stir bars. In the lab, air is evacuated from the septum tubes after loading with phosphoric acid and magnetic stir bars. In the field, water samples are introduced into the septum tubes using a syringe. Sample collection and injection takes less than 30 seconds. The acid-water reaction begins immediately upon injection. The evolved CO2 is extracted in the laboratory, where yields are measured and isotope ratios determined.

10.3.1.1 Procedure Sample container preparation List of supplies~equipment needed 9 16 x 100 or 16 x 165 mm glass septum tubes (VACUTAINER| Serum Tubes, Becton Dickson & Company, Franklin Lakes, NJ 07417) 9 85 % phosphoric acid 9 13 x 8 mm magnetic stir bars (Fisher brand| cat. No. 14-511-62)1 9 Extraction needle (consisting of an inner luer ground joint joined to a piece of 9mm diameter tube by glass blowing. A 26 gauge needle (small diameter to prevent septum coring) is affixed to the luer end 9 Laboratory vacuum extraction system To prepare the septum tubes for DIC extraction, the septum of each tube is removed and a magnetic stir bar inserted into each of the tubes. If the CO2 is removed from the septum tubes using ultrasonic water bath, the magnetic stir bar is omitted. Approximately 0.5 ml of 85% phosphoric acid is dispensed into each tube and their septa replaced. Each septum tube is evacuated of air introduced into the tubes while loading with phosphoric acid and magnetic stir bars using an extraction needle attached to the laboratory vacuum line by way of a Cajon union (Figure 10.2, insert A). Evacuation of air takes a few minutes and is complete when the vacuum gauge on the system has recovered to proper vacuum conditions. During this step, most of the moisture and any CO2 in the phosphoric acid is removed.

1. A modification of the DIC extraction technique where CO2 is removed from the septum tubes using an ultrasonic water bath does not require use of magnetic stir bars (see DIC extraction section).


207

Figure 10.2- Simplified schematic showing a laboratory v a c u u m extraction system. Insert (A) shows an extraction needle used to evacuate septum tubes pre-loaded with phosphoric acid and a magnetic stir bar. Insert (B) shows a DIC extraction set-up used to transfer CO2 from water-acid reaction in septum tubes into the laboratory vacu u m system. This set-up consists of a heater~stirrer, a water bath (or an ultrasonic water bath), an extraction needle and an external U-tube moisture / ice trap.

Water sample collection

List of supplies~equipment needed 10-ml plastic syringes with a luer lock tip (Becton Dickson & Company, Franklin Lakes, NJ 07417) 9 26-gauge needles (Becton Dickson & Company, Franklin Lakes, NJ 07417) 9 Prepared septum tubes (Pre-evacuated septum tubes loaded with phosphoric acid and magnetic stir bars)

9

In the field, a known volume of water (usually 5 or 10 ml) is introduced into the prepared septum tube using a plastic syringe fitted with the 26-guage needle. Before collecting a water sample, the needle is removed from the syringe in order to facilitate rapid uptake of water into the syringe and to prevent possible vacuum degassing of the sample due to the small needle size. After filling, the needle is quickly reattached to the syringe. Air bubbles, if present, are removed by tapping the syringe gently and excess water subsequently expelled to obtain the desired volume. Filtration of the water before injection in the septum tube can be accomplished using syringe filter with a female luer-lok inlet and male luer slip outlet. The sample is then injected into the prepared septum tube. The tube is shaken for a few seconds to mix the acid and water. The reaction between the water and the phosphoric acid starts instantaneously.

208

Chapter 10 - E.A. Atekwana& R.V. Krishnamurthy

Reacted samples are transported as such to the laboratory for CO2 extraction. No special storage conditions are needed except care to prevent breakage of the glass sample tubes.

Laboratory extraction of C02 from the water~acid reaction mixture List of supplies~equipment needed 9 Septum tubes with reacted samples 9 50~ water bath (water in a beaker on a hot plate/stirrer is sufficient) or water in an ultrasonic agitator (e.g. Fisher Scientific Catalog 15-335-6) 9 -80~ ice slush (mixture of dry ice and organic solvent) ~ Liquid nitrogen 9 External U-tube trap (9 mm Pyrex glass tube blown to fit vacuum system; see Figure 10.2, insert B) 9 Extraction needle (consisting of a 23-gauge needle attached to an inner luer ground joint, joined to a piece of 9 mm tube by glass blowing) 9 Laboratory vacuum extraction system Evolved CO2 in septum tubes is extracted using an external modification to the vacuum system (Figure 10.2, insert B). During extraction, the septum tubes with reacted samples are placed in a water bath at about 50~ (or in an ultrasonic water bath). A removable external U-tube trap is utilized to ensure that the bulk of moisture is trapped prior to gas entry into the extraction line. A Cajon union is used to connect the external U-tube trap to the vacuum system. The extraction needle is attached to the external U-tube trap by way of a threaded vacuum seal. Prior to CO2 extraction, the needle attached to the external U-tube trap (Figure 10.2, insert B) is partially inserted into the septum of the tube containing the reacted sample, thereby covering the needle opening. This step is required to evacuate the dead volume between the stopcock on the vacuum line and the needle-tip (Figure 10.2). After evacuation, the needle is pushed through the septum and the CO2 is extracted dynamically while the reaction mixture is continuously stirred. This step is carried out for 5 to 10 minutes. It is noteworthy to mention that stirring was found necessary in this procedure. Without stirring the extraction efficiency for standard solutions varied between 85 and 93 %. If the ultrasonic water bath is used, sonic agitation is sufficient to effectively transfer all CO2 from the sample/acid mixture in the septum tube to the vacuum line. The extracted CO2 is trapped in the vacuum line with liquid nitrogen (-180~ Residual non-condensable gases are pumped away and the trapped CO2 is purified cryogenically by substituting the liquid nitrogen in the trap with-80~ ice slush. The CO2 released is then re-trapped in a calibrated cold finger using liquid nitrogen and its yield measured manometrically. The CO2 is collected from the cold finger and introduced into the mass spectrometer for carbon isotopic measurements.

Extraction efficiency and r

precision of the Vacutainer Gas Evolution Technique

Several experiments were conducted to verify the accuracy and precision of the vacutainer gas evolution technique. Initially, C02 evolved from "solid-form" samples,


209

each representing different amounts of solid Na2CO3, were reacted with 100% phosphoric acid and extracted using the method of Krishnamurthy et al. (1997). The amount added, the experimental yield and percentage yield as CO2 and the corresponding 613C of the extracted CO2 are presented in Table 10.1. The data shows that the extraction efficiency (% yield as CO2) of the solid standard is better than 99% with a ~13C of 7.5 + 0.1%o. Next, a solution was prepared using the solid Na2CO3 for DIC extraction. Evolved CO2 from 5-ml aliquots of the standard and 5 ml of natural waters were extracted using the above DIC extraction procedure. Natural waters included laboratory tap water, stream water, runoff from a wetland and spring water downgradient of a landfill, and water obtained from a landfill leachate collection system. The purpose was to determine the optimum time required for CO2 extraction from a reaction mixture and the extraction efficiency of the procedure. Individual extractions were carried out for 5, 10, 15, and 20 minutes. The results of this time series extraction of the Na2CO3 solution and natural water samples are shown in Table 10.2. It is observed that 100 % of the theoretical CO2 yield from the Na2CO3 solution is obtained from the reaction mixture within five minutes. Additionally, the ~13C of the Na2CO3 solution (-7.5 + 0.1%o) was identical to that of the Na2CO3 solid (-7.5 + 0.1%o, Table 10.1). The natural waters extracted by this technique also yield highly reproducible isotopic results. With a standard deviation of CO2 yields of less than 1 ~mole, at least 98% of the DIC was extracted from the natural samples within five minutes. Except for the stream water with a 613CDIC range of 0.5%0, the rest of the natural waters do not vary by more than 0.2%0 for different extraction time intervals. The lower precision of the 613CDIC for the stream water could be due to natural variability of DIC concentrations in flowing systems. Precision and accuracy of the technique was further verified by extracting a batch of 10 samples of Na2CO3 solution at 50~ for 10 minutes. The results are also presented in Table 10.2. The results show that the extraction process is 99 + 1% efficient, and the ~13C is measurable to better than 0.1%o. Therefore, an extraction time of 5 to 10 minutes at 50~ with constant stirring provides highly reproducible DIC concentrations and 613C with this technique. Modification of the vacutainer technique where water samples are collected in septum tubes without magnetic stir bars and CO2 is removed from the septum tubes and transferred to the vacuum line by ultrasonic agitation also produced precise and accurate results, a total of 10 samples of Na2CO3 Table 10.1 - CO2 yield and h13C from solid Na2CO3 standard. Run # 1 2 3 4 5

Amount Added (m moles CO2)

Experimental Yield% (m moles CO2)

Yield as CO2

613C

167.7 121.7 109.1 86.9 75.7

167.1 119.9 107.4 86.2 76.4

99.8 98.5 98.4 99.2 100.9 99.4 + 1.0

-7.6 -7.6 -7.5 -7.5 -7.4 -7.5 + 0.1

210

Chapter 10 - E.A. Atekwana & R.V. Krishnamurthy

Table 10.2 - Time series extraction of Na2CO3 solution and selected water samples at 50~ Sample

Extraction Time (minutes)

CO2 Yield (/~ moles)

% Yield as CO2 (~ moles) 100.0 99.8 98.5 99.7 99.5 + 0.7

~)13CDIC (%o)

Na2CO3 1 2 3 4

5 10 15 20

101.3 101.1 99.8 101.0 100.8 + 0.7

Tap Water 1 2 3 4

5 10 15 20

25.4 25.4 25.6 25.9 25.6 + 0.2

-10.6 -10.7 -10.5 -10.9 -10.7 + 0.2

Stream Water 1 2 3 4

5 10 15 20

22.3 22.0 21.8 21.9 22.0 + 0.2

-10.8 -11.1 -11.3 -11.0 -11.1 + 0.2

Wetland Runoff 1 2 3 4

5 10 15 20

96.2 95.6 95.2 94.3 95.3 + 0.8

4.8 5.0 5.0 5.0 5.0 + 0.1

Spring 1 2 3 4

5 10 15 20

101.0 101.5 101.3 101.5 101.3 + 0.2

7.7 7.9 7.9 8.0 7.9 + 0.1

Landfill Leachate 1 2 3 4

5 10 15 20

131.7 131.0 131.3 130.6 131.2 + 0.5

3.6 3.8 3.7 3.7 3.7 + 0.1

Na2CO3 (n = 10)

10

98.9 + 1.0

-7.4 -7.5 -7.5 -7.5 -7.5 + 0.1

-7.5 + 0.1

s o l u t i o n e x t r a c t e d u s i n g t h i s m o d i f i c a t i o n w a s 99 + 1% efficient, a n d t h e (~13C w a s m e a s u r a b l e t o b e t t e r t h a n 0.1%o ( D a t a n o t s h o w n ) .

10.3.1.2 Advantages of the vacutainer gas evolution technique compared to other gas evolution technique T h e v a c u t a i n e r g a s e v o l u t i o n t e c h n i q u e of A t e k w a n a & K r i s h n a m u r t h y (1998) d e s c r i b e d a b o v e has a d v a n t a g e s in that the p r o c e d u r e o v e r c o m e s m a j o r p r o b l e m s a s s o c i a t e d w i t h t h e o t h e r g a s e v o l u t i o n t e c h n i q u e s . A l l e v i a t e d p r o b l e m s are d i s c u s s e d below.


211

Sample reaction in C02-free atmosphere Prior to sample collection, the septum tubes are preloaded with magnetic stir bars and phosphoric acid and then evacuated. This not only removes the ambient air introduced during loading but also removes any CO2 from the acid, consequently, preventing sample contamination.

Transfer of aliquot of water sample with representative DIC to the vacuum system during extraction The problem of selecting a suitable container for water collection that allows for complete transfer of evolved CO2 from the sample container (Hassan, 1982) is also solved by this technique. For example, one difficulty with water samples collected via other techniques is in assessing the CO2 partitioned in the headspace relative to the water phase. This is especially true for water with high CO2 content. The septum tubes do not suffer from headspace CO2 partitioning because all the CO2 evolved in the septum tube via the acid reaction is from the parent-sample and is effectively transferred to the laboratory vacuum line in this procedure.

Post sampling alteration of DIC Upon water collection and injection into the prepared septum tube, DIC is immediately converted into CO2 and confined therein. The highly acidic condition of the reaction mixture (pH 99.9999%) 9 10 ml gas tight syringes (Dynatech Precision Sampling Corp.) 9 CO2-free 6N HC1 (prepared by bubbling with ultrapure He gas for at least 3 h)

Extraction of Dissolved Inorganic

Carbon

(DIC) in Natural Waters for Isotopic Analyses

225

The vapor phase equilibration is carried out at constant room temperature (23~ A headspace of known volume (5 ml) is created in each sample container by injecting ultra-pure He using a gas tight syringe. A similar volume of water is withdrawn with another syringe (Figure 10.5). Another plastic syringe is used to acidify the water sample by injecting 0.5 ml of CO2-free HC1 solution (6N) through the septum. The samples are shaken vigorously by hand and left upside down and in the dark for at least 40 hours to equilibrate. After equilibration a known amount of headspace gas is withdrawn with a 1-ml gas tight syringe and injected into the G C / C / I R M S for CO2 concentration and isotopic ratio determination. 10.6.1.3 Determination of DIC concentration and c513CDIC Before outlining the procedure for determining DIC concentrations, a simplified description of the GC/C~ IRMS would be appropriate. The G C / C / I R M S is a gas chromatograph interfaced with an isotope ratio mass spectrometer by a combustion furnace. The GC is used to separate the constituents in an injected gas mixture in a packed column using a carrier gas (usually He). The separated constituents are passed through the combustion furnace. The combustion furnace, which operates at temperatures of 800 ~ - 900~ is used for combusting organic carbon constituents. However, in DIC analysis, it is maintained at the proper operating temperature to stabilize analytical conditions even though no combustion is needed. The separated CO2 is then passed into the IRMS where isotopic ratios are determined by comparison to a reference CO2.

Equilibrating CO2 between the headspace and the liquid phase in glass serum bottles containing acidified water sample with headspace (A). Withdrawing a portion of the headspace gas with a gas-tight syringe (B). (Modified from Miyajima et al., 1995). F i g u r e 10.5 -

226


The output from the GC/C/IRMS consists of the calculated (~13C, along with some chromatographic parameters such as retention times, peak width, and peak area. The ion current signal for M/e (molecular weight per electronic valence) of 44 (12C1602+) of the mass spectrometer is used to estimate the CO2 concentration. As the fraction of 12C1602+ exceeds 98% in natural CO2 samples, the peak area for M/e is closely related to the amount of CO2 injected into the GC/C/IRMS (Miyajima et al., 1995). This allows the concentration of aliquots of CO2 analyzed from the headspace gas to be determined from the peak area for M/e - 44. The ~)13Cvalue of XCO2 (DIC) for the sample is calculated from aliquots of the headspace gas injected into the GC/C/IRMS using the following equation (Miyajima et al., 1995)" a

~)13C(YCO2) - [V h x n + (V b - V h ) x ~ x (n + ~g)] + [V h + (V b - V h ) • ~]

[10.3]

w h e r e Vh is the volume of the headspace, Vb is the inner volume of the serum bottle, n is the ~)13Cvalue of the headspace CO2, ~ is the Ostwalt solubility coefficient and ~ga is the isotopic difference between the gas and aqueous CO2. Both 13and ega are tempera-

ture dependent.

10.6.1.4 Precision and accuracy of the technique The precision and accuracy of determining both the concentration and ~)13CDIC using the vapor phase equilibration method depends on the assumptions used in the headspace equilibration and the analytical precision of isotopic analysis. Equilibration between the CO2 in the headspace and aqueous phases is assumed which can differ, however, depending on the nature of water sample. Factors that affect this equilibrium in water, such as salinity, ionic strength, and concentration of humic substances, are usually not measured or are difficult to determine. This could potentially lead to error in DIC analyses. No experiments have been conducted to determine how these parameters affect the precision or accuracy of the DIC analysis possibly restricting this technique to dilute waters where these effects are considered minimal (Miyajima et al., 1995). Optimization of analytical conditions and correction for isotopic discrimination of headspace CO2 injected into the GC/C/IRMS, as well as, the CO2 sample size affect precision and accuracy of DIC concentration and g13CD~C determination using this technique. A detailed discussion related to these and other parameters in this technique is found in Miyajima et al., (1995). 10. 7 General considerations and conclusions

Widespread use of DIC in research has been hampered by the technical difficulties associated with the DIC extraction and mass spectrometric procedures. In terms of the extraction procedure, ease of sampling and reliability/reproducibility of the DIC extraction procedure is paramount. The salient points of the main DIC extraction procedures are summarized in Table 10.8. The vacutainer gas evolution technique of Atekwana & Krishnamurthy (1998) is simple, versatile, requires less technical sophistication and provides accurate DIC concentration and 613CDIC. Compared to the pre-

Table 10.7 -Summary comparison of the methods of DIC extraction from natural waters Gas Evolution

Sample volume collected

Precipitation

5 or 10 mL

Precipitation and Gas Evolution

20 m L - 1 L

Ca

Vapor Phase Equilibration

300 ml - 1 L

o 9

68 mL

9 f0r...r0utine...D!C...~y.se.s .< Water sample preservaNot required for the method Not required Not required Saturated HgC12 for tion of Atekwana and long term storage ....................................................................................K r ! s ~ a m u r t h y ( ! 9 9 8 ) O Simultaneous determination Routinely determined Not routinely determined Routinely determined Routinely determined of DIC and 6 1 3 C D I C Ca ........................................................................................................................................................................................................................... Extraction of DIC from small Possible, limited by amount Difficult to process on a Difficult to process on a Posslbiewi~:hmodificafions) ....... r sample volume riers Shipping of extracted CO2 to Possible Possible Possible Not routine. Possible with other laboratories for analysis significant procedural modifications Time required for routine sam- Hours Two or more days Two or more days Two or more days pie collection to isotopic measurement of extractedCO2 Allows for 14C analysis Possible Possible Possible Not routine. Possible with of DIC significant procedural tO modifications .

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

.

..................................................................

.............................................................................................................................

tO --.a

228


cipitation, precipitation and gas evolution or the vapor phase equilibration techniques, it is experimentally less involving. Using the vacutainer gas evolution technique, water samples can be obtained from many regions of the world with little difficulty and without the need for the researcher to travel to remote sites if cost of travel is a problem. For example, prepared sample tubes have been shipped to scientists in Central America and southern Africa where water samples were collected and shipped to the USA for DIC extraction. Although possible with the other techniques described here, both the logistics and the level of expertise required to conduct similar remote sampling are daunting. Using the vacutainer technique, water samples can be collected and extracted in a matter of a few hours. In a single vacuum set-up, more than 30 samples can be extracted in an 8hour day. Commercial laboratories can easily adapt this technique for routine DIC analysis. These laboratories can prepare the sample tubes, ship them with appropriate instructions to field sites for water sampling. The tubes with water samples can be shipped back to the laboratories for later extraction and isotopic determinations since no special storage conditions are necessary and the length of storage is not problematic. Although the vapor phase equilibration technique offers similar advantages for remote sampling, the DIC extraction and analysis procedure is also technically demanding relative to the vacutainer gas evolution technique. In addition, sample preservation is recommended for long periods of storage requiring additional chemical expense. The method is less amenable to water samples with low DIC concentration since only a small portion of the gas phase is sampled for isotopic analysis. In addition, it does not allow for use of the extracted DIC for possible 14-C analysis if needed. The precipitation and gas evolution techniques provide accurate determinations of t h e 13C of DIC. Accurate determinations of DIC concentration are possible with the precipitation and gas evolution technique. These techniques are less attractive due to the elaborate procedure and precautions needed to obtain accurate results. In addition, the large volumes of water sample (approximately 20 ml to 1L) used in current techniques make it less attractive for analyses involving smaller volumes of water extracted from sediment cores, for example. The main advantage of the techniques lies in the fact that sufficient CO2 can be produced for both conventional and AMS 14C determinations.


CHAPTER 11 Compound Specific Isotope Analysis of the Organic Constituents in the Murchison Meteorite M. A. Sephtonl & I. Gilmour Planetary and Space Sciences Research Institute, Open University, Milton Keynes, Buckinghamshire, MK7 6AA, United Kingdom e-mail: 1 [email protected]

11.1 Introduction Most meteorites are pieces of asteroids that have been propelled towards the Earth by collisions in the asteroid belt (Wetherill & Chapman, 1988). The parent asteroids have escaped the extensive geological recycling prevalent on the planets and, consequently, are composed of extremely ancient materials. As a result, meteorites constitute a record of the early solar system and to some extent the interstellar environment from which it formed. Of all the meteorites, the carbonaceous chondrites are the most primitive and attract particular attention from the organic cosmochemist because they contain up to several percent carbon, the majority of which is present as organic matter. Less than 25% of the organic matter in carbonaceous chondrites is present as low molecular weight free compounds which can be extracted with common organic solvents, while the remaining 75% or so is present as high molecular weight macromolecular materials which are insoluble in common organic solvents (Hayatsu & Anders, 1981). Meteoritic organic matter may have been created in a number of extraterrestrial environments associated with the transformation of an interstellar cloud, through a solar nebula, into a star and planetary system (Cronin et al., 1988). Subsequently, once on the asteroidal parent body, the organic matter may have undergone modification by secondary thermal and aqueous alteration. Scrutinising the record contained within meteoritic organic matter can lead to an improved understanding of the inputs and processes contributing to the early solar system environment. Inexorably, the study of meteoritic organic matter also provides valuable information on prebiotic molecular evolution and the origin of life.

1. Correspondence should be adressed to this author.

230

Chapter 11 - M.A. Sephton & I. Gilmour

Much of our current understanding of meteoritic organic matter has come from investigations of the Murchison meteorite, approximately 100 kg of which fell in Australia in 1969. The indigenous nature of the organic matter within Murchison was quickly established (Kvenvolden et al. 1970) and over the last three decades it has been the focus of intensive research using the most modern techniques. The Murchison organic inventory has become a valuable reference to which all other extraterrestrial organic matter may be compared.

11.2 Stable isotopes The abundances of stable isotopes are generally expressed as 6 values. These indicate the difference, in per mil (%o), between the relevant ratio in the sample and the same ratio in an international standard as follows: %o - ( Rsample- Rstandard / x 1000 Rstandard

[11.1]

The stable isotopic composition of extraterrestrial materials may be determined by nucleosynthesis, nuclear processing and chemical fractionation processes (Penzias, 1980). By establishing the isotopic composition of meteoritic organic matter the significance of each extraterrestrial source region in its production and modification may be constrained. Early studies of this kind obtained isotopic measurements on broad assemblages of meteoritic compounds in the form of either the bulk macromolecule, acid residue or evaporated unfractionated solvent extracts (Smith & Kaplan, 1970; Krouse & Modzeleski, 1970; Kvenvolden et al., 1970; Belsky & Kaplan, 1971; Becker & Epstein, 1982; Robert & Epstein, 1982; Yang & Epstein, 1983). Later work achieved higher resolution and measured isolated fractions containing compounds with broadly similar chemical structures. Such crude fractions were isolated by standard preparative chromatography and then subjected to direct combustion-mass spectrometry (MS). Fractions analysed in this way have included hydrocarbons (Krishnamurthy et al., 1992), amino acids (Epstein et al., 1987; Pizzarello et al., 1991), carboxylic acids (Krishnamurthy et al., 1992; Epstein et al., 1987; Cronin et al., 1993) and polar hydrocarbons (Pizzarello et al., 1994). These studies represent significant advances in confirming the indigenous nature of organic matter by detecting noticeable enrichments in the heavy isotopes of C, N and H (D) when compared to terrestrial values. These measurements also provide valuable clues as to the extraterrestrial environments that contributed to the various classes of meteoritic organic matter, with perhaps the most notable of these being Denrichments characteristic of interstellar space. However, these isotopic measurements can only be averages of a wide variety of organic compounds. Isotopic analysis at the molecular level, known as compound specific isotope analysis (CSIA) represents a further stage in the deconvolution of the structural and isotopic constitution of meteoritic organic matter.

231

Compound Specific Isotope Analysis of the Organic Constituents in the Murchison Meteorite

Table 11.1 - Results from carbon isotopic measurements performed on individual molecules from the Murchison meteorite.

Compound Amino acids glycine D-alanine L-alanine glycine + alanine c~-aminoisobutyric acid c~-aminoisobutyric acid + c~-aminobutyric acid aspartic acid L-glutamic acid isovaline isovaline + valine glycine + alanine

613C %o

Ref.

+22 +30 +27 +41 +5 -1

(1) (1) (1) (2) (1) (2)

+4 +6 +17 +30 +41

(2) (1) (1) (2) (2)

Carboxylic acids, CO and C02 CO -32.0 +2.0 CO2 +29.1 +0.2 acetic acid +22.7 +0.2 propionic acid +17.4 +0.2 isobutyric acid +16.9 +0.2 butyric acid +11.0 +0.3 isovaleric acid +8.0 +0.2 valeric acid +4.5 +0.2

(3) (3) (3) (3) (3) (3) (3) (3)

Sulphonic acids methyl suphonic acid ethyl suphonic acid propyl suphonic acid isopropyl sulphonic acid

(4) (4) (4) (4)

Light hydrocarbons methane ethane propane isobutane butane ethene Normal alkanes n-C12 n-C13 n-C14 n-C15 n-C16 n-C17 n-C18

+29.8 +9.1 -0.4 -0.9

Compound Aromatic compounds benzene toluene naphthalene 2-methylnaphthalene 1-methylnaphthalene phenanthrene fluoranthene pyrene chrysene benzo[ghi]fluoranthene benzo(e)pyrene benzo(j)fluorenthene

(3) (3) (3) (3) (3) (3)

-35.5 -28.5 -33.2 -28.2 -25.3 -31.0 -30.6

(5) (5) (5) (5) (5) (5) (5)

Ref.

-28.7 +0.2 -28.8 +1.1 -12.6 +2.3 -5.8 -11.1 -7.5 +1.2 -5.9 +1.1 -13.1 +1.3 -14.5 +2.2 -14.2 +2.2 -22.3 +4.1 -15.4 +3.3

(3) (6) (6) (6) (6) (7) (7) (7) (7) (7) (7) (7)

Fragments of macromolecular material toluene -24.6 +0.2 -1.3 +2.0 ethylbenzene m-xylene p-xylene methylethylbenzene benzaldehyde phenol 2-methyl phenol 3-methylphenol naphthalene

+9.2 +1.0 +3.7 +0.1 +1.2 +0.1 +4.4 +0.1 +2.4+0.1 +0.1 +0.4

~)13C%o

benzothiophene 2-methylnaphthalene 1-methylnaphthalene acenaphthene

-5.4 -21.9 +1.4 -19.6 -21.7 +0.4 -20.3 -17.8 -18.1 -23.7 +0.6 -24.0 +0.3 -24.1 -10.3 -10.4 -13.9 +0.9 -18.5 +1.9 -6.5 +2.5 -5.5 +0.5 -6.4 +0.8 -15.8 -5.6 +2.1 -6.4 +0.8 -7.2 +2.0 -7.1 -5.9 +1.7

(6) (8) (8) (6) (6) (8) (8) (6) (6) (8) (8) (6) (6) (6) (8) (8) (6) (8) (8) (6) (6) (8) (6) (8) (6)

References: (1) Engel et al., 1990; (2) Pizzarello et al., 1991; (3) Yuen et al., 1984; (4) Cooper et al., 1997; (5) Sephton et al., 2001; (6) Sephton et al., 1998; (7) Gilmour & Pillinger, 1994; (8) Sephton & Gilmour, 1999.

232


Table 11.2 - Results of nitrogen isotopic measurements performed on individual amino acids from the Murchison meteorite (Engel & Macko, 1997).

Compound R-aminoisobutyric acid sarcosine isovaline glycine f~-alanine D-alanine L-alanine L-leucine D,L-proline D,L-aspartic acid D-glutamic acid L-glutamic acid

615N %o + 184 +129 +66 +37 +61 +60 +57 +60 +50 +61 +60 +58

Table 11.4 - Results of hydrogen isotopic measurements performed on individual amino and sulphonic acids from the Murchison meteorite. Parenthical values are corrected for hydrogen exchange during the water extraction procedure. Compound

8D %o

aspartic acid glutamic acid c~-aminoisobutyric acid isovaline + valine glycine + alanine methyl sulphonic acid ethyl suphonic acid n-propyl sulphonic acid i-propyl sulphonic acid

Ref

+214 (+630) +523 (+1013) +67 (+149) +713 (1014) +1072 (+2448) +483 +787 +536 +852

(1) (1) (1) (1) (1) (2) (2) (2) (2)

References: (1) Pizzarello et al., 1991; (2) Cooper et al., 1997.

Table 11.3 - Results of sulphur isotopic measurements performed on individual sulphonic acids from the Murchison meteorite (Cooper et al., 1997). Compound methyl sulphonic acid ethyl suphonic acid n-propyl sulphonic acid i-propyl sulphonic acid

633S %0

(~34S %0

636S %0

A33

+7.63 +0.33 +0.20 +0.32

+11.27 +1.13 +1.20 +0.68

+22.5 +0.8 +2.1 +2.9

+2 -0.24 -0.40 -0.02

11.3 Compound Specific Isotope Analysis (CSIA) Determining the isotopic compositions of individual molecules in complex organic mixtures has been a long-held goal of stable isotope mass spectrometry. In recent years this type of analysis has produced a significant data set for meteoritic organic components (Tables 11.1-4). Such measurements are particularly useful as they can expose the reaction mechanisms and possibly source environments from which the organic constituents have been formed. Some attempts at CSIA involve the use of preparative chromatography to isolate fractions containing simple mixtures or individual molecules. Subsequently, in a separate procedure, the fractions are then combusted to CO2 or N2 for carbon and nitrogen isotopic measurements respectively or transformed to H2 or SF6 for determining the hydrogen or sulphur isotopic compositions. These gases are then introduced into an isotope ratio mass spectrometer and the isotope ratios measured. Measurements obtained by this direct combustion-MS method include the carbon and hydrogen isotopic compositions of C2 to C5 amino acids (Pizzarello et al., 1991) and the carbon, hydrogen and sulphur isotopic compositions of C1 to C3 sulphonic acids (Cooper et al., 1997).

Compound SpecificIsotope Analysisof the Organic Constituentsin the MurchisonMeteorite

233

Recent analytical advances have led to the combination of separation, combustion and isotopic measurement as a single procedure. The resulting gas chromatographyisotope ratio mass spectrometry (GC-IRMS) system allows the relatively rapid determination of the isotopic compositions of individual molecules within complex organic mixtures (Mathews & Hayes, 1978). As GC-IRMS relies on the GC-amenability of the analytes, volatile meteoritic compounds with hydrocarbon skeletons and few functional groups have been the focus of several studies. Using GC-IRMS, carbon isotopic compositions have been determined for CO, CO2, C2 to C5 aliphatic hydrocarbons (Yuen et al., 1984), C1 to C5 carboxylic acids (Yuen et al., 1984), C6 to C20 aromatic compounds (Yuen et al., 1984; Gilmour & Pillinger, 1994; Sephton et al., 1998) and C12 to C18 n-alkanes (Sephton et al., 2001). GC-IRMS analyses of the less volatile polar compounds are more problematical. These compounds need to be derivatised to increase their volatility and transform them into compounds amenable to analysis by GC (see Chapter 8 for a detailed description of derivatisation and GC methods). The 613C value of the derivative can then be determined by GC-IRMS and the true 613C value can be obtained following a correction for the carbon added during derivatisation. Using this procedure, Engel et al. (1990) and Engel & Macko (1997) have successfully measured the carbon and nitrogen isotopic composition of C2 to C6 amino acids from the Murchison meteorite as trifluoroacetic acid-isopropyl ester derivatives. Compound specific isotope analysis of the macromolecular material represents a formidable challenge. Although by far the most abundant organic component in carbonaceous chondrites and, therefore, arguably the most important meteoritic organic fraction, its study by CSIA has only been attempted recently. This neglect is presumably due to the analytical difficulties associated with a solvent-insoluble high molecular weight organic network. In its natural state the macromolecular material will not yield combined structural and isotopic information at the molecular level by either combustion-MS or GC-IRMS methods. The intractable nature of the macromolecular material has been overcome by the use of off-line hydrous pyrolysis to liberate C7 to Cll aromatic and alkyl aromatic fragments of the macromolecular material for carbon isotopic analysis by GC-IRMS (Sephton et al., 1998). However, although the small-scale nature of these experiments reduced the conventional hydrous pyrolysis sample size requirements from several hundred grams to one or two grams (Sephton et al., 1999), this is still too demanding for many of the meteorite samples available for analysis. Recently, in an attempt to accommodate this restriction, on-line pyrolysis-GC-IRMS has been used effectively on milligram-sized samples of the Murchison meteorite, providing carbon isotopic measurements for C7 to Cll aromatic and alkyl aromatic fragments of the macromolecular material (Sephton & Gilmour, 2001).

234


11.5 CSIA of organic compounds from the Murchison meteorite 11.5.1 Carbon

Figure 11.1 shows that the C1 to C5 aliphatic hydrocarbons, amino acids, carboxylic acids and sulphonic acids from the Murchison meteorite appear to follow a common trend when their 613C values are plotted against carbon number. 613C values generally decrease as the amounts of carbon in the molecules increase. This trend has been interpreted as the result of a kinetic isotope effect during the stepwise formation of higher molecular weight compounds from simpler precursors (Yuen et al., 1984). The more reactive 12C is preferentially added during the synthesis of the carbon skeleton of these compounds. T h e C6 to C20 free and macromolecular aromatic compounds from the Murchison

meteorite also appear to follow a definite trend when their ~13C values are plotted

40 30

J

20 10

A ,,']:,

m I

r-1 El i r-1

A

0

@

,.~ I

0 A A

-10 -20

$8e

o

0 0

C'

/'5, I~

-30

00

V

X X

-40

o

0

' 5

X X

' , 10 15 carbon rtumber

9

O (3,

X X

X

, 20

Figure 11.1 - Carbon isotopic compositions plotted against carbon number for individual compounds from the Murchison meteorite. Standard deviations and references are listed in Table 1. Assignments are as follows: (A) free aliphatic hydrocarbon, (11) free amino acid, (~) free sulphonic acid, (rl) free carboxylic acid, (O) free aromatic compound, (O) macromolecular aromatic compound, (A) macromolecular oxygen-containing compound, (V) free CO, (V) macromolecular benzothiophene, (x) free normal alkane.

25

Compound Specific Isotope Analysis of the Organic Constituents in the Murchison Meteorite

235

against carbon number. This trend is expressed as an arch where C6 to Cll compounds have ~13C values that increase with additional carbon atoms while Cll to C20 compounds have ~i13Cvalues that decrease with increasing carbon number. The Cll to C20 trend is consistent with the C1 though C5 compounds from Murchison and indicates an origin by a synthetic process (Gilmour & Pillinger, 1994). The C6 to Cll trend is the opposite of that seen for the C1 to C5 compounds and, if this trend is also the result of a kinetic isotope effect, suggests that bond breaking or 'cracking' has produced these compounds (Sephton et al., 1998). During cracking the more reactive 12C bonds are preferentially broken leading to a concentration of this isotope in the products. Therefore the aromatic entities in Murchison seem to have been produced by a process in which both bond formation and destruction was significant. A further noticeable feature of the C6 to C20 aromatic compounds in Murchison is the similarity in 613C values for both free and macromolecular moieties suggesting a genetic relationship. Close inspection of the cases where 813C values are available for both free and macromolecular compounds reveals a consistent enrichment in 12C in the free compounds. This indicates that free aromatic compounds have been released from the macromolecular material in a pre-terrestrial generation event (Sephton et al., 1998). The ~13C values for the C12 to C18 n-alkanes from Murchison form a distinct group

in which none of the compounds exhibit the 13C-enrichment characteristic of indigenous extraterrestrial organic matter. In fact the ~13C values are identical to those of petroleum products or other terrestrial fossil hydrocarbons (Sephton et al., 2001). The Murchison n-alkanes also lack the distinctive increase or decrease in 813C value with carbon number common to the other compound classes. These features confirm the long-held suspicion that these molecules are contaminants from the terrestrial environment added to the meteorite following its fall to Earth (Cronin & Pizzarello, 1990; Sephton et al., 2001). The possibility that terrestrial contamination has produced another meteoritic organic feature, an apparent excess of the L-forms in amino acids, was investigated by Engel et al. (1990) who obtained carbon isotopic compositions of individual enantiomers. The similar 13C-enrichment of both D- and L-forms of alanine from Murchison implies that the L-excess is an indigenous extraterrestrial property.

11.5.2 Nitrogen The 615N values for individual amino acid enantiomers in the Murchison meteorite have been used to determine the origin of these compounds. Engel & Macko (1997) established that both L- and D-enantiomers of alanine and glutamic acid have significantly heavier ~15N values (ca. +60 %0) than their terrestrial counterparts (ca.-10 to +20 %o). On this basis, they argued that the excess of L-forms over D-forms is an extraterrestrial feature and not the result of terrestrial contamination. Furthermore, the 15N-enrichments in the C2 to C6 amino acids in the Murchison meteorite suggested an interstellar source for these compounds or their precursors (Engel & Macko, 1997). It

236


should be noted, however, that Pizzarello & Cronin (1998) questioned this work and believed that other meteoritic amino acids were coeluting with L-alanine and contributing to the L-excess and 815N determinations.

11.5.3 Sulphur As sulphur contains four stable isotopes it is possible to construct a three isotope plot on which normal mass-dependent fractionations lie on a line of 633S = 0.5 634S. Deviations from this line indicate mass-independent fractionations (expressed a s A33S = 633S - 0.5 ~)34S).Cooper et al. (1997) discovered that the sulphur isotopic composition of methyl sulphonic acid from Murchison implied a mass-independent enrichment in 33S. These isotopic features were attributed to the interstellar gas-phase UV-irradiation of symmetrical CS2 molecule leading to the production of the methyl sulphonic acid precursor.

11.5.4 Hydrogen Compound specific hydrogen isotopic measurement for amino and sulphonic acids display a substantial enrichment in D when compared to terrestrial ratios (Pizzarello et al., 1991; Cooper et al., 1997). These deuterium enrichments indicate the formation of the hydrocarbon skeleton of the amino and sulphonic acids in a low temperature environment such as is found in interstellar clouds. Furthermore, it has been proposed that the relatively constant D / H ratios of the different sulphonic acids imply that the hydrogenation of their unsaturated precursors occurred within a pool of nearly uniform D-enrichment.

11.6 History of meteoritic organic matter It is becoming widely considered that significant amounts of meteoritic organic matter, or its precursor materials, are synthesised in an interstellar environment (Cronin & Chang, 1993). It is also becoming clear that, although the feedstock for meteoritic organic matter may predate the solar system, its final molecular architecture is strongly determined by the effects of aqueous alteration on the meteorite parent body. Organic precursors synthesised in interstellar space are hydrolytically transformed to water-soluble polar organic compounds (Bunch & Chang, 1980). Portions of the macromolecular material are liberated to become soluble compounds (Sephton et al., 1998) and the chemical composition and degree of condensation of the residual macromolecular material is directly controlled by extent of aqueous processing (Sephton et al., 1999). The influence that the interstellar and asteroidal environments appear to exert on the final constitution of meteoritic organic matter has led to this model being called "the interstellar-parent body hypothesis" (Cronin & Chang, 1993). Finally, following the arrival of a meteorite at the Earth's surface, its organic assemblage may be compromised by the addition of terrestrial contaminants.

Acknowledgements The authors are grateful for the constructive reviews of Dr J.R. Cronin and Dr F. Robert.


CHAPTER 12 A New Method for the Isotopic Examination of Sub-Milligram Carbonate Samples, Using Sulphamic Acid (NH2.SO3H) at Elevated Temperatures H. Le Q. Stuart-Williams Research School of Biological Sciences, Australian National University, GPO Box 475, Canberra City, ACT 2601, Australia e-mail : [email protected]

Abstract High-concentration polyphosphoric acids used for the reaction of carbonates for isotopic analysis have equilibria which are complex and difficult to quantify, resulting in reactions with carbonates having rather uncertain isotopic values. This method of reacting sulphamic acid with carbonate was developed in an effort to find an alternative technique for producing isotopically representative CO2 from small carbonate samples without using an automated device. The method is also interesting because it is entirely anhydrous and takes place at a relatively high temperature. The fixed stoichiometry of sulphamic acid potentially results in excellent reproducibility, producing gas with an isotopic composition dependent only on the make-up of the carbonate and the reaction temperature. Finely ground carbonate material is mixed with sulphamic acid (NH2.SO3H) in evacuated borosilicate glass tubes and heated to 220~ to 235~ in an isothermal furnace. Dry CO2 is produced which can be passed directly into a mass spectrometer. Typical analytical precision for the method is 0.04%0 (813C) and 0.06%o (8180), based on a test run of 14 samples. Recommended, tested, processing parameters are 200 mm by 6 mm borosilicate tubes, sample sizes of 100-350/~g matched to 3.5 to 4.5 times as much dry sulphamic acid, reacted for 30 minutes at 220~

12.1 Introduction This paper is presented in three main parts. This, the first section, is an introduction to the aspects of carbonate reaction addressed by this study and is a brief review of the chemistry of phosphoric acid, examining why the reaction is difficult to control in small scale experiments. The second part (section 12.2) describes the reaction of sulphamic acid with carbonates and quantifies preferred reaction conditions while the third part (section 12.3) provides the experimental background for the parameter selection, based on approximately 200 analyses. Different subsets of the data are used in different figures. These subsets were selected so that as many variables as possible were well controlled to demonstrate the effect of varying one particular parameter.

238

C h a p t e r 12 - H. L e Q. S t u a r t - W i l l i a m s

T a b l e 12.1 - D a t a sets u s e d in t h e F i g u r e s 12.1-14. Figure number

Material

Mass range (~g)

R a n g e of sulphamic

Reaction temperature

L e n g t h off reaction

acid: calcite

(~

(minutes)

masses 1

BangC

52 - 4187

3.7 - 4.4

220

30

2

BangC

88 - 4187

3.7 - 4.4

220

30

3

BangC

200 - 246

3.82 - 4.28

220 - 450

30 - 36

4 5a 5b

BangC BangC BangC

160 - 356 52 - 4187 4 4 - 1281

3.18 - 4.29 3.47 - 4.36 3.56 - 5.0

220 - 450 220 235

20-60 20-60 30

6

BangC

118 - 410

2.23 - 3.94

220

5 - 60

7 8 9

BangC BangC BangC

179 - 646 179 - 646 52 - 4187

1.51 - 4.36 1.51 - 4.36 3.47 - 4.36

220 220 220

20 - 60 20 -60 20 - 60

10 11

BangC BangC

52 - 4187 52 - 4187

1.51 - 21.05 3.47 - 3.46

220 220

5 - 60 20 - 60

12a & 12b

BangC

52 - 4187

3.47 - 3.46

220

20 - 60

13

BangC

161 - 235

3.83 - 4.24

235

30

14

(3 s i e v i n g s ) Synth-0% n=l

210

4.15

235

30

Synth-25% n=3

227 - 235

3.85 - 4.17

235

30

Synth-50% n=3

232 - 236

3.92 - 4.16

235

30

Synth-75% n=2

204 - 239

4.02 - 4.19

235

30

Synth-100% n=2 BangC n=18 NBS-18 n=4

191 - 246 170 - 242 2 0 7 - 249

3.99 - 4.29 3 . 8 0 - 4.33 3.89 - 4 . 1 9

235 235 235

30 30 30

NBS-19 n=4

217 - 248

3.99 - 4.17

235

30

These data sets are summarised in Table 12.1. Attempts to analyse sub-milligram samples of carbonate manually using phosphoric acid (McCrea, 1950) often produce isotopic results which vary considerably from analyses of larger samples of the same material. Automated devices of either the common- or separate- reaction-vessel type have good analytical precision for samples as small as 50 gg or less, but are not available in all laboratories. Due to the complexity of the phosphoric acid reaction and the variability of the methods used to formulate the acid, it may be desirable to have a reaction based on a reagent supplied in the condition in which it will be used, for inter-laboratory comparisons. To limit the impact of variables other than reaction temperature, the reagent should be non-aqueous, preferably reacting only above 100~ to reduce the effects of water produced in the reaction, and inert at room temperature. If an acid reactant is used then the production of H20 is inevitable (making opportunities for isotopic exchange), by the reaction CaO + H2X CaX+ H20. Consequently, in these experiments emphasis was placed on materials that contain a radical such as SO3 which might combine rapidly with produced water SO3 + HaO ---, 2H + + SO4a-. Sulphamic acid (NHz.SO3H) was selected as it has a melting point of about 210~ and is widely available. -

A New Methodfor the Isotopic Examinationof Sub-MilligramCarbonate Samples ...

239

Off-line analyses of carbonates by reaction with phosphoric acid are typically made using versions of the method published by McCrea (1950) in which the carbonate is reacted with concentrated phosphoric acid under vacuum. Usually this process takes place at room temperature or, nowadays, somewhat warmer to make temperature regulation easier. After an extended reaction at a constant temperature the reaction vessels are then opened and the CO2/H20 mixture is passed through a cold-trap to remove water vapour before the sample is admitted to the mass spectrometer. Automated methods are generally similar in principle except that the reaction temperature is normally close to 90~ and the CO2 is commonly frozen into a liquid nitrogen cooled cold finger as it is evolved. At the high concentration of phosphoric acid used for this reaction the major component is pyrophosphoric acid (Jameson, 1959). As the first dissociation constants of pyrophosphoric acid are greater than those of orthophosphoric acid, the reaction of carbonate with these mixtures may be approximated by: CaCO3 + H4P207 ~ CaH2P207 + H20 + CO2

[12.1]

Acid for the off-line process is typically made up by mixing 85% phosphoric acid with P205 with dehydration by heating until either no more P205 can be dissolved in the heated mixture or the relative density has reached 1.9 to 1.92 (Coplen et al., 1983). The product is commonly termed 102% phosphoric acid or similar. The recipe used in the laboratory where this research was conducted (University of East Anglia (UEA), Environmental Sciences, RF. Dennis, pers. com.) is rather different from the method used by Coplen et al. (1983) and consists before heating of 1 litre of 85% H3PO4 and 1.25 kg of P205. The UEA method achieves the same density as stated above but with less boiling. The resulting acids, once they have been cooled and stood, probably have similar compositions, except for minor kinetic effects relating to polymerization of the phosphate. This is not the place for a detailed discussion of the chemistry of phosphate but it is important to emphasize that phosphoric acid mixtures are to all practical intents and purposes "black boxes". For example it is clearly not possible to have 102% orthophosphoric acid (H3PO4) and the product produced is in reality more than 70% pyrophosphoric acid (H4P207)(Jameson, 1959) plus orthophosphoric acid and minor amounts of more highly polymerized phosphates, such as H5P3010. Even '100%' H3PO4 would contain approximately 10% of pyrophosphoric acid (Jameson, 1959). The first dissociation constants for all these acids are different; for example pyrophosphoric acid is a much stronger acid than orthophosphoric acid. Re-equilibration rates subsequent to deprotonation are also probably different as a result of different rates of hydrolysis and polymerization. In short, it is difficult to be certain which acid is actually reacting with the sample in what proportions. Perhaps more importantly it is also difficult to know how long water is available for equilibration with the produced CO2, although McCrea (1950) did not consider this to be a significant problem. Acid variability is not so much of a concern in automated systems where great reproducibility of conditions is possible but it could greatly reduce precision in offiine preparation, especially when

240

Chapter 12 - H. Le Q. Stuart-Williams

very small samples are being processed. In addition to poorly constrained acid chemistry, variables such as grain size must also influence the reaction rates. Despite these apparent difficulties, some off-line methods for small samples have been published (e.g. Ball et al., 1996; Wada, 1988). Common acid-bath automated systems use elevated temperatures to speed the reequilibration of the acid between samples, to reduce the viscosity of the acid and to speed the release of trapped gas. Additionall~ freezing the evolved CO2 as it is produced reduces problems associated with solution of the CO2 and probably removes the water from the system before the acid equilibrium is substantially perturbed" H4P207 + H20 --~ 2H3PO4 is slow at room temperature. This sulphamic acid method attempts to avoid these difficulties by using fast reactions in a non-aqueous environment. 12.2 Sulphamic acid reaction of calcium carbonate

12.2.1 Equipment A tube furnace with accurate temperature regulation is required. For these experiments a furnace from the author's phosphate reaction line (Stuart-Williams & Schwarcz, 1995) was used. The furnace consists of a 5 cm diameter nickel plated copper rod about 12 cm long. The rod (which serves as a large thermal mass with a very even temperature) has a 13 mm hole bored lengthwise through it, lined with a 12 mm O.D. quartz tube into which the 6 mm borosilicate sample tubes are inserted. For phosphate reactions the furnace is heated by three windings but for the low temperatures required by the sulphamic acid reaction only the central Nichrome wire winding is used. The temperature is sensed by a "J" type thermocouple inserted into a small boring in the side of the rod. The plated copper rod assembly is mounted in a stainless steel casing and insulated (electrically) with split mica sheets and (thermally) by vermiculite. The temperature of the furnace is regulated by a Eurotherm 2116 PID controller, using a solid-state relay to switch a 15 volt current from a variable transformer. The temperature varies less than 1~ from the target 220~ or 235~ except for a brief period when a cold sample tube is first inserted. The furnace was not built for these experiments and is not ideal. For example, an aluminium block with multiple cartridge-heater inserts and borings to hold several 6 mm sample tubes would permit the processing of a number of samples simultaneously, so that standards and samples could be reacted together. Samples tubes are about 200 mm long and are made in the laboratory from 6 mm thin-wall borosilicate glass tube. A carrier tool was made for loading the powdered samples into the tips of the glass tubes without smearing powder on the sides, which would result in low yields and incomplete reactions. The tool consists of a 400 mm long, 3 mm diameter stainless steel rod with a small trough milled into it close to the end. The trough is 1.3 mm wide, about 1.5 mm deep and 10 mm long.

A New Method for the Isotopic Examination of Sub-Milligram Carbonate Samples ...

241

A small, glass tube cracker was used to break the tubes and release the gas into the VG SIRA II IRMS.

12.2.2 Chemicals Most testing was p e r f o r m e d using a UEA internal standard marble, BangC (short for Bangor Carrara). X-ray diffraction of the marble detects only calcite. The isotopic composition of the marble has been d e t e r m i n e d to be: ~13C 1.995 + 0.006%0 VPDB and 6180 -1.312 + 0.020%0 VPDB. This is largely based on approximately 25 comparisons with another UEA internal standard, UEACMST. UEACMST has been extensively c o m p a r e d (hundreds of analyses) with reference material NBS-19 and has an isotopic composition very close to 613C 1.988%o and 6180-2.044%0 VPDB. The majority of the BangC used was sieved to the size fraction 63-125 ~m but is u n w a s h e d and contains some finer powder. Other materials used are noted in Table 12.2. NBS-18 (carbonatite) and NBS-19 (limestone) were also analysed, as they are well characterised and have widely separated isotopic values. To fill in the gaps and extend the tested range, additional synthetic carbonates were prepared. Distilled water was evaporated on a hotplate to 15% of its original v o l u m e and then mixed in varying proportions with u n e v a p o r a t e d distilled water. Carbonates were p r e p a r e d by leaving dissolved s o d i u m carbonate in the p r e p a r e d waters for two days and then a d d i n g calcium chloride to precipitate calcium carbonate rapidly. The materials are n a m e d Synth-X%, where X% is the proportion of 180 enriched water p r e p a r e d by evaporation. These p r e p a r e d carbonates are extremely fine-grained and could therefore not be w a s h e d or sieved. The fineness of the material was not regarded as a problem as there w a s insufficient time for isotopic exchange before the carbonate was reacted. X-ray diffraction d e m o n strated only the presence of pure calcium carbonate. The sulphamic acid (NH2.SO3H) used is from 2 batches of reagent grade chemical m a n u f a c t u r e d by BDH Chemicals Ltd., Poole, England. The reagent is p r o v i d e d as dry, coarse crystals. Table 12. 2 - Standard and reference materials used in this study. Material

NBS-18 Synth-0% Synth-25% Synth-50% Synth-75% Synth-100% NBS-19 BangC

Sulphamic analysis

Sulphamic analysis

Phosphoric analysis

Phosphoric analysis

Number of 618OPhossamples 618Osulp (sulphamic / phosphoric)

613C

8180

~13C

6180

n

A180

- 5.22 - 6.19 - 6.61 - 6.37 - 6.39 - 6.29 1.81 1.91

- 17.03 - 4.00 - 0.57 3.10 6.92 10.56 4.09 5.12

- 5.08 - 6.34 - 6.38 - 6.35 - 6.39 - 6.27 1.91 2.00

- 13.07 - 0.80 2.88 6.58 10.29 14.15 8.04 8.92

4/ 3 1/ 3 3/2 3/ 3 2/ 2 2/ 3 4/ 3 24/5

3.96 3.20 3.45 3.48 3.37 3.59 3.95 3.80

242


12.2.3 Method

12.2.3.1 Sample Preparation The 6 mm borosilicate sample tubes are placed in batches in a tube furnace and heated to 500~ for at least 30 minutes before use. They are then covered with polythene sheeting to prevent contamination by dust and stored in racks until required. Samples of carbonate (mainly BangC, see above) are weighed out on aluminium foil boats on a Sartorius Supermicro $4 balance. Optimum masses for 200 x 6 mm tubes range from 100- 350/~g. The weighed powders are carefully tipped into the carrier described above and placed into the tips of the glass tubes, striking the carrier gently to remove adhering powder. The carrier is cleaned with de-ionized water and dried with compressed air between samples. Once loaded the tubes are kept vertical until reacted to prevent spreading of the sample. Sulphamic acid is prepared by grinding the rather coarse crystals in an agate mortar. The mortar with the ground sulphamic acid is placed in the glass drying oven at 80~ when not in use. This helps to keep the chemical dry and to prevent clumping, making it easier to weigh out small amounts. The sulphamic acid is weighed out in a similar manner to that used for the samples and placed in the tubes on top of the samples using the stainless steel carrier. The masses of the sulphamic acid 9sample powder need to be matched at a ratio of 3.5 to 4.5 91. Once the acid is added to the powder the tubes should be evacuated fairly promptly: typically groups of 4 are loaded and then put on a vacuum line and pumped down to less than 6 x 10-6 mbar. Pumping down usually takes 5 to 10 minutes, so that each group completes pumping while the next group is weighed. Preheating the samples in the tubes to 100~ to drive off water had no apparent effect on the analyses and is probably unnecessary. Once the tubes are evacuated they are cutoff and sealed using a gas torch. The tubes can then be stored but must be kept upright.

12.2.3.2 Sample reaction and analysis The furnace is preheated to 220~ Before reaction the sample and reactant are mixed by holding the tube at 45 ~ from the vertical and gently rolling and tapping it. The end of sample tube containing the carbonate and sulphamic acid is then inserted into the furnace, leaving the other half of the tube projecting. The sample is reacted for approximately 30 minutes, after which it is placed in a rack at room temperature to cool. Once cool, the sample tubes are scored and inserted into the cracker on the mass spectrometer. No removal of water from the gas is required. The tubes are then run manually on the VG SIRA II mass spectrometer or as single-sample auto-runs, using a 6 minute cold finger freeze-down. Very approximate relative yields are measured by recording the beam strength and valve configuration when the bellows were not in use. The gas was not measured manometrically for this study to avoid possible contamination but more accurate yield assessment would be desirable for routine analyses.


243

12.2.4 Results

The results presented in this section apply only to data that were produced precisely by the method stated above. The majority of the remaining data (> 180 points), associated with testing and refinement of the method, were obtained with somewhat different parameters and are presented in the discussion section that follows. The model data here include 18 samples of BangC ranging in mass from 52 ~g to 259 ~g and 2.17 mg to 4.19 mg, plus three samples of NBS-18 carbonatite ranging from 177 ~g to 210 ~g. The 2 smallest samples (52 ~g and 88 ~g) had to be run at ion-beam strengths significantly less than the optimum 5E-9 Amps (mass 44). The mean CO2 analyses and standard deviations for the 18 BangC samples are 613C 1.941 + 0.058%o and 6180 5.407 + 0.095%o, compared with the accepted values from phosphoric acid reaction of 613C 1.995 + 0.006%o VPDB and 6180 - 1.312 + 0.020%o VPDB. One sample (the most isotopically enriched) has greater than 99% probability of being an outlier (American Society for Testing Materials Tn test) and was rejected. Without this sample and the other > 1000 ~g samples the standard deviations become 0.042%o and 0.056%o for carbon and oxygen respectively. The means and standard deviations for the 3 NBS-18 samples a r e ~13C - 5.121 + 0.012%o and 6180-16.417 + 0.030%o. The isotopic composition of CO2 from NBS-18 was determined (using phosphoric acid, see below) as 613C - 5.081 + 0.016%o and 6180 - 13.074 + 0.06%o. Using these numbers, a laboratory-standardised calcite/CO2 fraction factor of 10.229%o and the previously determined values for BangC, A18OSulphamic-Phosphoric c a n be determined for NBS-18 and BangC using phosphoric and sulphamic analyses. A18OSulphamic-Phos phoric(NBS-18) - 3.596%o and A18OSulphamic-Phosphoric(BangC) - 3.510%o. From this, the approximate fractionation factor for CO2-calcite using sulphamic acid reaction at 220~ is determined as c~ - 1.00669. The analytical precision is excellent, especially for the smaller samples. One other strength of this method should be emphasised: there is every evidence that the composition of the CO2 produced by the reaction is highly reproducible. Different batches of sulphamic acid off-the-shelf produce results which are statistically identical (see below): there is no individual variability such as may occur in the mixing of phosphoric acid. The only significant variable is reaction temperature which can be reproduced very reliably. This means that not only can repeatable values be obtained from reference materials but that both phosphoric and sulphamic methods can be compared for the same material. Despite the somewhat limited variation in the parameters used for these runs, there is some spread in the data which should be examined. Two particular types of analytical error are shown in Figures 12.1 and 12.2. Figure 12.1 shows what may be a kinetic isotope effect as each of the three sub-sets of data (large, normal and tiny) shows a very approximate fit to a slope of A180 - 2A13C. The origin of this effect is uncertain but may be related to yield (see below). The samples (with the exception of the outlier: 613C 2.105%o) have ~)13C that clusters around the probable true value for BangC" 1.995%o VPDB. This is to be expected as all carbon in the sample should be

244

Chapter 12 - H. Le Q. Stuart-Williams Figure 1 2 . 1 - The correlation between 813C and 6180. Very small samples lie slightly off the main trend, while large samples show increased kinetic fractionation with this combination of parameters. Regression of samples > 100 ~tg shows the relationship : 8180 = 1.56 813C + 2.39, with R2 = 0.88. Data parameters in Tablel2.1.

released in the reaction and converted to CO2. The 2 smallest samples may be kinetically fractionated from a gas with a slightly different initial composition. It is, unfortunately, difficult to separate mass spectrometer and reaction effects from each. Figure 12.2 shows the relationship between ~180 and relative yield calculated as beam I strength divided by sample mass (Amps-mass 44 initial beam)/(gg of sample) for the sub-milligram samples. The outlier (noted above) was removed from this diagram to produce a better graphical spread of the other data. Measurements using a variable capacitance manometer show that samples of 200 ~g produce about 70% of the theoretical maximum yield of CO2. Higher relative yields tend to produce isotopically heavier gas over a small range of variation, although larger variation is shown Figure 12.2- The relationship between relative yield and 6180. Data set as for Figure 12.1 but the smallest sample (52 /2g), which had a very low yield, was removed to improve graphic spread. Relative yield is calculated as the mass 44 beam strength in amps divided by the sample mass in /2g. Samples with higher yields are typically less isotopically fractionated. Data parameters in Table 12.1.


245

below to produce the opposite effect, perhaps due to the introduction of extraneous oxygen.

12.3 Experimentation: The effects of changing reaction variables 12.3.1 Introduction The experimental section is based on the entire data set. The experimental variables are highly correlated with each other, so a summary of the relationships in this paragraph is a good introduction to the more detailed discussion that follows. Much of the variation in isotopic analyses by sulphamic reaction can be attributed to changes in yield, but yield changes with a number of other parameters such as sample size, reaction temperature, reaction time, ratio of sulphamic acid:carbonate and grain size. Most of the isotopic variation probably results from the reaction of the carbonate with the sulphamic acid but non-systematic errors include fractionation during the introduction of the samples into the mass spectrometer. The introduction of organic dust may cause outliers, such as seen in samples where substantial errors in ~13C are not mirrored in the 6180. The relatively high reaction temperature may cause problems with gas being produced from carbohydrate and hydrocarbon contaminants. At lower temperatures, increasing yield is correlated with increasing reaction temperature, rising to a peak at 235 ~ - 250~ and then decreasing at higher temperatures. If these yield effects at different temperatures are removed by attempting an oxygen isotopic correction based on the 613C of the carbonate obtained by phosphoric acid reaction and the 613C obtained by the reaction with sulphamic acid, then an approximate 6180/temperature function of-0.013%o ~ is obtained. The precision of analyses gets worse at temperatures higher than 220~ for that reason 220~ is the preferred reaction temperature despite the slightly lower yield. The reaction is relatively insensitive to increases of reaction period over 30 minutes but periods from 10 to 20 minutes show increasing yield and increasing 613C and 6180. Yield is also sensitive to the sulphamic acid 9calcite ratio - increasing amounts of sulphamic acid give higher yields but the analytical precision is worse. Larger grains (for example the coarse NBS-18 grains) result in decreased yields and depleted isotopic values. Sample size is not correlated with isotopic ratio over small ranges but very large and very small samples show changed yields and isotopic ratios. This may be related to the partial pressure of the CO2 in the reaction tube. It is possible to construct functions to correct for the yield effects but this is unnecessary if the range of reaction conditions is restricted.

12.3.2 Temperature effects Substantial temperature variation produces two particular effects" changing yield and changing isotopic fractionation. The variation of yield (calculated from mass 44 beam strength, as discussed above) with reaction temperature is shown in Figure 12.3. Temperatures below 220~ produce a negligible yield as the sulphamic acid does not fuse completely. The highest yields are found at 235 ~ - 250~ Above those temperatures the yields decrease. The reasons for this decrease are unknown but may relate to either the way in which the molten acid spreads up the tube or to the production of different chemical products. Once the acid is molten it starts to diffuse up the walls of the tube and the reaction ceases either when the carbonate has been completely reacted or when it is no longer in contact with the molten acid. At 220~ some acid

246

Chapter 12 - H. Le Q. Stuart-Williams Figure 12.3 - Relative yield (see Figure 12.2) versus reaction temperature. M a x i m u m yields w e r e obtained at 235 ~ to 250~ b u t better precision w a s achieved at 220~ The grey line is not a statistical fit a n d represents the a u t h o r ' s interpretation. Data par a m e t e r s in Table12.1.

always stays at the bottom of the tube but this is not the case at higher temperatures. The exact reactions involved are unknown: the chemical products are currently unidentified: x-ray diffraction of the glassy, water soluble residues shows only the remains of the two compounds initially i n t r o d u c e d - sulphamic acid and calcite. Blanks containing just sulphamic acid produce no gases when fused at 220~ The temperature dependent fractionation is especially interesting as this is a truly anhydrous process. The results of tests made at 220~ to 450~ are shown in Figure 12.4. Three sets of points are plotted: 613C, ~180 and corrected 5180. The carbon isotopic composition should remain constant as the single carbon atom per molecule should not be temperature fractionated: any fractionation that does occur should therefore be as a result of other process associated with partial yields. The "uncorrected" 5180 includes all normal corrections to relate the gas composition to VPDB but does not include partial yield correction. The corrected points putatively have partial yield effects removed, based on the assumption that a) the carbon composition should be constant, b) that points resulting from a temperature dependent fractionation should plot close to a straight line over the temperature range tested and c) that the carbon and oxygen fractionation should be proportional. An average carbon composition for samples up to 275~ was calculated as 513C 1.93%o. The oxygen was then "corrected" according to the relationship: ~18Ocorrected = ~ 1 8 0 +

((1.93 -

~13C ) x

2)

[12.2]

assuming a mass dependent fractionation with a slope of A13C - 2 A 1 8 0 . Slopes close to 2 are indicated by other results in this study (see below). Gradients of changing fractionation factors with varying temperature are presented in Table 12.3.

247


IPtqlI 4b

2

IA~mL

>

o0

9

cX~

L)

d ~A

Ak

o all

-2

~)13C E3

cr)

O

~)180

-4

E3

9

6180 (Corrected)

-6

I 200

250

300

1-350

400

450

Reaction temperature ~ Figure 12.4 - The temperature dependence of 613C and 6180. Reaction periods were mostly 20 or 30 minutes but some samples were reacted for up to 60 minutes. The filled circles represent 6180 plots which have been approximately corrected to account for fractionation due to kinetic isotope effects: 618Ocorrected = 618Ouncorrected + 2 (1.93 - 613C). Data parameters in Table 12.1.

Table 12.3 - The relationship of temperature variation and changing fractionation factors. The data are subsets of the data presented in Figure 12.4. The data m a r k e d as 'Corrected' have been modified for mass d e p e n d e n t fractionation, based on variations in 613C as described in equation [12.2] in the text. Full temperature range (Corrected) Slope (%o / ~ 25~ intercept Fit (R2)

220 ~ to 300~ (Corrected)

220 ~ to 300~ (Uncorrected)

- 0.0104 + 0.001%o

- 0.0125 + 0.0005%0

- 0.0129 + 0.0005%0

7.34%o

7.9%o

7.93%o

0.92

0.95

0.96

248


The agreement of the corrected and uncorrected results for the temperature range from 220~ to 300~ is good and not substantially different from the corrected results for the full temperature range. The 25~ intercept is calculated only for comparison with phosphoric acid data: the sulphamic reaction does not occur below 200~ It is noteworthy that the projected intercept to 25~ is less than 2.5%o different from the carbonate/carbon dioxide fractionation for phosphoric acid at that temperature. The results suggest that a furnace designed to stay within 3~ of the target temperature would have a maximum thermally induced error of less than 0.04%0. The variation of sample gas ~180 with temperature is probably one of the main reasons that reaction at the lower temperature (220~ gives better precision than higher temperatures (235~ to 250~ even though the yields are greater for the higher temperatures. At 220~ the reaction can only occur between the melting point of the sulphamic acid (N 210~ and the target tem5.0 p e r a t u r e - a range of only A r d ~ 10~ whereas the range of L NL possible reaction tempera2.5 . L . I tures is much greater when a T:~U L L the furnace target temperar,/3 ~ L ture is higher. 0 . 0 immmiRiam.,aimm ,m.mm. aa--,,. ~.-_ m L

n

D n

~m ,.~ m[m O

C, D n~mu=u

I0

12.5- A) Reproducibility of analyses at 220~ For description of reaction time length see Figure 12.4. Sample masses are indicated in the top histogram. Two 6180 series are shown. The top series (180) shows the raw data. The bottom series (18Ob)is partially corrected for kinetically induced fractionation: ~)18Ocorrected = ~)18Ouncorrected + 2 (1.93 - ~ ) 1 3 C ) - 1. The subtraction of 1%o is to prevent superimposition of the data in the display. The data were acquired during two periods separated by several months. B) Reproducibility of analyses at 235~ from a single batch of analyses. Sample masses are indicated in the top histogram. The two series of 6180 analyses are as described in Figure 12.5A. Data parameters in Table 12.1.

Figure

b,~>

4

3 ~

9 m.m

._

~

,oooo, o~176

9 mm-m

m.m

m

,

m-,O

m

~

mI:m

.

'~

18Ob

o

CO ~

o

2

m.m. n m . m

ram-ran

luuEuh.mm I

-

mk.m.Eu

ml..mmdl I

mD411Vdl m ,.m.4ndllm

Samples (in order created)

13C


249

Figure 12.5A shows both ~)13C and 5180 for 33 samples of BangC run at 220~ over a period of about 5 months. A third trace shows "kinetically corrected" 5180 based on equation [12.2], which tends to reduce the g180 of the isotopically most enriched results. One permil was subtracted from the resulting values to separate the plots. The resulting plot is remarkably uniform for most samples, suggesting that almost all apparent isotopic variation results from kinetically related effects. There is little variation related to sample size, except that non-systematic errors increase for samples less than 100 g g mass. Figure 12.5B illustrates the isotopic data for 34 samples of BangC run at 235~ over a period of 2 weeks. The traces represent the same parameters as in Figure 12.5A. All analyses show conspicuously more variation than was found for reactions at 220~ Samples of less than 100 gg in particular show much greater variance with one conspicuous outlier in the centre of the diagram.

12.3.3 Reaction time and proportion of material reacted The effects of varying the length of the reaction at a temperature of 220~ are shown in Figure 12.6. Insufficient gas was obtained for an analysis for reaction times < 5 minutes. Reaction times of less than 20 minutes show low yields and correspondingly more depleted isotopic compositions. Essentially similar analyses are obtained for reaction times e 20 minutes although there is some scatter in the data resulting from a lack of a standardised sulphamic acid:carbonate ratio which was adopted later in the testing. The limiting factor for the length of reaction would appear to be the rate at which the molten sulphamic acid migrates up the walls of the tube away from the sample, resulting in little increased yield for longer reactions. Less than 100% yields can be produced in several ways, including incomplete reaction, incomplete recovery of "free" gas from a complete reaction and loss of gas by back reactions. In the case of grains of carbonate being reacted at their surfaces by an acid environment it seems probable that there can be little isotopic fractionation as

t )K C~

5

~

4

O

3

~

8

~

N

6180 > 613C

,/

2

i! 2 r

,t

Yield 0

o

':j

0

10

20

a ~CU

i

30

40

50

length of reaction (minutes)

60

0

Ox~ Z

Figure 12.6 - The effect of the length of reaction on 613C, 6180 and yield. The sulphamic acid: calcite mass ratio is lower and more variable than was employed in later tests. This has led to some relative yield variation as seen in the 30 and 45 minutes reactions. Yields are shown as e m p t y triangles, related to the scale on the right of the graph. The isotopic analyses are barely time dependent for periods > 20 minutes. Data parameters in Table 12.1.

250


there is no possibility of re-equilibration of the gas with the carbonate. Partially reacted material should therefore produce a gas of representative composition, but surprisingly it has been shown that this is not the case: the isotopic composition of CO2 evolved by the reaction of phosphoric acid with carbonate does change during the reaction (Fritz & Fontes, 1966; Walters et al., 1972). The identification of traces of carbonate by XRD in the residue of sulphamic reactions indicates that partial reaction of the carbonate may also produce fractionated gases in this case. If the cause of the low yields were incomplete recovery of an unfractionated gas then purely mass dependent fractionation could be expected, with A180 - 2A13C but in most cases the gas composition corresponds to A180 < 2A13C. Despite this, the general approximation to &180 - 2A13C probably indicates partial gas recovery as a major factor. Gas may be dissolved or lost as bubbles which are frozen into the melt as it cools to a rigid, impermeable glass. To test for back-reactions a group of sample tubes was treated in three different ways following heating: some tubes were held at 110~ for 30 minutes before cooling, some were allowed to cool by standing at room temperature and some were quenched rapidly in cold water. There was no obvious difference between the isotopic results. In summary, partial reaction certainly occurs and may influence the results, partial recovery of a fractionated gas may also be important but it is improbable that the composition is influenced by back-reactions. There is no significant exchange of oxygen between the sulphamic acid and the evolved CO2 or longer reactions would probably show changing 6180, as the 6180 of all the sample gases and the sulphamic acid cannot have been in equilibrium. As a further test of the possibility of exchange between the evolved CO2 and the sulphamic acid melt, some BangC samples were processed using another, very much older, bottle of the chemical from a different manufacturer. The analyses were essentially identical (6180 5.18%o at 235~ apart from a single outlier that may represent contamination of the reagent. Figure 12.7- The effect on the yield of increasing the sulphamic acid : calcite ratio in the reaction. Yields increase with increasing ratios, although other tests (Figure 12.10) show that isotopic analyses are less good at higher ratios of sulphamic acid : calcite. Data parameters in Table 12.1.

3.5 ~, • :Va

~A

3.0 2.5

N

"~

2.0

9 6

1.5

~

5.5 2 mg failed to produce any ODa measurable amount of CO2. 5.4 0 Due to these slightly ~ 5.3 unpredictable effects it is 6 advisable to control the sul5.2 I e phamic" carbonate ratio fairly 5.1 closely. Between mass ratios 1.80 1.90 2.00 2.10 of approximately 3.5" 1 and ~13C (%o V P D B ) 4.5" 1 there is no appreciable correlation with the isotopic Figure 12.9 - Correlation of i~13C t o 6 1 8 0 . Regression : 6 1 8 0 = composition of the CO2 pro1.59 613C + 2.34. R2 = 0.64. There appears to be some fractionduced, either with or without ation of the gas which is not entirely mass controlled, resulta correction for the apparent ing in a slope of < 2. Data parameters in Table12.1. .m. w

I

.m. .m.

252


kinetic isotope effect (Figure 12.11).

6.0 i

5.5

I

i_JIi

12.3.5 Sample size > There is very little rela5.0 tionship between sample size and analytical error for O 4.5 oo either 6180 (Figure 12.12A) oo or 613C (Figure 12.12B) 4.0 within the range of masses 0:1 5:1 11):1 15:1 20:1 25:1 recommended (100 - 350 Ratio of mass sulphamic acid: ~g) for use in 200 x 6 mm calcium carbonate glass tubes. Larger samples are enriched in both 13C Figure 12.10 - The effect on the 6180 of increasing the sulphamic and 180 by the KIE noted. acid:calcite ratio. Figure 12.7 shows that increasing the Sulphamic acid : calcite ratio over a narrow range improves the yield Very small samples show a (and isotopic analytical accuracy) but larger increases in the more substantial enrich- amount of sulphamic acid produce poor results. This effect may ment in 13C than 180. This result from fractionation of produced CO2 (by entrapment of may imply minor contami- gas in the acid) or from the introduction of extraneous oxygen. nation by an oxygen-free Data parameters in Table12.1. source of fractionated carbon, either in the processing or in the mass spectrometer. ,II

5.8 9 ~)180

5.7

j

O 5180

(corrected)

"~

5.6

>

5.5

0

~b

5.4

O

oo oO

C: 9

o_~~

~

5.3

OI

@

5.2 5.1

3.4

3.6

3.8

4.0

0

0

I

0

4.2

4.4

Ratio of mass sulphamic acid: calcium carbonate Figure 12.11 - The influence of minor variation in the sulphamic acid:calcite ratio on 6180. When the range of ratios is restricted to 3.47 to 4.36 there is no significant relationship between the variables: R2 = 0.00 for uncorrected 6180. Data parameters in Table 12.1.

12.3.6 Grainsize effects X-ray diffraction, as noted above, has failed to identify any of the reaction products in the resulting glassy mass. Compounds produced must therefore either be soluble in the melt or essentially amorphous, which increases the probability that the reaction is not hampered by coating of the grains with reaction products. Approximately twice as much sulphamic acid as is required for a stoichiometric relationship with the carbonate is present in the tube, so incomplete reaction of the grains is probably a consequence of the thermally induced migration of sulphamic acid away from the reac-

A New Method for the Isotopic Examination of Sub-MilligramCarbonate Samples ... The effect of sample size on 613C and 6180. Very large (> 1000 ~g) and very small (< 100 ~g) samples are often relatively enriched in 13C and 180. Data parameters in Table 12.1.

Figure 12.12.A,B -

253

:a~ 5000 4000 .m.,

8000 tion site. This effect is likely r,,o r,,rj to be more apparent with larger grains, which will take ,~ 2000 longer to react. Figure 12.13 shows analyses of three different grain sizes of BangC, 1000 coarse (125 - 180 ~m), medium ( 6 3 - 125 ~m) and @ ~~Jm edr 9 0 fine (< 63/am). The data sug1,80 1 ,,.,,., '~ 1 ,~1:1 1 ,95 '2,Ll(I 2,05 2,10 2,15 gest lower relative yields for (~13C (%0 V P D B ) the coarser sizes with slightly lower isotopic values as a consequence. The same effect B 50130 was noted with NBS-18 which has a grain size of 4000 about 250 ~m: before grinding relative yields averaged 2 mA/~g-1, rising to 2.3 mA Hg8000 1 after grinding in a mortar, o'J o'J despite the coarser material having been reacted at 235~ 2000 which usually produces a higher yield. Very high relative yields of about 2.5 mA 1000 Hg -1 were obtained from the o 9 9 9 9 ed~ q ~ j d ee'e'i 9 extremely fine grained synO 4 a thetic carbonates produced 5,1 F, 9 5 ,P, 5,4 5 ,F, 5 ,~, ~ 7 5,8 by direct precipitation 6180 (~o V P D B ) (Synth-0% to Synth-100%, Table 12.2), with close to 100% true yield. This results in a smaller ~)18OPhos - 618OSulph. Grain-sizes of 100 ~m or less are desirable to obtain good isotopic analyses and high yields, which is compatible with results from studies using phosphoric acid (Swart et al., 1991). ..

lie

.

.

.

.

.

.

".11

12.3.7 Experimentation - the overall validity of the results It has been shown that high precision analyses can be obtained using this technique by controlling variables such as the sulphamic 9carbonate ratio, but for the method to be useful the relationship between the analyses by sulphamic acid and the

254

Chapter 12 - H. Le Q. Stuart-Williams Figure 12.13 - More finely ground calcite tends to have a higher yield of gas and produces a more accurate analytical result. Coarsely ground (125 to 180/am) BangC marble has a lower yield and produces a more depleted isotopic analysis than medium (63 to 125/am) or fine (< 63 /am) sievings of the same material. Data parameters in Table 12.1.

5.4 Finb

5.3

Fin

w

q~

Fihe Coarse Mediumq}

C~ 5.2

0

O

5.1

Medium

9 5.0

O

Coarse

4.9 4.8

O

Coarse

oo

O

1.6

1.8

2.0

2.2

2.4

2.6

Yield Amp (mass 44) / ~g x 1E-n

true isotopic values must be predictable. To test this eight isotopically varied calcium carbonates, ranging from NBS-18 to an isotopically enriched synthetic carbonate prepared for this study (Table 12.2), were analysed by both conventional phosphoric acid analysis and sulphamic acid reaction. The sulphamic acid analyses for these comparisons were performed at 235~ so the oxygen isotopic values are about 0.2%o depleted relative to analyses by the preferred method at 220~ and the precision is less good than at the lower temperature. Despite the reduced precision, regression analysis of the results shows the oxygen isotopic relationship to have a slope of 1.007 and R2 = 0.999. Sulphamic analyses are on average 3.63%o depleted relative to phosphoric analyses.

15 ~

I

75%~'"

10

o

"l,

9

0

~ ~

-5

50%~" NBS-19 \9 25%4" '~/'"

Bar gC /D & ,- NBS-19

0%,,.o"/" /-,'

~

~ /BangC

8180

I-i

~~

J"'d"~

Figure 12.14- The relationship of calcite analyses using sulphamic and phosphoric acids. Carbon isotopic analyses between the two methods are very similar, with a crossplot of the two methods having a slope of unity and a zero/zero intercept. The relationship of 6180 analyses also has a slope of unity but sulphamic acid analyses are depleted by approximately 3.7 1 relative to phosphoric acid analyses at 235~ Data parameters in Table 12.1.

D 100%

,,,,."

/ / .,~ _~'NBS-18 ' II Synthetic

--813 C

-10

d"" 1

-15 -20

9 NBS-18 I

-15

-10 -5 0 813C & 8 1 8 0 (%o

5

VPDB) Sulphamic acid reaction

10

15

A New Methodfor the IsotopicExaminationof Sub-MilligramCarbonateSamples ...

255

The final data column in Table 12.2 shows that Asulphamic-phosphoricis larger for the natural carbonates than for the synthetic ones. This is almost certainly, as noted above, the result of the extremely fine grain-size of the Synth- samples and their very high yields. 12.4 Conclusions

These tests have shown that if the method described in section 12.2 is followed that isotopic analyses of 100 to 350 ~g calcite samples can be made off-line with precisions of 0.04%o (813C) and 0.06%o (6180) respectively. With good control of the reaction furnace very similar replicate analyses should be achievable in any laboratory. While loading clean tubes with weighed ingredients is somewhat time consuming, the preparation of phosphoric acid is avoided and there is no need to pass the gas through a transfer line to remove water vapour. The method could be improved by the construction of a purpose built furnace which would facilitate the reaction of numerous samples and standards simultaneously. Sulphamic analysis of carbonates is also interesting because it offers the opportunity to compare isotopic effects encountered in conventional phosphoric acid analyses with those in a high temperature anhydrous environment.

Acknowledgements

This work was conducted in the Stable Isotope Laboratory of the School of Environmental Sciences at the University of East Anglia and was funded by the laboratory. I would like to thank Paul Dennis (Head of the Laboratory), for his support and discussion, and Alex Etchells (Research Technician) for his assistance with equipment. Karen (my wife) encouraged me when things did not work! Paul Kennedy of The University of Wales, Bangor, kindly supplied the Carrara marble used for BangC. I would like to thank H. Wada and an anonymous reviewer for their help in making this paper more concise and correct.


CHAPTER 13 Determination of the Abundance and Stable Isotopic Composition of Trace Quantities of C and N in Geological Samples: The Practice and Principles of Stepped-Heating at High Temperature Resolution Stuart R. Boydl CRPG-CNRS, 15 rue Notre Dames des Pauvres, B.P. 20, 54501 Vandoeuvre-les-Nancy Cedex, France

13.1 Introduction Carbon and nitrogen can both be present within geological samples in a variety of unrelated components that may have markedly different isotopic compositions. A simple example would be a metasedimentary feldspar crystal, containing 10 ppm of biologically fixed ammoniacal nitrogen, whose surfaces were 'contaminated' by gaseous N2 and air-borne organic nitrogen. Other samples can contain several components that are all indigenous to the sample. For example, carbonaceous meteorites contain trace quantities of nano-diamonds (ppm level) co-existing with high concentrations of organic matter (% level). Clearly, bulk isotope analyses of these samples would lead to meaningless results, and for carbonaceous meteorites, the isotope signal from the nano-diamonds would be totally masked by the organic matter. In such cases, stepped heating techniques (+ various pre-treatments) can be used to release C or N from the various components in a sequential manner, allowing the content and isotope composition of each element to be determined separately. Samples are heated to progressively higher temperatures in a step-wise manner (e.g., 100~ 150 ~ 200~ etc.), either under vacuum ('pyrolysis') or in an oxygen atmosphere ('combustion'). The duration of each temperature step is usually between 30 and 40 min. At the end of each step, the resulting gases are collected and purified, prior to quantification of the element of interest, and the determination of its isotopic composition. For more details on historical aspects of the technique, see Wright & Pillinger (1989) and Robert & Halbout (1990) and references therein. Components which decompose or oxidise at different temperatures will release their gases over different temperature intervals, allowing for their separation. Components occurring in trace (ppm) quantities can be concentrated by chemical/physical pre-treatments prior to the determination of the abundance and isotopic composition the of elements present. For example, Russell et al. (1991) and Hough et al. (1995) 1. Deceased in November 2001.

Determination of the Abundanceand Stable Isotopic Compositionof Trace Quantities ...

257

extracted fine-grained diamonds from whole-rock samples by a series of acid treatments. During these procedures, the components risk becoming contaminated by organic matter, either related to the chemical reagents themselves, or present within the atmosphere. Such contamination, usually rich in elements such as C and N, would interfere with any attempt at a bulk analysis. However, since the contaminants usually oxidise at low temperatures (< 500~ they can be removed during the early stages of a stepped-heating experiment (or by a precombustion; Ash et al., 1990), to leave the component of interest in a pristine state, from which the gases can be liberated at higher temperatures. For most studies, stepped-heating has only been used to isolate different components and measure their isotopic compositions. However, in the last few years, it has become apparent that stepped-heating techniques have other applications. For example, they can be used to characterise the thermal decomposition of materials, to aid in the development of new techniques (Boyd et al., 1993a), although what is perhaps of more general interest is that, by increasing the number of temperature steps (i.e. decreasing the size of each temperature increment), the shape of the release profile can aid in the identification of the component in which the element is present. This feature is advantageous when only a small amount (< 1 mg) of material is available for study, as is often the case for extra-terrestrial samples. However, by increasing the number of temperature steps, the length of the experiment may become prohibitively long for manually operated systems, hence the possibility of automation of stepped-combustion is being explored (Grady et al., 1996; Verchovsky et al., 1997). Automation will also allow for strict control of operating conditions which is imperative for subse-

Figure 13.1 - Apparatus used for high resolution stepped-combustion mass spectrometry of carbon-bearing samples: see text. E, extraction; P, purification; Q, quantification; M, mass spectrometry; HV, high vacuum; CM, capacitance manometer; MS, mass spectrometer; A-D, valves; SF, sample furnace; LNT, liquid nitrogen trap; CuO-Cu20, copper oxide furnace; VCT, variable temperature cryogenic trap.

25 8

Chapter 13 - S.R. Boyd

quent kinetic analysis of the data. 13.2 Technical aspects Figure 13.1 shows the basic apparatus for a system used for the stepped-combustion of carbonaceous samples (Boyd et al., 1998). A system for nitrogen has been described by Boyd et al. (1988). For both carbon and nitrogen, the systems that are used have four aspects in common: extraction (E), purification (P), quantification (Q) and finally mass spectrometry (M), to determine the isotopic composition of the element.

13.2.1 Extraction (E) Extraction involves the heating of the samples in a step-wise manner in regular temperature intervals (e.g. 50~ resulting in steps of 250~ 300~ 350~ etc.) for a fixed duration of 30 minutes, for example. It is important to adhere strictly to the chosen duration, otherwise a subsequent kinetic analysis of the data would be impossible. Samples are either pyrolysed under vacuum, or combusted in an oxygen atmosphere, the latter being provided by an on-line CuO-Cu20 furnace (Figure 13.1), in which wire-form copper oxide is wrapped in platinum foil and partially depleted in oxygen by pumping on the furnace with the temperature at around 900~ (see Boyd et al., 1988, 1995). Such a furnace, when operated at 850 to 920~ can provide oxygen pressures of between 10 and 90 torr, although the oxygen pressure needs to be verified periodically, since the oxygen content becomes depleted with use. At the end of the period of combustion, excess oxygen can be resorbed by changing the copper oxide temperature to 450~ A liquid nitrogen trap (LNT#1;-196~ is also present within the extraction section to remove condensible species such as CO2, which would otherwise build up in the gas phase and inhibit further reaction between the sample and the oxygen. Such a trap is also necessary, in the case of the pyrolysis of carbonates, to minimise 'backreaction' between evolved CO2 and the residual metal oxide. If it is necessary to separate trace quantities of NOx from a large amount of CO2, then a liquid nitrogen trap can be replaced by a CaO-Cu furnace (Boyd et al., 1994; 1995). When nitrogen is being extracted from samples, a CaO trap can be used to retain all condensible gases in the extraction section of the line, since there is little point in transferring them to the purification section.

13.2.2 Purification (P) For carbon, this involves the purification of CO2 for which there are numerous papers (e.g. Sakai et al., 1976; DesMarais, 1978a) and only the simplest example is covered here. After the combustion/pyrolysis period, valve A (Figure 15.1) is opened and C O 2 , SO2 and H20 are condensed onto a variable temperature cryogenic trap (VCT; Figure 13.1), held at-196~ The VCT is contained within the purification section and can

Determination of the Abundanceand StableIsotopicCompositionof Trace Quantities ...

259

operate between -196 ~ and about +150~ the latter being used for overnight degassing. The VCT is essentially a tube of glass surrounded by a heating element, the whole unit being immersed in liquid nitrogen. With no current flowing through the element, the trap is at -196~ and will condense CO2, SO2 and H20. After condensation of the gases, valve A can be closed, the next temperature step started, and valve D (Figure 15.1) opened to pump non-condensible gases such as N2. After the non-condensible gases have been removed, valve D can be closed and the temperature of the VCT raised to -130~ allowing CO2 to expand into the purification section, whilst retaining SO2 and H20; see also Miller & Pillinger (1997). A system for nitrogen purification has been described recently by Boyd et al. (1995).

13.2.3 Quantification (Q) and Mass Spectrometry (M) After extraction of the gas and its purification, valve B (Figure 13.1) can be opened and the gaseous species of interest transferred to the quantification section of the line. A gauge, such as a capacitance manometer (see Boyd et al., 1995; Boyd, 1997 for their operation and performance), is commonly used for both carbon and nitrogen although, for very small samples of nitrogen, the intensity.of the major ion beam (as recorded by the mass spectrometer) has to be used (Boyd et al., 1993b). Ideally, one would like a range of international standards that contained ppm amounts of C and N, which were homogeneous to better than + 5%, e.g. a carbon content of 250 + 10 ppm. Also it would be desirable that this result could be guaranteed for aliquots weighing about 1 mg. At present this is not the case. For example, consider the standard steel SRM - 368 which has a quoted N content of 100 + 10 ppm. In order to achieve this level of accuracy and precision, it is recommended that at least 1 g of material be used, yet the stepped-heating lines are designed to study samples having a maximum weight of typically < 50 mg. Although international standards are not yet available, Boyd & Pillinger (1991) developed a technique for producing small amounts of nitrogen standards which, for milligram amounts of material, gave errors of400 mL) is required, and significant quantities of liquid nitrogen are consumed in the process of analysis. Barring improvements or breakthroughs in the precision of GC-IRMS methods or optical absorption techniques, this method will remain a benchmark method for measuring the carbon and oxygen isotopes in atmospheric CO2. 14.3 Gas chromatography-isotope ratio mass spectrometry (GC-IRMS) measurements of atmospheric CO2 14.3.1 GC-IRMS introduction Three problems are experienced with traditional dual-inlet IRMS methods used to perform isotopic measurements of atmospheric CO2" 1) co-extraction of contaminating N20 such that post-measurement correction is required 2) the requirement for relatively large samples of CO2 (a few tenths to several retool), and 3) the requirement for coolant use.

To avoid the problems and sources of error that arise due to the inclusion of N20 in the CO2 sample during dual-inlet analysis, either a standardized correction procedure must be used by all laboratories, as recommended by Allison et al. (1995), or N20-free analysis of samples must be achieved.

Stable isotope measurements of atmospheric CO 2 and CH 4

283

To address these shortcomings in the dual inlet technique, this paper describes a GC-IRMS technique that enables N20-free, isotopic analysis of atmospheric CO2 that has been recently developed at National Institute of Water and Atmospheric Research (NIWA) in New Zealand (Ferretti et al., 2000). 14.3.2 Experimental 14.3.2.1 GC-IRMS inlet system The layout of the GC-IRMS system is shown in Figure 14.6. Aliquots of air samples are introduction into an evacuated sample loop and pressure-equilibration-volume (PEV) and are described in section 14.3.2.3. A 10-port valve (Valco Instruments Co.

F i g u r e 1 4 . 6 - Gas chromatographmisotope ratio mass spectrometry (GC-IRMS) schematic. The 10-port valve and adjacent plumbing are enclosed in a temperature controlled box (TCB) and automatically switch between load/back flush (solid loops) and inject (dashed loops). Pressure equilibration of sample gas before injection to the GC column occurs in the pressure-equilibration-volumn (PEV). After GC separation, gas effluent flows through a Nation drier and enters the open split. The Nation drier is cooled by a thermoelectric cooler (TEC) and is purged by a countercurrent flow of clean dry "zero"air. To enable gas stream switching, the sample and purge helium capillaries alternately move into the immediate proximity of the tip of the transfer capillary, which is permanently positioned at the bottom of the open split. The open split has two positions: (1) "open split in," (sample capillary extended and helium capillary retracted, as shown in Figure 14.1), where the Finnigan MAT 252 IRMS receives undiluted GC sample effluent and, (2) "open split out" (sample capillary retracted and helium capillary extended), where the Finnigan MAT 252 IRMS, receives pure helium. The shutoff valve (SOV) is closed to enable operation of the IRMS in dual-inlet mode.

284

Chapter 14 - B.H. Vaughn, J. Miller, D.F. Ferretti & J.W.C. White

Inc., Houston, Texas) is used to direct sample air into the GC, an HP5890 Series II (Hewlett Packard, Avondale, Pennsylvania, USA) where gas separation is performed with a packed column described in section 14.3.3.1. Switching between sample/reference gases and vacuum on/off is performed with pneumatic valves (Nupro) controlled by LabView (National Instruments, Austin, Texas, USA). After GC separation, sample gas enters the open split interface via a Nation MD Series Gas Drier (Perma Pure Inc., Toms River, New Jersey,) that is discussed further in section 14.3.3.2.1. Two independent air-actuated pistons switch between the sample capillary (deactivated quartz glass, 0.32 mm ID flow rate 3.0 mL/min) and the helium purge gas capillary (deactivated quartz glass, 0.32 mm ID flow rate 5 mL/min). Transfer of sample gas or purge helium from the open split to the IRMS is through a deactivated glass capillary, 0.11 mm ID, 1 m long, (SGE, International Pty Ltd., Ringwood, Victoria, Australia) at a flow rate of 0.3 mL/min. Obtained precision was found to be critically dependent on the open split design, which is further discussed in section 14.3.3.2.2. Isotopic analysis of CO2 is performed by a Finnigan MAT 252 IRMS with an accelerating potential of 10 kV. This IRMS utilizes a M u l t i e l e m e n t - Multicollector (MEMCO) system with Faraday cups to simultaneously measure m / e 44, 45, and 46 ion currents, respectively. The ion currents for the m / e 44, 45, and 46 beams are integrated at 4 Hz as the Gaussian peaks elute from gas chromatographic column and the Finnigan software calculates ~)13C and 6180 values. The IRMS was tuned for a compromise between maximum sensitivity ( 40 ~M (Voerkelius, 1990; Durka et al., 1994). As detailed in the previous section, distillation and diffusion methods attempt to remove ammonium and DON from the solutions prior to distillation/diffusion, and provide further purification by the selective uptake of ammonium with an acidic solution, disk, or cation resin. Ion exchange methods try to eliminate non-nitrate species by adsorption onto resins. Procedures that isolate only the nitrate include the chemical reaction method of Johnston et al. (1999) and the microbial denitrification methods of Sigman et al. (2001) and Casciotti et al. (2002).

15.4.6.1 Evaporating NO3- samples Methods suited for samples with variable inorganic-N concentrations The earliest methods for concentrating sample waters for 615N analysis were evaporation and freeze drying. Sample preservation by acidification, addition of mercuric chloride, or freezing is required when there is a delay between collection and preparation. The resulting dry solids include nitrate and nitrite. A popular modification to the original Dumas process is the addition of CaO to the reagents to eliminate CO2 and water vapor so that the resulting gas can be admitted directly into a mass spectrometer (Kendall & Grim, 1990; Fiedler & Proksch, 1972). The limitations to simply drying water samples and combusting the residues are that 1) freeze-drying of large water samples requires considerable time and, at some volume, becomes impractical, 2) samples with high concentrations of dissolved solids produce large quantities of dried salts which are cumbersome to load into combustion tubes, 3) the quartz or Vycor combustion tubes often fail because of reaction of alkali metals, particularly Na, and 4) dissolved organic nitrogen (DON) is included in the 615N analysis. The nitrate procedure is essentially the same as that for ammonium except that the nitrate must first be converted to ammonium by use of Devarda's alloy (Bremner & Keeney, 1965; Keeney & Nelson, 1982; Mulvaney, 1996). Prior to nitrate reduction, native ammonium is purged from the sample by raising the pH with the addition of MgO, which converts sample ammonium to ammonia gas. Steam is then bubbled through the sample to remove the ammonia. If necessary, nitrite (NO2) may also be eliminated by addition of sulfamic acid. Following the removal of ammonium, nitrate is quantitatively reduced to ammonium by the addition of finely divided Devarda's alloy and the ammonium is distilled and trapped as previously described. The Devarda's alloy may be prepared by ball-milling until it passes a 100-mesh screen and 75% passes a 300-mesh screen (Bremner & Keeney, 1965); however, Mulvaney (1996) mentions that a satisfactory product is produced by Merck (Darmstadt, Germany). Various degrees of N contamination associated with Devarda's alloy have been reported in the literature in reference to diffusion techniques (Liu & Mulvaney, 1992; Stark & Hart, 1996; Sigman et al., 1997; Goerges & Dittert, 1998; Johnston et al., 1999). Therefore, reagent N contamination should be considered, particularly for samples of low NO3- concentration.

Preparation and Analysis of Nitrogen-bearing Compounds in Water for Stable Isotope Ratio ...

331

15.4.6.2 Ion exchange resins for NO3The use of anion columns to collect and concentrate nitrate was an important development, allowing sample concentration to occur in the field. This meant that samples could be collected from access-limited sites. It also made possible the sequestration of nitrate from fresh waters with extremely low nitrate concentrations, typical of snow and rain. Two anion exchange methods that were widely used in the late 1990's were those of Chang et al. (1999) and Silva et al. (2000). Both methods allow for analysis of both 615N and 6180 and nitrate is converted to AgNO3. The methods differ in that the Silva et al. (2000) method is suited for samples in which NO3- concentrations are relatively high, requiring only a few liters to obtain N 100 ~moles of nitrate and in which the total organic load is < 30 mg. In contrast, the Chang et al. (1999) method was designed to obtain samples from snow and streams in forested watersheds. In these samples, NO3- concentrations are low, (down to 0.7 ~M), larger volumes of water are required (up to 70 L), and the cumulative load of DOC may approach 30 mg of more. Both methods use anion exchange resin (BioRad AGIX8 or AG2X8). Water samples are first filtered through a 0.45 gm filter. An apparatus of flexible design is used to gravity-drip sample water through the columns at up to 2 L/hr. A resin column is suspended on tubing far enough below an appropriately sized vessel for the head pressure to force the sample through at a proper rate. Alternatively, a pump may be used to establish flow through the column. Before passing the sample through the anion column, it is important to calculate the volume of water that can be passed through the column to obtain a minimum of 100 gmoles of NO3-, without exceeding the exchange capacity of the column. This is important because if the exchange capacity is exceeded, NO3- will bleed out of the column with the isotopically heavy NO3- exiting first, resulting in fractionation (Silva et al., 2000). The volume of sample that can be passed through the exchange column is a function of the total exchange capacity of the column and the total load of anions in the sample. The exchange resin has an exchange capacity of 1.2 meq/L, therefore 2 mL columns can retain 2.3 meq (Silva et al., 2000) and 5 mL columns can retain 6 meq (Chang et al. 1999). Silva et al. (2000) described the approximate capacities and limits of small, prefilled columns containing 2 mL of resin, in terms of nitrate, chloride, sulfate, and dissolved organic matter (DOM) concentrations. The modifications made by Chang et al. (1999) were made to accommodate the large sample volume (10 to 70 Ls) and to prevent DOM accumulation on the anion column. DOM can occupy sites otherwise available for nitrate, can impede flow through the column, and can contaminate the 6180 portion of the sample if ~18ODOM a r e transferred. In the Chang et al. (1999) method, the sample is filtered through a filter capsule (Gelman, with pore size 0.45 gm) with a large surface area (600 cm2). Placement of a cation exchange (AG 50 W; mesh size 100 to 200) column in front of the anion

332

Chapter 15 - C.C.Y. Chang, S.R. Silva, C. Kendall, G. Michalski, K.L. Casciotti & S. Wankel

exchange (AG-2X mesh size 100 to 200) column minimized clogging of the anion column by DOC accumulation and served to minimize transfer of unwanted 618ODOC to the sample. The cation column absorbs cationic DOM and may protonate some anionic DOM thereby allowing it to pass through the anion column as a neutral species. Other modifications made to accommodate the large sample volume were to increase the mesh size of the exchange resin to 100 to 200 (the Silva et al. (2000) method used 200 to 400 mesh) and to use an anion resin (AG-2X) that was less aggressive in adsorbing DON (the Silva method used AG 1X). This increased the sample flow and decreased the possibility that DON would adsorb to the anion column.

Elution of nitrate from the anion column After samples are concentrated on columns, the columns are chilled and returned to the lab for processing. Nitrate is eluted from the columns with several successive aliquots of 3 M hydrochloric acid (HC1). The nitrate-bearing acidic solution is neutralized with silver oxide (Ag20), which precipitates chloride in the form of silver chloride (AgC1) and leaves the nitrate in solution as silver nitrate (AgNO3). The solid AgC1 and excess Ag20 are filtered (Whatman no.5) from the solution. If 6180 analysis is desired, then all oxygen bearing species other than nitrate must be removed. To remove barium sulfates and barium phosphates a saturated solution of barium chloride is added to precipitate sulfates and phosphate. The sample is refrigerated overnight to allow the precipitate to form. The next day, the sample is filtered to remove the precipitate, the sample is put through a cation column to remove excess barium and neutralized with a small amount of Ag20. After neutralization, the sample is filtered to remove the solid silver, after which the sample can be split for separate 615N and ~180 analysis. The 615N portion is freeze-dried and rehydrated in progressively smaller volumes of water as necessary to transfer the sample to either a quartz combustion tube for sealed tube combustion (Kendall & Grim, 1990) or to silver capsules for EA-IRMS (see below). In addition to automated analysis, EA-IRMS reduces the size of sample and therefore the volume of water that must be processed. For EAIRMS, silver capsules are used rather than the more common tin capsules because AgNO3 reacts with tin. A small amount of table sugar (about 2 mg) is added to the sample to aid in combustion for 615N analysis (Silva et al., 2000); however, Stickrod & Marshall (2000) show good precisions without sugar.

15.4.7 Hybrid approaches for nitrate The methods of Garten (1992) and Downs et al. (1999) combine ion exchange methods and distillation or diffusion. Garten (1992) used 2.5 g of anion exchange resin (Dowex 1-X8, 50 - 100 mesh) to sequester nitrate from rain and throughfall samples. The nitrate was eluted with 40 mL 2 M KCL. The method was similar to that previously described for NH4 + (see section 15.4.5) except that Devarda's alloy was added to the eluant directly prior to distillation. Note that neither of these methods recommend eluting the nitrate off the column with sequential aliquots of eluant as do Chang et al. (1999) or Silva et al. (2000). The reader is referred to section 15.4.5 for a description of the Downs et al. (1999) method, which marries exchange resins and diffusion to NO3- and NH4 + for 615N analysis.


333

15.5 Analysis of nitrate for ~180 There are currently three methods for obtaining 6 1 8 0 from nitrate: off-line combustion, in-line combustion and the denitrifier method. Amberger & Schmidt (1987) developed the first reliable off-line combustion method which has since been refined (Voerkelius, 1990; B6ttcher et al., 1990; R6v6sz et al., 1997; Silva et al., 2000). In this method nitrate (10's of/~moles of KNO3 or AgNO3) is combusted at 520 to 850~ with different C-bearing reducing agents in quartz or Vycor tubes to produce CO2. The CO2 is cryogenically transferred to sample tubes for ~180 analysis. In the on-line combustion method (Kornexl et al., 1999a) only a few gmoles (~ 3) of nitrate is required. Combustion occurs in a reaction tube at ~ 1400~ and a continuous flow of He transports the gas to a GC column where CO is separated from other gases (N2 is also produced). In the denitrifier method (Sigman et al., 2001) NO3- is microbially converted to N20 and as little as 20 nmoles of N can be analyzed. The gas passes through a cryogenic trap, into a gas chromatograph for separation and is analyzed by continuous flow spectrophotometery. The three approaches for obtaining ~180 from nitrate are described in more detail, below.

Off- line combustion The method developed by Amberger & Schmidt (1987) does not concentrate NO3and is suitable only for samples with high NO3- concentrations, as described in section 15.4.2. This method used activated carbon to remove DOM from the sample. While this may be acceptable for high concentration samples, the use of activated carbon for samples with low nitrate concentrations is not advised. Chang et al. (1999) determined that the amount of DOM removed from activated carbon is not only highly variable, but that nitrate is also indiscriminately removed. These authors also observed that nitrate was removed from waters in which the nitrate concentration was 0.7 gM (unpublished data). Amberger & Schmidt (1987) passed sample through a cation exchange resin column (Amberlite IR 129 in the H + form) then neutralized the sample with KOH. Sulfate and phosphate were precipitated by addition of an excess of BaC12 and filtered out of solution. The remaining solution was evaporated to dryness. An aliquot of the resulting solid KNO3/BaC12 mixture was combusted in a closed tube with mercuric cyanide (HgCN) to form CO2, which was then extracted cryogenically and analyzed for 6180. The drawbacks of this method are the toxicity of HgCN and low yields of CO2. Wassenaar (1995) found that silver cyanide gave more complete yields of CO2 and better precision than HgCN. Several other methods have been used for the preparation of nitrate for 6180 analysis. R6v6sz et al. (1997) combusted nitrate in the form of potassium nitrate (KNO3) with catalyzed graphite (C + P d + Au) in a sealed tube to form CO2, N2, and K2CO3. The fractionation caused by incomplete yields of CO2 is accounted for using the fractionation factor between the CO2 and K2CO3. With this method, both 615N and 6180 may be determined from the same preparation, with excellent precisions. Br~iuer & Strauch (2000) combust nitrate to CO2 in sealed tubes using guanidine hydrochloride.

334


This method is cheap, safe, and gives yields ~ 100%, comparable to the silver cyanide and graphite methods; however, ~15N cannot be measured. The graphite, silver cyanide, catalyzed graphite, and guanidine hydrochloride methods all give similar precisions, better than 0.5 %0. Br/iuer & Strauch (2000) suggest that more variation arises through collection and isolation of nitrate than through conversion to CO2, a conclusion shared by the authors of this chapter. On-line combustion

The ability to do on-line combustion for 6 1 8 0 (Kornexl et al., 1999a) was a significant advancement because continuous flow mass spectrophotometers needed much smaller samples (~ 3 to 15 ~moles) compared to dual inlet mass spectrophotometers. In this method the sample is combusted at high temperature, 1400~ a process referred to by these authors as pyrolysis, but is also referred to as Thermal Conversion with an Elemental Analyzer (TC/EA) (R6v6sz & B6hlke, 2002). The difference in nomenclature derives from the use of different manufacturers (GVI versus Finnigan MAT). One of the advantages of this method is that samples can also be analyzed for 615N and 613C, and 6180 at the same time. Both inorganic and organic substances can be analyzed, with some caveats (see below). Unlike previous methods (R6v6sz et al., 1997), smaller samples are required and the primary product is CO and N2. This eliminates the need for a second reaction step to convert sample 02 to CO2 for 6180 analysis. Kornexl et al. (1999a) initially described doing analysis on samples containing 50 to 100 ~moles, but R6v6sz & B6hlke (2002) report using less than 10 ~mole. In the high temperature combustion method, a glassy carbon tube is encased within a ceramic A1203 reaction tube. Encasement of the glassy carbon tube is necessary because sample oxygen can exchange with oxygen from the A1203 at high temperatures. The glassy tube is filled with glassy carbon grit, nickelized graphite, and nickel powder (see Kornexl et al., 1999a for details). Downstream from the tube is an ascarite trap (for CO2 removal) and GC column. The reaction furnace is heated to 1400~ and the system is flushed with He carrier gas. The standard deviation for 615N values for inorganic substances was somewhat higher than for conventional techniques (up to 0.8 %o as compared to 0.3 %o). The method appears less suited for organic substances because of the relatively small amount of N2. In general, samples with high CO/N2 ratios were less suited. For samples with C / O ratios ~ 1 ~13C values were in good agreement with conventional methods, though somewhat higher than for more traditional methods. For organic compounds such as caffeine, positional ~13C values for the oxygen bearing carbon atoms were found. Recovery for nitrates and ammonium sulphates were similar to the expected stoichiometric values. Some compounds (certain amino acids) higher CO/N2 ratios indicated that other nitrogen products beside N2 were formed. Recovery of silver phosphate and some carbonaceous materials was incomplete (90 and 70%). The incomplete recovery of the carbonaceous material indicated that only two of the three oxygen atoms were reduced to CO. The third O atom remains with the tube, presumably


335

in the form of an oxide.

Denitrifier method for nitrate ~15N and ~180 analyses The latest method for isotopic analysis of nitrate utilizes denitrifying bacteria to quantitatively convert NO3- to N20 for isotopic analysis (Sigman et al., 2001). The two bacterial strains used, Pseudomonas chlororaphis and Pseudomonas aureofaciens, lack nitrous oxide reductase and thus naturally produce N20 as the final denitrification product. The bacterial strains (obtained from the American Type Culture Collection) are initially cultured in tryptic soy broth amended with ammonium sulfate and potassium nitrate for 6 - 10 days, in which time all the initial nitrate is consumed. The bacteria are then concentrated by centrifugation, resuspended in 4 mL of the decanted supernatant, and divided into headspace vials. The sealed vials are degassed using pure N2 at ~ 15 m L / m i n for a minimum of 2 hours to remove N20 produced during initial culture growth and atmospheric 02 that would inhibit denitrification. Aqueous samples containing 10 - 20 nmoles of nitrate are then injected into the prepared vials and incubated overnight to allow for complete conversion of the NO3- to N20. The extraction and purification of N20 can be performed off-line (Sigman et al., 2001) but may also be performed on-line with automated sample extraction and isotopic analysis using a modified Finnigan GasBench or other appropriate apparatus (Casciotti et al., 2002). Mass interference by CO2 is eliminated in each case by using a chemical trap and chromatographic separation from N20, and the purified N20 passes into a continuous flow mass spectrometer. Masses 44, 45, and 46 are measured, and the 615N is calculated from the 45/44 mass ratio after applying a correction for 170 interference (assuming ~ 1 7 0 - 0.52 * ~180). The 6180 of the N 2 0 analyte is obtained from the 46/44 mass ratio. The denitrifier method can also be used for ~ 1 8 0 analysis of nitrate in seawater and fresh water samples (Casciotti et al., 2002). While mass balance for nitrogen requires that the final 615N of N20 equals the 615N of the nitrate sample, oxygen atoms are not quantitatively carried over into the N20 analyte. The calculation of 6180 of nitrate from 6180 of N20 thus involves additional considerations for the effects of fractionation during conversion and possible exchange of oxygen atoms with water during denitrification. Fractionation due to oxygen loss during conversion of nitrate to N20 is constant for a given batch of samples and therefore does not affect the 6180 of samples relative to a known nitrate standard analyzed in parallel. Water exchange contributes less than 10%, and frequently less than 3%, of the oxygen atoms in the N20 product for Pseudomonas aureofaciens. Because this exchange is constant for a given batch of samples, the analysis of appropriate standards can be used to accurately correct the measured 180/160 ratio of samples for the low levels of exchange catalyzed by P. aureofaciens. Therefore, the 6180 of nitrate can be reached by measurement of the 6180 of N20 using this method. The denitrifier method offers a number of distinct advantages over other conventional nitrate isotope analysis methods. First, it reduces the sample size by two orders of magnitude. This is helpful in watershed studies where sample volumes required

336


for conventional methods reach 10 - 20 liters, and is vital in applications where limited volumes of sample are available, such as ice core or sediment porewater analyses. Second, it eliminates problems associated with waters of high ionic strength. This is especially true for oceanic nitrate studies, where the concentration of nitrate in seawater using anion exchange resins is virtually impossible due to the high concentration of C1- that competes with nitrate for absorption on the resin. Third organic contaminants are avoided because the bacteria works exclusively on nitrate. Fourth, the preparation time is greatly reduced over other methods. Fifth, both 515N and 5180 analyses can be obtained from the same sample. There are aspects of the denitrifier method that still require further refinement at this point, namely a persistent blank (up to 5% of sample size for 10 nmol samples) and the correction of 515N for 1 7 0 in atmospherically-derived nitrate. The blank appears to be consistent among vials in any given run and is believed to arise from N20 absorbed onto the bacteria from the initial growth step. Because there is little variation in 515N of the synthetic nitrates that are used in the growth medium and because most 615N values for nitrates fall in a limited range, the blank appears to be of minimal concern and a correction can be applied. This correction becomes more significant as the sample c515N move farther away from "normal" values such as in strong denitrification zones or Antarctic nitrate aerosols (515N ~ 100 %o). The analysis of 515N in atmospherically-derived nitrate using this method is currently hampered by the large mass-independent fractionation (A170 ~ 2 0 - 30 ~ o ) that occurs in atmospheric nitrate, where 5170 = 0.52 * 6180. The assumed relationship of ~ 1 7 0 = 0.52 * 6180 used for correcting the ~15N of such samples can lead to miscalculation of their 615N values by I - 2 %o. Approaches for overcoming this limitation are currently being developed (M. Galanter Hastings, in prep.).

15.5.1 Comparisons between on-line, off-line and the microbial method for ~)180 As shown by R6v6sz & B6hlke (2002), off-line methods for 6180 of nitrate are subject to 5180 exchange between the sample and the combustion tube. Off-line combustion methods typically rely on one point calibration, and use a reference with an oxygen isotope composition close to that of atmospheric 02:5180 - 23.8 %o. However, as shown by R6v6sz & B6hlke (2002), when multiple reference samples are used, with 5180 values above and below that of atmospheric oxygen, there is an offset. The offset is more pronounced as sample size decreases, as the surface area of the tube increases, and as the 6180 of nitrate becomes heavier. The offset is negligible for samples with 6180 values between 10 to 20 %o. The isotope ratio scale factors for 50 ~mole samples were more compressed relative to 100 ~mole samples when compared to off-line techniques. This finding shows the importance of doing multiple point calibration with reference standards that have 5180 values that are below and above ~ +20 %o. The microbial method of Casciotti et al. (2002) should also give values similar to the offline combustion method because there is no contact with quartz or vycor tubes. This does not invalidate the conclusions of investigations that used off-line combustion methods, which were used by most studies published before 2003. The 6180

Preparation and Analysisof Nitrogen-bearingCompoundsin Waterfor Stable IsotopeRatio ...

337

values reported by these studies should be viewed as being operationally defined by the off-line combustion method. Since the offset is affected by sample size, and because few studies reported sample size, it would be difficult to correct the 6180 values of all off-line studies. 15.6 Analysis of nitrate for ~170 Only recently has an experimental method been developed to measure 6170 (A170) in nitrate (Michalski et al., 2002). Because of isobaric interference with 13C, standard conversion of NO3- to CO or CO2 is unsuitable for ~170 measurements and 02 is the requisite analytical gas. Conversion to 02 using well developed CO2 fluorination techniques are both time consuming, expensive and are impractical for routine analysis. Recently, the thermal decomposition of AgNO3 into 02 and NOx has been developed for both 6170 and 6180 measurements. The decomposition (AgNO3 ~ 02 + NOx (NOx - NO2 + NO)) is carried out under vacuum at 520~ in a quartz reaction tube and the NOx is cryogenically trapped as the gases evolve in two Pyrex traps at liquid nitrogen (as N203). The quartz reaction tubes are pre-cleaned in HNO3 and HF then annealed at 800~ to reduce labile O atoms on the quartz surface. AgNO3 samples are isolated and purified as described by Silva et al., (2000), then loaded into silver foil boats. The boats are slightly crimped to prevent volatilization of the silver salt at the decomposition temperature. Failure to crimp the boats can result AgNO3 vapor migrating out of the heated section of the reaction tube and re-condensing in cooler sections. This generates high ~180 values since AgNO3 with low 6180 values preferentially volatilizes, leaving AgNO3 with high 6180 values behind to decompose. The 02 generated is measured for yield, collected in a molecular sieve sample tube and transferred to the mass spectrometer for analysis. The fractionation during the partitioning of the oxygen atoms into 02 and NOx is fairly consistent (std. dev. 6180 + 1.1%0) and therefore 6180 values can be obtained by applying a correction factor after the decomposition line has been calibrated using known nitrate reference materials. Calibration is important because the 6180 values can vary widely depending on distance of the traps from the decompositions chamber and whether the NOx is directly trapped or allowed to equilibrate and therefore consistency of procedure is crucial. Since the decomposition is a kinetic process, and therefore mass dependent (e.g. 6170 = 0.52 * ~180) the fractionation leaves the A170 unaffected. This is reflected in a smaller standard deviation for A170 (+ 0.11%o) even for sample sizes as small as 5 ~mol. Sample sizes between I - 7 mg AgNO3 gave the most consistent result. The size of experimental blanks becomes relevant below ~ 5 ~mol sample sizes. The authors estimate the blank to be approximately 50 nmol due to silver oxides on the boat surface. The blank increased to 100 nmol when samples ~1 gmol are purified by ion chromatography because of impurities in the Ag20 and its solubility. The authors note they are currently exploring better purification techniques including Ag + exchange membranes to replace the Ag20 neutralization. Although this technique

338


was developed as an offiine method, it is readily adaptable to continuous flow/auto sampler technology.

15.7 Contamination by organic matter An unresolved issue in the preparation of nitrate for 6180 and (~170 analysis, and to a lesser extent 615N, is the contribution of O or N from organic matter. DOM contains roughly 30 to 50% oxygen compared to a few % of nitrogen. Silva et al. (2000) reported that most of the DOM measured in sample waters was eliminated during the collection and preparation process. About half of the remaining DOM was eliminated by shaking the sample with Norit* G-60 activated carbon. However, Durka (personal communication) and Chang et al. (1999) found that the adsorption of DOC by activated charcoal was highly variable, even within sample replicates, and sometimes incomplete. Furthermore, nitrate was also adsorbed by the charcoal, which would greatly affect samples with low nitrate concentrations. Given that different types of activated carbon have quite different behaviors for adsorption of organic compounds, the use of activated carbon to remove DOM should be done with extreme caution, if at all. Haberhauer & Blochberger (1999) described the use of poly(vinylpyrrolidone) for DOM removal and compared activated carbon and the solid phase extraction column material Isolute C18. The sample preparation procedure involved evaporating the sample to 50 mL, passing the sample through a cation resin column, adding poly(vinylpyrrolidone) (included heating and 12 hour contact time), filtering, and removal of sulfate with BaC12. The method showed excellent removal of DOM and recovery of nitrate for samples of high nitrate concentration (10 - 60 mg/L) requiring fairly small volumes. There are other methods to remove organic matter" 1) XAD-type resins have been used to adsorb organic matter (eg. Wassenaar, 1995). Their effectiveness varies greatly depending on the DOM composition. Initial cleanup procedures are rather lengthy but they do not adsorb nitrate. 2) Dialysis. 3) ultrafiltration. 4) ultra centrifugation (RNA thimbles). Johnston et al. (1999) took a slightly different approach to minimize DOM contamination. These authors developed a method that converts nitrate to 1-phenylazo-2napthol (Sudan-I) for combustion using EA-IRMS. This method was a modification of similar methods designed for 15N tracer studies (Schell, 1978; Kator et al., 1992). Sample sizes of 80 - 100 gg N per sample were required for adequate precision using EAIRMS. This is a very complicated method, requiring several chemical conversions and a final product Sudan-1 that contains one reagent-derived N and one N from nitrate. In brief, water samples are filtered, and the humic acids are removed by passage through prepacked, conditioned, reverse-phase columns (C18 resin) using a solidphase extraction manifold that holds 10 columns. Samples were adjusted to ~107 gM nitrate, buffered, and then slowly passed through columns packed with specially pretreated granular cadmium to quantitatively convert the nitrate to nitrite. The nitrite is then converted to Sudan-1 by a complicated reaction of aniline dye, naphthol, metha-

Preparation and Analysis of Nitrogen-bearingCompoundsin Water for Stable IsotopeRatio ...

339

nol, and assorted buffers. The Sudan-1 is then concentrated by passage through C18 columns, eluted with acetone, and then dried gently overnight. The Sudan-1 is then resuspended in acetone, pipetted into tin boats, freeze-dried, and analyzed by EAIRMS. This method provided accurate and precise 615N values for test samples. Spectrophotometric analysis showed complete conversion of nitrite-N to Sudan-1. About 93% of the dye formed could be put into the boats. Although the Sudan-1 contains one N from the aniline and one N from the sample, the resultant 615N is not the mean of the two values. Instead, the 615N was ~ 9 %o lower than the m e a n ~)15N, probably a result of a kinetic isotope effect involving an excess of the reagent (Johnston et al., 1999). However, the effect was consistent and by analyzing a set of samples with a range of known c515N values along with the samples, a regression line (r2 - 0.9999) was obtained that corrected the Sudan-1 615N values to provide precisions of 0.2 %0.

15.8 Sequestering dissolved organic nitrogen Dissolved organic nitrogen (DON) is usually operationally defined as the N in dissolved organic matter (DOM). Fractions of DOM rich in N include amino acids, proteins, and phenols. Because a wide variety of dissolved organic molecules contain nitrogen, exactly what is analyzed for DON-615N depends primarily on how the water sample is processed after particulate N is removed. Specific types of organic nitrogen can be converted to ammonium using modifications of the Kjeldahl method (see section 15.4.2), and then converted to N2 for ~15N analysis, using various methods. Methods for isolating the DOM for later combustion to N2 and analysis for ~15N range from the very simple freeze-drying of the sample, to the addition of several possible methods prior to freeze-drying, including removal of DIN by ion chromatography, ultrafiltration to isolate a specific molecular weight fraction followed by ion chromatography (Bronk & Glibert, 1991), roto-evaporation followed by dialysis to remove DIN (Feuerstein et al., 1997), and passage through various XAD resins to separate different fractions by their chemical properties (Aiken et al., 1979; Thurman &Malcolm, 1981). Procedures for isolating DOM that result in solid residues can be analyzed for 615N as well as 613C. Automated instruments for the analysis of DIC/DOC for 613C using a DOC analyzer connected to a mass spec are now commercially available (Gilles St. Jean, personal communication, 2003); these could be potentially be modified for 615N of DIN/ DON in the near future.

15.9 Methods for producing N2 from nitrogen-bearing species Sealed tube combustion

Solids resulting from any of the above concentration methods, including ammonium and nitrate, can be converted to N2 gas by Dumas combustion, in which the sample plus Cu and CuO are heated in an evacuated and flame-sealed tubes at various elevated temperatures (Kirshenbaum et al., 1947; Fraser & Crawford, 1963; and many others).

340


A commonly used modification to the original Dumas process involves the addition of CaO to the reagents to eliminate CO2 and water vapor (Fiedler & Proksch, 1972), with combustion at 850~ in quartz tubes followed by slow cooling, so that the resulting pure N2 gas can be admitted directly into a mass spectrometer with no further purification (Kendall & Grim, 1990). Some caveats include: 1) large fluffy samples are cumbersome to load into combustion tubes, 2) any powder remaining on the surfaces later flame sealed can weaken the glass, causing cracking of the tubes during cooling, 3) the combustion tubes often fail because of reaction of alkali metals, particularly Na, with the walls of the tubes, 4) the tubes must be allowed to cool slowly to produce pure N2 (Macko, 1981), and 5) if the CaO is not dried carefully to remove H20 and CO2, tubes often explode during combustion. Precisions better than 0.1%o are common (see Kendall & Grim (1990) for details). Off-line sealed tube combustion methods have been automated using elemental analyzers (see below). Nitrate solids resulting from the methods above can also be analyzed for ~ 1 8 0 using silver cyanide (Amberger & Schmidt, 1987), catalyzed graphite (R6v6sz et al., 1997), guanidine hydrochloride (Br/iuer & Strauch, 2000), and graphite (Silva et al., 2000). These methods are described in the previous sections. More recently such samples are reduced to CO using a glassy carbon furnace in an automated high temperature elemental analyzer (EA-IRMS), and analyzed for 6180 (see section 15.5.3).

Rittenberg method The "classic" Rittenberg method (Rittenberg et al., 1939; Rittenberg, 1948; Sprinson & Rittenberg, 1948, 1949) for analyzing ammonium produced by distillation (before or after Kjeldahl digestion) involves conversion of ammonium to N2 by oxidation with alkaline sodium hypobromite in an evacuated vessel. The reaction is performed in a special "Rittenberg" Y-shaped tube (see section 15.4.3) that can be tilted to bring the hypobromite in one arm, in contact with the liquid sample in the other arm. The tube must be carefully degassed prior to letting the liquids react. The reaction is usually represented as 2NH3 + 3OBr- ~ 3Br- + 3H20 + N2

[15.3]

However, the reaction is not quantitative and small amounts of N20 and NO3- are formed (Bremner, 1965a). The N2 produced must be purified on a vacuum line to remove hypobromite and water vapor. This is performed by freezing the sample container (Y-tube or vial) in a low-temperature bath, and passing the N2 through a liquidnitrogen filled dewar before admission into the mass spectrometer. Although this technique has been applied extensively in the agricultural and biological sciences, it is generally regarded as a tedious and labor-intensive procedure because of the time required to prepare the unstable reagent, clean the Y-tubes, degas the solutions in the Y-tubes, control effervescence during reaction, prevent leaks, and purify the resulting N2 (Mulvaney, 1993). Many modifications of this method have been published and are reviewed in Mulvaney (1993). Some particularly notable modifications are: use of vacuum stopcocks


341

and o-ring seals, to eliminate air leaks (Ross & Martin, 1970); use of lithium hypobromite instead of sodium hypobromite, especially for dry samples, because it is more stable (Ross & Martin, 1970); and the use of disposable reaction vials (Ross & Martin, 1970; Porter & O' Deen, 1977). The Rittenberg method has been successfully automated and commercial "Automated Rittenberg Analysis mass spectrometery" (ARAMS) units have been used by soil scientists for over 20 years (McInteer & Montoya, 1981; Mulvaney et al., 1990; Mulvaney & Liu, 1991). These ARA-MS system could handle small water samples with ammonium amounts as small as 25 gg N, and could analyze several hundreds of 15N tracer samples a day by automatically adding hypobromite to each sample, purifying the gas, and analyzing it for 615N (Mulvaney, 1993).

Automated combustion and analysis using an elemental analyzer Any of the above methods for preparing N-bearing samples for 615N analysis can be automated by placing the dry products into tiny foil boats, and loading them onto an automated C / N elemental analyzer connected to a continuous flow stable isotope mass spectrometer where the samples are combusted, purified, and the N2 is analyzed for 615N. Other combustion products, i.e. CO2, water, and SO2 are either trapped or separated from the N2 by the GC column. Such instruments, originally developed for 15N tracer samples (especially soils) and sometimes referred to as Automated Nitrogen and Carbon Analysis Mass Spectrometry (ANCA-MS) units, were developed almost 20 years ago (Otsuki et al., 1983; Preston & Owens, 1983; and many others), and have been commercially available for > 15 years. For a nice review of the history of ANCA-MS, see Barrie (1991). This general category of automated combustion is more typically referred to as "Elemental Analyzer- Isotope Ratio Mass Spectrometry", or EA-IRMS. Since the procedures are discussed in detail elsewhere (Chapter 8), only a brief description will be given here. The procedure for loading samples into foil boats depends on sample type. Ground and homogenized powders (from POM or biological samples) are weighed into small foil boats, which are crimped shut. Ammonium sulfate samples on glass filters are usually folded into larger boats, which are compressed and crimped. Tin boats are usually used, but silver boats are preferred for corrosive acidic samples. Samples and reference materials are placed into the wells of a carousel, which is mounted in an autosampler on an elemental analyzer. Under computer control, the samples are purged of air by a flow of helium (He), the samples are dropped one by one into a quartz tube, typically half filled with Cr203 and silvered cobalt oxide CO3/Ag generally between 120 and 1050~ The silver acts as a trap for halogens and sulfur. Helium continuously flows through the system. A measured pulse of oxygen is timed to enter the helium stream when the sample is dropped. The tin capsules oxidize at ~1700~ adding heat to the sample combustion. Combustion of organic samples generally produces N2, NOx, SOx, CO2, and H20. He carrier gas sweeps the gases through a reduction tube filled with Cu wire at ~ 600~ where the NOx is reduced to N2 and excess 02 removed. A dessicant removes water vapor. The N2 and CO2 peaks are separated by a GC column, and the sample flows into the mass spectrometer for analysis of 615N and 613C. Precisions of 0.1 to 0.2%0 for 615N and ~13C samples are routinely achieved, for sample sizes > 1 gmole of N or C.

342

Chapter 15 - C.C.Y. Chang, S.R. Silva, C. Kendall,G. Michalski, K.L. Casciotti& S. Wankel

If the ammonium ions have been collected on a sufficiently small volume of exchange resin or zeolite, then such samples can simply be added directly into the tin capsules. If samples have been trapped in acid and dried to form a salt, then the sample vials should contain a few mg of dry ammonium sulfate or other salt as a film at the base. If quantity allows it may be possible to scrape this off with a micro-spatula, and weigh it into tin capsules directly but some analysts have found it easier to add 100 ~L of distilled water to the vial to re-dissolve the salt and then dispense it with an adjustable (up to 25 mL) micro-pipette. The 6 x 4 mm tin capsules are prepared by adding I - 2 mg of a nitrogen free cross-linked dextran such as 'Ultrodex' (Pharmacia Biotech, Uppsala, Sweden), then an appropriate aliquot of the sample solution is pipetted in. The Ultrodex forms a paste with the sample and makes subsequent handling easier. If less than 20 mL of sample is added then it is possible, with care, to crimp the capsules closed over the wet absorbent but if more liquid is added then it is dried down into the capsules in an oven at 75~ Capsules may then be crimped and run as normal. Samples prepared using the diffusion techniques of Sorensen & Jensen (1991) and Stark & Hart (1996) are enclosed in the PTFE parcels. Once removed, the paper strips were of such a size that they could be simply rolled up and placed into a 6 x 4 mm tin combustion cup, and then loaded directly onto the EA-IRMS. However, as the paper strips had been wetted with an excess of sulfuric acid, they were transferred to the tin capsules immediately prior to analysis as the acid attacks the tin. The encapsulated samples cannot be stored for more than an hour without the capsules falling apart. Combustion of these samples seems to attack and shorten the life of the quartz tube forming the oxidation furnace and this will need changing more frequently. Use of silver boats avoids corrosion but the resultant mass of melted silver will eventually clog the combustion tube. 15.10 Methods for producing CO and CO2 from nitrates and other oxygen-bearing species Many groups are currently analyzing the resultant AgNO3 (or KNO3) using automated high temperature carbon reduction systems (sometimes erroneously called "pyrolysis systems") that produce CO and N2, which are separated by passage through a GC, and analyzed for 6180 and 615N, respectively. This will undoubtedly soon become the preferred method because it provides simultaneous analysis for 6180 and 615N. The first pyrolysis method of Koziet (1997) provided for simultaneous measurement of CO for 6180 and N2 for 615N using an automated elemental analyzer. Organic samples were pyrolyzed in a reactor with nickelized graphite and vitreous (glassy) carbon. This and similar procedures developed by Farquhar et al. (1997), Kornexl et al. (1999a), and others were described earlier in this chapter. 15.11 New Frontiers Since the late nineties significant progress in analytical techniques has been made for stable isotope analysis of materials since the original developments almost half a century ago. As a result, the number of stable isotope laboratories in the world increased rapidly, and an exploding number of isotopic analyses are being reported in


343

the literature. Overall quality control of isotopic data (e.g., sample documentation and tracking, information management for sample preparation, mass-spectrometric measurements, and data reduction) is becoming an important issue. These advances have included: 1) the ability to simultaneously determine more than one element in a single analysis, 2) the ability to processe samples in the field, 3) the ability to measure smaller sample sizes, 4) to shorten the sample preparation time, and 5) the ability to differentiate between different species of organic compounds and isotopomers of N20. The challenge will be to maintain accuracy and precision as the sample size decreases. For instance, recent developments of continuous-flow mass spectrometry are making it possible to determine both 615N and 6180 values of nitrate samples simultaneously (high-temperature carbon pyrolysis method, also the denitrifier N20 method). Such techniques are particularly useful for small-size samples. With decreasing sample size, possible isotopic fractionations associated with extraction from the matrix, contaminants in the reagents, and leaks will become more significant. Simultaneous measurement of 615N, ~170 a n d ~180 of nitrate requires some means for separating the isobaric interference of 170 within the analyte gas or gases (CO2, CO or N20). When analyzing compounds of multiple elements with multiple stable isotopes (e.g. carbon and oxygen in CO2) there is a certain amount of 'overlap' in the ion beams measured by the mass spectrometer- 'isobaric interference'. For example, when measuring M / Z 45 one cannot distinguish between 130160-160 and 120160170. While for mass dependent samples the effect of this interference is easily removed by measurement of the 6180 and the application of the "Craig correction," (Craig, 1957) this is not the case for mass independent fractionations (MIF; containing anomalous 170). A few recent approaches to measuring the ~170 of nitrate have been developed. Combined thermal decomposition and electrical discharge has been used effectively (Michalski, 2002) for measuring 6170 of NO3. Simultaneous measurement of N20 generated via the denitrifier method and N2 (reduced from the N20) as a means for solving for the isotope ratios seems to lack the precision necessary for most environmental research - but could potentially be improved (Wankel et al., 2002). With the analytical advance of the denitrifier method for analysis of the 615N and 6180 of nitrate, new doors have been opened for the analysis of atmospheric and porewater nitrate, as both are typically either low in concentration and/or small in volume. Furthermore, the merging of other new techniques should allow both the ability to simultaneously m e a s u r e ~15N, ~170 and 6180 of nitrate as well as a reduction in the amount of nitrate required for analysis of all three isotope ratios. This should open the way for research into the three-isotope ratio tracer system of nitrate, especially useful in atmospheric studies and in tracing N sources and cycling. Using purified non-microbial reductases to reduce nitrate also shows some promise. Such reductases that are unaffected by the presence of oxygen offer the advantage over oxygen sensitive microbial reductases and circumvent the need to grow the

344


microbes. For instance, soluble nitrate reductace (NaR), purified from corn leaves has been used to irreversibly catalyze the reduction of nitrate to nitrite with high specificity (Patton et al., 2002 and references therein). The measurement of isotopomers of N20 (the distribution of the 15N atom within the linear N-N-O molecule) carried out with multi-collector mass spectrometry is another rapidly advancing field (Yoshida & Toyoda, 2000; Chapter 19). While this method has proven to be useful in understanding the distributions and sources of N20, multi-collector technology may spawn other techniques useful for tracing nitrate as well. Multicollector IRMS systems are becoming increasingly common and the isotopic analysis of N20 produced via the denitrifier method along with the secondary fragments produced in the ion source of the mass spectrometer (N2+ and NO +) should allow measurement of the 1515N, ~1170, and 6180 of nitrate simultaneously. Finally, because of the interest in distinguishing different species of organic-N compounds, the diffusion method has been applied to differentiate types of organicN (Mulvaney & Khan, 2001). These efforts include procedures to determine urea615N in soil extracts (Marsh et al., 2003, in press), and the recovery of ~14C and 615N from alpha-amino acids in soil hydrolysates (Mulvaney, personal communication). This is of interest because there is some indication that cornfields may be insensitive to fertilizer application if amino acid sugar concentrations in the soil are elevated (Mulvaney et a12001). Clearly, new advances in stable isotope analysis (smaller sample sizes, multiple analysis of more than one isotope, automation and the ability to differentiate between species) are providing new tools for investigating sources, patterns, and transformations (denitrification, NOx to NO3- ). The ability to analyze water samples, in conjuction with atmospheric, biological, and soil samples will aid our understanding of how biologically sensitive elements such as C, N, O, and S cycle, are transformed, and stored.

Acknowledgements We thank David Velinsky, Richard Mulvaney and Bernhard Mayer for their comments and reviews. Use of firm, brand, and trade names in this manuscript is for identification purposes only and does not constitute endorsement by the U.S. Geological Survey. Thanks also to Darren Sleep for his contribution to an earlier version of this draft.


345

Appendix 15.A1 - H o w to distill a sample a soil solution for ammonia, as described by Mulvaney (1996) Pipet about 10 - 20 mL of the NH4 + containing solution into the distillation flask. Buffer the sample to p H 9.5 with (0.2 g) MgO and commence distillation. W h e n the distillate reaches 35 mL the process is stopped and the solution can be titrated to determine the a m o u n t of NH4 + recovered. To titrate, 5 mL of H3BO3-indicator solution is a d d e d to a 50 mL beaker, m a r k e d to indicate a volume of 35 mL prior to distillation. The beaker is placed u n d e r the distillation apparatus. After distillation the sample is titration with 0.0025 M H2SO4. The endpoint is reached w h e n the color changes from green to faint pink. Boric acid indicator solution: Dissolve 400 g of reagent grade boric acid H3BO3 is added to 18 L of deionized water in a 20 L bottle. Add 400 mL of indicator solution (0.495 g of bromocresol green and 0.33 g of methyl red in 500 mL ethanol) and bring the volume to 20 L with deionized water. With continuous stirring, adjust the pH to 4.8 or 5.0) or until the solution assumes a reddish purple, by careful addition of I M NaOH or single NAOH pellets. If excess NaOH is added, lower the pH with dilute HCL.

15.A2- How to diffuse a sample with high NH4 + concentrations onto a disk as described by Khan et al. (1998) 1 - A water sample or 2 M KCL soil extract (5 - 100 mL) containing 50 to 150 ~g of NH4 + is put into a 1 pint (473 mL) m a s o n jar. For fresh water samples, add e n o u g h KCL to make a 2 or 4 M KCL solution. 2 - A d d ca. 0.2 g of MgO, swirl, then place the acidified disk(s) into the disk holder. In this method, the disk is not encased b e t w e e n Teflon envelopes, rather it is susp e n d e d above the solution. See paper for details. 3 - Incubate sample at 20~ or on a heating plate at 45 - 50~ If the sample is heated, then a 4 M solution should be used. The a m o u n t of time to diffuse a sample d e p e n d s on the sample volume. A 20 mL sample heated at 45 - 50~ will take 5 hours, whereas a 100 mL sample will take 14 hours. 4 - Dry the disks in a v a c u u m desiccator containing Drierite. 15.A3- Summary of experiments by to Mulvaney & Khan (1999) to investigate the affects of organic solutions on ammonia diffusions Solutions contained 714 ~moles (10 mg) of a m m o n i a and ~ 11.4 mg of soluble organic-C. The lowest N concentration determined by this m e t h o d was ~ 15 mg L-1 or 1.07 mM. The sample volume was 10 mL and contained ~g of N labeled as (NH4+)2 SO4, KNO3, or NaNO2. Three grams of KC1 was a d d e d to each sample to give a 4 M KC1 concentration. The solutions were put into the m a s o n jars. Each m a s o n jar contained 7 mL of the boric acid indicator solution (see section 15.A1 or Kahn et al., 1997 for instructions on h o w to prepare the solution). The solution was placed in a small petri dish, s u s p e n d e d above the sample solution by a clamp (Khan et al., 1997). The samples were heated for 1.5 hours at 45 - 50~ or allowed to stand for 18 hours at room temperature.

346


To evaluate whether there were interferences caused by DON, a concentrated solution with (10 g N L-l) containing eight ammonia acids and two amino sugars was made. Organic interference was more pronounced for large sample volumes, samples that were heated, had Devarda's alloy) and in which the concentration of KC1 was less than 4 M. Small samples requiring shorter diffusion periods, and with low surface-to volume ratios had the least carryover. Carryover was minimized in samples with 4 M KC1, an effect attributed to the ability of salt promoting liberation of NH4+-N thereby minimizing the time needed to diffuse the sample. This is because larger samples take longer to diffuse. Organic carryover also occurred in samples with Devarda's alloy, presumably because the samples were heated. These results are consistent with those of Kahn et al. (1997) and Mulvaney et al. (1997) in which alkali-labile nitrogen compounds (300 gg of N as glucosamine or glutamine) were used as test compounds. No organic interference occurred during diffusion of NH4+-N, from 10 or 20 mL samples, but organic N was detected on 50 to 100 mL samples. Organic interference was also more pronounced in distilled samples because of the sample heating than samples diffused into H3BO3. 15.A4- Recovery of 815N-NO3- in waters with low nitrate concentrations (~ 2 to 5 /IM) as described by Sigman et al. (1997) 1 - Measure out a volume of water that will give the minimum amount of N need for mass spectrometry, - 2 to 5 ~mol N. For freshwater samples, 50 g NaC1 ~ L-1 (ashed 4 hours at 450 ~ must be added to increased to increase the ionic strength to ~ 35 ppt. This prevents the filter packs from bursting because of osmosis (Holmes et. al. 1998). 2- Add 300 mg MgO per 100 mL of sample (precombusted at 650 ~ to raise the pH to 9.7. 3- Reduce the sample volume by boiling or evaporation. Heating the sample promotes breakdown of DON. Reducing the volume will minimize the time needed to diffuse the sample. Since DON hydrolysis is promoted by heat, evaporating the sample at 65~ rather than boiling may be preferable. 4- Make diffusion packets out of I cm GF/D filter disks (Whatman #1823 010) or Whatman quantitative paper no. 41. Kahn et al. (1998) recommends the latter, because it is ashless, avoiding accumulation of glass in the combustion tube, and because the paper remains flexible after drying. The disks are acidified with 20 -30 ~L of 4 N H2SO4, H3PO4, or KHSO4 (the latter is less corrosive to tin capsules that the traps are put in) and sandwiched between two 2.5 cm diameter 10 ~M pore sized Teflon membranes. The reader is referred to the paper for details on making the diffusion packets. 5- Boil or evaporate the sample to reduce the sample volume (15 - 25%) and remove dissolved ammonia. 6 - Add the diffusion packet to the sample bottle. Add Devarda's alloy (75 mg/100 mL sample). Tighten cap immediately. 7- Swirl bottle and incubate at 65~ for 4 days. 8- Remove bottles from oven and place on reciprocating shaker for up to 7 days, depending on sample volume.


347

9 - Remove diffusion packet and dip into 10% HC1, then distilled water. Place packets in a dessicator in the presence of an open container of sulfuric acid (to remove trace ammonia). Leave packets for 1 - 2 days to dry. Alternatively, the samples may be dried in an oven, but this increases the chances of a m m o n i u m loss or fractionation. 10 -To store, place packets in individual air-tight vials with sealing caps. 11 - On the day of analysis, remove the filter from the Teflon membrane and put into tin boats immediately before analysis to minimize corrosion of the tin by the sulfuric acid in the filter. 12- Correct isotopic measurements for the blank effects caused by Devarda's alloy and DON.


CHAPTER 16 815N Analyses of Ammonium-Rich Silicate Minerals by Sealed-Tube Extractions and Dual Inlet, Viscous-Flow Mass Spectrometry Gray E. Beboutl & Seth J. Sadofsky2 Department of Earth & Environmental Sciences, Lehigh University, Bethlehem, Pennsylvania 18015, U.S.A. e-mail: 1 [email protected], 2 [email protected]

Abstract

In this paper, we report progress on the development of techniques for the routine, quantitative extraction and isotopic analysis of trace amounts of nitrogen in micas and whole-rock silicate samples. Using relatively straight-forward sealed-tube extraction techniques and standard cryogenic purification methods and dynamic-vacuum, dualinlet, viscous-flow mass spectrometry, it is possible to routinely obtain highly accurate and precise 815N and N concentration data for such materials. For 815N values of samples containing > 100 ppm N, lo for 2-10 duplicate analyses typically approaches 0.1%o, and N concentration data nearly always have < 5% uncertainty. Blanks of these extraction techniques are < 0.025 ~mol N2, affording analyses of small N2 samples to 1.0 ~mole. The precision of these analyses is typically higher than that obtainable by static mass spectrometry techniques, with the latter methods commonly producing data with l a nearer 0.5%o. However, the requirement of viscous-flow conditions (in turn requiring maintenance of sufficient pressure in the inlet volume) and the geometry of the sealed-tube experiments (placing limits on the maximum sample size that can be fused without rupture of the tubes) impose a practical limit of ~ 10 ppm as the minimum N concentration that can reliably be analyzed by these techniques (for samples containing < 100 ppm N, lo ~ 0.30%o, using the same techniques). In many crustal igneous and metamorphic rocks, white-micas and biotite, both of which strongly partition N as NH4 + into their interlayer sites, contain the majority of the whole-rock N and commonly have N concentrations in the range of 10 - 1000 ppm. Despite the trace concentrations of N in most silicate rocks, the analytical precision obtainable by sealed-tube extractions and dual-inlet, viscous-flow mass spectrometery, and the degree of isotopic variation for N within individual suites, rival those for 8180 data in the same suites. This, and the particular ability of N isotope data to trace sediment-derived components in fluid-melt-rock systems and potassic metasomatic alteration in mafic and ultramafic settings, should make the N isotope system an attractive, more routine addition in field studies of crust-mantle mass transfer.

815N Analyses of Ammonium-Rich Silicate Minerals by Sealed-Tube Extractions ...

349

16.1 Introduction

Despite early pioneering work in the 1950s (Hoering, 1955; Mayne, 1957; Scalan, 1958), N isotopes have been underused in crustal petrologic studies, presumably largely due to the trace quantitites commonly present in silicate systems and the analytical difficulties inherent with the large potential for atmospheric contamination. After this early reconnaissance-style work on N isotopes in crustal silicate rocks, further development of the N system as a routine tool applied to petrologic problems appeared to have "fallen by the wayside" somewhat (with only a few exceptions; Becker & Clayton, 1977), as research on O, H, C, and S isotopes provided exciting new information in an increasingly broader range of earth-science disciplines. Over the last 30 years, N isotopes have been employed extensively in studies of biogeochemical pathways, meteorites and mantle materials (primarily diamonds), largely due to the ease of analysis in the relatively N-rich organic systems (see recent work by Williams et al., 1995; Ader et al., 1998) and the common use of high-vacuum, low-blank extraction (in some cases, stepped-heating) systems in laboratories conducting the meteoritical and mantle petrologic work (e.g., Kerridge, 1985; Boyd et al., 1987; Hashizume & Sugiura, 1995; Boyd, this volume). The common analysis of Ar isotopes in micas and feldspars by stepped heating (e.g., Wijbrans & MacDougall, 1986; Grove & Bebout, 1995), and the shared interlayer residency of Ar and NH4 + in micas, have provided the potential for some interesting intellectual synergy among Ar and N isotope geochemists, in particular, in considering release during incremental heating experiments (and implications for diffusive and other mechanisms of release; see discussions for N by Boyd et al., 1993a) and in assessing closure during cooling of higher-T rock suites (see Sadofsky & Bebout, 2000). More recent work indicates that the underexploited N isotope system shows great potential for elucidating processes such as fluid-rock interaction and ore formation, magma provenance, crystallization, and degassing, and crust-mantle cycling (see recent work by Bebout & Fogel, 1992; Boyd et al., 1993a; Mattey et al., 1994; Bebout, 1995, 1997; Boyd & Philippot, 1998; Bebout et al., 1999a, b; Sadofsky & Bebout, 2000). High concentrations of NH4 + in igneous and metamorphic minerals (Honma & Itihara, 1981; Duit et al., 1986; Hall, 1999) suggest that N may in some cases be an important fluid constituent. Nitrogen species (particularly N2) are commonly found in fluid inclusions (Kreulen & Schuiling, 1982; Kreulen et al., 1982; Bottrell et al., 1988; Darimont et al., 1988; Andersen et al., 1989, 1995). Nitrogen is a common constituent of volcanic gases (e.g., Fischer et al., 1998; Sano et al., 1998), is abundant in some magmatic/hydrothermal deposits (Junge et al., 1989; Hall, 1999; Krohn et al., 1993; Bebout et al., 1999a), and occurs in trace amounts in various mantle-derived rocks (see Javoy et al., 1986). It is an important constituent of natural gases from some sedimentary basins (Jenden et al., 1988) and deep crystalline rock environments (Sherwood et al., 1988). Previous N extraction techniques for isotopic analysis of rocks and minerals have, in general, involved wet chemical techniques (dissolution/distillation) or high-temperature combustions and fusions of samples, and have employed either dynamicvacuum/viscous-flow or static-vacuum mass spectrometry. The methods of Mayne

350

Chapter 16 - G.E. Bebout& S.J. Sadofsky

(1957; largely adopting the techniques of Rayleigh, 1939) employed long-duration (a 15 hours) heating of large samples (up to 30 grams) at 950~ in open-ended quartz tube linings in stainless steel vessels, and the elegant experimental study by Scalan (1958) tested a wide variety of extraction techniques including R. F. induction heating, resistance-furnace heating in stainless steel or Ni vessels with or without CuO, NaOH-fluxed fusions, and HF dissolutions/Kjeldahl determinations. The time-consuming Kjeldahl distillation techniques, which have continued to see some use in silicate studies (Haendel et al., 1986; Junge et al., 1989), involve numerous chemical steps and may not provide complete yields (Minagawa et al., 1984; Rigby & Batts, 1986), resulting in possible isotope fractionation and relatively poor precision (Rigby & Batts, 1986; Haendel et al., 1986). The fluxed-fusion techniques developed and applied by Zhang (1988; see description of earlier work with fluxes by Scalan, 1958), involving fusion of ~ 1.5 grams of sample using LiBO2-V205 flux mixtures, and dual-inlet, viscous-flow mass spectrometry, are relatively time-consuming but afford analyses (with reduced precision) of extremely low-N materials (e.g., basalts containing < 1 ppm N; blanks for this technique are ~ 0.01 gmoles). Scalan (1958) and Prombo & Clayton (1993) employed on-line heating in molybdenum and tungsten-wire crucibles, respectively, by R.F. induction (achieving temperatures of > 1500~ and obtained reliable N-isotope data for silicate and iron meteorite materials (Prombo & Clayton, 1993, report reproducibility of ~0.4%0 for NBS SRB 1098 steel standard). Static mass spectrometry (e.g., Boyd et al., 1993a; Boyd, this volume), in some cases coupled with lasers (for analyses of diamonds; Boyd et al., 1987) and stepped-heating experiments, has afforded analyses of extremely small N2 samples; however, the somewhat lower precision of these techniques (~ 0.5%o) relative to that obtainable by dual-inlet, viscous-flow mass spectreometry (~ 0.1-0.2%o; this study) can be problematic in studies of suites showing only several %0 overall variation. Combustion-gas chromatography-isotope ratio monitoring techniques (GC-C-IRMS; see Brand, 1995b), perhaps ultimately involving laser ablations, constitute a "next frontier" for N-isotope analyses of rock systems but have not yet been applied to analyses of silicates. Over the last nine years, at the Geophysical Laboratory (1990-1991; Carnegie Institution of Washington, Washington, D.C.) and at Lehigh University (1992 to the present), we have modified the sealed-tube combustion techniques initially used in biogeochemical studies (Macko, 1981; Minagawa et al., 1984; Rigby & Batts, 1986; Kendall & Grim, 1990; Boyd & Pillinger, 1990; Ader et al., 1998) and applied these modified techniques to the analysis of N isotopes in silicate minerals and whole-rock silicate samples. These techniques, which can be employed in practically any stable isotope laboratory (i.e., any laboratory operating off-line extraction vacuum lines and dynamic-mode, dual-inlet, viscous-flow mass spectrometers with reasonable sensitivity), with little or no additional extraction-line construction, involve heating of mineral separate and whole-rock samples to 910 ~ - 1250~ in sealed quartz tubes with CuO wire, Cu metal, with or without CaO (to remove H20 and CO2; Kendall & Grim, 1990; Bebout & Fogel, 1992; Boyd& Pillinger, 1990; Sadofsky & Bebout, 2000). The N2 gas samples we analyze are in the size range of I - 50/~moles and can easily be analyzed using the variable volume (bellows) and microvolume (cold finger) inlets on a viscous-flow mass spectrometer (in our case, on the Finnigan MAT 252 at Lehigh Uni-

815NAnalysesof Ammonium-RichSilicate Minerals by Sealed-TubeExtractions ...

351

versity). Nitrogen is known to strongly partition as NH4 + into micas (particularly biotite), in rocks containing these minerals, relative to other minerals capable of incorporating it into their structures (e.g., K-feldspar and plagioclase; see Honma & Itihara, 1981). Honma & Itihara (1981) reported that, in a rock containing biotite, muscovite, K-feldspar, and plagioclase, the muscovite contains on the average ~ 40% of the N concentration in the coexisting biotite, the K-feldspar ~ 40%, and the plagioclase (depending on Ca content) on the order of 10%. The tendency of muscovite to contain -- 40% of the N concentration in coexisting biotite has been confirmed in more recent studies by Boyd & Philippot (1998) and Sadofsky & Bebout (2000); however, the latter two authors have identified significant variation within individual suites that could be attributed to differential closure or varying effects of retrogradation reactions during cooling of higher-grade metamorphic suites. Our field-based studies (Bebout & Fogel, 1992; Bebout, 1995, 1997; Bebout et al., 1999a, b; Sadofsky & Bebout, 2000) have focussed primarily on metasedimentary suites lacking K-feldspar and in which biotite and/or muscovite are present, and on metamafic and metaultramafic suites in which K-feldspar is absent and one or both of the micas occur as metasomatic products (e.g., fuchsite in alkali-metasomatized ultramafic rocks; muscovite and biotite in K-metasomatized metabasalt; see Bebout, 1997). In such suites, N is extremely concentrated in the mica phases (> 95% in micaceous rocks lacking K-feldspar and containing minor plagioclase and/or hornblende). Our development of extraction techniques has largely focussed on the quantitative extraction of N from biotite with varying compositions and white-mica, the latter ranging from nearly endmember muscovite in higher-grade metamorphic rocks (Sadofsky & Bebout, 2000) and pegmatites to extremely celadonitic, in some cases, Cr-rich white-micas (fuchsites) in lower-grade suites (Bebout & Fogel, 1992; Bebout, 1997; Bebout et al., 1999a). 16.2 Discussion of analytical methods Clean, unweathered rock samples are crushed in a jaw crusher (or, for fine-grained samples, in a large mortar and pestle) and then (for some samples) in a disk mill to a size of 0.25 to 0.05 mm (smaller than most mica crystals in medium- to high-grade metamorphic rocks). These samples are then washed in deionized water to remove any adhered powder, and the minerals of interest are separated from one another by standard magnetic and gravitational techniques. Mechanically separable chlorite is removed from the biotite samples, however, it is possible that some chlorite is interlayered at a very fine scale and remains in the biotite samples (see Veblen & Ferry, 1983). No attempt has yet been made to separate white-mica phases (muscovite and paragonite); however, we have for the most part avoided sample suites in which paragonite is known to occur as a separate phase. Purity of mineral separates is determined by examination under a binocular microscope or by x-ray powder diffraction techniques, and only mica samples of greater than 95% purity are analyzed. Some samples are run as whole-rock powders, specifically, relatively fine-grained metasedimentary rocks (e.g., Bebout & Fogel, 1992; Bebout et al., 1999a, b) and rocks such as Kmetasomatized ultramafic rocks (e.g., fuchsite-bearing talc schists; see Bebout, 1997) and metamafic rocks (e.g., metaconglomerate cobbles, with gabbroic protoliths with

352 pseudomorphs of white-mica after plagioclase; Bebout & Barton, 1993; Bebout, 1997) in which mica is known to be the only likely mineral host for N. Samples of 20 to 1000 mg (occasionally larger for low-N samples) are loaded into quartz tubes (9 mm O.D., 7 mm I.D., 23 cm long) with Cu and CuO (2.5 g CuO, 4 g Cu). Some samples with higher N concentrations are loaded into quartz tubes with 6 mm O.D. and containing somewhat smaller amounts of reagent. The tubes are then evacuated for two hours while being warmed intermittently with a hightemperature heat gun and sealed under vacuum. The ends of the tubes containing the samples are heated in a resistance tube furnace (Deltech Inc. Model DT-28-HT, with heating elements capable of achieving > 1500~ for thirty minutes at 1200 ~ - 1250~ in order to fully fuse the micas and extract all N. Heating of the samples at these temperatures for periods longer than 30 minutes does not result in the release of Figure 16.1 - Calculated yield (in p p m N) of N from biotite mineral separates (all from Townshend Dam, Vermont) combusted to a variety of temperatures. Samples analyzed at 910~ were heated for three hours at that temperature; samples heated to higher temperatures were heated for thirty minutes in a tube furnace and then heated to 850~ with CuCuO reagents to ensure proper speciation of all gases. Heating at the higher temperatures for longer periods produces no additional N yield. The small numbers on the figures show the 615N values for each analysis (note the summaries of these data, comparing the results for the low-T and higher-T extractions, in Tab?e 16.1).

Chapter 16 - G.E. Bebout & S.J. Sadofsky

lO-ui_-u

II

II

I IIII

I II

IIII

I II

IIII

II

E

800 ...-"

"=n,.~,.n."

Z

2.7

II

I

.P

j"

...P-'

--'---

94 . 4

4.5

'--' J8 ,l

403

VT 9G 7a Bi ot ite

203 1";J30

....

1C00

:3.4 32

,,I.n.L.

E

III

,j-"l

1 ....""

>-

I II

....-"

8~n

n i n

I IIII

4.4

..N,.I.L.

n ....

I ....

I ....

I~1 ....

,

n nn n

'~" .3

___-8

.-

n l n nn n lip

04,5

..----'4 ..'5_,, 4 . 2 , 4 . 3

9

....-'-

45

800 3"2

t~uu

/

>-

400

Z

200

:3 2

i n

VT 96~ 71:., Bi otite

1400 .,=.1,..

E

.-,.~..,. /

~3

>-

,.

,:

I ,.

in

lln

Ii

In

I,

Inn

n.

il

i in

nl

,.n

nl

,

.in

O

_.._----~, , , - - " - - - 4 . 5

1200

e4. 3

._.~'

1000

=,--

2B

8C8 600

i /

Z

400

. . . I.I . .I. .I I.I I. . . . .I . .n l. . l.l l.l . . . . I. .I.I . .i l l l

II

12f'n 0 . A .

E

81~1_-I

e.--

I|

I III

...-,f-

4

.7

III I. .I I. . .I I.I . . |

4.5,

4B,

II

9

413

I "z

,

4.6 -.

+

_B

4.5

.=.==""=

6Ell

>-

4CEI

Z

20-1

n l

4.3/"

1Cn-uO

~

n

VT 96- 7c Biotite

21313

0

:-: .3

VT , 1200~ have been determined as necessary in order " ~ 2Ci3 r 94 to achieve full release of N from higher- ._. grade metamorphic and igneous mus- ~-. covite and biotite separates (see Fig- Z 1 CO VT c ,-. 7d rn us; oov ite ures 16.1 and 16.2 and Table 16.1). This finding is consistent with the results of O_m_ 101ZlI-I 1100 1 :;130 1. ~ 0 vacuum stepped-heating studies of N release from micas (Boyd et al., 1993a; T ernp erat u re (C) see Figure 15.4 in Boyd, this volume), in which the bulk of the N release has been documented to be in the 900 ~ or 1000~ steps below the tem-perature at which sintering of the micas began (~ 1100~ Smykatz-Kloss, 1974) but significant amounts of N are retained to higher temperatures. This higher-temperature (> 1200~ procedure does not appear to be necessary for analyses of whole-rock, mica-rich samples and sam-ples containing more celadonitic and Cr-rich white micas from low-grade suites. For these materials, we have found that heating at 910 ~ - 950~ for three hours quantitatively releases the N and produces isotopic and concentration data indistinguishable from those obtained by the higher-T (1050~ - 1250~ heating methods (see discussions of yields for the 910~ experiments in Bebout & Fogel, 1992). The similarity in yield in our experiments on the samples containing more celadonitic white. . . .

.

. . . .

,

. . . .

,

. . . .

.

. . . .

.

. . . .

,

. . . .

|

. . . .

354

Chapter 16 - G.E. Bebout & S.J. Sadofsky

Table 16.1 -Demonstration of N-Isotope results discussed in text. Sample mean

815Nair (per mil) (std. dev.)

Concentration (ppm) mean (std. dev.)

SL-1 (fuchsite; see Figure 16.3) 910 (n = 5) ___1100 (n = 9)

2.24 2.32

0.13 0.23

2039 2052

45 (2%) 101 (5%)

WE-2 (fuchsite) 910 (n = 11)

1.63

0.10

2197

147 (7%)

0.15 0.18

624* 635

22 (4%) 20 (3%)

T54 (Silurian/Devonian Graptolite shale whole-rock) 910 (n = 7) 0.56** 0.14

4183"*

259 (6%)

20019 20168

715 (4%) 590 (3%)

WS-1 (metasedimentary whole-rock)f 910 (n = 40) 2.93* >1050 (n = 2) 2.83

Buddingtonite-2 910-950 (n = 29) ___1100 (n = 6)

-1.59 -1.22

0.23 0.21

Mica Separatestt- Townshend Dam, Vermont (from Sadofsky & Bebout, 2000) VT96-7a biotite 1250 (n = 6)

4.50

0.14

873

41 (5%)

VT96-7a white mica _>1200 (n = 2)

4.46

0.02

256

17 (7%)

VT96-7b biotite ___1100 (n = 6)

4.38

0.12

1134

30 (3%)

VT96-7c biotite _ 1200 (n = 2)

4.37

0.12

1286

44 (3%)

VT96-7d biotite ___1200 (n - 6)

4.54

0.05

1117

28 (3%)

VT96-7d white mica _ 1100 (n = 2)

4.25

0.07

299

12 (4%)

* K. Brauer reports +2.9 per mil mean ( l o - 0.1) and 550 p p m (lo = 30 ppm). ** K. Brauer & K. Hahne report +0.95 per mil (lo = 0.25) and 4000 p p m (lo = 300 ppm). f Combining the data from this lab with additional analyses obtained at the Geophysical Laboratory (during the period of 1990-1993). t data are presented in Figures 16.1 and 16.2.

355

815N Analyses of Ammonium-Rich Silicate Minerals by Sealed-Tube Extractions ... ;-II-II-II-I

I

i

I

-

E -,~,-9

I

i

"]

n

,n

2052 + 101 I:,F.rn I~%)

-

..- JUU

i

[" SL-1 (Fu ct]si te E,I:r~,datd)

275n-I ,-I1~

I

Io;i

Iq = L:! 9

,T...Ol-llI

.

..........................................................................

. . . . . . . .

~,)

1750

,

1.'__ql-ll-I

"1

% i,

3-hour headng

i

3 0 - r n i n u t e FiealJ rig, fol Io,~~ed h y 8 5 0 ' C : I"l_-II.,I ~1"18

1250 1 I_-II-II-I 91:10

I

I

1000

4.13

n

I

1100

I

n

"

3.5

12110

I

n

1300

I

n

'l

SL- 1 ,[R.lcil.si re_.E,t;~ldard)

., i

,=

3.0 i"

II

,::.. '-' 5

l~.,

.........................................................

2 .U

9

9 ............................ o

+2 ' 2 +

01.

'='1

II tl

...... ...o.....o......

+2.3+

.

.

.

0

9

. nC,

1.5

,0

900

i

i

i

1000

1100

1200

1300

Teml:.,erature (~C) of E.xperimer~t Figure 16.3 - Calculated yields (a) (in ppm N) and 615N values (b) for SL-1, a sample of fuchsite (Cr-rich, phengitic white-mica from the Catalina Schist; see Grove & Bebout, 1995) combusted to a variety of temperatures. Samples analyzed at 910~ were heated for three hours at that temperature; samples heated to higher temperatures were heated for thirty minutes in a tube furnace and then heated to 850~ with reagents to ensure proper speciation of all gases. The small numbers on the figures show the 815N values for the each analysis (note the summary of these data, comparing the results for the low-T and higher-T extractions, in Table 16.1).

356


micas at the lower (910~ and higher temperatures (a 1100~ is demonstrated in Table 16.1 for sample SL-1, a fuchsite-bearing sample of ultramafic schist (also see Figure 16.3), and for WS-1 (two higher-temperature extractions, at 1050~ and 1100~ a medium-grade metasedimentary sample provided by K. Brauer (University of Leipzig; see her data in a footnote in Table 16.1) and analyzed in her laboratory using the dissolution/distillation techniques employed by Haendel et al. (1986). Consistent with this observation, more celadonitic white-mica is known to release Ar at lower temperatures than those for Ar release from muscovite (M. Grove, personal communication, 1999; see discussions and Figure 5 in Wijbrans & McDougall, 1986, who documented enhanced stability of muscovitic white mica relative to celadonitic white mica during "in vacuo" incremental heating; also see Grove, 1993; Grove & Bebout, 1995). Fine-grained, whole-rock samples of micaceous metasedimentary rocks (even those containing more muscovitic white-mica) are (based on visual examination of the tubes following the extractions) believed to experience melting (in part reacting with the quartz tubes) at 910 ~ - 950~ thus promoting enhanced N release at the lower temperatures. At the Geophysical Laboratory, reaction products from the 910~ experiments for several whole-rock metasedimentary rocks (all showing obvious evidence of having undergone melting; samples 6-2-27a, 6-4-68H, and 6-3-41' with a wide range in their white-mica compositions from phengite to muscovite) were reloaded into 9 mm quartz tubes with reagents and reheated at 1000~ (90 minutes then slow cooling) to test for incomplete reaction in the initial experiments. The resulting amounts of N2 were similar to only slightly larger (all ~ 0.3 ~moles, most too small to analyze for their 615N) than those of the blanks (~ 0.1 ~moles at that time; note that the N2 samples extracted in the first experiments were all > 40 ~moles). One of the N2 samples from these reruns (for 6-3-41') barely large enough (~ 0.3 ~moles) to obtain reliable isotopic data yielded a 615N of +3.6%o similar to, but slightly lower than, the value obtained in the initial extraction (+4.0%o; see Table I in Bebout & Fogel, 1992). Interestingly, buddingtonite (an ammonium feldspar; see Voncken et al., 1993) does appear in our experiments to yield somewhat differing concentrations and ~15N values in extractions employing combustions at 910~ (for three hours) and in those employing combustions at > 1100~ (see data for Buddingtonite-2 in Table 16.1). The differences are consistent with the retention in the samples of a small amount of isotopically fractionated (likely isotopically heavy) N to temperatures exceeding 910~ as was indicated by Boyd et al. (1993a; also see Figure 4: in Boyd, this volume). In the stepped-heating runs for K-feldspar of Boyd et al. (1993a), less release of N occurred at temperatures below the melting point of the minerals (~ 1150~ and the bulk of the N was retained in the feldspar samples to melting temperatures. Above the melting temperature, N continued to be released with further heating (with all N apparently being released by ~ 1250~ as the viscosity of the feldspar melt decreased (discussion by Boyd et al., 1993a). For both the mica and the K-feldspar extractions, N release below the mineral melting temperatures is thought to be diffusive or perhaps more complex, involving a combination of diffusive and instantaneous release over a wide temperature range (M. Grove, personal communication, 1999; cf. Boyd et al., 1993a). Diffusive release, in general, favors the release of relatively light N and retention of heavy N in the residual solids - this retention to higher temperatures of an isotopically

815NAnalysesof Ammonium-RichSilicate Minerals by Sealed-TubeExtractions ...

357

heavy N component is demonstrated by the trends of N concentration and isotopic composition for micas showing varying degrees of N degassing in Figures 16.1 and 16.2. After combustion, N2 is cryogenically separated from the CO2 and H20 in the gas samples, using standard cryogenic techniques (at liquid nitrogen temperature; e.g., Macko, 1981; Bebout & Fogel, 1992; Boyd et al., 1993a). Purified samples of N2 gas are transferred by molecular sieve from a glass vacuum line to the inlet of the Finnigan Mat 252 mass spectrometer. Large N2 samples (> 7 ~moles) are transferred into the variable-volume (bellows) of the standard inlet system and run though the standard procedure. Small samples (0.5 - 2 ~moles) are frozen into a microvolume (cold-finger) inlet containing silica gel. Blanks for this technique have been reduced to < 0.025 ~moles N2 for routine large samples in the 9 mm quartz tubes and are thought to be related to static leakage during extractions (primarily on O-ring fittings; blanks are lower on the average for runs utilizing 6 mm quartz tubes and smaller, better-fitting 1/4-inch diameter O-rings) on the glass vacuum line. Variations in the isotopic composition of N are defined as: ~15N - I(15N/14N)spl - (15N/14N)std1103 (lSN/14N)st d

[16.1]

where the standard is atmospheric N2. Accuracy of our measurements of 615N is maintained by routine, repeated analyses of atmospheric N2 (615N - 0%0), two international N-isotope standards (NBS N2-ammonium sulphate with 615N = +20.41%o; USGS 32 potassium nitrate with 615N - +179.1%o), and one secondary, laboratory standard TCH-1 (ammonium sulphate from the Geophysical Laboratory with 615N = 0.15No). Low-grade, metasedimentary whole-rock, powdered sample WS-1 (with 615N - +2.9%o, provided by K. Brauer) has also been used, and 40 analyses of this sample over the last eight years, using the very different extraction lines and mass spectrometers at the Geophysical Laboratory (with the modified, double-focussing DuPont 491 mass spectrometer) and in the Lehigh University laboratory (using a Finnigan MAT 252) produced mean 615N - +2.93%0 (lcJ- 0.15; mean N content - 624 ppm, with lo of 22 ppm or 3.5%; see Table 16.1). Precision in 615N of the analyses of relatively high-N biotite and muscovitic whitemica separates (samples with >100 ppm N) ranges from lo - 0.02 to 0.14%o (n - 2 to 6; Table 16.1), but lo is < 0.35%o for analyses of lower-N samples (10 - 100 ppm N). Reproducibility of isotopic data (with lo nearly always ___0.2%o and commonly 0.1%o) and concentration data for several laboratory silicate standard materials and some higher-N micas (from an amphibolite-facies metasedimentary exposure at Townshend Dam, Vermont; Sadofsky & Bebout, 2000) are demonstrated in Table 16.1. For samples SL-1, WE-2, and WS-1, sample sizes used in the extractions varied by nearly an order of magnitude (e.g., for sample SL-1, 19-168 mg) and both the variablevolume (bellows) inlet and microvolume inlet on the mass spectrometer were used (each inlet with its own calibration for determinations of concentrations). Concentra-

358


tions are obtained by measurement of voltage on the m / z 28 peak for calibrated volumes in the mass spectrometer (either a variable-volume/bellows or microvolume inlet); voltage vs. gmoles N2 is calibrated by extractions and analyses of non-silicate standards with known N contents (usually combinations of ammonium sulphate and phenylalinine yielding calibration lines with r2 > 0.98). Uncertainties (again, expressed as lo) for N concentration data are nearly always < 5% (typical concentration data for various sample types are presented in Table 16.1). We have also constructed an all-metal, low-blank extraction line which is mounted on a cart and can be attached directly to the inlet of the mass spectrometer. This system is used for analyses of low-N materials for which experiments (fusions with and without addition of fluxes; cf. Zhang, 1988) involve large amounts of low-N sample (> 2 grams) and for extractions of N2 from fluid inclusions. Sealed-tube extractions of N from fluid inclusions in quartz veins, employing heating to 910 ~ - 1000~ for up to three hours (with Cu and CuO reagents), have been extremely successful in producing reproducible isotopic compositions (see earlier N-isotope work on fluid inclusions in rocks from the Dome de l'Agout, France, by Kreulen et al., 1982). In collaboration with J. Touret, M. Moree (both at Vrije Universiteit, Amsterdam) and M. Sintubin (Katholieke Universiteit Leuven, Belgium), we have performed analyses of the 615N of the generation of N2-rich fluid inclusions (a 93 mole % N2 with small amounts of CO2 and / or CH4) and coexisting biotite ("bastonite" with 2280 ppm NH4 +) in a quartz vein from these exposures. Three sealed-tube decrepitation experiments yielded 615N values of +3.7, +3.9, and +4.0%o for the quartz-vein fluid inclusion N2, and a single biotite analysis (using the techniques described in this paper) yielded a 615N value of +6.8%o. The difference in 615N of the N2 and biotite (mean vein ~15N is +3.9%o, thus making the inclusion-biotite difference ~ 2.9%o) is extremely similar to the A15NNH4+N2 calculated by Hanschmann (1981) for the petrologically inferred temperatures of formation of ~ 400~ (~ +2.8%o; indicated on Figure 16.4). 16.3 Outlook for future applications of N isotopes in studies of high-T silicate systems As in many other subfields of stable isotope biogeochemistry, upcoming advancements in both extractions and mass spectrometry will further widen the range of potential applications of N isotopes to an increasing number of petrologic problems. In particular, the application of lasers in heating experiments (Hashizume & Marty, Chapter 17 of this volume; Humbert et al., 2000) and the use of gas chromatographcombustion systems interfaced with carrier gas systems (see Brand, 1995b) should afford more routine and rapid applications of N isotopes (see early experimentation with elemental analyzer-continuous flow-IRMS methods by Jia & Kerrich, 1999; Kyser et al., 2000). The techniques we describe in this paper (and similar high-temperature "sealed-ampoule" extraction techniques employed by Boyd & Pillinger, 1990, and Boyd, 1997, in analyses employing static mass spectrometry) can easily be applied to a variety of other N-isotope petrologic pursuits, with (for example) extensions to facilitate extractions of N from feldspars, cyclosilicates (containing channel N2; Scalan, 1958, reported beryl with up to 234 ppm N and cordierite containing up to 125 ppm N; also see analyses of channel fluids by Damon & Kulp, 1958; Giuliani et al., 1997),

815N Analyses of Ammonium-Rich Silicate Minerals by Sealed-Tube Extractions ...

359

Figure 16.4 - Calculated fractionations (A15N - h15NNH4+- •15NN2,NH3 ) among N fluid species and NH4 +, the latter of which is bound structurally in silicate minerals in metamor-phic rocks (fractionations are from Hanschmann, 1981). The stippled horizontal and vertical lines indicate the excellent match of the calculated fractionations with the measured A15N for the "bastonite" (NH4+-rich biotite) and fluid inclusion N2 in quartz veins at Bastogne, Belgium (see description of veins and fluid inclusion chemistry in Darimont et al., 1988) for estimated vein formation temperatures.

and fluid inclusions (see Figure 16.4 and Kreulen et al., 1982). The precision obtainable for N isotopes using our methods (with lo commonly approaching 0.1%o; see Table 16.1) is similar to that obtainable for O in silicate systems (lo of 0.1 to 0.2%0 for n ~ 4; see 8180 data for garnet and quartz from various laboratories for laser probe methods and methods involving externally heated Ni reaction vessels in Valley et al., 1995), and the degrees of isotopic variation for the two systems in a given rock suite can also be quite similar (on the order of 5 - 10%o overall variation; see comparison of the O and N isotope systematics in the Catalina Schist, California, by Bebout, 1997; ~ 8%o 815N range in the Skiddaw aureole and granite, Bebout et al., 1999a; ~ 10%o 815N range at the Townshend Dam metamorphic locality, Sadofsky & Bebout, 2000). Two studies (Richet et al., 1977; Hanschmann, 1981) have published calculated Nisotope fractionation factors based on spectroscopic data. Of the two theoretical studies, only Hanschmann (1981) calculated fractionations involving NH4 + in solid phases. The calculated N-isotope fractionations among various N molecules based on spectroscopic data (see Figure 16.4) predict substantial N-isotope fractionation as a result of devolatilization, metasomatic alteration, and magmatic volatile release leading to 815N variation of the magnitude observed in recent integrated petrologicgeochemical studies employing N isotopes (see Javoy et al., 1986; Bebout & Fogel,

360


1992; Boyd et al., 1993a; Bebout, 1997; Bebout et al., 1999a,b; Sadofsky & Bebout, 2000). Conspicuously lacking for applications of N isotopes are experimentally determined fractionation factors for appropriate fluid-melt-mineral systems. The experimentally derived fractionation data and international N-isotope silicate standards will both be required in future applications of N isotopes to petrologic systems, regardless of the future direction of the analytical methods. However, one should consider applying N isotopes, particularly in studies of relatively potassic and organic-rich rock systems (e.g., metasedimentary and felsic-intrusive systems; K-metasomatized metamafic and metaultramafic rocks), to complement other geochemical data, just as one would consider the other more commonly used O, H, C, and S isotope systems in settings for which compositions are appropriate and there is some expectation that N isotopes will yield unique, useful constraints.

Acknowledgements

GEB acknowledges early support of this research by M. Fogel, the late T. Hoering, and the Geophysical Laboratory (Carnegie Institution of Washington, Washington, D. C.). Continued work has been supported by the National Science Foundation (grants EAR-9206679, EAR-9220691, EAR-9405625, EAR-9406135, and EAR-9727589), with some support also coming from the American Chemical Society Petroleum Research Fund (grant #25246-G2). GEB thanks M. Fogel, T. Hoering, S. Boyd, M. Grove, B. Idleman, and P. Zeitler for helpful discussions, and we thank Karin Brauer for providing several samples for interlaboratory comparison and Pier de Groot for undertaking this ambitious and extremely useful publication project.


CHAPTER 17 Nitrogen Isotopic Analyses at the Sub-Picomole Level Using an Ultralow Blank Laser Extraction Technique Ko Hashizumel,2,4* & Bernard Martyl,3** 1 Centre de Recherches P6trographiques et G6ochimiques, 15 Rue Notre-Dame des Pauvres, B. P. 20, 54501 Vandoeuvre-16s-Nancy Cedex, France 2 Department of Earth & Space Sciences, Osaka University, Toyonaka, Osaka 560-0043, Japan 3 Ecole Nationale Sup6rieure de G6ologie, Avenue du Doyen Roubault, 54501 Vandoeuvre-16s-Nancy Cedex, France e-mail: * [email protected]; ** [email protected]

Abstract We describe our recent achievement in determining the isotopic composition of nitrogen contained in extremely small samples. The aim of our development is to enable isotopic analysis of single mineral grains using the smallest quantity of nitrogen with a precision sufficient to resolve isotopic variations of nitrogen in extraterrestrial samples. The limiting factors are (i) residual nitrogen generated during extraction and purification of nitrogen, and (ii) mass interferences from hydrocarbons, N2H and CO at masses 28 and 29. The use of a defocused CO2 laser as a heating source allowed circumventing the first problem. Hydrocarbon (C2H5) is mass-resolved with our mass-spectrometer. Contribution of N2H was reduced dramatically by optimizing the ion source setting. CO is efficiently removed during purification and its residual contribution is corrected numerically using measurements at mass 30 in addition to masses 28 and 29. We are able to perform isotopic analyses of sub-picomole (< 30 picogram) quantities of N2 with a precision typically of + 10%o. The required amount of samples for nitrogen isotopic analyses, in the case of the lunar regolith, is reduced by a factor of-10-5 compared to the pioneer's works in the 1970's.

17.1 Introduction Nitrogen trapped in extraterrestrial materials exhibits large variations in the 15N/ 14N ratio, allowing to set strong constraints on the origin and evolution of the solar system and of the planets (e.g., Hashizume et al., 2000). The 615N values (which stand for the permil deviation of the 15N/14N ratio relative to terrestrial atmospheric nitrogen) range between-300 to +100%o among lunar grains (e.g., Kerridge, 1993; Hashizume et al., 2002), -200 to +1600%o among bulk meteoritic samples (see references in Hashizume et al., 2000), or even wider (log-scale variation of the absolute ratio, i.e., 5 4. Correspondence should be adressed to this author

362

Chapter 17 - K. Hashizume & B. Marty

< 14N/15N < 104) among the presolar grains contained in primitive meteorites (e.g., Zinner, 1998). Systematic studies of the nitrogen isotopic composition among the extraterrestrial materials started in the 1970's with the analyses of the regolith samples returned back from the Moon (e.g., Becker & Clayton, 1975), followed by bulk analyses of meteorites (Kung & Clayton, 1978). At that time, analyses were performed using dynamic type mass-spectrometers, which required 0.1 - 3 grams of the precious samples for each analysis. Later on, technical improvements were dedicated to reduce the sample size, by increasing the sensitivity of the mass-spectrometry system, as well as by reducing the blank levels, and to better resolve the indigenous nitrogen components from the terrestrial contamination. The sensitivity for nitrogen analyses increased dramatically following the development of static-type mass-spectrometers (e.g., Frick & Pepin, 1981; Wright et al., 1988; Hashizume & Sugiura, 1990; Murty & Goswami, 1992; Boyd et al., 1993b; Craig et al., 1993; Marty et al., 1995; Takahata et al., 1998; Yamamoto et al., 1998; Humbert et al., 2000; this study). The deconvolution of different nitrogen components has been improved by the introduction of the stepwise combustion technique (e.g., Frick & Pepin, 1981; Boyd et al., 1988; Boyd, Part L Chapter 13), or by using non-heating extraction methods such as vacuum crushing (e.g., Marty et al., 1995). In this paper, after presenting the general picture of the nitrogen isotope analysing system in CRPG-Nancy, we highlight our recent improvements concerning our ultralow blank extraction system and the mass spectrometry of sub-picomole quantity of nitrogen. 17.2 Overview of the N analyzing system The outline of the system and the analytical procedures are described here, although most of the details are described elsewhere (Humbert et al., 2000). The nitrogen isotope analysing system is described diagrammatically in Figure 17.1. It consists of a laser heating unit, an all-metal purification line, and a high-sensitivity static-type mass spectrometer.

Samples in the laser chamber are loaded in small pits machined in a stainless steel. The samples are heated by a defocused CO2 laser, working in a continuous mode with a wavelength of 10.6 gm. In case samples require stepwise analyses, several steps are performed before the final fusion step. The heating temperature is adjusted by inserting variable length of interval (e.g., 20 - 200 gsec) between laser emission (which last for 100 - 1000 ~sec per cycle). The temperature of the heated sample is estimated from its color and brightness. Samples are heated in oxygen atmosphere (P02 > 0.1 Torr) generated from CuO heated at > 720~ (combustion mode), or in vacuum (pyrolysis mode). Organic compounds released from the heated sample together with the sample gas (N2 and rare gases) are dissociated to oxide molecules (CO2, SO2, H20 etc.) in contact with the hot CuO, then are removed from the sample gas using the cold trap, a U-shaped glass tube cooled at -183~ The partial pressure of 02 is then lowered by slowly cooling down the CuO to 400~ Sample gases are normally split into two, N2 and Ar fractions, by volume dilution. These fractions are purified and introduced into the mass spectrometer sequentially. The impurities in the argon fraction are chemi-

Nitrogen Isotopic Analyses at the Sub-Picomole Level Using an Ultra-low Blank Laser ...

sorbed by two Ti-getters (GT1 and GT2) heated at 600- 700~ For the nitrogen fraction, no trap is used else than the CuO/ cold traps. Before introducing the gases into the mass-spectrometer, the gas pressure is monitored with an ion gauge having a tungsten filament working at a low filament current of 200 ~A. (The filament does not seem to pump out or emit measurable amount of nitrogen gas under the current condition, which is verified by the linear correlation between the nitrogen intensity measured by the massspectrometer and the pressure of standard air, adjusted by splitting the gas between known volumes.) The amount of nitrogen to be introduced into the mass-spectrometer is adjusted in order to get a signal compatible with the mass-spectrometer dynamic range. A rare gas mass spectrometer (VG 5400, Micromass 9 working in a static mode is used for the

363

Figure 17.1 - Diagram of the N / rare gas analytical system developed at CRPG-Nancy. The system is evacuated by ion pumps (IP) and turbo molecular pumps (TMP) connected to rotary pumps (RP). PG, IG, CT and GT denote Pirani gauge, ion gauge, charcoal trap and Ti-getter, respectively. The introduction line which includes the vacuum crushers is used for other extraction procedures and is not described in this contribution. Charcoal trap and helium cryotrap are used when analyzing light rare gases (He and Ne).

364


analyses. It has a Nier-type ion source, a bent tube of 0.54 m radius, across which a magnetic field is delivered by an electro-magnet stabilized with a Hall probe, and two collectors, a Faraday cup and an electron multiplier working at 2.5 kV. Its extended geometry allows a high mass resolution of 650 at 2% valley on the electron multiplier collector, whereas the resolution obtained on the Faraday collector is 250. On the Faraday collector, the output voltage is measured through a 1011 f~ resistor, whereas the count rates of electron burst output from the electron multiplier (dead time: 13 nsec) in response to the incident ions are measured by a pulse counting system (Orted~ 996). Comparing the count rates of ions detected by the two collectors (ion currents divided by the charge per ion, in the case of the Faraday collector), the yield of the electron multiplier collector was usually observed to be around 0.80 - 0.85 times that of the Faraday collector, which is normal for this type of collection. The source settings were 10 - 100/~A for the trap current, and 60 eV for the ionization energy. The trap current was reduced when we analyzed small quantities (e.g., sub-picomole) of N2, in order to avoid production of interfering N2H (See discussion in the "Nitrogen Mass-Spectrometry" section). Peak heights of N2 isotopes (14N14N, 14N15N and 15N15N) and interfering species are measured 1 0 - 15 times repeatedly. The total time required to complete the nitrogen analysis is 15 - 20 minutes. The mean sensitivity for N2 (the current measured on the collector divided by the partial pressure of the gas in the mass-spectrometer) ranged from 1.8 x 10-5 A/Torr at a trap current of 10 gA to 2.6 x 10-4 A/Torr at 100 gA. The half-life of N2 in the ion source was 350 and 25 min. for trap currents of 10 gA and 100 gA, respectively. In the standard procedure, hot blank (discussed in section 17.3) and standard gas are measured at least once a day, respectively before and after beside the sample gas analyses. Standard measurements are performed following the same procedure as for the sample gas analyses. A pipette (4 x 10-10 mol N2) of standard air, prepared from atmospheric air and stored in a vial (STD2), is taken for the measurement. The amount of the standard gas to be introduced into the mass spectrometer is adjusted by volume dilution to be comparable to that of the sample gas. 17.3 Low blank gas extraction using a laser A critical issue which controls the quality of analyses of small amounts of nitrogen is the hot blank level, which is the amount of gas released when the extraction and purification procedures are performed without a sample. Several sources can contribute to the blank, among which are the "hot spots" in the vacuum line, e.g., the extraction chamber, where the nitrogen contained in samples is extracted by heating, copper oxide, or the vacuum gauge. After the vacuum line had been well degassed by cyclic baking under high vacuum, the sample gas extraction part became the largest blank source. In order to reduce the blank generated during the extraction procedure, we used a defocused laser beam as a heating device. Franchi et al. (1986) also developed a laser system to extract nitrogen and other volatiles. They used a laser beam to heat a small area of a large sample, whereas we use it to homogeneously heat a small sample. The CO2 laser beam is absorbed by the silicate sample, while the part of the beam arriving at the surrounding stainless-steel surface is reflected and scattered. Thus only


365

the silicate sample is heated, minimizing the source of hot blank. The variation of the hot blank level of our system is s h o w n in Figure 17.2. Periods w h e n a series of analyses were p e r f o r m e d after the clean-up procedure (described below) are s h o w n by the hatched bars. The sample chamber, after fusing all the samples loaded in the pits, is opened and w a s h e d with acids (which consist mainly of fluoric acid and a small a m o u n t of nitric acid), distilled water and acetone before reloading new samples. It normally requires 2 - 4 days, after the installation of the chamber to the v a c u u m line, to reduce the blanks to acceptable levels. The installed chamber is baked at 120~ at least overnight, often for two days. We often performed a "pre-combustion" procedure, introducing ~ 1 Torr of 02 into the hot chamber, connected to the cold trap (Figure 17.1). This procedure promotes removal of adsorbed air, organic contamination and water from the samples and from the stainless steel apparatus (At the baking temperature of 120~ most of the organic matter is not com-

Figure 17.2 - Hot blank level of nitrogen plotted as a function of time. Hot-blank analyses are carried out following exactly the same procedure as that applied to the samples. Data plotted on days within a hatched bar represent a series of analyses performed after the clean-up procedure subsequent to the reloading of samples. Intervals between analyses series are taken arbitrarily. The h15N values of the blank nitrogen were constantly lower than- 30%o during the first five series of analyses, as a probable result of contamination from nitric acid used in the cleaning procedure. Subsequently, the use of additional cleaning treatments (see text) reduced the N blank level, and resulted in blank h15N values close to 0%0.

366


busted. Nevertheless, we believe that removal of organic matter weakly adsorbed on the surface of the stainless steel is facilitated by the pre-combustion procedure). We verified that the major source of the blank nitrogen in combustion mode is not the CuO, which provides the pure oxygen. Indeed the "cold" blank level, the amount of nitrogen released by the purification procedure itself, was lower than 5 picogram, well below blank values observed during combustion experiments (Figure 17.2). We observed that the blank level was primarily determined by the cleanliness of the sample chamber on the one hand, and the mode of extraction, vacuum pyrolysis or combustion, on the other hand. Initially, we performed analyses in the combustion mode, heating the sample under an oxygen atmosphere ranging from 0.05 to 0.2 Torr. The blank level in this mode was generally higher than that obtained by heating the sample without 02 (pyrolysis mode), probably because the oxygen molecules mobilize residual nitrogen-bearing compounds adsorbed on the internal surface of the laser chamber. One of the problems encountered during the first five series of analyses was that ~515Nvalues of the blank nitrogen showed constantly negative values, -30%0 or lower. We suspect that it originates from the nitric acid added to the liquid to wash the sample holder. From the sixth series, the sample holder was boiled in distilled water several times after the acid treatment, and was washed again with a solution of an organic acid (COOH)2 to remove a possible chemisorption layer involving nitrogen. These additional treatments lowered the blank level by at least a factor of two (Figure 17.2). Subsequently, the 615N values of the blank nitrogen were observed to be constantly around 0%0. The blank nitrogen produced during vacuum pyrolysis was within 5 - 1 0 picogram (Figure 17.2). The amount of residual N obtained during sample heating may be higher than the blank level obtained when shooting the laser at an empty pit, because the heated sample can heat the sample holder by conduction. Though it is not easy to quantify the "true" blank level which involves sample heating (since no sample is granted to be perfectly nitrogen-free), we give an example suggesting that the radiation from the sample does not dramatically increase the blank level. Among the lunar grains we measured, several grains possessed extremely small amounts of nitrogen. When fusing such a sample in vacuum, the amount of nitrogen (including the blank) was 17 picogram, which is not dramatically higher than the hot blank range (5 - 10 picogram). In summary, we have achieved the lowest blank level among the systems used to analyze nitrogen isotopes. The literature hot blank levels range between 200 - 5000 picogram, while the typical blank level of our system is 5 - 10 picogram during pyrolysis, and 2 0 - 40 picogram during combustion, which enables nitrogen isotopic analyses of N2 in the picomole to sub-picomole range.

17.4 Nitrogen mass-spectrometry The nitrogen isotopic ratio of the extracted gas is determined basically by comparing the peak heights of N2 at masses 28 (14N14N) and 29 (14N15N). However, contribution of CO, N2H and hydrocarbons (C2Hx) interfering at these masses is critical, especially when the amount of N2 introduced in the mass spectrometer is small. The


367

mass resolving powers (M/AM) required to discriminate these compounds from N2 at mass 29 are 5903 (13CO), 2688 (N2H) and 807 (C2H5). The best reproducibility for the 29/28 isotopic ratio is obtained when the N2 pressure exceeds 10-7 Torr (~ 300 picogram), i.e., when the contribution of the interfering species can be neglected. From the peak heights of N2 at masses 28 and 29 measured using the Faraday collector, the external reproducibility of the ~515Nvalue of standard air is typically 1 +_0.5%0 (Marry & Humbert, 1997; Dauphas & Marry, 1999). Below we describe the methods we developed to avoid or numerically subtract mass interferences at very low levels of N2.

Hydrocarbons: We are

able to separate partly the C2H5 peak from the 14N15N peak at a mass resolution of 650 using the electron multiplier collector. Figure 17.3 shows a mass-scan profile at mass 29 using a small amount of sample gas (which is actually ~ 3 picogram of blank N2 gas from the laser chamber). When > 3 x 10-9 Torr (7 picogram) of N2 is introduced into the mass-spectrometer working with a trap current of 100 gA, however, the ion counting system is saturated when analyzing N2 at mass 28. There are two countermeasures to cope with the problem. 1) Use of the Faraday collector for signals at masses 28 and 29, combined with additional analyses of two peaks, N2+CO+N2H and C2H5 at mass 29, using the ion

Figure 17.3 - Peak profile at mass 29 scanned over an interval of 0.1 amu for a blank N2 a m o u n t of -~ 3 picogram. The C2H5 peak is partly separated from 14N15N peak at the mass resolution p o w e r of 650 characterizing the electron multiplier collecter.

368


counting system. The ratio (28/29)N2+CO+N2H, corrected for interferences from hydrocarbons, is calculated as: (28 / 29)N2+CO+N2H - (28 / 29)Faraday X {1 + (29)CH/ (29)N2+CO+N2H}

[17.1]

where (29)CH and (29)N2+CO+N2H a r e the peak heights at mass 29 obtained in ion counting mode. 2) Reducing the trap current, for example to 10 gA, to enable the analysis of all N2 peaks with the electron multiplier. In the < 30 picogram range, the reproducibility of the nitrogen isotopic ratio of standard air analyzed in this mode was better than that obtained by using the first mode (The reason is described in the next subsection on N2H). The drawback of this mode is that it sometimes prevents simultaneous analyses of rare gases when their concentrations are low.

N2H: This species, along with CO, cannot be resolved from N2 by our mass-spectrometer. A small amount of N2H seems to be generated in the mass-spectrometer, although its occurrence depends largely on the ion source conditions. In this subsection, evidence for its occurrence is discussed in detail, because, unlike the interferences by hydrocarbon or CO, it is impossible to precisely correct for its contribution. In Figure 17.4, the 30/28 ratios of standard air N2 are plotted against the 29/28 ratios. This diagram exhibits a mixing relationship between N2, CO and N2H. With the trap current set at 10 ~A, all data lie on a mixing line between N2 and CO, suggesting that the contribution of N2H is negligible. In constrast, significant amounts of N2H are generated when the trap current is set at 100 ~A. The protonation rate can be reduced presumably by lowering the H2 partial pressure (e.g., Frick & Pepin, 1981). The dependence of the protonation rate on the trap current may be due either to the increase of the hydrogen partial pressure around the ion source, or to the increase of the reaction rate at 100 ~A, which are both likely to be a function of the filament temperature. Two observations suggest that the N2H observed in the standard air is produced around the ion source in the mass-spectrometer, not in the purification line or in the standard tank. First, the amount of N2H increases with increasing trap current, as stated above, and second, the 29/28 ratio is growing rapidly after introduction of the gas into the mass-spectrometer. The average growth rate of the 29/28 ratio was observed to increase with decreasing pressure of N2 in the mass-spectrometer (The growth of the ratio does not seem to be primarily due to mass-dependent isotopic fractionation processes, since the rate was observed to be much higher than the one expected by fractionation processes such as nitrogen consumption by the source filament). The average growth rate A(ln29/28)/dt was roughly proportional to the inverse square root of the N2 pressure, ranging from 1.5%o/min at PN2 - 2 x10-7 Torr to 15%o/min at PN2 - 2 xl0 -9 Torr. We interpret the linear correlation between PN2-1/2 and A(ln29/28)/dt as due to a rate determining process for the nitrogen protonation which involves the atomic form of nitrogen (e.g., N + H ---, NH, N + 1 / 2 H2 ---, NH or

HANDBOOK OF STABLE ISOTOPE ANALYTICAL TECHNIQUES VOLUME 1

Pier A. de Groot editor Economic Geology Research Institute, School of Geosciences, University of the Witwatersrand, Johannesburg, South Africa and Delta Isotopes Consultancy, Pastoor Moorkensstraat 16, 2400 Mol-Achterbos, Belgium (present address)

2004

ELSEVIER Amsterdam - Boston - Heidelberg - L o n d o n - New York - Oxford Paris - San Diego - San Francisco

- Singapore

- Sydney

- Tokyo

370


If the production of N2H was a linear function of residence time in the mass-spectrometer, the contribution of N2H could be corrected by taking the 29/28 ratio at time zero (introduction of N2 into the mass-spectrometer). Nevertheless, the contribution of N2H is still seen at time zero, which is demonstrated in Figure 17.4 where the ratios 29/28 and 30/28 extrapolated to time zero are plotted. Such effect seems to be due to the non-linear production of the N2H after introduction of the gas. The 29/28 ratio generally grew faster at the beginning, slowing down with time. The slower growth at longer time is possibly due to equilibration of generation and dissociation processes of N2H. The apparent protonation rate of N2 at time zero, represented by the N2H/N2 ratio, depends on the N2 pressure in the mass-spectrometer (Figure 17.5). At the N2 pressure range of ~ 10-8 Torr (~ I picomole), the protonation rate is variable and large enough to largely degrade the quality of nitrogen isotopic determination. For exam-

'

'

'

I

'

'

'

'

I

'

'

'

'

I

'

'

0

7-I

'

'

I

'

'

'

'

STD N2 (Under 100gA Trap Current)

@ 6 - ~@ 5

•

~

t'-,l

O

3 _0

Z

O O O

O O

0

,

0

,

,

,

I

5

,

,

,

,

I

10

,

,

,

,

I

,

i

i

15

i

I

20

,

,

,

,

2

N2 Pressure (10-8 Torr) Figure 17.5 - The protonation rate expressed as the N2H/N2 ratio, under the 100 ~A trap current condition, plotted against the nitrogen pressure in the mass-spectrometer. The amount of N2H in the standard air is calculated assuming that variations of the isotopic ratios 29/28 and 30/28 are caused by admixture of three components (N2, CO and N2H) with known isotopic ratios (i.e., obtained from the mixing diagram shown in Figure 17.4). Under the 100 ~A trap current condition, the contribution of the interfering species N2H cannot be neglected when the nitrogen pressure in the mass-spectrometer is below 10-7 Torr.


371

ple, ambiguity in the protonation rate of + 2 x 10-4 results in an uncertainty of 30%o in the nitrogen isotopic ratio. Therefore, we applied a source condition of 10 gA trap current to avoid N2H interference when the amount of N2 to be measured was in the picomole to sub-picomole range. CO" The amount of CO interfering at masses 28 and 29 can be estimated by analyzing the N2+CO peak at mass 30, and solving the following mass-balance equations: Mass 28 = 14N14N + 1 2 C 1 6 0 Mass 29 = 14N15N + 13C160 (+ 12C170) Mass 30 = 15N15N + 12C180 (+ 13C170)

[17.2] [17.3] [17.4]

Species in parentheses are the ones which are normally negligible compared to the primary CO isotopes at each mass (They are ignored in the following formulae for simplicity). Assuming that isotopic equilibrium is achieved among the N2 molecules, the abundance ratios between their isotopes are: 14N14N:14N15N:15N15N=

1 : 2r : r2

[17.5]

where r = 15N/14N (Note that this assumption do not always hold in the case of extraterrestrial samples that contain extremely anomalous isotopic components, such as cosmogenic nitrogen. See Hashizume & Sugiura, 1992; Sugiura et al., 1995 for details). Then the mass balance equations are: Mass 28 = N2 + CO Mass 29 = 2r x N2 + 1 3 C / 1 2 C x CO Mass 30 = r 2 x N2 + 180/160 x CO

[17.6] [17.7] [17.8]

Assuming that the 1 3 C / 1 2 C and 180/160 ratios of the interfering CO are known, the remaining unknowns are (1) r (= 15N/14N), (2) the amount of N2 and (3) the amount of CO. Therefore, we obtain a unique solution by solving the three equations [17.6/8]. CO originates chiefly in the mass spectrometer ion source, since CO is efficiently removed during the purification procedure (Humbert et al., 2000), and because the calculated amount of CO is nearly constant, regardless of the nature of the sample gas. Therefore, we assume the 13C/12C and 180/160 values of the CO are equal to their standard terrestrial values even when we analyze extraterrestrial samples. However, to be conservative, we assume uncertainties for these ratios to be _+50%0, exceeding the range observed among bulk meteorites. These uncertainties are propagated on the uncertainty for the nitrogen isotopic ratio. When we introduce 10 picogram (0.3 picomole) of standard N2 in the mass-spectrometer, the contribution of CO at mass 28 is normally around 2%, and the ~15N is corrected typically by 10%o.

Nitrogen Isotopic Ratio: The overall performance of our nitrogen isotope analyses at the sub-picomole range is shown in Figure 17.6, which represents the 14N14N/14N15N ratios of various amounts of standard air N2, corrected for interference of CO. Since

372


Figure 17.6 - The 14N14N/14N15N ratios of sub-picomole quantity of standard air as a function of N2 amount in the mass-spectrometer (CO interference is corrected). This series of analyses is performed under 10 gA trap current condition, where the N2H interference can be neglected. The ion counting system was used for analyses of all peaks including N2 at 28 and 29. In this range of N2 amount, statistical counting error during measurements of the 14N15N peak height (plotted as dotted curves) appears to be the main cause of errors.

these analyses were performed with the trap current set at 10 t~A, we assume that the N2H interference is negligible. The absolute value of the 14N14N/14N15N ratio in this series was systematically higher by ~ 3% compared to the commonly accepted air ratio of 136, which we regard as a mass-discrimination effect. When the magnet and/ or the source conditions are changed, the absolute nitrogen isotopic ratio can vary from 136 to 141, however, the ratio is stable once these parameters are fixed. We adopt the observed mean ratio, e.g., 140.3 in this case, as the reference air value to calculate the 815N values. The 815N is formally defined by comparing the isotopic abundance ratio between 15N and 14N atoms relative to the atmospheric air value. However, we obtain an equivalent value by comparing the abundance ratio between 15N14N and 14N14N molecules with the corresponding reference value, provided that the isotopic equilibrium is achieved among the nitrogen molecules. 515N __ ((15N/14N)observed - 1 ) x 1000 - ( (15N 14N/14N 14N)~ (15N14N/14N14N)AIR (15N/14N)AI R

1) x 1000

[17.9]


373

Error bars in Figure 17.6 represent lo e r r o r ((Jspectrometry) comprising the counting statistic error (Ocounting), the uncertainty (Ocorrection) caused by correction of interfering CO, and the e r r o r (ofitting) caused by fitting ratios to a line as a function of time and extrapolating it to time zero (i.e., when the gas is introduced into the mass spectrometer). We obtain the final uncertainty (Ospectrometry) concerning the mass-spectrometry by the following formula; (Ospectrometry) 2 = (Ocounting) 2 + (Ocorrection) 2 + (Ofitting) 2

[17.10]

The counting statistic e r r o r (Ocounting) for a given amount of nitrogen in the mass spectrometer is calculated to be; ((Jcounting) 2 - ((Jcounting(o)) 2 x [No/N]

[17.11]

where, N is the amount of nitrogen introduced in the mass spectrometer. The No and Ocounting(o) a r e the reference amount of nitrogen and its corresponding statistical counting error, respectively. When 10 picogram of nitrogen (No) is introduced in the mass spectrometer, the counting e r r o r (Ocounting(o)) is 4.4%0. The dotted curves in Figure 17.6 represent the counting statistic e r r o r (Ocounting). It is therefore shown here that the error on the isotopic ratio of sub-picomole quantities of nitrogen obtained by our procedure is limited primarily by counting statistics. 17.5 Total Performance- A case study

In the above sections, we described the technical options chosen for isotopic analyses of sub-picomole quantities of N2. Here we synthesize the performance of our analyzing system. Although practical results are shown elsewhere (e.g., single grain analyses of lunar grains; Hashizume et al., 2002), the grain size and the N concentration assumed here are typical of a natural sample analysis. We consider the case of a silicate sample weighting 0.03 mg (a 200 ~m-sized grain) loaded in the laser chamber. We assume that this sample contains 1 ppm of surface-correlated N, which is extracted at a heating temperature of ~ 1000~ and 1 ppm of volume correlated N, which is mostly extracted by fusion of the sample (Such a situation is often seen among the lunar grains. See for example, Kerridge (1993) for the release profile of lunar nitrogen by stepwise heating). First, the sample is heated at 1000~ in the vacuum pyrolysis mode (If the sample appears to be dirty with a significant amount of organic compounds at the surface, these contaminants can be removed by combusting the sample at lower temperature, e.g. at 600~ We monitor the color and brightness of the heated sample using a CCD camera (Humbert et al., 2000), which enables us to control the heating temperature with a precision of + 100~ Following the purification and analyses of nitrogen and rare gases of the 1000~ temperature step, the sample is melted at a temperature of 1600~ and the extracted gas is purified and analyzed in the same way. In each heating step, ~ 30 picogram of N2 is extracted. During the extraction procedure in vacuum pyrolysis mode, the blank contributes 5 - 10 picogram of N2 to the sample gas. We observed that the hot blank level is variable by + 20% relative to the

374


blank level measured on the same day. After the purification procedure, the sample gas is introduced into the mass-spectrometer. Since we split part of the gas for rare gas analysis, the amount introduced in the analyzer for nitrogen isotopic analysis is around 10 - 20 picogram. When analyzing 15 picogram of N2 in the mass-spectrometer in ion counting mode, the 615N value for the introduced gas (= sample + blank) is usually determined with a precision of + 7%o. Subtracting the blank contribution (which normally accounts for 15 - 25% of the total gas, in this case), the 615N value of the nitrogen gas is finally determined with a precision typically of + 10%o (1~). The 615N values with such uncertainties are still well informative to discriminate various isotopic components seen among extraterrestrial samples. 17.6 Conclusion

In this contribution, we have presented a new analytical procedure aimed to measure the isotopic composition of sub-picomole quantities of nitrogen. With respect to standard analytical procedures, two major points have been improved, the level of the hot blank and that of interfering species, especially that of N2H. The hot blank has been lowered by minimizing the hot area during the extraction procedure using a defocused laser beam as a heating device. The amount of the interfering species N2H arising from the mass-spectrometer ion source was lowered by optimizing the ion source condition. Having done that, our mass-spectrometry system performs nitrogen isotopic analyses within limits imposed only by the counting statistics. Using this system, we are able to measure the N isotopic ratio with a precision of + 10%o for a - 200 ~m sized single grain containing I ppm of nitrogen. Such system is useful for untangling the nitrogen isotopic puzzle exhibited among the extraterrestrial materials, that is, to identify fine-grained anomalous nitrogen isotopic carriers.

Acknowledgments

We are indebted to Laurent Zimmermann, Pascal Robert, Gregory Sauder and Jean-Claude Demange for their technical assistance. Discussions with Nicolas Dauphas, Frank Humbert, Stuart Boyd and Rainer Wieler greatly improved the quality of this work. We thank Ian A. Franchi and Haraldur R. Karlsson for their detailed and constructive reviews. This work was supported by Programme National de Plan6tologie/Institut National des Science de l'Univers. K.H. acknowledges support from the Japanese Ministry of Education, Science, Sports and Culture, and the Institut National des Sciences de l'Univers - Centre National de la Recherche Scientifique for his stay in Nancy. This work is CRPG-CNRS contribution 1510.

Handbook of Stable Isotope Analytical Techniques, Volume 1 P.A. de Groot (Editor) 9 2004 Elsevier B.V. All rights reserved.

CHAPTE R 18 Mass Independently Fractionated Ozone in the Earth's Atmosphere and in the Laboratory Jeffrey C. Johnston1 & Mark H. Thiemens2 Iterations, P.O. Box 590805, San Francisco, California 94159, USA Department of Chemistry and Biochemistry, University of California, San Diego, La Jolla, California 92093-0356, USA e-mail: 1 [email protected],2 [email protected] 1

2

18.1- Introduction The isotopic analysis of gases in the Earth's atmosphere provides a wealth of information regarding sources, sinks, chemical transformation, and transport processes that govern the budgets of atmospheric species. A non-random distribution of isotopes is observed in many atmospheric species because of a variety of kinetic and equilibrium isotope fractionation effects. This chapter examines the mass independent isotope fractionation observed in ozone (03), both in the Earth's atmosphere and in the laboratory. Many of the isotopic fractionations observed in nature are termed mass dependent because the magnitude of these fractionations scales with the relative mass difference between the nuclides of a specific element. As discussed in detail elsewhere, a mass dependent fractionation of oxygen isotopes will result in an isotopic value that closely follows the relation 6170 - 0.526180 (Thiemens, 1999). Equilibrium chemical processes - such as isotope exchange, diffusion, evaporation and condensation- are sensitive to isotopic mass, and different isotopes may thus be incorporated into the molecules at different rates. The oxygen isotope exchange between H20 and CaCO3, and vapor pressure isotope effects are two examples of mass dependent, equilibrium chemical processes. The sensitivity of a chemical rate constant to isotope substitution, termed a kinetic isotope effect, can generally be understood in terms of transition state theory. Isotopic substitution affects many terms in the transition state theory expression for the reaction rate, but the resulting fractionations are always mass dependent (Bigeleisen & Wolfsberg, 1958). In a system such as for oxygen isotopes, the isotopic fractionation for the 170/160 is always half that of that for 180/160. This is a direct result of the mass dependency of isotope effects. In the case of 170/160, the mass difference is 1 1. Correspondence should be adressed to this author.

376

Chapter 18 - J.C. Johnston & M.H. Thiemens

amu while for 1 8 0 / 1 6 0 is 2 amu. As a result, for example, a 5 per mil fractionation in the 1 7 0 / 1 6 0 ratio is accompanied by a 10 per mil effect in the 1 8 0 / 1 6 0 because of the dependency upon relative mass differences. Such isotope effects are termed mass dependent. In addition to these mass dependent processes, there are a significant number of gas phase fractionation processes that do not scale with differences in nuclidic mass. These so called mass independent fractionations (MIF) have been observed in a variety of chemical systems, including 03 formation (Morton et al., 1990; Thiemens & Heidenreich, 1983), $2F10 formation (Bains-Sahota & Thiemens, 1989), the photopolymerization of CS2 (Coleman et al., 1996), the O + CO reaction (Bhattacharya & Thiemens, 1989a; 1989b), OH + CO reaction (R6ckmann et al., 1998b), and H + 02 recombination (Savarino & Thiemens, 1999b). To date, mass independent isotopic compositions have been observed in meteorites, atmospheric 02, 03, CO, CO2, N20, H202, aerosol sulfate and nitrate, as well as solid samples from the Namibian desert, Antarctic dry valleys, Miocene volcanic ash and Death Valley varnishes (Thiemens et al., 2001). Thiemens (1999) and Weston (1999) have provided thorough reviews of mass independent fractionation processes. Although the mechanisms responsible for producing them are not completely understood, the identification of MIF in natural samples can provide information not available from concentration measurements and single isotope analysis alone (Cliff & Thiemens, 1997; R6ckmann et al., 1998a; 1998b). To date, most laboratory research on MIF has involved molecules incorporating the three stable isotopes of oxygen (160=99.762%, 170=0.038% and 180=0.200%). The Earth's atmosphere is an ideal place to search for MIF due to the wide variety of oxygen containing species and the wide range of isotopic fractionation processes occurring there. The deviation from a purely mass dependent fractionation is conveniently defined by the value A170, where A170=~5170-0.52~5180. Thus, when looking at oxygen isotopes, a mass independent fractionation is one in which A170 ~ 0. Sulfur (32S=95.02%, 33S=0.75%, 34S=4.21%, 36S=0.02%) is another isotopic system in which MIF have been found (Coleman et al., 1996; Cooper et al., 1997; Farquhar et al., 2000b), although to date less work has been done to search for these fractionations in the Earth's atmosphere. This chapter reviews the MIF observed in 0 3 in the Earth's atmosphere and in the laboratory. Chemically produced MIF processes were first observed in 03 (Thiemens & Heidenreich, 1983), and over the past 20 years a significant amount of research in many laboratories has been devoted to developing an understanding of this process. While there remain gaps in our understanding of the mechanisms responsible, recent experimental and theoretical work (discussed in section 18.4) has provided interesting new insights into the process involved. Section 18.2 briefly reviews the chemistry of 03 in the atmosphere, and section 18.3 discusses what is known about the isotopic composition of atmospheric 03. Section 18.4 examines what is known about the source of the MIF in 03 from laboratory experiments and theoretical models; conclusions and references are in sections 18.5 and 18.6.

Mass IndependentlyFractionated Ozone in the Earth's Atmosphereand in the Laboratory

377

18.2 - O z o n e in the Earth's a t m o s p h e r e

While 03 is found throughout much of the atmosphere, it has a very uneven distribution; a consequence of the fact that 03 is not emitted directly from the ground in any significant amount, but is formed photochemically in the atmosphere. Ozone is produced when an oxygen atom combines with an oxygen molecule:

(R1)

O + 02 + M --. 03 + M

where M symbolizes a third body species which is required to remove excess energy from the nascent 03 molecule. In the stratosphere, the oxygen atoms are produced largely by UV photodissociation (X ~ 240 nm) of 02, while in the troposphere the oxygen atoms are produced primarily by photodissociation (K ~ 430 nm) of NO2. Most of the atmosphere's 03 is located in the stratosphere between approximately 25-40 km in altitude. 03 in this region of the atmosphere is critical due to the fact that 03 absorbs UV radiation between 200-310 nm very strongly. Were it not for the stratospheric 03 layer, this biologically harmful radiation would penetrate to the Earth's surface. In contrast to the troposphere, which is characterized by decreasing temperature with increasing altitude, temperature increases with altitude in the stratosphere due to the presence and photolytic activity of 03. Only about 10% of atmospheric 03 is located in the troposphere, and its concentrations there are extremely variable. In very polluted regions it can rise to several hundred parts per billion by volume (ppbv), while in remote regions it can drop below 10 ppbv. The oxidizing potential of the Earth's atmosphere is largely maintained by the hydroxyl radical (OH), the nitrate radical (NO3), and 03. Reaction with OH is the predominant sink for a large number of reduced trace gases including carbon monoxide, methane and nonmethane hydrocarbons (NMHC), hydrofluorocarbons (HFCs), and hydrochlorofluorocarbons (HCFCs). The most important tropospheric source of OH begins with the ultraviolet photolysis of 03 (K < 320 nm). This produces oxygen atoms in the first electronically excited state, O(1D), 7-10% of which then react with water to form OH, the primary daytime oxidizing agent of the troposphere. 03 + hv --* O(1D) + 02 O(1D) + H20 ---, 2OH

(Z. < 320 nm)

[18.2] [18.3]

Reaction between 03 and NO2 leads to the formation of the nitrate radical (NO3), the primary oxidizing agent in the nighttime troposphere. For a number of biogenic and anthropogenic hydrocarbons, mostly alkenes, reaction with 03 represents the major loss process. Thus, to a large extent, 03 controls the oxidation state of the atmosphere. A direct and indirect influence on climate is the second important role played by 03. The direct effect is due to the fact that 03 is a greenhouse gas. 03 is optically active, absorbing solar radiation in the UV and visible regions, and absorbing and emitting terrestrial IR radiation in the 8-10 mm region. This fact means that a change in the distribution of 03 in the atmosphere will disrupt the radiative energy budget and possi-

378


bly disrupt climate. The indirect climatic effect of 0 3 in the troposphere is a consequence of the strong chemical tie between 03 and several other important greenhouse gases, particularly CH4, HFCs and HCFCs which are removed from the atmosphere primarily through reaction with OH. As distributions of 03 change, distributions of OH also change, thus influencing the lifetimes of these greenhouse gases. Tropospheric 03 is also important because of its phototoxic effects on plants and animals. 03 is one of the primary components of photochemical smog, millions of people live in regions that regularly experience episodes of ground level 03 concentrations greater than what the U.S. Environmental Protection Agency considers safe. Effects on human health include reductions in lung function and increases in respiratory symptoms, airway reactivity, permeability, and inflammation. 03 exposure is also responsible for billions of dollars in crop losses each year (Adams, 1985). According to current understanding two tropospheric 0 3 sources, transport from the stratosphere and in situ photochemical production, are balanced by two loss processes, deposition at the Earth's surface and in situ chemical destruction. While these production and loss terms are all estimated to be of the same order of magnitude, spatially and temporally quantifying them is difficult. As with many other atmospheric components, the isotopic characterization of source and sink terms could help to provide important, and otherwise unobtainable, information about the budget of atmospheric 03. 18.3 - The isotopic c o m p o s i t i o n of atmospheric o z o n e

Cicerone & McCrumb (1980) published the first paper on a possible non-statistical distribution of isotopes in stratospheric 03, suggesting that 3402 (160180) may be preferentially dissociated in the stratosphere and mesosphere due to selective absorption in the Schumann-Runge bands (175-205 nm) when compared to 3202 (160160). The proposed effect was an isotopic self-shielding arising from the separate and differential adsorption of the isotopically substituted molecular isotopic species. The individual absorption lines coupled with the differences in natural abundances of the isotopes was suggested to produce a variation with altitude in the photolysis of molecular oxygen, leading to increased production of 180, which would in turn produce more 5003. In 1981, the first measurements were published of stratospheric 0 3 between 22 and 38 km using a balloon borne mass spectrometric system (Mauersberger, 1981). The balloon payload consisted of a high-speed differential pumping system that formed stratospheric gases into a molecular beam. This beam traversed the ion source of a Mattauch-Herzog magnetic instrument, with a mass range of 11-90 amu, without wall collisions. The ion detector was a low noise counting multiplier which permitted long sampling intervals (Mauersberger, 1977; Mauersberg & Finstad, 1979). The results showed a pronounced enhancement in 5003 (160160180 and 160180160) reaching a maximum of over 400%o (relative to the expected statistical distribution of natural 160

Mass IndependentlyFractionated Ozone in the Earth's Atmosphere and in the Laboratory

379

and 180 in the atmosphere) at 32 km and decreasing toward higher and lower altitudes. As only a single isotope ratio was measured in this work, it was not possible to distinguish between mass dependent and mass independent isotopic compositions. Kaye & Strobel (1983) showed that the rapidity of the isotopic exchange between O and 02 would dilute any increased concentration of 180 from preferential 3402 photolysis, and thus concluded that a paradox existed between the measurement of large 5003 enrichments in the stratosphere, and theory which predicted no enrichment. The rapid isotopic exchange between O and 02 has important implications on the isotopic composition of atmospheric 03 and makes it important to distinguish between the chemical and isotopic lifetimes of the 03 molecule. While the chemical lifetime can be on the order of months (Liu & Trainer, 1988), the isotopic lifetime will generally be much shorter. The photochemical lifetime of 03 is extremely variable, but an average value of on the order of hours can be calculated using standard actinic flux data. Because the dominant fate of the O(3P) produced during 03 photolysis is to recombine with 02 to produce 03, the effective chemical lifetime is longer than the lifetime against photolysis. However, because isotopic exchange between O and 02 is approximately 300 times faster than recombination to form 03, the isotopic composition of the of the original 03 has been lost. Thus, rather than providing an integrated picture regarding a variety of sources and sinks, as is found when looking at the isotopic composition of relatively long lived species like CO2, N20, or CH4, the isotopic composition of 03 should be viewed as providing information regarding processes that are typically on the order of hours old. Results from two additional flights of the balloon-borne in situ mass spectrometer payload confirmed the large and variable enrichments in 5o03, and showed that the concentration of 4903 (160160170 and 160170160) is also enhanced (Mauersberger, 1987). Due to a very low abundance, quantitative analysis of 4903 profiles from the in situ mass spectrometer data was not possible. Thus, Mauersberger and co-workers developed a balloon-borne cryogenic sampler capable of returning samples for laboratory analysis (Schueler et al., 1990). Cryogenic sampling utilizes the differences in the thermodynamic properties of gases to separate them. A gas will condense on a surface when the partial pressure of the gas is significantly higher than the equilibrium vapor pressure of the gas, which is determined by the temperature of the surface. By flowing a gas stream through a series of cold traps at progressively colder temperatures, the components of the gas stream can be isolated as a function of vapor pressure. Thus, more abundant gases like H20 and CO2, which condense at higher temperatures, can be separated from 03. Equilibrium vapor pressures of most atmospheric gases are well known, and precise vapor pressure measurements were made over liquid and crystalline 03 by Hanson & Mauersberger (1986). The stratospheric collection system (Schueler et al., 1990; Stehr et al., 1996) involved a series of liquid nitrogen cooled traps, the first at 80 K to condense H20 and CO2, and the second at 63 K to condense 03. These two traps were followed by a small zeolite pump which is used to remove the last traces of non-condensable gases (primarily N2, 02, and Ar)

380


after the collection had ended. The trap temperatures were controlled by regulating the N2 pressure above the liquid. At 63 K, over 95% of the 03 passing through the system condensed, while the major atmospheric gases did not. For stratospheric flight, the balloon platform carried three 03 samplers connected by a common air inlet line, allowing multiple samples to be collected on a single flight. Prior to analysis, the 03 samples were allowed to warm to room temperature to convert the 03 into 02. The samples were analyzed with a mass spectrometer attached to a beam system, as described in (Anderson & Mauersberger, 1981; Mauersberger, 1977; Mauersberger & Finstad, 1979) using atmospheric 02 as the isotopic standard. Results from three flights were published by Schueler et al. (1990), and results from a fourth flight of the same system in Mauersberger et al. (1993). More recently the results from a second series of flights were reported in Krankowsky et al. (2000) and Mauersberger et al. (2001). This data showed that stratospheric 03 is mass independently enriched in 180 and 170; see Table 18.1. The relationship between 5o03 and 4903 is complex; at times the enrichments are nearly equal while at other times 4903 enrichments are significantly less than those in 5o03. While there is a significant amount of variability in the data, the enrichments measured with the sample return system never approach the very high enrichments measured with the in situ system (Mauersberger, 1981; 1987). Mauersberger et al. have attempted to resolve this discrepancy by stating that the older, high values of 5o03 are suspect, and should thus be disregarded (2001). The source of the errors has not yet been identified, however. There have been a variety of ground, atmosphere, and space-based studies of the isotopic composition of stratospheric 03 using infrared emission and absorption spectroscopy. Optical spectroscopy has an advantage over mass spectroscopy in that it is able to distinguish between the different isotopomers of 03, i.e., 160160180 from 160180160, but a disadvantage in that the measurement uncertainties are generally large. Rinsland et al. (1985) reported column enrichments of 50 + 50%o in the symmetric isotopomer and 110 _+ 110%o in the asymmetric isotopomer using ground based FTIR solar absorption measurements. Abbas et al. (1987), using a balloon-borne far infrared spectrometer, reported a larger enrichment in the symmetric isotopomer (690 +_200%0 at 33 km for example) than in the asymmetric isotopomer (200 _+200%0 at 33 km for example). This was a surprising result considering the laboratory (Anderson et al., 1989) and atmospheric data (Rinsland et al., 1985) which indicated that the enrichment was primarily in the asymmetric species. This result is also in disagreement with subsequent measurements of stratospheric 03. Goldman et al. (1989) used a balloonborne FTIR absorption spectrometer to measure isotopic enrichments in the 03 column above 37 km. The first flight, over Fort Sumner, New Mexico (34 ~ N, 104 ~ W), showed enrichments of 200 +_140%o and 400 _+180%o for the symmetric and asymmetric isotopomers, respectively. A second flight over Palistine, Texas (32 ~ N, 96 ~ W), showed slightly smaller enrichments, 160 + 80%0 and 250 _+ 120%o in the symmetric and asymmetric isotopomers, respectively. Meier & Notholt (1996) used a groundbased FTIR instrument to measure solar absorption at Spitsbergen, Norway (79 ~ N, 12~ E), and found a column enrichment of 112 + 14%o and 154 + 9%0 in the symmetric

T

a

b

l

e

1

8

.

1

-Summary

of the isotopic composition

5 0 0 3 (%o) 160180160

of atmospheric

03

(see text and original references for details)

4 9 0 3 (%o)

160160180

Total

Total

Method

Altitude

Location

Reference

t~

(km) ................................................................

-

-

-

-

-

-

-

-

-

-

Str

-

400+115"

-

-

410+70"

-

-

141.7+7.0

!0

phere

..............................................................................

balloon borne mass spec.

102+11

sample

return mass spec.

S c h u e l e r e t al., 1 9 9 0

"

33.5-29.0 30.6 37.1-29.8

-

88.6+3.6

80.1+5.7

"

-

76.6+3.0

71.8+4.8

"

-

88.6+3.3

80.5+5.2

"

-

96.3+3.5 89.5+4.1

82.3+5.5 82.5+5.2

" "

-

-

balloon borne emission

based absorption

200+140

400+180

-

-

balloon borne absorption

160+80 112+14 85+25 f

250+120 154+9 90+33 f

-

_

ground

" based absorption ,. .

100+70

150+60

-

space based absorption

-

ground

130+50

170+4

-

-

102+43

73_+60

airborne emission

- ............

91~--2 . . . . . . . . .

7i!-3 ........

collection, mass spec.

-

82+7

69+7

.

-

86+6

66+6

.

.

-

90+4

78+5

.

.

enrichment

observed

measurements.

during

IR IR

.

Sweden

Krankowsky

28.9-20.9

France

"

30.7-25.6

Sweden

"

31.8-23.3 33.2-25.9

France "

" Mauersberger

column

Arizona

Rinsland

33 k m

Texas

A b b a s e t al., 1 9 8 7

IR .

> 37

.

New

Mexico

Texas Norway .

25-41

global

column

California

18-37

various

~,,L.

o

.

22.4-21.6

" column .

IR

.

" Mauersberger

column

IR

.

New Mexico "

IR

IR

based absorption

TroFosphere

* Maximum

"

" "

200+200*

f Night-time

33.9-27.5

86.5+18

690+200*

-

1987

83.0+8.0 110+29

ground

-

Mauersberger,

85.0+14

-

-

1981

"

91.0+5.0 158+28

-

.... - .............

Texas

-

110+110

-

Mauersberger,

42.5

-

50+50

-

32

"

e t al., 1 9 9 3 e t al., 2 0 0 0

e t al., 2 0 0 1

e t al., 1 9 8 5

Goldman

> 9

e t al., 1 9 8 9

" Meier & Notholt,

m" t~

1996

. I r i o n e t al., 1 9 9 6 " Johnson

e t al., 2 0 0 0 a

t9

..............................................................................

.

Heidelberg

Krankowsky

.

La Jolla

Johnston

.

.

Pasadena

"

.

.

New

"

.

ground

flight, see original reference for enrichment

level

profile.

9 N O

Mexico

e t al., 1 9 9 5

& Thiemens,

1997

O

382


and asymmetric isotopomers, respectively. These authors also used lunar spectra to measure the enrichment within the polar vortex during the polar night. These results show slightly lower enrichments, 85 + 25%0 and 90 + 33%0 for the symmetric and asymmetric isomers, respectively. These results are interesting in that the lower enrichments are consistent with the known temperature dependence of the 03 formation process (Morton et al., 1990). However, during the polar night no 03 can be formed, so the significance of this correlation between enrichment and temperature is unclear. The fact that the difference between symmetric and asymmetric enrichments is lower in the dark is also interesting, and currently unexplained. The ATMOS (Atmospheric Trace Molecular Spectroscopy) FTIR spectrometer has also been used to measure distribution of 5003 in the stratosphere, with measurements being made both from space and from the ground (Irion et al., 1996). The results from four Space Shuttle missions show the globally averaged 5003 enrichment between 2.6 and 26 mbar (~25-41 km) is 130 + 50%0, and the enrichment for 160160180 and 160180160 are 150 + 60%o and 100 + 70%o, respectively. A series of ground-based measurements from Table Mountain, California (34.4 ~ N) resulted in an average total column 160160180 enrichment of 170 + 4%o, with no discernable seasonal variation. The Smithsonian Astrophysical Observatory far-infrared spectrometer (FIRS-2) has been used to measure the isotopic composition of stratospheric 03 from a balloon altitude of around 37 km down to 18 km. The average enrichments from 7 flights (1 flight at 68 ~ N, 6 flights between 30 and 35 ~ N) are 102 + 43%0 and 73 + 60%0 for 5003 and 4903, respectively (Johnson et al., 2000a). Every technique used to measure the isotopic composition of 03 in the stratosphere has detected a substantial and variable enrichment, see Table 18.1. There does not appear to be any correlation between enrichment and time of day, season or altitude, although the larger enrichments tend to be at the higher altitudes (Mauersberger et al., 2001). Laboratory measurements (see section 18.4, below) show that at stratospheric temperatures and pressures, stratospheric 03 should be enriched by about 8090%o, slightly less in 4903 than in 5o03. While the current understanding of isotopespecific stratospheric 03 chemistry is able to explain the bulk of the stratospheric measurements (provided the very large enrichments are disregarded as suggested by Mauersberger et al. (2001)), it is possible that there remain questions regarding the variability in both enrichment magnitude and in ~170/~180 values. Additional measurements of the relative reaction rates of other isotopically substituted species, particularly those of 170, along with their parameter dependencies, will be important. In contrast, the isotopic enrichment of 5003 and 4903 in the troposphere is close to what is expected based on laboratory measurements (see section 18.4, below). Two groups have used cryogenic collection techniques to measure the isotopic composition of ground level 03 (Johnston & Thiemens, 1997; Krankowsky et al., 1995). The operational premise behind the collection is the same as discussed above for the stratospheric collection system, however, the requirements of the system are more

Mass Independently Fractionated Ozone in the Earth's Atmosphere and in the Laboratory

383

extreme due to the fact that mixing ratios of 03 in the troposphere are about 100 times lower than in the stratosphere. The Krankowsky et al. (1995) collection method utilizes a preliminary trap at 77 K for H20, CO2, and N20, followed by the 03 trap at a temperature of 54.9 + 0.1 K and a pressure of about 6 mbar. Following the collection, the 03 is converted to 02 and separated from Xe, prior to analysis on a Mattauch-Herzog magnetic instrument (Krankowsky et al., 1995; Stehr et al., 1996). Johnston & Thiemens (1997) collected 0 3 at a temperature of 55.0 + 0.1 K and a pressure 7.67 mbar (5.75 Torr). At 55 K the vapor pressure over crystalline 03 is lx10 -u bar (8x10-9 Torr) (Hanson & Mauersberger, 1986). To prevent 02 from condensing along with 03, the 02 partial pressure must be kept below 2.0 mbar (1.5 Torr), its vapor pressure at 55 K. At a total pressure of 7.67 mbar, the partial pressure of 03 will be higher than ~1x10-10 bar (lx10-7 Torr), and the partial pressure of 02 will be 1.6 mbar (1.2 Torr). Thus, by maintaining a total pressure of less than 7.67 mbar in the 03 trap, 03 should trap effectively, without any concomitant trapping of 02. In this temperature and pressure regime, 03 and Xe (87 ppbv in air) are the only atmospheric components collected in the trap. The 03 trap is preceded on the collection system by four liquid N2 traps at 77 K to remove H20, CO2, and N20. Following the collection, the 03 trap is warmed to ~82 K and the 03 and Xe recondensed on molecular sieve at 77 K. Following the transfer, the sample is allowed to warm to room temperature, quantitatively converting the 03 to 02 for mass spectroscopic analysis using a Finnigan MAT 251 isotope ratio mass spectrometer. The fact that Johnston & Thiemens (1997) did not separate Xe from the 0 3 derived 02 prior to mass spectroscopic analysis is a potentially important difference between their method and that of Krankowsky et al. (1995), who did perform this separation. There is some evidence that the presence of Xe can enrich the measured ~180 values by more than 10%o (S. Chakraborty, unpublished data, 1998). However, as discussed below, the mean values reported by Krankowsky et al. (1995) fall within the range of values reported by Johnston & Thiemens (1997). The effect of Xe remains to be quantified. Both groups reported the enrichment of tropospheric 03 relative to air 02 (6180 23.5%o, ~170 = 12.2%o V-SMOW). Krankowsky et al. (1995) reported mean enrichment values of 91 + 2%0 in 5003 and 71 + 3%o in 4903 (2 standard deviations) for 47 measurements of ground level 03 from an urban environment near Heidelberg, Germany (49 ~ N, 8 ~ E). The enrichments vary by 27%o and 46%o in 6180 and 6170 respectively, but show no systematic variation. The scatter is statistical in both axes with variances compatible with statistical errors. Johnston & Thiemens (1997) reported the isotopic composition of ground level 03 from three environments in the western United States: La Jolla, California (33 ~N, 117~ W), Pasadena, California (34 ~ N, 118~ W), and White Sands Missile Range (WSMR), New Mexico (32 ~ N, 106 ~ W). The mean values of the enrichments for 5003 and 4903, respectively, in La Jolla (n=29) were 82 + 7 and 69 + 7; for Pasadena (n=6) 86 + 6 and 66

384


+ 6, and for WSMR (n=7) 90 + 4 and 78 + 5, where the stated uncertainty is the standard deviation in the measurements. Significant isotopic variability was observed at each location, in addition to potentially important differences between the sampling locations. While the measured isotopic variability in ground level 03 showed no correlation with 0 3 or NOx mixing ratios, meteorological parameters, or time of da35 there did appear to be a correlation between the pattern of isotopic fractionation and degree of photochemical control over the local 03 budget at each sampling location (Johnston & Thiemens, 1997). Before these differences can be really understood, a much larger data set is required, both of the isotopic composition of 03, as well as many other species that could be interacting with 03. The MIF originating in atmospheric 0 3 has been shown to work its way into a wide variety of other important atmospheric species, including CO (R6ckmann et al., 1998a), CO2 (Johnston et al., 2000; Yung et al., 1997), 0 2 (Luz et al., 1999), H 2 0 2 (Savarino & Thiemens, 1999a), and sulfate deposits on the Earth's surface (Bao et al., 2000b). A recent review by Thiemens et al. provides a thorough review (2001). Lyons utilized a photochemical equilibrium model to demonstrate that the mass independent isotopic composition of 0 3 c a n be transferred to HO2, NO2, and a variety of other species (2001). 18.4 - The origin of mass independent enrichment in 03: Experiment and theory

The fractionations associated with the formation and decomposition of 03 have undergone intensive experimental and theoretical scrutiny in a variety of laboratories since the mass independent enrichment in 03 was first reported (Thiemens & Heidenreich, 1983). Initially~ the mass independent enrichment in 03 was believed linked to the 02 dissociation process (Thiemens & Heidenreich, 1983). Optical shielding by the major isotopic species 1602 will result in the preferential dissociation of 170160 and 180160 with subsequent formation of isotopically heavy 03. As pointed out by Kaye & Strobel (1983) and Navon & Wasserburg (1985), however, this mechanism cannot be responsible because the isotopic exchange between O and 02 (equation [18.4]) is significantly faster than 03 formation. 180 + 1 6 0 1 6 0 160 + 1 6 0 1 8 0

[18.4]

Thus, any increase in 170 or 180 atom concentrations relative to 160 will be removed before they can be incorporated into the 03 molecule. This exchange is mass dependent and, based on the value for the reduced partition function, resets the isotopic composition of the O(3P) to 6180 = -81.0%o at 298 K, relative to the 02 (Urey, 1947). Since the initial experiments in which 02 was dissociated via electron impact in the radio frequency region (Thiemens & Heidenreich, 1983), experiments in the microwave region (Bains-Sahota & Thiemens, 1987), and with UV light (Morton et al., 1990; Thiemens & Jackson, 1987; 1988; 1990) have all produced 03 with A170 ~ 0, provided the formation occurs in the gas phase. At low pressures (< ~3 Torr), wall effects become important; 03 formed under these conditions is mass dependently depleted in heavy isotopes (Bains-Sahota & Thiemens, 1987; Morton et al., 1990).


385

Using a "photolysis/recycling" technique, Mauersberger and coworkers demonstrated that the mass independent enrichment in 03 arises during the O(3P) + 02(3~g) (ground state), recombination reaction (Mauersberger et al., 1993; Morton et al., 1990). These experiments show that the enrichment is not tied to chemistry of excited states, thus eliminating a mechanism of Valentini and coworkers that involved nonadiabatic collisions between different electronic states of symmetric linear molecules (Valentini, 1987; Valentini et al., 1987). While not relevant for explaining MIF in 03, the mechanism observed by Valentini and co-workers remains an interesting symmetry-induced isotope effect for electronically excited species. Bates considered the 0 3 isotope anomaly in several publications. In the first, symmetry numbers were used to conclude that the O + 02 recombination reaction is faster for the substituted than the unsubstituted 03 (Bates, 1986). Anderson & Kaye (1987) pointed out that isotopic exchange and recombination reactions share a common energized collision complex, which can dissociate to products that are isotopically distinct from the original reactant, thus invalidating Bates' contention. It was then suggested that the role of symmetry in the 03 + energized complex is linked to the process of energy randomization and dissociation (Bates, 1988). However, this mechanism produced a pure 180160160 enrichment, which is inconsistent with experimental observations (Anderson et al., 1989). To accommodate this deficiency, an intermolecular "flip" was invoked to rearrange the metastable 03 + complex (Bates, 1990). This rearrangement occurs during a bond stretch that takes place before the excess energy of the 0 ~ 0 2 association is completely randomized. The bond to the terminal atom is stretched sufficiently to initiate bond formation with the other 03 terminal atom, thus converting 160160180 into 160180160. A problem with this mechanism is that in order to produce the enrichments observed in the laboratory (~85%o), the randomization frequency for the molecular rearrangement must be on the order of 1013 s -1 (Bates, 1990). This is equivalent to the fastest vibrational frequency, so is physically unrealistic. When a more plausible randomization frequency was used, the predicted enrichment increased to ~430%o. Experimental artifacts were invoked to explain this discrepancy between model prediction and observation (Bates, 1990). As pointed out by Thiemens (1992), however, experimental artifacts are insufficient to explain these differences. A mechanism involving highly vibrationally excited 0 2 w a s developed by Miller et al. (1994) and Houston et al. (1996) to address both the 03 deficit problem (the fact that model predicted 03 concentrations in the upper stratosphere were lower than observed 03 concentration levels, until satellite observations were revised downward and the rate of OH+C10 was measured to be faster than previously believed) and the large isotope enrichments in stratospheric 03. In this scheme, (1) a fraction of 03 photodissociation events produce 02 (v a 26) + O, (2) the 02 (v >__26) reacts with ground state 02 to produce 03 + O, and (3) each O atom recombines with 02 to form 03. Step (1) is more likely for heavy 03 than for normal 03, and since one of the original 03 oxygen atoms is incorporated into a new 03 molecule through step (2), heavy oxygen is distilled into the 03 pool. As 03 is cycled by photodissociation many times (ranging

386


from about 50-400 times at altitudes of 50-32 km) before destroyed by other processes, even a very small enrichment factor can be strongly amplified. This mechanism can explain only about 3% of the heavy 03 enrichment in the stratosphere, however. Another theory regarding the source of the mass independent enrichment in 0 3 is that it is related to the symmetry of the different isotopically substituted 03 molecules, 160160160, 170160160, and 180160160 (Heidenreich & Thiemens, 1986). The rate of a three-body recombination reaction is very sensitive to the lifetime of the metastable species, 03 + in this case. The greater this lifetime, the greater the probability that the metastable molecule will be collisionally quenched by the bath molecule (M). The fact that 170160160 and 180160160, both Cs, will have a slightly longer lifetime than the C2v species, 160160160, due to the appearance of alternate rotational states in the asymmetric species is taken into account by standard 3-body reaction rate theory. Thus, for this lifetime argument to be valid, some additional factor must be responsible. Additionally, the lifetime mechanism requires the isotopic enrichment to be independent of pressure until the falloff region of the O + 02 recombination is approached. Measurements of this falloff curve show that the deviation from low pressure, third order behavior does not occur until about 8 atmospheres, and that the high pressure limit is not reached until greater than 400 atmospheres (Croce de Cobos & Troe, 1984; Hippler et al., 1990). Laboratory measurements show that the isotope enrichments are quite sensitive to pressure and do not obey the kinetics of the observed falloff curve. The room temperature enrichment reaches a maximum at approximately 10 Torr, with 5 1 8 0 - 130%o and 5 1 7 0 - 100%o, and begins to fall off quickly at pressures greater than ~ 1/2 atmosphere (Morton et al., 1990; Thiemens & Jackson, 1988). By 56 atmospheres the enrichment has disappeared entirely, and between 56-87 atmospheres the 03 shows a small mass dependent depletion in 180 and 170 (Thiemens & Jackson, 1990). It is therefore unlikely that theories involving differential lifetimes of the metastable state are the source of the observed mass independent fractionations. As discussed below, however, Marcus and co-workers have recently proposed a mechanism by which a slightly higher density of states in the asymmetric vibrationally excited 03 molecule can reproduce the observed MIFs in a way consistent with observed pressure dependencies (Gao & Marcus, 2001; 2002; Hathorn & Marcus, 1999; 2000). It is established that absolute rate coefficients for 0 3 formation are strongly influenced by the bath gas composition, and vary by as much as one order of magnitude, depending on the quenching efficiency of the third body (Steinfeld et al., 1987). Guenther et al. have shown, however, that the composition of the third body does not alter the relative rates of the isotope specific formation reactions (Guenther et al., 2000). While the enrichment mechanism based on differential lifetimes appears to be invalid, experimental observations and theoretical models continue pointing to the importance of molecular symmetry. Anderson et al. (1989) used a tunable diode laser system to measure the isotopomeric distribution in 03 produced in an electric discharge and found that 80% of the enrichment is in the asymmetric molecule, while

Mass IndependentlyFractionatedOzonein the Earth's Atmosphereand in the Laboratory

387

only 20% is in the symmetric molecule. A purely statistical argument predicts that only 66% of the enrichment should reside in the asymmetric isotopomer. In natural oxygen, the abundances of 180 and 170 are small enough that concentrations of multiply substituted species, such as 180160180 and 180170160, are too low to contribute to the observed enrichments. However, by forming 03 from 02 that is artificially enriched in 180 and 170, all isotopomers of 03, from 4803 to 5403, can be formed. Mauersberger and co-workers produced 03 via the photolysis/recycling method and found that, relative to 4803, the symmetric isotopomers (170170170, 180180180) are slightly depleted (Mauersberger et al., 1993; Morton et al., 1989). This result agrees with the 03 isotopic formation theory developed by Kaye & Strobel (1983), that predicts a small mass dependent depletion. For the asymmetric isotopomers, large enrichments are observed. A 180%o enrichment is seen in 160170180, and approximately 2/3 of that in the other asymmetric species. These results seem to indicate that the enrichment lies in the asymmetric molecules alone, and that the symmetric molecules show the depletion predicted from recombination, theory (Kaye & Strobel, 1983; Kaye, 1986). Gellene (1996) applied a theory of symmetry induced kinetic isotope effects (SIKIEs) to the O + 02 recombination reaction and was able to quantitatively explain the enrichment and depletion pattern in isotopomers 4803 - 5403 observed by Mauersberger and co-workers (Mauersberger et al., 1993). The general theory (Gellene, 1992) was developed from observations of a variety of ion-molecule SIKIEs, including the formation of 04 + (Griffith & Gellene, 1992), He2 + (Gellene, 1993), and (CO2)2 + (Yoo & Gellene, 1995). In the case of 03, symmetry restrictions arise for homonuclear diatomics (i.e., 160160 and 180180) involved in the O + 02 collision because only a fraction of their rotational states correlate with those of the corresponding 03 molecule. In contrast, all of the rotational states of the heteronuclear (i.e., 160180) 02 molecules correlate with those of the resulting 03 molecules. While the theory developed by Gellene and co-workers remains relevant to other SIKIEs, the rate coefficient measurements of Mauersberger et al. (1999), (discussed below) show that this mechanism is not relevant to 03. A number of studies have focused on characterizing the adiabatic energies, binding energies, lifetimes and geometries of the low-lying metastable electronic states of 03 in an effort to isolate the source of the mass independent enrichment (Anderson et al., 1992; 1993; 1995; Anderson & Mauersberger, 1995). The motivation being that these metastable states may be populated during 03 formation, and subsequent symmetry selective pathways to the ground state could then enhance the asymmetric molecules over the symmetric ones. The spectroscopy of 03 is very complex, eight electronically excited states with adiabatic energies below 8 eV are known, at least half of which lie below 2 eV. Despite these efforts, no connection between the isotope effect and the metastable states has been made to date.

388


Kinetic studies have provided a great deal of information about the mechanism of 03 formation and recently, the source of the MIF. Studies using 160 and 1802 showed that 03 is formed via a simple end on addition, rather than through a more complicated insertion process (Larsen et al., 1991). Recent work by Mauersberger and coworkers indicates that the MIF occurring during the O + 02 reaction is determined by the nature of the collision rather than by molecular symmetry. Anderson et al. (1997) used isotopically pure 02 ( 1 6 0 1 6 0 and 1 8 0 1 8 0 ) , a filtered deuterium lamp, and a molecular beam mass spectrometer gas analyzer to investigate the kinetics of four 03 formation channels: 160 180 180 160

+ + + +

160160 160160 180180 180180

+ M --* 160160160 + M + M --* 180160160 + M + M --* 180180180 + M 4- M --* 160180180 + M

[18.5] [18.6] [18.7] [18.8]

Reactions [18.5 - 18.7] all have similar rates of formation, while [18.8] has a rate coefficient 50% faster than the first three. If molecular symmetry was the controlling factor, reactions [18.6] and [18.8] should display equal rate constants. Using a slightly different technique, Mauersberger et al. (1999) confirmed the results of Anderson et al. (1997) and extended the studies by including reactions involving 170. By combining these measured rate coefficients with other laboratory observations (i.e., Anderson et al., 1989 and Mauersberger et al., 1993), it is possible to infer information about the rate coefficients, which cannot be directly measured For example, atmospheric 5003 is formed by reaction [18.6] and by: 160 + 160180 + 160 + 160180 +

M M

--* 1 6 0 1 6 0 1 8 0 + M ~

160180160 + M

[18.9] [18.10]

The rate coefficient determined by Mauersberger et al. (1999) indicates that [18.6] cannot be responsible for the large enrichment in atmospheric 03. Using a diode laser Anderson et al. (1989) showed 80% of the enrichment is carried in the asymmetric 5003 isotopomer. Thus, the rate coefficient for reaction [18.9] must be ~43% faster than the rate coefficient for reaction [18.5] (Mauersberger et al., 1999). In describing the observed pattern in isotopic enrichment, a collision between a light atom and a heavier molecule will result in a rate constant that is higher than the rate constant from reaction involving only one isotope, while a collision between a heavy atom and a lighter molecule will not yield the rate enhancement (Mauersberger et al., 1999). Janssen et al. (1999) used tunable diode laser and mass spectrometry to measure the relative formation rate coefficients of each of the four channels that can form 5003 and 5203 from mixtures of 160 and 180. The results, consistent with those of Anderson et al. (1997) and Mauersberger et al. (1999), show that molecular symmetry plays no apparent role in the 03 formation process, and that the isotopic enrichment in 5003 arises from an enhanced rate of one formation reaction [18.9] (Janssen et al., 1999).


389

Following up on the earlier work of Hathorn & Marcus (1999; 2000), Gao and Marcus developed a statistical (RRKM)-based model with a hindered-rotor transition state that explains most of the laboratory observations (Gao & Marcus, 2001; 2002). The theory assumes (1) an "~l-effect" which can be thought of as a small deviation from the statistical density of states for the symmetric versus the asymmetric isotopomers, and (2) weak collisions in the deactivation of the vibrationally excited 03 (Gao & Marcus, 2002). A partitioning effect, arising from small differences in zero-point energies of the two exit channels of dissociation of an asymmetric 03, controls the ratios of the recombination rates. These small differences are magnified into large differences in numbers of states in the two competing exit channel transition states. In isotopically unscrambled systems this second effect dominates, leading to the large, unconventional mass-dependent effects in the rate constants reported by Mauersberger and coworkers (Anderson et al., 1997; Janssen et al., 1999; Mauersberger et al., 1999). In contrast, in the scrambled systems the partitioning factor disappears exactly, leaving the ~l-effect responsible for the observed fractionations. The calculated isotopic enrichments are consistent with a wide variety of laboratory experiments examining 03 formation (Bains-Sahota & Thiemens, 1987; Heidenreich & Thiemens, 1986; Morton et al., 1990; Thiemens & Heidenreich, 1983; Thiemens & Jackson, 1987; 1990; Yang & Epstein, 1987). 18.5 - Conclusions

Ozone carries a large and variable mass independent isotope fractionation, which has been studied extensively in the atmosphere and in the laboratory for more than 20 years. Measurements of MIF in tropospheric 03 (Johnston & Thiemens, 1997; Krankowsky et al., 1995), and a recent assessment of MIF in stratospheric 03 (Mauersberger et al., 2001) indicate that the atmospheric variability is largely in agreement with laboratory measurements. The recent work of Gao & Marcus (2001; 2002) has provided a theoretical model to understand a wide variety of laboratory results. The development of this model is timely, as it will undoubtedly contribute to the interpretation of the wide variety of natural species that have now been shown to possess a mass independent fractionation.

Acknowledgements The National Science Foundation and NASA is gratefully acknowledged for their support for the many facets of work reported in this paper.


CHAPTE R 19 Site-specific Nitrogen Isotope Analysis in N20 by Mass Spectrometry Sakae Toyodal,4* & Naohiro Yoshidal,2,3,4** Department of Environmental Chemistry and Engineering, Frontier Collaborative Research Center, and Department of Environmental Science and Technology, Tokyo Institute of Technology, 4259 Nagatsuta, Midori-ku, Yokohama 226-8502, Japan, and SORST, Japan Science and Technology Corporation (JST), Kawaguchi, Saitama, Japan 1 Department of Environmental Chemistry and Engineering, Tokyo Institute of Technology 2 Frontier Collaborative Research Center, Tokyo Institute of Technology 3 Department of Environmental Science and Technology, Tokyo Institute of Technology 4 SORST,JST (Japan Science and Technology Corporation) e-mail'* [email protected]; ** [email protected]

Abstract A novel method has been developed for site-specific nitrogen isotope analysis in N20, which is an asymmetric linear molecule and an important trace gas in the atmosphere. The method makes use of mass analyses of the molecular (N20 +) and fragment (NO +) ions of N20 in an isotope-ratio mass spectrometer equipped with a special ion collector system. The fragmentation of N20 in the electron impact ion source is stable, and the precision of isotope ratio measurements of the fragment ion is better than 0.1%o for pure N20 samples introduced from a conventional dual-inlet system. Although the observed isotope ratio of the fragment ion is affected by rearrangement reactions in the ion source, a correction can be applied using an experimentally determined rearrangement fraction. This technique has been shown to supply useful information on environmental N20 when it is coupled with a continuous-flow technique to accommodate highly sensitive analyses.

19.1 Introduction N20 is one of the important trace gases in the atmosphere that affect radiative balance and atmospheric chemistry. Since it has a long lifetime of ca. 120 years, and absorbs infrared radiation, it is a major greenhouse gas, following CO2 and CH4 (IPCC, 2001). In the stratosphere, it is decomposed by ultraviolet light, and chemically active species (atomic oxygen and nitrogen oxides) are produced, which then react with ozone. The tropospheric concentration of N20 is increasing by 0.2 - 0.3% per year presumably due to human activity, but the global budget and cycle of N20 has not been well resolved because it has a variety of sources and sinks. Isotopic studies of N20, which are essential to the understanding of its origins, chemical/physical processes, and fate, have not been extensively conducted, while there have been a num-

Site-specific NitrogenIsotopeAnalysisin N20 by Mass Spectrometry

391

ber of extensive works based on concentration or flux measurements in various fields. One of the reasons that it is difficult to analyze isotope ratios of N20 is its low concentration in nature" for example, ~50 L of atmospheric or water sample are required for the conventional dual-inlet isotope analysis. However, recent developments in instrumentation and analytical techniques allow us to measure isotopic ratios of N20 with a smaller sample size and a less time-consuming procedure. In earlier studies, N20 was converted to N2 and 0 2 / C O 2 / H 2 0 , and then introduced into the conventional dual-inlet system of an isotope-ratio-monitoring mass spectrometer (Yoshida & Matsuo, 1983; Wahlen & Yoshinari, 1985; Yoshinari, 1990; Kim & Craig, 1990, Thiemens & Trogler, 1991). Later, direct introduction of N20 into the ion source of the mass spectrometer was performed, and found to be applicable to obtaining both nitrogen and oxygen isotope ratios simultaneously, although separation and purification of N20 from isobaric CO2, which is much more abundant in the environment, are critical for the accuracy of the measurement (e.g. Kim & Craig, 1993; Tanaka et al., 1995; Rahn & Wahlen, 1997). This technique can be combined with a continuous-flow technique for on-line analysis of smaller size samples (Yoshinari et al., 1997; Dore et al., 1998). Although previous studies for the isotopic characterization of N20 were based on the bulk element contained in the molecule (nitrogen and oxygen), N20 has more isotopic information owing to its asymmetric molecular structure (N-N-O), i.e. the intramolecular distribution of nitrogen isotopes. Since formation and cleavage of N-N and N-O bonds of N20 take place in naturally occurring processes, an intramolecular site preference for nitrogen isotopes is expected (e.g. Yung & Miller, 1997). We (Toyoda & Yoshida, 1999) attempted to monitor the isotope ratio of the fragment ion of N20, as well as the molecular ion formed in the ion source of a mass spectrometer, and reported for the first time that the site-specific nitrogen isotope analysis in N20 can be performed with high precision. Brenninkmeijer et al. (1999) also reported a mass spectrometric technique basically the same as ours, and an infrared spectroscopic one has been developed (Esler et al., 2000a; Uehara et al., 2002), although the precision of the latter technique is not as high as that of the mass spectrometric technique at present. In this manuscript, a notation for the intramolecular distribution of isotopes is briefly discussed, and instrumentation, experimental techniques, and applicability are described.

19.2 Notation for the isotopomers of N20 Let us define an isotopomer as one of a set of molecules that contains isotopically substituted atoms. The number of isotopomers of a certain compound depends on the number of elements that constitute the compound, the number of isotopes of each element, and the symmetry of the molecular structure. In the case of N20, 12 isotopomers can exist, but only the following five are significant at the natural abundance level: 14N14N160, 15N14N160, 14N15N160, 14N14N170, and 14N14N180.

392

Chapter 19 - S. Toyoda & N. Yoshida

If the nitrogen atoms at the center and end positions are denoted as N~ and N~, respectively, the nitrogen isotope ratio for each site is expressed as the isotopomer ratio as follows (Toyoda & Yoshida, 1999): 15Ra = 14N15N16 O / 14N14N16 O 15R~- 15N14N16 O / 14N14N16 O

[19.1] [19.2]

The conventional nitrogen isotope ratio, which is hereafter referred to as the bulk isotope ratio, corresponds to the average of the two isotopomer ratios, since it does not distinguish N-isotopomers: 15Rbulk = (15Rc~+ 15R~) / 2

[19.3]

In delta notation 615Ni- {15Ri / 15Ri (std)- 1} 1000

[19.4]

for i = a, [3, and bulk, where std means the standard or reference material. If atmospheric N2 is chosen for the standard, equations [19.3] and [19.4] lead to the following equation: 515Nbulk- (615NR + 615N~ ) / 2

[19.5]

Note that equation [19.5] does not strictly hold when N20 is used as the reference (Toyoda & Yoshida, 1999).

19.3 Experimental methods 19.3.1 Natural and 15N labeled N20 samples Commercial N20 in a cylinder (Showa Denko Co., Ltd., purity > 99.999%) was used as the working standard gas without further purification. The isotopomer/isotope ratios for the gas were determined to be -4.0, 0.2, and 23.3%o for 515N% 515Nf~, and 6180, respectively, by our nitrogen-isotopomer calibration technique (Toyoda & Yoshida, 1999) and by conversion of N20 to N2 and CO2 in the presence of graphitic carbon (after Yoshinari, 1990), where the standards for nitrogen and oxygen are atmospheric N2 and SMOW (standard mean ocean water), respectively. Labeled samples of 15N14NO and 14N15NO, whose 15N purity at the labeled and unlabeled positions are > 99% and 0.37% (natural abundance level) by atomic fraction, respectively, were purchased from Icon Services Inc., USA. From each of the materials three samples were prepared by static dilution with the working standard so that they have different 15N concentrations at the labeled position. Their 15N concentration was calculated from the manometrically determined dilution factor: 0.582, 0.758, and 1.06% for a-labeled N20 and 0.524, 0.749, and 1.09% for the 13-labeled N20. Accuracy was estimated to be better than 1% (relative error against calculated 15N concentration), although uncertainty of the 15N content of the working standard gas was not

Site-specific Nitrogen Isotope Analysis in N20 by Mass Spectrometry

393

included because it canceled out in the relative measurement between the labeled and working standard gases.

19.3.2 The principle of site-specific nitrogen isotope analysis for N20 by mass spectrometry In conventional mass spectrometric isotope analyses for CO2, N2, H2, etc., molecular ions produced by electron-impact ionization are separated and detected. Fragmentation of the molecule also takes place in the ionization chamber of the mass spectrometer, although their fraction is smaller than for molecular ions. For N20, the formation ratio of N20 + to NO +, one of the fragment ions, is about 3 91. Considering the asymmetric structure of the N20 molecule, one would expect that the nitrogen atom in NO + originates mostly from the R-N atom of the parent molecule (there is a minor contribution from the f~-N atom as shown later), while N20 + contains both c~and ~-nitrogen atoms. Therefore, mass analyses of both ions make it possible to deduce the intra-molecular distribution of nitrogen isotopes. In this study a sector type, isotope-ratio mass spectrometer (Finnigan MAT 252, Thermo Quest K. K.) was used throughout the experiments. The multi-collector system was specially designed and modified to measure the isotope ratios for both molecular and fragment ions of N20 without changing the Faraday collector cup configuration and the amplifiers associated with the cups (Figure 19.1). The modified collector system consists of five cups, three of which (cups no. 1, 3, and 5) have larger collector slits than the standard cup in order to allow the measurement of other gases, such as N2, 02, etc., as well as N20 and NO. In the N20+-measuring mode, cups no. 1, 2, and 4 were used to monitor ions of m / z 44, 45, and 46, respectively, and in the NO+-measuring mode, cups no. 1, 3, and 5 monitor m / z 30, 31, and 32, respectively. Taking account of the isotopomers of N20 that contribute to the m / z 44, 45, and 46 molecular ion beams and the m / z 30, 31, and 32 fragment ion beams in the mass spectrometer, the 45R, 46R, 31R, and 32R isotope ratios for molecular and fragment ions relative to the most abundant ones, are given by: 45R = 46R = 31R = 32R =

15R~ + 15R~ + 17R 18R + (15R~ + 15R~) 17R +15Ra15R~ 15R~ + 17R 18R +15R~ 17R

[19.6] [19.7] [19.81 [19.9]

Since the observed reproducibility of 32R is much worse than for the other ratios, probably because of an unstable instrumental background level of 02 and the 02 + produced from ionizing reactions of N20, we assumed the mass-dependent fractionation of oxygen contained in the sample N20 and used the following equation (Craig, 1957) instead of equation [19.9] to obtain 15R~, 15R~, 17R, and 18R from 45R, 46R, and 31R of the sample" 18R / 18R(std)= {17R / 17R(std)}2

[19.10]

394


Figure 19.1 - Schematic diagram of the measurement of N20 isotopomers on a mass spectrometer (A modified version of Figure 1 in Toyoda & Yoshida, 1999). IC: ionization chamber, C: cathode, ET: electron trap, EP: extraction plate, AN: analyzer, MC: multi-collection system.

This assumption is not valid when a mass-independent oxygen isotope fractionation occurs. Such an 170 anomaly is observed in stratospheric and tropospheric N20 (Cliff & Thiemens, 1997; Cliff et al., 1999). However, the magnitude of the anomaly is about 1%o for 6170 and equation [19.10] brings about an overestimate of about 0.1 and 0.05%o for 615N~ and 615Nbulk respectively. These differences are small compared with the current analytical precision for atmospheric N20 by a continuous-flow technique (Yoshida & Toyoda, 1999), although corrections will be needed if the 170 a n o m aly is independently determined for the same sample. 19.4 R e s u l t s a n d d i s c u s s i o n

19.4.1 Fragment pattern of N20 and precision of the measurement The critical point of the mass fragmentation analysis is that the fragment ratio, which refers to the ion beam intensity ratio of fragment ion to molecular ion, and isotope fractionation in the fragmentation, should be constant in a series of measurements in which sample and reference are introduced alternately into the ion source. Figure 19.2 shows the variability of the fragment ratio when the electron energy is


395

Figure 19.2 - Relationship between the fragment ratio of N20 and ion source conditions. The parameter'ext.' is the scale on the potentiometer that modulates the extraction plate voltage. A larger number corresponds to a lower potential difference between the ionization chamber and the extraction plate.

intentionally changed over a wide range. Typically, experimental parameters of the ion source of the mass spectrometer were optimized so that high linearity (isotope ratio is not dependent on sample pressure) can be obtained, although this tuning does not give the maximum sensitivity that the machine can produce. For the instrument used in this study, the optimum setting values were between 0 and 5 for the extraction plate voltage parameter, and between 60 and 100 eV for the electron energy. The fragment ratio was about 0.3 under typical conditions, and its variability was within _+0.2% over several hours. The effect of the total pressure, or matrix effect, in the ionization chamber was evaluated by introducing He continuously using the interface for isotope-ratio-monitoring gas chromatography-mass spectrometry. As shown in Figure 19.2, pressure

396


dependence of the fragment pattern was small under optimum conditions. The precision of the ten sets of measurements and variation among nine independent measurements between two identical samples obtained using the dual inlet system are shown in Table 19.1. For the molecular ion, precision was nearly equal to that for the other gases such as CO2 and N2, although in the case of the fragment ion it became a little worse because the beam intensity of the latter was weaker than that of the former.

19.4.2 Rearrangement reaction of N20 In order to check the stability of isotope fractionation during electron impact fragmentation of N20, 15N-labeled samples were analyzed relative to the working standard. The analytical precision of the samples was better than 0.2 - 1.2%o, depending on the 15N concentration. Bulk nitrogen isotope ratios of the six samples observed in the N20-measuring mode agreed well with calculated values (~15Nbulk(obs) - 1.01615Nbulk(calc) - 3.78, R = 0.99998), indicating that there is no isotope fractionation in the sample preparation and that the fractionation factor for the formation of N20 § is constant within a range of 1000%o (Toyoda & Yoshida, 1999). However, the isotopomer ratios obtained from the fragment NO measurements were not identical with the calculated values, although a linear relationship was found. As shown in Figure 19.3, 615N was lower than the calculated value at the labeled position and higher at the unlabeled position. From the magnitude of the 6 deviation and 15N purity of the purchased material validated by near-infrared absorption spectroscopy (Uehara et al., 2002), we attributed this result to the rearrangement reaction of N20 under electron impact ionization. Here, we define "rearrangement" or "scrambling" as inclusion of the [3-N atom Table 19.1- Analytical precision of the isotopomers' determination. Experiment

Average + l o

615Nbulkair (%o)

615N%ir (%o)

-1.896• -1.895• -1.887• -1.888• -1.915• -1.921• -1.899• -1.963• -1.882• -1.905+0.025

-3.981• -4.020• -4.034• -3.980• -4.021• -4.037• -4.027• -4.009• -4.015• -4.014+0.021

615Nf~air (%o)

618OsMow (~o)

0.188• 0.229• 0.261• 0.204• 0.191• 0.195• 0.229• 0.082• 0.251• 0.203+0.053

23.308+0.038 23.313+0.040 23.336+0.045 23.343+0.024 23.342+0.038 23.297+0.062 23.341+0.033 23.219+0.024 23.330+0.019 23.314+0.040

About 2 ml STP of N20 identical with the working standard is introduced into the mass spectrometer through a conventional dual inlet system. Average of 10 set of relative measurements is used for 6-calculation and the standard deviations of molecular and fragment ion analysis for ~)15Nbulk and ~)15Ncz are listed, while square mean errors are estimated for 615Nfi (= 2 x 615Nbulk - 615Ncz).


397

Figure 19.3 - Relationship between the observed nitrogen isotopomer ratio of N20 and the calculated value (A modified version of Figure 2b in Toyoda & Yoshida, 1999). Calculated values refer to the labeled position. in the parent N20 molecule in the fragment NO + ion, irrespective of what the mechanisms are. The formation mechanisms of NO + after electron impact of N20 are considered to be spontaneous and collision induced dissociation of the N20 + (Begun & Landau, 1961) and ion-molecule reaction between O + / N + and N 2 0 (Derwish et al., 1964; Ryan, 1972), and the formation of N~O + has been observed from the ionization of highly 15N-enriched N20 (Friedman & Bigeleisen, 1950; Begun & Landau, 1961). Although the relative contributions of these reactions may depend on the geometry of the ionization chamber, gas pressure, etc., their effect on the overall isotopic fractionation of NO + should cancel out when the sample and reference gases are analyzed under the same instrumental conditions.

398


If the r e a r r a n g e m e n t fraction y is defined as the ratio of N O + b e a r i n g 13-N of the p a r e n t N 2 0 to total N O +, the o b s e r v e d isotope ratio of N O + can be e x p r e s s e d by the following e q u a t i o n 15R%bs - (1 - y) 15R~ + y 15RI3

[19.11]

F r o m e q u a t i o n s [19.3], [19.4], a n d [19.11], y is expressed as a function of the u n s c r a m bled isotopomer, b u l k isotope, a n d o b s e r v e d i s o t o p o m e r ratios" y - A(a15N~ - 815NC%bs) / [2{(1 - A) 815N~obs - 815N bulk + A815N~}]

[19.12]

where A ~ (15R~ / 15Rbulk )std

[19.13]

In the p r e s e n t s t u d y A w a s d e t e r m i n e d to be 0.998 by c o m p a r i n g the w o r k i n g stand a r d to the calibration s t a n d a r d p r e p a r e d from thermal d e c o m p o s i t i o n of NH4NO3, w h i c h forms N 2 0 w h o s e c~-N a n d f3-N atoms are d e r i v e d from nitrate a n d a m m o n i u m ions, respectively (Toyoda a n d Yoshida, 1999). U s i n g the calculated a n d o b s e r v e d i s o t o p e / i s o t o p o m e r ratios of 15N-labeled N 2 0 , y w a s e s t i m a t e d to be a b o u t 0.08, w h i c h is in g o o d a g r e e m e n t w i t h other r e p o r t e d values ( F r i e d m a n & Bigeleisen, 1950, B e g u n & L a n d a u , 1961). A l t h o u g h this p a r a m e t e r m a y d e p e n d on the i n s t r u m e n t used, the variation of y w a s f o u n d to be v e r y small u n d e r typical e x p e r i m e n t a l conditions, even if m e a s u r e m e n t s w e r e p e r f o r m e d on difTable

19.2

-

Reproducibility of the rearrangement fraction y.

...Exper!m.ent . a..........................................................Ionsourc.e..p.a!~.ametersb ........... .............................................................Va!ue....of.yC ............ ..................... Trape (V) Extractiond Electron energy (V) 1 2 3f 4 5 6 7

4.64 2.46 2.46 0.00 0.00 0.00 0.00

54.9 37.4 37.4 50.0 50.0 50.0 50.0

66.9 86.6 86.6 86.5 86.5 86.5 102.1 Average & lo

0.0803 0.0810 0.0813 0.0813 0.0805 0.0812 0.0823 0.0811+0.0006

All experiments were performed on different days except for no. 4 and 5, which were performed at a 6-hour interval on the same day. b These parameters were optimized so that both high linearity and high sensitivity could be obtained, but optimum values were different from day to day. c Calculated from eq 12 in the text. Sample was b-labeled gas (15N14NO, 15N = 0.744 at.%). d See the caption in Figure 19.2 e Potential difference between the ionization chamber and the electron trap. f Sample was introduced into the ion source with He. [A modified version of the Table 2 in Toyoda & Yoshida, 1999]. a


399

ferent days and the pressure in the ionization chamber was changed by introducing He (Table 19.2). Once the rearrangement rate is determined, the nitrogen isotopomer ratio of the sample can be obtained from the observed ratio: 615N~ = ~)15Naobs + 2y(615N%bs- 615Nbulk) / {A(1- 2y)} 615N~ = 1515Nbulk+ A(615Nbulk - ~)15Na) / (2 - A)

[19.14] [19.15]

In the above equations y is assumed to range between 0 and 0.5, otherwise one could not obtain the intramolecular nitrogen isotope distribution by mass fragment analysis.

19.5 Summary and applications of the technique Site-specific nitrogen isotope analysis for N20 has been made possible by mass fragmentation analysis on a conventional isotope-ratio mass spectrometer for a sample size at the micromolar level. The precision of the analysis is almost as high as that of the bulk isotope analysis. This technique, combined with continuous-flow analysis, reduces the required sample size to the nanomolar level, which enables us to analyze N20 in the natural environment. For instance, tropospheric N20 has been found to have the ~-site preference for 15N (~)15N~- ~115Nf3)of about 20%0, while it is as much as 90%o in the stratosphere (Yoshida & Toyoda, 2000; Toyoda et al., 2001). Measurements of the N20 isotopomers from various sources and sinks are now in progress and several results have shown that the 15N-site preference is characteristic of each production/ consumption process (e.g. Yamulki et al., 2001; R6ckmann et al., 2001; Toyoda et al., 2002; Sutzka et al., 2003). Another approach to site-specific nitrogen isotope analysis, which utilizes infrared absorption spectroscopy of the N20 isotopomers, has also been developed (Esler et al., 2000a; Uehara et al., 2002). Although the sensitivity and precision of the measurement is not as high as that of mass spectrometric determinations at present, it will serve effectively as an alternative technique in the near future, since it has some advantages over the mass spectrometric technique (e.g. non-destructive analysis, less need for purification of N20 from other species).


CHAPTER 20 Fluorination Methods in Stable Isotope Analysis Bruce E. Taylor Geological Survey of Canada, Ottawa, Ontario KIA OE8, Canada e-mail: btaylor@nrcan, gc.ca

20.1 Introduction Fluorine, is the most oxidizing element known (one of the few elements more oxidizing than oxygen), and possesses the highest electronegativity of all elements (e.g., Pauling, 1964). Consequently, fluorine or fluorine-bearing compounds have been the principal means by which to successfully facilitate the extraction of oxygen from oxygen-bearing substances for stable isotope analysis. Among most earth scientists, silicates and oxides have historically dominated the materials of interest treated by fluorination. However, fluorination techniques have also been successfully extended to other minerals, including phosphates, sulfides and selenites, in addition to elemental S, Se and U, among others). The purpose of this chapter is to review the most common fluorination methods currently applied in stable isotope geochemistry, incorporating both conventional and the more-recently applied laser-induced heating techniques, and to describe the construction and utilization of the basic apparatus used for fluorination employing fluorine (F2), or interhalogen fluorides (e.g., BrF5, C1F3) with or without the hydrogen halide, HF. It is not feasible to cover all possible variations in the details of customconstructed apparatus or of their application. Rather, this paper attempts to summarize the basic components and methods deemed reliable and functional. Although primarily intended for those not previously familiar with the fluorination techniques used in stable isotope geochemistry, I hope that this contribution will also be of use to the more-experienced. Fluorination is a robust technique for oxygen isotope analysis of silicates, oxides, and phosphates, as well as for sulfur isotope analysis of sulfides, one that is likely to continue to be used into at least the near future, whether for macro- or micro-analysis. As a result, it is important to incorporate health and safety issues in our discussion, as well as provide as much guidance as possible, based largely on our own experience, for the actual application of fluorination methods in isotope extraction.

Fluorination Methods in Stable Isotope Analysis

401

20.2 Background and overview

Early efforts by Manian et al. (1934) employing CC14 at 1000~ and producing CO as the analyte, for oxygen the isotope analysis of silicates were not encouraging, largely due to the lack of sufficient sensitivity (+ 2.5%o) to detect isotopic variation among the samples analyzed. Subsequently, two principal techniques were developed to extract oxygen from silicates and oxides: (1) reduction by graphite (carbon); and (2) oxidation by fluorine (+ HF), or one of its interhalogens. Carbon reduction of silicates and oxides was attempted before the adoption of fluorination as the method of choice, and met with some success in selected applications. Fluorination procedures dominated through the second half of the 20th century, but recent developments in automated, on-line reactors employing high-temperature furnaces coupled with continuous-flow mass spectrometric inlet techniques, have prompted some renewed interest in carbon reduction. Accordingly, a brief review of carbon reduction is also given below for completeness. Recent attempts at high-temperature pyrolysis of samples for oxygen isotope analysis (e.g., Kornexl et al., 1999a; Werner et al., 2001) indicate that replacement of silicate/oxide fluorination by carbon reduction, or a carbon-reduction related process, however tantalizing, may not be on the immediate horizon. 20.2.1 Carbon Reduction

Following the seminal work of Urey (1947) on the thermodynamic properties of isotopically-substituted compounds, several investigators sought techniques for the isotopic analysis of minerals (especially oxygen-bearing minerals) and water. Early attempts at carbon reduction of silicates produced CO, but determined the 180/16 0 ratio in one case by mass spectrometry on electrolytically recovered 02 (Manian et al., 1934), and, in the other case, by the measurement of the density of H20 formed upon reaction of the CO with H2 (Vinogradov & Dontsova, 1947). The reported uncertainties of the relative 6-values were, respectively, 2.5%o and 1.5%o. Subsequently, Adams (1949-1950) reduced quartz to CO for determination of its 180/16 0 ratio directly by mass spectrometry. However, accurate isotopic analysis was prevented by the presence of N2 and hydrocarbons, and, consequently; Adams converted the CO to CO2 by reaction with copper oxide at 455~ to facilitate cryogenic purification. Schwander (1953) published the first detailed description of the carbon reduction method, sources of uncertainty, and analytical results for various materials, from waters to different rock types. A graphite resistance furnace, mounted in vacuo and externally water-cooled, with power supplied by a 220V (primary)/6V, 3.5-4.0kW (secondary) transformer in combination with a variable transformer (secondary), permitted out gassing and carbon reduction at temperatures up to ca. 2200~ (Schwander, 1953). Isotopic measurements were made mass spectrometrically on CO, reportedly to an uncertainty in the relative &value of 0.4%o. The nominal C-reduction reaction is: Metal oxide(s) + C(s) ~ Metal carbide(s) + CO

[20.1]

Which, in the case of quartz, can be written as" SiO2(s)+ 3C(s) ~ SiC(s) + 2CO

[20.2]

402

Chapter 20- B.E. Taylor

Throughout this chapter, reactants and reaction products are assumed to be in a gaseous state during the reaction as written, unless indicated to be either liquid (1) or solid (s). Typically, sub-equal weights (ca. 60mg) of finely pulverized sample and graphite were mixed together and placed in a cylindrical, capped graphite container of ca. 0.4 cm3 volume. Following high-temperature out gassing of the furnace, and simultaneous heating of the sample to ca. 900~ the sample was inserted further into the furnace so that it out gassed at temperatures up to ca. 1600~ (measured by optical pyrometer) to remove all water and any traces of carbonate (Schwander, 1953). The temperature was then raised, and carbon reduction began generally above about 1700~ (from thermodynamic data for silicates, Clayton & Epstein (1958) estimated ca. 1727~ for quartz and 1200~ for iron oxides, and was usually complete in about 10 minutes, yielding virtually no more CO (from quartz) above ca. 2000~ The measured 6180 value was found to increase with yield, and that yields of at least 70-80% were required for reliable isotopic measurement. In cases where a silicate melt phase formed upon heating, and separated from the graphite, Schwander (1953) noted that yields were greatly reduced (10-30%) and c5180 values were too low. This was avoided by adding more graphite. Cryogenic purification using liquid air readied the sample of CO for mass spectrometry. Clayton & Epstein (1958) used a radio frequency (RF) induction furnace to heat a capped, 2g graphite crucible in vacuo surrounded by a platinum shield, to simplify the heating process. Use of the RF furnace avoided the presence of reactive, heated components in the vacuum system other than the graphite capsule and sample. To aid purification for mass spectrometry, product CO was converted to CO2 by means of nickel powder at 450~ as a catalyst. A pressed pellet containing ca. 10mg of sample and 100mg of graphite powder was prepared to ensure intimate contact between graphite and sample, and out gassed at a temperature, between ca. 1000-1250~ depending on sample composition. The tendency of graphite to absorb gases required pre-reaction out gassing of the crucible as well. For yields of ; 97%, Clayton & Epstein (1958) indicated a reproducibility of ca. 0.5%0. Dontsova's (1959) procedures and apparatus closely followed those of Schwander (1953). Reactions proceeded measurably above about 1000~ and were complete by 1900~ after 15-18 minutes. At this temperature, yields of ~ ca. 95% for a range of rock types (granites to dunites) and minerals (quartz and olivine) were achieved. Reproducibility was indicated to be between 0.2 and 0.4%o. Schwander (1953) noted that alkali metals and some metal carbides, volatile under the conditions of the carbon reduction, often condense as metallic films on the cooler walls of the reaction vessel and vacuum line. Clayton & Epstein (1958) suggested that reaction between released CO and such metal films can cause a marked reduction in yield (e.g., reduced to 70% in one case), with an attendant isotopic fractionation that depletes the residual CO in 180. Indeed, they found poor comparison (differences in 6180 of up to 5.2%o) between analyses for the same orthoclase by C-reduction and flu-


403

orination, and attributed this to reaction between CO and metal films precipitated from volatized sample. The isotopic effect of the CO-metal film reaction was opposite to that reported by Schwander (1953) for cases where yields were low due to incomplete C-reduction. In contrast, Dontsova (1959) indicated in her study that CO-metal reaction was insignificant owing to the presence of dispersed graphite. Franchi et al. (1986) demonstrated the possibility of preparing CO by carbon reduction/ laser heating of small aliquots of silicates. They used a Nd-glass laser to heat a mixture of quartz and graphite pressed into a pellet. The product CO had a carbon isotope composition similar to that produced by combustion of the same graphite in an oxygen atmosphere, but the oxygen isotope composition of the CO was not determined. Using a Nd-YAG laser, Sharp & O'Neil (1989) heated small (3mm) pressed pellets consisting of sub-equal weights (< 1 to 3mg) of sample and graphite to temperatures above 1727~ in a glass vacuum line. Product CO was converted to CO2 by platinumcatalyzed, high-voltage discharge. They found that, with the exception of potassium feldspar, isotopic analyses were as precise as those by fluorination, and independent of yield. Erratic results for feldspar noted by Clayton & Epstein (1958) were also encountered by Sharp & O'Neil (1989), but Sharp & O'Neil (1989) found less positive, rather than more positive measured values of 6180 compared to accepted values. Again, a kinetic reaction may be the cause. Since CO was absorbed onto molecular sieve during the laser heating, selective decomposition rather than reaction between CO and precipitated metals was suspected (Sharp & O'Neil, 1989). The success of fluorination procedures described below in facilitating oxygen isotope analysis of even the most resistant of minerals with acceptably low uncertainty soon led to the abandonment of the carbon reduction method. Nevertheless, the apparent simplicity of reaction [20.1], safety, and the apparatus required, would seem to make this type of extraction still attractive in certain applications, providing that the sources of uncertainty (e.g., temperature of reaction; metal volatility; extent of reduction; etc) noted in Schwander (1953) and Clayton & Epstein (1958) could be overcome. With regard to modern laser-heating techniques, and on-line, continuous flow preparation procedures, it is encouraging that Schwander (1953) noted that rapid heating, although resulting in a low yield of CO (probably from the reaction of primarily grain surfaces), made no detectable difference in isotopic composition of the product CO. Consequently, the potential of this reaction, especially when assisted by a small amount of a fluorination reagent (M. Gehre, pers. comm., 2001) to assist in metal-oxygen bond breakage, has, perhaps, not yet been fully tested using recently developed, high-temperature, glassy-carbon furnaces (cf., Kornexl, et al., 1999a). High-temperature reaction with carbon in sealed quartz tubes (O'Neil et al., 1994) provides a convenient, precise alternative to fluorination for the oxygen isotope analysis of purified macro phosphate samples that, if properly calibrated in each lab, is accurate. This method is relatively straightforward, and requires simple vacuum apparatus found in many isotope laboratories. Moreover, a number of samples may

404

Chapter 20 - B.E. Taylor

be processed at one time. The sealed-tube method of O'Neil et al. (1994) can, by extension, be readily adapted to automated techniques employing high-temperature pyrolysis, with reduction in the size of the analyzed sample (e.g., 0.1mg) and an increase in throughput. An additional advantage of this sealed-tube method is the apparent lack of dependency of 6180 on reaction yield (O'Neil et al., 1994). Yet, as with silicates, oxides, and sulfides, analysis of phosphates by fluorination still provides the best accuracy and precision. Nevertheless, oxygen isotope analysis of phosphates by fluorination still provides the best accuracy and precision. Sulfates have proven to be less amenable to fluorination, although Pickthorn & O'Neil (1985) have demonstrated that a correction, required owing to incomplete, selective fluorination, may be used for improved accuracy. Carbon reduction, either by classical resistance heating (e.g., Nehring, et al., 1977), or external, radio frequency (RF) heating methods using macro samples of ca. 20mg, combined with spark discharge conversion of CO to CO2 (Nehring, et al., 1977), typically results in a combined accuracy and precision of ca. 0.1 to 0.2%o. High-temperature (1400~ pyrolysis of sulfate to CO, aided by nickelized graphite, can be carried out in an automated fashion considerably faster than by traditional methods (e.g., Nehring, et al., 1977), but at a cost of comparatively lower accuracy and precision ( ~ 0.5%0; Kornexl et al., 1999a). As noted in some detail below, fluorination is also the method of choice for the analysis of silicon isotope ratios, and for sulfide minerals, where precision and accuracy are of principal concern. The earliest isotopic analyses of selenium were carried out by fluorination (see below), but very recently instrumental methods have taken over, proving to be comparable in speed and precision. 20.2.2 Fluorination 20.2.2.1 Silicates and oxides

Successful and reliable oxygen isotope analysis of silicate and oxide minerals by fluorination was first demonstrated by Baertschi (1950), who found that sedimentary rocks contained more 180 than did igneous rocks. Detailed description of the fluorination method and apparatus by Baertschi & Silverman (1951) comprised the foundation of the fluorination technique, on which later refinements were based. Baertschi & Silverman chose C1F3 (+ HF) as the fluorinating reagent due to " ... the omission of certain necessary precautions .... " which prevented a successful test of F2 (+HF). Silverman (1951) used C1F3 to analyze a variety of rocks and meteorites. Others (e.g., Tudge, 1960; Taylor & Epstein, 1962; Clayton & Mayeda, 1978; Borthwick & Harmon, 1982) soon followed suit refining the fluorination method and expanding the range of reagents and applications. Two principal methodologies were developed: one method relying on fluorine gas (F2) derived from a high-pressure cylinder, and another method utilizing an interhalogen fluoride (e.g., BrF5 or C1F3) derived from a low-pressure cylinder. Though similar in many respects, each of these methods has its own advantages and disadvantages, as noted below, and both are in use today for macroand micro-analysis.


405

Both conventional vacuum lines ("macro'-systems by today's standards) and laser-based ("micro"-) vacuum systems are used to release oxygen by fluorination. Each system offers particular advantages, and many laboratories find these "macro-" and "micro-" scale analytical tools complement each other. Conventional systems typically require 5-20mg of mineral or whole-rock powder, reacted in externally-heated vessels, generally for ca. 12-16hrs, whereas laser-assisted fluorination utilizes considerably less sample, typically on the order of 0.1mg (e.g., several grains or small fragment from a thin-section), or involve small in situ reaction volumes ("spots" or craters) some 150-500 mm in diameter, and of similar depths, and occurs often within seconds or fractions of a second. Thus, in some cases, the desired scale of analytical resolution, nature of the sample, or purpose of the analysis may guide the choice of analytical system. The diameter of the incident ion beam and the depth of sputtering determine the volume of a sample analyzed by ion-beam analytical methods (SIMS, or Secondary Ion Mass Spectrometry; see Ireland, T., Chapter 30, this volume). In contrast, volumes of minerals analyzed in situ by laser-assisted fluorination may be considerably larger than the nominal laser beam diameter, owing to the fact that the resultant volume is the result of a fluorination reaction. The size of the volume (i.e., the spatial resolution of analysis) depends upon a number of factors, including duration of reaction, temperature reached (in those cases where absorption of laser radiation produces the requisite heating), and pressure of fluorinating reagent. Isotopic analysis by SIMS can be accomplished at a higher spatial resolution relative to that by laser-assisted fluorination. However, one must accept a lower analytical precision and accuracy (owing to drift), and present intractability of some minerals to analysis (e.g., S: Eldridge et al., 1987; Chaussidon & Lorand: 1990; O: Ricuputi & Paterson, 1994), although the use of multi-beam collection, analytical precision for some SIMS analyses is nearing that capable by fluorination-based methods (Valley, pers. commun., 2003; see also Ireland, T., Chapter 30, this volume). As even many early SIMS-based investigations showed, however, the isotopic variations on a microscopic scale may readily exceed the analytical uncertainty (e.g., O, C, and H: Valley et al., 1998; S: Eldridge et al., 1987 and McKibben & Riciputi, 1998; Si: Zinner et al., 1987). In these cases, the high spatial resolution can be of great advantage.

20.2.2.2 Silicates for silicon isotopes Silicon isotope analysis of silicate minerals requires their fluorination to silicon tetrafluoride (SiF4) as the analyte used for the measurement of 30Si/28Si ratios (e.g., Taylor & Epstein, 1970). Samples containing a1% of carbonates, carbon, phosphates, boron-bearing minerals, sulfides and sulfides must first be purified using HC1 (carbonates, sulfides), high-temperature oxidation (carbon), or fluxing with NaOH and Na202 followed by dissolution in HC1 and firing at 1000~ (e.g., Ding et al., 1996). Allenby (1954) was the first to measure silicon isotope ratios in rocks, preparing SiF4 by fluorinating rock samples with HF at room temperature to 100~ Although his analytical procedure and the reporting of silicon isotope ratios as 28Si/30Si are no longer followed, his observation is still valid that rocks exhibit a comparatively smaller range in silicon isotope composition than found for oxygen isotopes. Success-

406


ful fluorination procedures for the isotopic analysis of purified silica or silicates used today employ either F2 (Taylor & Epstein, 1962; Douthitt, 1982; De La Rocha et al., 1996), or BrF5 (Ding et al., 1988; 1996; and 2003), and generally follow the procedures for oxygen isotope extraction from silicates and oxides. Purification of the SiF4 for mass spectrometry, as discussed in a later section, differs slightly from routine methods in oxygen isotope analysis. Reynolds & Verhoogen's (1953) early investigations of silicon isotope variations in rocks, sinters, and cherts made use of a rather involved wet chemical sample purification procedure that culminated in the precipitation of the stable compound BaSiF6. Compared to other methods of fluorination (discussed later), their procedure has the distinct advantages of producing not only a form of silica that is pure and decomposes under vacuum heating according to the reaction: BaSiF6(s) ~ BaF2(s) + SiF4

[20.3]

but also obviates the need for fluorination by gaseous or liquid fluorinating reagents, as is currently done. This simple reaction, in spite of the time required to purify and convert the sample to BaSiF6 obviates the need for an elaborate vacuum fluorination apparatus and purification protocol for the SiF4 produced. In fact, reaction [20.3] can be performed in sealed tubes, individually or as a batch, and stored for subsequent mass spectrometry (Reynolds & Verhoogen, 1953). Reynolds & Verhoogen found that the ca. 2.3% silicon isotope exchange between the SiF4 and the walls of the glass sealed tubes had a negligible effect on the measured isotopic composition of the sample. As found by subsequent workers, Reynolds & Verhoogen (1953) reported a small (ca. 3%0) variation of 630Si (reported in terms of the 30/28 mass ratio) among natural samples, which, they suggested was related to the temperature and mode of formation and warranted further investigation. Tilles (1961a), using the Reynolds-Verhoogen technique, extended the range of natural variations in 630Si to 5.3%0 (reported in terms of the 30/(28+29) mass ratio), and published the first measurements of permil silicon isotope fractionations between quartz and feldspar from a pegmatite that suggested a potential for silicon isotope geothermometry, and also indicated (Tilles, 1961b) a tendency for the order of enrichment (in 30Si) among common rock-forming minerals. Some minerals (e.g., olivine) were found to give low yields (ca. 80%), however, and, accordingly, results for these minerals are not reliable by the Reynolds-Verhoogen technique. Silicon isotope studies have been limited in number (but increasing recently), and largely focused on extra-terrestrial material because the variation in reported terrestrial 30Si/28Si ratios is comparatively small (Douthitt, 1982). Recent isotopic determinations of dissolved riverine silica Ding et al. (2002 and 2003) have demonstrated that the terrestrial variation of 630Si is as large as 3.0 (0.4 to 3.4%0), compared to that of dissolved marine silica (+0.06 to 2.2%0; De La Rocha et al., 2000), or to biologically-deposited silica (-1.2 to -3.7%o; De La Rocha, 2003). The routine measurement precision of 0.1%o nevertheless allows for the investigation of a number of phenomena of interest, especially among surficial processes, despite the fact that the terrestrial variation of


407

silicon isotopes is not large by comparison with isotopes of hydrogen, carbon, nitrogen, oxygen and sulfur. 20.2.2.3 Phosphates Fluorination as a process for isotopic analysis is, of course, not restricted to oxygen isotopes in silicates and oxides. As is also the case for sulfides, fluorination of phosphates provides the most accurate and precise means of (oxygen) isotope analysis (Vennemann et al., 2002; L6cuyer et al., 1993; see L6cuyer, Ch., Chapter 22, this volume)). Fluorination of phosphates can be employed in both macro- and micro-techniques. However, the fact that oxygen can reside in several sites (PO4 -3, CO3- and OH), in biogenic apatite (Ca5(PO4)3-x(CO3)xOHyFl-y), and that additional, organic compounds, are present requires the isolation and purification of the phosphate-oxygen (as Ag3PO4 or BiPO4) from biogenic apatite prior to fluorination. The purification step imparts no isotopic fractionation to the results from biogenic phosphate (e.g., Karhu and Epstein, 1986), and is not required for analysis of abiogenic (e.g., hydrothermal or magmatic) apatite (Conway & Taylor, 1969; Fortier & L~ittge, 1995; Rhodes & Oreskes, 1999). Unlike Ag3PO4 or BiPO4, however, fluorination of abiogenic apatite [Ca5(PO4)3(OH,F,C1)] by BrF5 requires high temperatures (e.g., 650~ for complete yields (Rhodes & Oreskes, 1999; B. Taylor & Mirnejad, unpub.).

Micro-analysis of biogenic phosphate by direct, laser-assisted fluorination (e.g., Kohn et al., 1998; Jones et al., 1999; Lindars, et al., 2001), without purification of contained phosphate-oxygen, limits the accuracy and precision of the results, but relative isotopic variations may still be of great value owing to the high spatial resolution afforded by laser sampling. The same may also be said of direct laser heating of tooth enamel in a He stream (Cerling & Sharp, 1996) which provides an alternative, albeit still relative, method of micro-analysis that does not involve fluorination. The principal issue in the fluorination of biogenic phosphate remains the choice between direct, laser-assisted fluorination of the phosphate mineral and classical fluorination of a purified phosphate (Ag3PO4 or BiPO4) prepared from the original phosphate mineral by dissolution and re-precipitation. The former, albeit on a microscopic scale, may yield only a relative analysis of lower accuracy~ whereas the classical technique is preferred for best accuracy and precision. Currently, the sampling resolution of biogenic phosphate for classical fluorination is ultimately limited by the techniques of purification. 20.2.2.4 Sulfides Although the sulfur isotope ratios of sulfide minerals have mostly been determined from SO2 produced by oxidation, using either gaseous 02 (commonly used in elemental analyzers) or a solid source of oxygen mixed together with the sample (e.g., V205 or mixture of CuO/Cu20), fluorination of sulfide minerals to produce SF6 can facilitate the most precise and accurate analysis. The reasons for this include both (1) mineral/reaction-specific factors, especially in micro-analysis (e.g., mineral composition dependent fractionation; Crowe et al., 1990; Kelley and Fallick, 1990), and (2) cross-contamination in the ion source of the mass spectrometer. In the former

408


instance, accuracy of the results is compromised by the mineral composition data. In the latter case, cross-contamination or memory in the ion source results in variable contraction of the 6 scale in different laboratories. Fluorination of sulfides requires apparatus similar to that used for oxides and silicates. However, owing to the mass of SF5+ (127), the principal ion measured in the mass spectrometry of SF6, an isotope ratio mass spectrometer that can measure to at least a mass of 130 is required, and above 131, if 636S is to be measured. Although high-mass capable models are readily available, only a few labs today utilize fluorination for sulfur isotope analysis and gas-source isotope ratio mass spectrometers more commonly purchased today have a mass range on the order of 2-80 and, with regard to sulfur isotope analysis, are used for measurement of SO2. Although conceived nearly 40 years ago (Hulston & Thode, 1965b), sulfur isotope analysis by fluorination has not, with few exceptions (e.g., Rees & Thode, 1977; Rees, 1978; Ding et al., 1985), commonly been employed, and has remained virtually dormant as a technique from the mid-70's to early 90's. The superior capabilities of SF6 in microanalysis (and for determination of 633S: Hulston & Thode, 1965b; Thode & Rees, 1971; Ding et al., 1987), however, have caused a recent renaissance in sulfide fluorination, almost exclusively by laser-assisted methodologies (e.g., Rumble et al., 1991, 1993; Beaudoin & Taylor, 1994; Farquhar & Thiemens, 2000; Farquhar et al., 2000a; Taylor et al., 2001, in press). 20.2.2.5 Water

The classic CO2-equilibration technique (Epstein & Mayeda, 1953) has been the method of choice because it is relatively easy and precise. On the other hand, fluorination offers several advantages, especially for the oxygen isotope analysis of microlitersized samples. Uncertainties of many types can be introduced in the isotopic analysis of water, from the calibration of internal laboratory (gas and water) standards and the experimentally determined mineral-water and CO2-H20 fractionation factors, to control of the conditions governing equilibration and analysis in an individual laboratory. Therefore, direct analysis of water would be preferable, at least theoretically, by either fluorination (O'Neil & Epstein, 1966b), reduction by carbon (Majzoub, 1966), or reaction with guanidine hydrochloride (Dugan et al., 1985).

20.3 Fluorinating reagents A number of physical and chemical properties of the commonly used fluorination reagents (HF, F2, BrF3, BrF5 and C1F3) have influenced both the choice of reagent and the development of different fluorination methodologies. Other sources of fluorine (both solid and gaseous) have been variously used either as fluorinating regents, or as an aid to laser ablation or pyrolysis (e.g., Kornexel et al., 1999). The solid compound XeF2 has, apparently, been used with some success as a source of fluorine by Dmitry Krylov, a Russian scientist (S. Hoernes, pers. commun., 2003). His procedure, in brief, comprises weight ratio of XeF2/sample of ca. 10"1 (e.g., 10-15mg of sample added to ca. 100rag of XeF2), loaded together (in a dry box) and reacted, as in the BrF5 procedure, in nickel tube at temperatures up to 600~ (above 600~ the XeF2 dissociates; it melts at 140~ This procedure has received little attention, but the ease of handling small amounts of the fluorinating reagent might prompt some interest in its applica-

409


tion in microanalysis. Other substances, such as Freon R-134a (CH2FCF3), for example, have been used recently to increase yield of the chemical and Pb-isotope analysis of zircon by laser ablation using the ICP-MS (Hirata, 2003). The application of such fluorocarbons in more traditional methods of stable isotope analysis has not been widely pursued, however, and is likely to be fraught with problems owing to mass spectrum interference and poor pumping characteristics (c.f., Rumble & Hoering, 1994). Differences in freezing points (Table 20.1), in particular, have played a large role in how the interhalogen fluorides on the one hand, and fluorine gas, on the other, have been utilized. Whereas, interhalogen fluorides can be transferred cryogenically in the vacuum line, fluorine gas cannot under normal circumstances. Cryogenic transfer of a fluorinating reagent facilitates the measurement of an appropriate aliquot size that may then be condensed into each sample tube without pre-mature reaction. The ability to cryogenically transfer the interhalogen fluorides also facilitates the post-reaction separation of residual reagent from extracted oxygen. Other physical properties may influence the choice of one interhalogen fluoride over another. For example, the vapor pressure (Table 20.1, cylinder pressure; psig, or pounds-per-square-inch gauge pressure) of BrF3 is sufficiently low as to render its use in typical vacuum fluorination apparatus rather impractical, and, for this reason, it is not commonly used today. Awareness of its properties is important, however, as BrF3 can occur as an impurity in BrF5 and create some difficulties.

Table 20.1 - Selected physical and chemical properties of fluorinating agents and fluorine compounds I Property

HF

F2

BrF3

BrF5

C1F3

CoF3

CaF2

KF

XeF2

State (S.T.P.)

gas

gas

liquid

liquid

liquid

solid

solid

solid

solid

Molecular Weight

20

38

136.9

174.9

92.5

1 1 5 . 9 3 78.08

58.1

169.29

Boiling Point @ 100 kPa (~

19.54

-188.1

135

41.3

11.3

1400

ca2500

1505

114.35

Freezing Point @ 100 kPa (~

-83.1

-219.6

8.8

-62.5

-83

ca1200

1423

858

129.03

Cylinder Pressure (kPa) @ 20 ~

6.21

2760

0.93

43.2

46.92

Critical Temperaure

188

Specific Gravity (H20 = 1.000)

0.99

(oc)

197

1.69

2.49

2.48

1.77

1 Lide, D. R. (ed.), 1996, Chemical Rubber Comapany Handbook of Chemistry and Physics: CRC Press, Boca Raton, Florida, U.S.A., 77th ed.

410


Although all fluorinating reagents pose potential (but manageable) health and safety risks, the use of F2 requires particular attention to safety and handling precautions as a commercially available tank or cylinder of F2 is typically under considerable pressure (e.g., ca. 400psig; see Table 20.1). The cylinder pressures of the interhalogen fluorides, on the other hand, are all less than one atmosphere. Minor leakage of a cylinder containing a halogen fluoride would initially result in migration of air into the tank, rather than halogen fluoride out of the tank. Purity also distinguishes the commercially available fluorinating reagents. Halogen fluoride compounds can be obtained in rather high purity, and further purified by vacuum distillation. A tank of commercially available F2, however, usually has a purity of only ca. 97%. Contaminants, among others, typically include oxygen. Its presence contributes, at least in a small way, to the uncertainty of each analysis, in spite of the fact that such a "blank" of oxygen can be quantified, and for which a corFigure 20.1 - Conventional fluorination line of metal and glass construction for the extraction of oxygen isotopes from silicates, oxides and phosphates utilizing either halogen fluoride reagents (e.g., BrF5 or C1F3) or F2 gas, and employing top-loading (see Figure 20.5), externally-heated reaction tubes that facilitate rapid, blank-free (i.e., without water) loading of samples. Components utilized with F2 gas, either commercial or laboratory generated and/or purified are shown in insets F and J. Details of construction may vary from lab to lab, but may include the following components: (A) Hg (or other) diffusion pump, backed by roughing pump to yield vacuum of ___10-3Torr; (B) Argon and drying agent (C; e.g., 4A mol. sieve) to ensure dry inert gas used primarily during sample loading; (D) KBr (or other salt), externally heated to 150-200~ used to passivate excess reagent (F2 or halogen fluoride compounds) by reaction to form KF and Br (trapped at LN2 temperature on P-style trap; see Figure 20.2); (E) halogen fluoride storage, and Kel-F and nickel tubes for distillation, visual inspection, and storage of working aliquot for 2-300 analyses; (F) commercial F2 supply and/or Asprey-type F2 generator for purification/production of pure F2; (G) nickel reaction tube and cooling collar fitted with rapid loading top (see Figure 20.3; one 12 tubes attached to manifold of stainless steel tubing); (H) furnace, thermocouple (tc) and temperature controller, or variable transformer, for the external heating of reaction tubes (G); (I) Pt-catalyzed carbon rod reactor to convert 02 to CO2 (see Figure 20.4); and (J) in-line reactors (2; not shown in detail) to convert F2 to Br (or other halogen) allowing separation of 02 from F2-based fluorination. Vacuum is monitored by G-1 and G-2 (capacitance manometers, e.g., Setra | 0-10psi absolute (G-l) and 0-25psi absolute), and thermocouples or other, similar range gauges, G-3 to G-5. Cold traps T-l, T-2, and T-3 in stainless steel portion of line are held at liquid nitrogen temperature and utilized during, respectively, sample loading, post-reaction conversion of excess halogen fluoride reagent, and during extraction of oxygen. Cold trap T-4, is held at liquid nitrogen temperature to insure against loss of CO2 formed during conversion of 02 (most of which is trapped in converter vessel; see Figure 30.4). U-trap T-5 facilitates measurement of the 02 reaction yield (as CO2) and cryogenic purification of the CO2 sample if required. All valves are high-vacuum valves. Those in the metal section of line are a 'diaphragm' style valves (e.g., Nupro| SSDLS-4 and SS-DLS-6, with Swagelock @ tube fittings) except for the reaction tubes, which utilize toggle-style bellows valves (Nupro| SS-4BKT; see Figure 20.3). Pyrex tubing (1 / 4in or 6mm) is attached to the vacuum line below G-5 using O-ring connectors. CO2 is collected in these tubes with liquid nitrogen then sealed in lengths of 5-6in with a torch. Note that the exhaust of both rotary (roughing) vacuum pumps is vented outside the laboratory. A copper wool-filled trap between the Hg diffusion pump and roughing pump acts to reduce transmission of any Hg vapor.


411

412


rection can be adopted (Taylor & Epstein, 1962). Alternatively, relatively pure F2 can be generated (or, purified from another primary source) in the laboratory at low pressure using the method of Asprey (1976). This method is in use with both conventional fluorination lines for macro scale analysis (e.g., University of Bonn) or, in laserassisted fluorination apparatus for micro-scale analysis (e.g., Rumble, et al., 1993; Taylor & Beaudoin, 1993; Beaudoin & Taylor, 1994), as described briefly below. Reagent cost and availability further separate F2-based from interhalogen fluoridebased techniques. Bromine pentafluoride (BrF5), a reliable reagent for many years, was once available rather inexpensively when it was used in industrial applications. More recently, BrF5 has been difficult to obtain, and then only at a cost of several thousand dollars (U.S.) per kg. Chlorine trifluoride (C1F3) is a cheaper alternative, and can be more readily obtained than BrF5. In contrast, a tank of F2 costs even less than either of these interhalogen fluorides, and is available from specialty gas companies. Be forewarned that the availability and shipping fluorine and some of its compounds may strictly controlled in some countries. The sealed integrity of reagent cylinders should be of utmost concern, and care should be taken that their shipment does not expose the cylinders to environmental extremes such as freezing conditions that could lead to valve leakage. Inspection of the cylinder valves should be made for stains or deposits (red-brown in the case of BrF5, for example) which would provide some indication of leakage.

20.3.1 Purification of fluorinating reagents The interhalogen fluorides BrF3, BrF5 and C1F3 can all be purified by vacuum distillation, although the low vapor pressure and relatively high freezing point of BrF3 renders its purification by this means more time-consuming. Trace amounts of noncondensable gases such as oxygen, nitrogen, argon, fluorine, SiF4, CF4, SF6, etc. can be readily separated by cryogenic, vacuum distillation using a -70~ trapping temperature (e.g., frozen/liquid ethanol mixture made by adding liquid nitrogen to ethanol). The process comprises cryogenic transfer of an aliquot of reagent from one small storage vessel to another, followed by pumping away of all gases not condensed with the reagent. Typically, this involves 1-2 distillations, although some investigators prefer triple distillation. Cylinders of reagent used infrequently may develop a headspace of non-condensable gases such that the first few aliquots of reagent removed may have relatively more associated contaminants. It is advantageous to be able to actually see the liquid reagents. For this reason, semi-translucent Kel-F, a fluorinated plastic, is often used to manufacture a storage vessel for a working aliquot of the halogen fluoride (see Figure 20.1). A vessel of the same dimensions as the nickel reaction tubes is convenient. In the case of BrF5, purity can, in part, be assessed from its color. A clear, pale yellow liquid is characteristic of pure BrF5, whereas the presence of a reddish or brownish hue indicates contamination by Br (and, in this case, possibly some BrF3 as well). Bromine has a freezing point o f 7.2~ allowing for its cryogenic separation from BrF5. Exposure of the working aliquot of liquid BrF5 to gaseous F2 apparently can be carried out in order to fluorinate any excess Br and BrF3. This was apparently routine in the laboratory of Prof. S.


413

Epstein, California Institute of Technology (G. Yoshiba, pers. commun., 2003). Fluorination of BrF5 consisted of overnight exposure of the BrF5 to F2 at ca. l atm pressure and room temperature. The fluorination reactions are exothermic, and it is advisable to work with small quantities (e.g., 20-30 ml) of BrF5, preferably in a Ni vessel. The efficiency can be visually assessed after excess fluorine has been pumped away and neutralized (see below), and the reagent transferred to a Kel-F vessel. For example, the conventional fluorination line illustrated in Figure 20.1 has both a Kel-F vessel and a Ni vessel, each of similar size to the Ni reaction tubes, and is constructed in such a way as to facilitate reagent 'pre-treatment' upon connection to a source of F2. Alternatively, a reaction tube could be utilized. Commercial F2 can be readily purified in small batches by means of the K2NiF6,KF- filled F2 generator ('fluorine pump') originated by Asprey (1976). When heated, K2NiF6.KF decomposes, releasing pure F2 according to the following reaction: 2[K2NiF6. KF](s)~ 2[K3NiF6](s) + F2

[20.4]

This reaction is reversible, and re-fluorination of solid K3NiF6 can be accomplished by exposure to F2 at about 125~ Below this temperature, impurities such as oxygen, nitrogen, argon, etc. can be pumped away. Asprey (1976) showed that thermal cycling of the F2 generator, followed by cooling below 125~ and additional evacuation, could result in the production of relatively pure F2 at a low pressure. The pressure of F2 that develops as the generator is heated should be closely monitored because, although relatively low, the pressure increases markedly with temperature as reaction [20.4] proceeds. The required K2NiF6. KF can be purchased commercially, or made in the laboratory. For example, the procedure followed at the University of Bonn (Prof. S. Hoernes, pers. commun., 2003) comprised mixing together 3mol NiF2 and 9mol KE and loading into a nickel (or, Monel) vessel sealed with a Cu-gasketed Conflat~ flange to construct a 'fluorine generator' suitable for a standard fluorination line (such as shown in Figure 20.1). K2NiF6.KF is quite hygroscopic and the 'fluorine generator' (and also any commercially-purchased K2NiF6.KF similarly contained) must be activated by several fluorination cycles as follows: evacuate and heat to at least 300~ to liberate any HF which formed that must be pumped away. The generator is exposed to a raw fluorine pressure of some 3-5b while heating to 500~ and which point the F2 tank is closed and the fluorine generator allowed to cool slowly to ca. 250~ during which the K2NiF6.KF is fluorinated. Any gases not absorbed by the K2NiF6.KF are then pumped away. Asprey (1976) recommended a F2 pressure of 10b. However, 5b is a typical output limit on commercial, 2-stage F2 tank regulators, and this is sufficient for the purpose. The cycle is repeated (perhaps several times) until the F2 pressure drops below the starting pressure. The 'fluorine generator' is then ready for service as described in the previous paragraph.

414


Figure 20.2- Pyrex glass Pstyle trap used to collect Br (if KBr is used to passivate a fluorinating reagent, or C1 in the case of NaC1, for example) at liquid nitrogen temperature and to evacuate the KBr reactor (component D, Figure 20.1) and vacuum line. After thawing, the Br is transferred to a vent line with the aid of a few psi of dry Argon (components B and C, Figure 20.1), or other inert gas. Vacuum seal between custom stainless vessel top (dimensions as in Figure 20.3) and metal/glass seal is made with Teflon ferrules. Valves are stainless steel, high vacuum valves (e.g., diaphragm type, Nupro | SS-DLS-6; Monel stem tips where adjacent to KBr reactor, otherwise Kel-F stem tips) that permit valve part replacement, if needed, without removal from line.

Asprey's (1976) method has garnered considerable attention and is becoming the method of choice for safe production or purification of very pure fluorine. Another method, described by Jacob & Christie (1977), requires the removal of 02 and other trace impurities from commercial F2 in two stages. Trace impurities (except N2 and Ar) having no measureable vapor pressure just above LN2 temperature (-196~ can be removed by distillation at this temperature. Others require colder trap temperatures (e.g., -210~ achieved only with some difficulty in standard stable isotope labs (e.g., by streaming He through LN2). Trace amounts of 02, on the other hand, can be effectively removed from F2 according to Jacob & Christie (1977) by the following reaction carried for 2 hours in a closed Monel vessel at ca. 190~ 0 2 + F2

nSbF5 ~ S b O 2 F 6 . (n-1)SbF3

[20.5]

The Monel reaction vessel is then cooled to-183~ and the O2-free F2 distilled to another vessel held at-196~ Reaction [20.5] might be used with F2 purified according to Asprey (1976) as an extra precaution in micro-analysis for oxygen isotopes.


415

Figure 20.3 - Rapid-loading reaction tube "chimney" caps from stainless steel. The vacuum seal of the Ni reaction tube to the "chimney" cap is made with Teflon ferrules and a 3/4" Swagelock | nut (not shown; see Figure 20.4). The top (0.375in O.D.) of the tube ("chimney") is closed with a Swagelock @ cap (SS-600-C), also fitted with Teflon ferrules (see Figure 20.4).

20.3.2 Reagent storage The storage of F2 and other fluorinating reagents often presents a problem for the researcher in carefully regulated working environments. The fear of an accident involving stored reagent can be acute among those charged with safety and regulatory responsibilities. This is not to downplay such concerns, which should be foremost in all of our minds, but the requirements placed upon laboratories can, at times, create other potential sources of worry. Typically, BrF5 and other interhalogen fluorides are kept in the cylinder in which they were received from the commercial source, and securely supported and attached to the vacuum fluorination line. This makes good sense inasmuch as the means of neutralization is at hand, and there is no concern regarding moving or relocation of such cylinders. In some cases, the cylinder of reagent is kept in a fume hood, and connected to the vacuum line via metal tubing. The connecting pipe work between such cylinders and the vessels used for storing small quantities (e.g., Kel-F vessels) should be as short as possible. Lengthy pipe work simply slows the process of reagent transfer and increases the possibilities of leaks via needed connections, etc. In rare cases

416


where the cylinders and valves received are poorly manufactured (as is the case from one supplier of BrF5 on the international market), cryogenic transfer to a pre-treated vessel having a reliable valve is most advantageous, if not absolutely necessary. As noted before, the less than one atmosphere cylinder pressure of BrF5 typically minimizes the danger of direct leakage into the laboratory. The presence of a reddish brown stain or white deposit around the cylinder valve, or its connections, however, provides evidence of leakage and transfer to a more reliable vessel is needed. Transfer should be carried out slowly on the vacuum fluorination line (where the means for reagent neutralization is accessible if need be), and monitored by a vacuum / pressure gauge. Steel, Monel| or nickel are suitable cylind e r materials for storage of fluorinating reagents, Such cylinders should be obtained new from a commercial source, and carefully leak-checked under vacuum before requisite pre-treatment with small amounts of

Figure 20.4 - (A) The Ni reaction tube, with attached "chimney" cap, is connected to the manifold of fluorination line (inset G, Figure 20.1) with a toggle-style, bellows high vacuum valve with Kel-F stem tip (Nup r~174SS-4BKT-KF). A 60mm stainless steel chromatographic filter (Swag el~174 p / n 696-530), inserted in the reaction tube side of the valve, prevents pieces of Ni- or other metal fluorides and sample particles from entering the valve seat. Note the direction of gas flow, as indicated by an arrow on the valve body, which ensures that the valve stem tip and not the bellows of the valve, is exposed to the fluorination reagent during the reaction. The toggle-design exerts sufficient pressure to guarantee a vacuum-fight seal, and insures against over tightening. (B) Stainless steel loading tool, with highly polished interior end, closed by a spring-loaded rod soldered to a stainless-steel bearing, permits rapid sample loading through the "chimney" of the reaction tube's "chimney" cap (see also Figure 20.3). The reaction tube and manifold are filled with slightly greater than latm Ar. The Ar flows continuously out of the "chimney' cap during loading, which prevents entry of atmospheric moisture to the reaction tube. A weighed sample is placed in the loading tool, and released at the bottom (only) of the reaction tube by depressing the spring-loaded central rod.


417

Figure 20.5 - Water-cooled, C-rod converter forms CO2 from extracted 02, which is condensed with liquid nitrogen at the bottom of the converter vessel. The platinum coil simultaneously heats the ca. 30mm x 3mm O spectrographic graphite rod, simultaneously, with the surrounding platinum shield, catalyzes the oxygen-limited reaction (based on a design by T. K. Kyser; pers. commun.). The temperature of the carbon rod can be visually monitored through the water-cooled glass reaction vessel. This design requires but a very small amount of liquid nitrogen, and facilitates conversion of ca. 100mm of 02 in less than 5 minutes. A socalled "back diffusion trap" (T-4, Figure 20.1) prevents loss of CO2 not condensed initially at the bottom of the converter.

reagent (at room temperature, followed by mild heating) until no m o r e 0 2 is evolved. Under no circumstances should cleaning and/or re-use of older cylinders be attempted- potentially dangerous residues of low vapor pressure compounds (e.g., BrF3) may reside in these cylinders. Similarly, no reagent cylinder should ever be opened later to the atmosphere, but disposed of properly (typically returned to its source or another qualified receiver). Leakage of cylinders of F2 is another matter. Where possible, the storage of highpressure (ca. 400psig, or 2760kPa; Table 20.1) cylinders of F2 in the laboratory itself should be avoided to minimize health and safety concerns. One solution is to pipe the F2 to the laboratory from a storage room. However, such piping must be accessible for

418


monitoring, protected and leak-free. Exposure to environmental factors (e.g., by extremes in temperature) could compromise the leak-free integrity of the gas-delivery system. An alternative, safe-storage solution for a cylinder of F2, particularly viable where smaller cylinders are used to recharge Asprey-type "fluorine pump", is to utilize a mobile, specialty gas storage. When not in use, the cabinet can be stored in an appropriate and safe location. At the Geological Survey of Canada (Ottawa), for example, a commercially available specialty gas storage cabinet, designed for use with hazardous gases and fitted with appropriate piping, regulator, valves, and HF getter (NaOH) is mounted on wheels and fitted with an exhaust port for connection to a vent system when in use in the laboratory. An HF sensor and auto-dialing mechanism provide unattended monitoring of the cabinet and its contents while in a separate, appropriately labeled storage room.

20.3.3 Neutralization and disposal of reagents With the same regard for safety in the storage and use of fluorine and fluorinating agents, the waste products of fluorination reactions and excess reagent require safe handling in their neutralization and disposal. In systems where F2 is the chosen reagent, on-line neutralization is required during the extraction and separation of oxygen. Neutralization of both F2 and halogen fluoride compounds can be readily accomplished by reaction with crystalline KBr (ca. 2-Smm grain size) at ca. 150-200~ to yield Br. The reagent passivation section of the fluorination line in Figures 20.1 and 20.2 illustrates one way in which this can be conveniently and safely accomplished. A stainless steel cylinder containing ca. 500g of coarsely crystalline KBr, closed at each end with Cu-gasketed flanges (e.g., Conflat~) and fitted with fine Monel screen at the inlet and outlet, is externally heated and insulated. A thermocouple affixed to the cylinder mid-way between flanges permits monitoring of the temperature. Coarse scrap KBr (cut-offs and broken single crystals) from the optics industry is particularly well suited as it is of high purity, contains a minimum of water, and can be appropriately crushed, and sieved to a narrow size range so as to maximize the life and permeability of the reactor. Too fine a grain size and/or too large a range in grain size can lead to clogging and inefficient throughput and pumping. The apparatus and methodology for BrF5 neutralization employed at the Geological Survey of Canada (Ottawa) involves on-line collection of excess reagent, neutralization (conversion to KF and Br) and storage of Br after each set of 12 fluorinations. The procedure entails collection of excess reagent and volatile waste products (e.g., Br) in a liquid-nitrogen cooled trap (%2, Figure 20.1) during an hour-long evacuation of the reaction tubes while heating (typically to 550-650~ depending on the next samples to be run; a slightly-higher-than-reaction temperature is preferred for out gassing). The collected excess reagent, waste products and KBr reactor are isolated together from the remainder of the vacuum line for ca. 20 minutes by closing valves 10 and 11 (Figure 20.1). The contents of trap T-2 are allowed to thaw naturally and react with the heated KBr, after which, the product Br is condensed into a glass P-trap maintained throughout the week at liquid nitrogen temperature and normally left

Fluorination Methodsin StableIsotopeAnalysis

419

open to the KBr reactor, which degasses Br slowly, via valve 11 (Figure 20.1). The small amount (< 10ml) of Br, accumulated from reaction of 48-60 samples in one week, is vented from the trap and lab area through valve 13 to power vented ducting by a ca. 10 psi stream of Ar. The Br is first thawed and then valves 13 and 14 are opened simultaneously. Warming the Br storage vessel (e.g., with a heat gun) while flushing with Ar ensures complete removal of Br. This procedure avoids exposure of the lab or its occupants to any Br vapor. In some labs, the Br is condensed into a removable vessel that is taken to a fume hood, under which the liquid Br may be collected and disposed of by others. In an alternative procedure, Ding et al. (1996) describe neutralizing residual BrF5 and Br by bubbling the thawed waste products through a Ca(OH)2 solution, using Ar as a carrier gas under a fume hood, to produce harmless CaF2, CaBr2 and 02. A similar method for passivation of Br (and C1 from NaC1 or KC1 reactions) that allows for their safe disposal via the drain has long been in use at the University of Bonn (S. Hoernes, U. Bonn, pers. commun., 2003). This method entails dissolving in 2N Na(OH) (80g NaOH + 21 H20), followed by neutralization of the alkalic solution with sodiumthiosulfate. The neutralized solution can then be safely flushed down the drain. On-line neutralization of F2 is necessary not only for the disposal of excess F2, but also to effect separation of F2 from oxygen. Because F2 is not condensed at LN2 temperature at the low pressures involved, a mixture of F2 and 02 is released from each reaction vessel after fluorination, while Br and other reaction products are retained in the reaction tubes submersed in liquid nitrogen. As in the above methods, KBr has traditionally been used to yield Br upon reaction with F2, which may then be cryogenically separated from 02 for later disposal. To overcome the additional reaction time necessar?4 and to facilitate an efficient collection and conversion of oxygen, Taylor & Epstein (1962) devised two heated KBr reactors (e.g., inset J, Figure 20.1), one of which has a small volume to allow efficient conversion of the final remaining F2 and accelerate the collection/purification of oxygen. 20.4 Fluorination reactions 20.4.1 Fluorination of silicates and oxides by F2 gas The F2-based technique and apparatus described by Baertschi & Silverman (1951) became the backbone procedure for a number of labs, especially those at the California Institute of Technology (Taylor & Epstein, 1962) and in laboratories subsequently established elsewhere by students of Professors H. P. Taylor, Jr. and S. Epstein. Compared to halogen-fluoride based techniques, widespread use of F2 as the fluorinating reagent of choice in conventional (macro-) vacuum lines was perhaps hindered by some of the analytical and safety. However, the advent of micro-scale analytical techniques employing lasers has seen a renaissance of the use of purified F2 as the preferred fluorinating reagent owing to the simplicity of the fluorination reaction (Table 20.2).

420


A typical reaction between an oxygen-bearing mineral (a metal oxide, where M metal) and F2 can be represented by" MxyOy(s) + yF2 ~ xyMF2/x(S) + Y / 2 0 2

[20.61

which yields solid metal fluorides (MF), oxygen gas (02) and, in the case of silicates, volatile fluoride compounds such as SiF4 (and others). For the mineral titanite (CaSiO5), for example, this becomes: 2CaTiSiO5(s) + 10F2 ~ 2CaF2(s) + 2TiF4 + 2SiF4 + 5 0 2

[20.7]

From the stoichiometry of reaction [20.7], and others involving halogen fluoride reagents discussed below, two moles of fluorine (F2; or halogen fluoride reagent) are required per mole of oxygen (02) for complete reaction. In practice, a fluorine/oxygen ratio of ca. 2.5-3.3 (or, 1.2 to 1.6 for the excess/stoichiometric fluorine ratio) was used by Taylor & Epstein (1962). Reaction of fluorine with the hot walls of the reaction vessel claims some of the loaded reagent, but this is not a major concern at the typical reaction temperature of 450~ (Haimson & Knauth, 1983). When insufficient reagent is present, the amount of oxygen released varies directly with the fluorine/oxygen ratio or, "fluorine excess" (Haimson & Knauth, 1983; Figure 20.6). The effect on 6180 values is variable, depending on the fluorinating reagent used. Epstein & Taylor (1971) and Haimson & Knauth (1983) found partial extracFigure 20.6 - Plot of fractional oxygen yield (as % of total) vs. fluorine tion of oxygen with excess (molar F2/02) in partial fluorination reactions at constant temperF2 gas occurred ature (450~ and time. A linear relationship between yield and fluorine without fractiona- excess for both quartz and opal A exists until F2/02 ratio is greater than tion of oxygen iso- 2 (after Haimson & Knauth. 1983). Notably the measured values of 6180 topes. This was true are constant for quartz-derived oxygen, but vary for opal A vary, owing whether the F2 pres- to progressive dehydration with fluorination. In contrast, isotopic analyses of oxygen derived from incomplete fluorination with BrF5 are inaccusure was varied at rate (Garlick & Epstein, 1967). constant temperatu-


421

re (Haimson & Knauth, 1983), or the reaction duration was reduced at constant temperature and F2 pressure (Epstein & Taylor, 1971). Oxygen yields from quartz may be inconsistent when the amount of fluorine present is


4~

4a

> Table 20.3 continued

tO

Pretreatment Sample Sample and chamber (pre-fluorination)

MS Analyte

Yield

Blank

% T.Y. Yield1

Blank/Prep. correction (%o)

{' 3 =r to o !

Table 20.3 continued > O

> Table 20.3 continued o +8X1(~o)

+8~8o(%o)

+6X2(%o)

MS 7

+6170 (%o) IRMS/irm-GCMS

Additional Notes

Reference 9

=r o

r 9 9

>,

Table 20.3 continued

> #a L,o LaJ


4a 4a

Sample Type

Powder / grains (P / G); in-situ (I)

Size Sample: (rag: P / G) (diameter,/am: I)

Laser type (pwr; ~. in tlm)l

Mode Pulse (P) Continuous (CW)

Fluence (J/cm2)10

Reagent


r

=r

b,a O !

o

o

> Table 20.3 continued 9

Pretreatment Sample Sample and chamber (pre-fluorination)

MS Analyte

Yield

Blank

% T.Y. Yield1

Blank/Prep. correction (%0)

e~

o

9

r

o o r~

>

v~


4~


_+6X1(%o) 4-6180(%o)

+6X2(%o) MS 7 +6170 (%o) I R M S / i r m - G C M S

4a Additional Notes

Reference


C3 r~ t,o O !

O


Sample Type

P o w d e r / grains (P / G); in-situ (I)

o Size Sample: (rag" P / G) (diameter, tim: I)

Laser type (pwr; X in ~m)l


Fluence (J / cm2) 10

Reagent

9 =r or

r o 9 >


#a ...1


Sample

Pretreatment Sample and chamber (pre-fluorination)

4~

MS Analyte

Yield

Blank

% T.Y. Yield1

Blank/Prep. correction (%0)

r r~ t,~ O !

Table 20.3 continued > o

>Table 20.3 continued +6X1(%o) q-&180(~o)

+6X2(%o) + 6 1 7 0 (%o)

=

9 3. MS 7

Additional Notes

Reference r

IRMS/irm-GCMS

o ~,,d.

o o

> v< ~,,Lo


ga

4~ 4~ o

> Table 20.3 continued Sample Type

Powder / grains (P / G); in-situ (I)

Size Sample" (mg: P / G) (diameter, tom" I)

Laser type (pwr; ;~ in/xm)l


Fluence (J / cm2) 10

Reagent

Table 20.3 continued > e3 to O !

O

>Table 20.3 continued +6X1(%o) _+6180(%o)

+6X2(%o) MS7 _+~34S(%0) IRMS/irm-GCMS

9

Additional Notes

Reference

~.

o S" o

r~ r 9 9 1_ pwr: power, in watts (W) 2_ Pwdr., w.r.: powder, whole-rock 3_ Mineral abbreviations: ol, olivine; g, garnet; mt, magnetite; ep, epidote; zr, zircon; tm, tourmaline; en, enstatite; sp, spinel; di, diopside; aug, augite; ky; ky; kyanite; mu, muscovitie; bt, biotite; pyx, pyroxene; p> pyrite; cp> chalcopyrite; gn, galena; spl, sphalerite; po, pyrrhotite; tr, troilite; acan, acanthite; grn, greenockite 4_ Absorbed on 5 ~ molecular sieve for transportation to mass spectrometer 5_ Indicated in Figure 20.4 of Farguhar & Rumble (1998) 6_ t: torr 7_ MS: IRMS, classical, dual-inlet isotope ratio mass spectrometry; irm-GCMS: isotope ratio monitoring-gas c h r o m a t o g r a p h y - m a s s spectrometry 8_ 0.1 atm = 10.13 kPa 9_ Ca. 100 m m in crater diameter 10_ j/cm2 on sample surface, but not reported by all (often laser power only is quoted, measured either above sample chamber or from exit port of laser) 11_ n.a., not available or not applicable

>

4~ 4x

442

Chapter 20 - B.E. T a y l o r

20.4.2 Fluorination of silicates and oxides by interhalogen fluorides Fluorination of silicates by interhalogen fluorides, principally BrF5 (Clayton & Meyeda, 1963) and C1F3 (Baertschi & Silverman, 1951; Borthwick & Harmon, 1982), can be described by the following example reactions (Table 20.2): KA1Si308(s)+ 8BrF5 ~ KF(s) + A1F3(s) + 3SiF4 + 402 + 8BrF3 3Mg2SiO4(s) + 8C1F3 ~ 6MgF2(s) + 3SiF4 + 4C12 + 602

[20.91 [20.10]

Similar to fluorination by F2, metal fluorides (some of which are volatile) can also comprise reaction products, in addition to oxygen (and, under certain conditions, new oxides; see below). In particular, fluorination by interhalogen fluorides produces BrF3, Br, or C1 and their fluorides as additional reaction products in the reaction tube(s). These are readily condensable, facilitating separation from the 02 gas in the reaction tube, but also add to the corrosion of the reaction tubes, or form hygroscopic deposits. Neutralization of excess of BrF5 (or C1F3) takes places after all oxygen has been removed from the reaction vessels. Consequently, less time is required for the extraction, collection, and conversion of oxygen from each sample when using BrF5 (or C1F3; e.g., often ca. _~10 minutes/sample) as compared with the F2 fluorination procedure. Moreover, the ability to condense excess BrF5, reaction-related waste products, and the CO2 converted from the released 02, permits cryogenic purification of the gas sample. Duration of the reaction is often determined by routine laboratory work schedules, and, accordingly, reaction times are typically on the order of 16 hours (i.e., overnight). This is a common time span between the end of the workday, when the reaction tubes have been loaded and furnaces turned "on", and the morning of the next day when the extractions are performed, and is often more a matter of convenience. Vennemann & Smith (1990) showed, for example, that fluorination with C1F3 at temperatures 450~ can be significantly shorter (e.g., 2hrs) for quartz, feldspar, biotite and pyroxene, and require only 4-8hrs at 600~ Nevertheless, from personal experience, it seems best (for a number of reasons related primarily to the reactivity of the reaction products) to keep the reaction vessels at their reaction temperatures until preparing the vacuum line for oxygen collection. Some minerals, as a reflection of their bonding and crystallographic structure, are especially resistant to fluorination, and their oxygen yields, which are not always 100%, depend on the fluorination reagent, and reaction time and temperature (Taylor & Epstein, 1962; Kyser et al., 1981; Vennemann & Smith, 1990). These include silicates such as high Mg-olivine, pyroxenes, garnets, and aluminosilicates, and oxides such as magnetite and corundum. For these minerals, reaction is promoted (but, not necessarily to 100%) by very fine grain size (few micrometers), longer reaction times, and higher temperatures (close to, but less than 700~ at which Ni begins to oxidize). In some cases (e.g., magnetite) higher oxygen yields are obtained when using BrF5 instead of F2. Unlike reactions with most minerals, however, a correct isotopic analysis of Mg-rich olivine can be obtained even with only an 80% theoretical yield of oxy-

Fluorination Methodsin StableIsotopeAnalysis

443

gen. In some cases (e.g., garnet) analysis is greatly facilitated by preparing a glass from the pure mineral (e.g., Taylor & O'Neil, 1977). Reuter et al. (1965; and also Garlick & Epstein, 1967) achieved 100% yields and accurate isotopic analyses from olivine by fusing known amounts of olivine and quartz (with a known isotopic composition) to form a pyroxene glass. Reuter et al. (1965) employed yet other methods, such as heating the sample to 550~ in a stream of H2, to remove oxygen and adsorbed H20 from slightly 'weathered' meteorites. Iron metal and sulfides were removed by reaction with dry C1 gas.

20.4.2.1 Yield and accuracy of ~180 values Garlick & Epstein (1967) found that the measured values of 6180 of quartz were lower when incomplete fluorination was due to insufficient reagent (BrF5). Although a 100% reaction yield from the sample as loaded is traditionally used as a criterion of the reliability of the reaction, there are exceptions, as noted above. This suggests that fluorination reactions do not all proceed in the same way. In the case of silicates (and quartz, in particular), partial, isotopically selective replacement of oxygen by fluorine in the SiO2 tetrahedral structure may be at work. Or, as suggested by Kohn et al. (1998) for micro-analysis of phosphate, some variability of 6180 may be due to formation of P-O-F compounds in cases when yields are low (e.g., 80%). In this instance, a correction must be applied. If reaction temperatures are allowed to exceed ca. 700~ NiO may also form in the reaction tube, reducing the yield, and (by virtue of the gasoxide fractionation) increasing the 6180 value of the residual 02. The formation of NiO has been suspected also at lower temperatures (S. Sheppard, pers. commun., 2003), with similar affects. 20.4.3 Additional fluorination procedures and comments 20.4.3.1 Step-wise fluorination, or "fluorination stripping" Partial fluorination reactions can be used to document isotopic inhomogeneity in minerals, whether of an intra-grain, or intra-crystalline nature, and to selectively analyze silicate-bonded oxygen. Step-wise fluorination, or fluorination "stripping", in a conventional fluorination line using F2, is accomplished by interrupting the fluorination reaction prior to completion. This technique has been used to detect, among others, grain-scale isotopic inhomogeneties and to infer isotopic zoning in quartz phenocrysts from hydrothermally altered rocks (Magaritz & Taylor, 1976).

Partial fluorination is more problematic with BrF5. Too little reagent, or too low a temperature, leads to 6180 values that are, respectively, too high or too low (Garlick & Epstein, 1967). These affects are apparently not so pronounced when F2 is used according to Epstein & Taylor (1971). However, Hamza & Epstein (1980) found that partial fluorination of kaolinite by F2-1imitation at 180~ yielded increasing 6180 values with extent of reaction. Using BrF5, Javoy et al. (1973) and Javoy & Fourcade (1973) found that initial 6180 values from partial extraction of oxygen from lunar fines were low, and interpreted this as due to contamination by adsorbed, low-180 atmospheric water. Similar results from a Hawaiian basaltic glass (control sample) led Javoy & Fourcade (1973) to suggest that partial fluorination could be accomplished

444

Chapter 20 - B.E. T a y l o r

using BrF5 without isotopic fractionation. Outgassing of the sample in the reaction tube, prior to fluorination, has been used (by Javoy et al. (1973), and others) in an attempt to rid the sample of adsorbed water. Inasmuch as nickel fluoride, which coats much of the inner wall of the reaction tube, and bromine, which may deposit at the cool end of the reaction tube and/or at the Swagelock ferrule on the top of the 'chimney' in the cap shown in Figures 20.3 and 20.4, are very hygroscopic, sample degassing in the reaction tube should be used with caution. Periodic checking and cleaning of the reaction tubes and caps is recommended. Reaction may proceed at any surface accessible during fluorination, including cracks, grain boundaries, crystal faces, as well as the grain surfaces, and, therefore, extreme caution should be used in interpreting the origin and isotopic composition of step-wise released oxygen. For example, the assumption of simple rim-to-core reaction (e.g., Epstein & Taylor, 1971; Magaritz & Taylor, 1976) was brought into question by in situ, high-spatial resolution isotopic analyses of quartz phenocrysts from Isle of Skye intrusions by Ion Probe, or Secondary Ion Mass Spectrometer (Valley & Graham, 1996) that demonstrated isotopic inhomogeneties caused by hydrothermal alteration followed inter-grain fractures. Where possible, knowledge of the textural context of any micro-scale isotopic variations is preferred, and, indeed, necessary to a correct interpretation.

20.4 3.2 Selective analysis by partial fluorination

Partial fluorination, which exploits the differences in reactivity, or reaction rate of various minerals, has been suggested as a useful technique for selective analysis of mineral mixtures. For example, Clayton & Mayeda (1963) noted that a one-hour reaction at 100~ between BrF5 and a difficult-to-separate quartz-feldspar mixture was sufficient to preferentially release the feldspar oxygen. Pumping away this oxygen while retaining the remaining BrF5 by cooling the reaction tube with LN2, allowed for reaction of the quartz in the usual way. The general procedure of separation by partial reaction seems to work best on simple mixtures of components with markedly different reactivity (e.g., analogous to procedures for mixed carbonates: A1-Aasm et al., 1990), but, again, interpretation must proceed with caution. Urey (1947) suggested, on the basis of a thermodynamic framework, that the oxygen isotope ratio of chemically and structurally different sites in minerals should differ. This raised the exciting possibility that intra-mineral isotopic fractionations could provide the basis for single mineral geothermometry. Several authors have investigated this potential (e.g., hydrous silicates: Savin, 1967; Hamza & Epstein, 1980; Bechtel & Hoernes, 1990; Girard & Savin, 1996; alunite: Pickthorn & O'Neil, 1985). Partial fluorination of pure minerals for such studies requires the use of F2. Savin (1967) and Hamza & Epstein (1980) attempted partial fluorination of hydrous minerals using F2 as a means for site-selective analysis in order to determine whether the OH-framework oxygen isotope fractionation could be measured. For kaolinite, Hamza & Epstein's (1980) procedure entailed an initial low-temperature (0~ reaction to extract OH-oxygen, followed by a high-temperature (100-450~

Fluorination Methodsin Stable IsotopeAnalysis

445

reaction to yield the remaining framework oxygen. Similar, two-step extractions were used for biotite, chlorite, muscovite, and phlogopite, differing only in the temperature of the second step (600~ in the case of phlogopite). Monitoring of recovered SiF4 assured that no framework oxygen was released during the first, 0~ reaction step. Intra-mineral oxygen isotope fractionations between OH- and framework oxygen were estimated to range from 5.2 (muscovite) to 12.6 (kaolinite), illustrating the expected increase with decreasing formation temperature and suggesting perhaps a fruitful field for further investigation (Hamza & Epstein, 1980). Bechtel & Hoernes (1985; 1990) devised a procedure for Fe-free minerals involving dehydroxylation and partial fluorination. Reduction of water released during vacuum dehydroxylation obviates this method for Fe-bearing minerals (Bechtel & Hoernes, 1990). Their methodology yielded an improved reproducibility over earlier results based on partial fluorination alone. Comparison of results from dehydroxylation and fluorination of the dehydrated residue and the original mineral indicated a consistent isotopic mass balance (and, therefore, complete analysis). Thus, using illite as single mineral geothermometer, one need only measure the 6180 values of the whole mineral and of the dehydrated residue. Internal fractionations of ca. 10.0 to 15.0%o were found for test samples. These fractionations correspond with temperatures of 200300~ (according to illite crystallinity studies), and are consistent with fractionations calculated from Sch~tze's increment method (Richter & Hoernes, 1988). Subsequently, Girard & Savin (1996) supported the results of Bechtel & Hoernes (1990) showing that the thermal dehydroxylation step provides more consistent results than partial fluorination for analysis of OH-oxygen. For reasons unclear, Girard and Savin (1996) were less successful in their attempt to determine and intra-mineral oxygen isotope fractionation, suggesting that factors such as crystal size, temperature of formation, etc. may be important to consider. In any case, caution in the application of partial fluorination and interpretation the results is advised. Acquisition of accurate and consistent oxygen isotope analyses of hydrated silica by step-wise extraction was used by Haimson & Knauth (1983), who demonstrated that step-wise fluorination successfully removed both adsorbed and absorbed water from hydrated silica. Yields of oxygen are, up to a fluorine excess of ca 2, linearly dependent on the abundance of fluorine present (Haimson & Knauth, 1983; Figure 20.6). Knowledge of the isotopic composition of different oxygen sites in single minerals can yield information on provenance or genesis. Supergene alunite (formed during weathering) can be distinguished from hypogene alunite (hydrothermal or magmatic) (e.g., Rye, 1993). As is the case with alunite, however, techniques other than partial fluorination are sometimes more suitable to site-selective, or mineral-specific analysis. For example, a simple, feldspar can be removed at room temperature from quartzfeldspar mixtures by pre-treatment with HF or fluorosilicic acid in Teflon beakers, followed by careful washing and separation of the residual, largely unreacted quartz. Although more time-consuming than the previously mentioned partial fluorination procedure of Clayton & Mayeda (1963), the accuracy of isotopic results from pure

446


minerals is assured.

20.4.4 Fluorination of phosphates and sulfides by halogen fluorides and F2 gas Tudge (1960) originally described the fluorination of biogenic phosphate (pre-purifled as BiPO4) for oxygen in PO42- using BrF3. The reaction: BiPO4(s) + 8 / 3BrF3 ~ BiF3(s) + 4 / 3Br2 + 202

[20.11]

was quantitative and took place for 30 minutes at ca. 100~ in an externally heated Teflon reaction tube. BrF3 is liquid below ca. 125~ and was considered the fluorinating reagent of choice because the reaction could be carried out at low temperature. Until BiPO4, was replaced by Ag3PO4 (e.g., O'Neil et al., 1994; see also Firsching, 1961) as a more stable compound, reaction [20.11] represented the classic technique. More commonly, however, fluorination was accomplished using BrF5, C1F3, or F2 (Longinelli & Nutti, 1973a; Vennemann et al., 2002). Fluorination of BiPO4 (and presumably using Ag3PO4) by BrF5 carried out at high temperature (e.g., 500~ 9Vennemann et al., 2002) yields no systematic differences from results obtained using BrF3 (Longinelli & Nutti, 1973a). The reaction likely proceeds as follows" Ag3PO4(s) + 4BrF5 ~ 3AgF2(s) + PF5 + 3BrF3 + 1/2Br2 + 202

[20.12]

indicating the need for a molar fluorine/oxygen ratio of at least 5 for complete reaction. Fluorination of abiogenic apatite [Ca5(PO4)3(OH, F,C1)] by BrF5 proceeds via a reaction analogous to [20.12], but requires high temperatures (e.g., 650~ to give complete oxygen yields (Rhodes & Oreskes, 1999; B. Taylor & Mirnejad, unpub.). Conversion of sulfide minerals and elemental sulfur to SF6 permits the most accurate determination of 634S, in addition to 633S and ~36S. Unlike SO2, which is polar and readily absorbed on surfaces, SF6 is non-polar and chemically inert. Consequently, its memory in the ion source, though finite, is minimal in comparison to that of SO2. Differences in absorption and pumping characteristics are largely to blame for discrepancies in analytical results using one or the other gas (e.g., Taylor et al., 2001a, b; Beaudoin & Taylor, 1994; Rees, 1978). Moreover, fluorine is mono-isotopic, unlike oxygen in SO2, which obviates any corrections for isobaric interferences in the ion spectrum. Hence, the four SF5+ ion currents measured as masses 127, 128, 129, and 131, are attributed to the stable sulfur isotopes 32, 33, 34, and 36, respectively. Puchelt et al. (1971) described the preparation of SF6 by conventional fluorination of sulfur and sulfide minerals using BrF3, following the reaction: 2FeS2(s) + 10BrF3 ~ 4SF6 + 2FeF3(s) + 5Br2

[20.13]

carried out in a nickel reaction tube, externally heated to 200~ overnight. Repeated


447

cryogenic distillation pre-purified the SF6 prior to final purification on a 5A mol sievepacked 5ft x 0.25in column. The measured 634S values were reproducible to 0.1%o. Whereas, elemental sulfur, greenockite (CdS), sphalerite (ZnS), galena (PbS), cinnabar (HgS), covellite (CuS), argentite (Ag2S), and pyrite (FeS2) reacted readily, giving yields of 96-98%, low yields were obtained from pyrrhotite (FeS), troilite (FeS), and digenite (Cu9S5). The latter minerals were converted to CdS for analysis. Thode & Rees (1971), Hoering (1990), and Gao & Thiemens (1991) have reported fluorination of sulfur compounds with BrF5. Gao & Thiemens (1991) used a conventional fluorination line, similar in many aspects to that shown in Figure 20.1, for overnight (16h) fluorination of sulfide minerals (previously converted to CdS) at 450~ with a 150X stoichiometric excess of BrF5. Thode & Rees (1971), on the other hand, found that a 300~ 16 h reaction with a 20X molar excess of BrF5 sufficed to quantitatively convert 3-7mg aliquots of Ag2S to SF6. In either case, accurate measurement of ~36S following cryogenic distillation requires further purification of the SF6 by gas chromatography. This process is time-intensive. For example, GC purification described by Gao & Thiemens (1991) required as much as an hour or more per sample, plus a 5h bake-out (100~ of the GC column (80-100 mesh Porpak Q in a 12ft, 1/8 in OD column) between samples. Hulston & Thode (1965b) prepared SF6 by fluorination of Ag2S (prepared from each sample) with F2 gas because it was found that fluorination with BrF3 by reaction [20.13] resulted in yields of only 75%, albeit without apparent isotopic fractionation. Although not noted specifically, yields were presumably improved with the use of F2 gas. The SF6 was purified on a 5~, mol sieve-packed column held in a temperature gradient (150~ to room temperature). Contaminants in the column were removed after ca. 6 samples by an overnight bake-out at 220~ Largely because of the time-consuming nature of sulfide fluorination in this manner, fluorination of sulfides did not gain wide use, in spite of the mass spectrometric advantages of SF6 (Puchelt et al., 1971; Rees, 1978) and the ability to measure 633S and 636S. Simpler, but less accurate techniques that produced SO2 for measurement of 634S have been more widely used (Taylor et al., 2001a,b). Unlike SO2-based laser methods, fluorination of sulfides to form SF6 occurs without isotopic fractionation, and, consequently, no sample-dependent correction factors are needed (Beaudoin & Taylor, 1994). The advent of laser-assisted fluorination (e.g., Sharp, 1990), combined with capability to safely generate pure F2 in the laboratory (Aspre~ 1976), has begun to re-vitalize fluorination of sulfides for isotopic analysis (e.g., Rumble et al., 1993; Beaudoin & Taylor, 1993; Taylor & Beaudoin, 1993; Beaudoin & Taylor, 1994). Fluorination by F2 may follow a simple reaction such as" FeS2(s) + 7.5F2 ~ FeF3(s) + 2SF6

[20.14]

448


that yields no halogen waste products such as form in reaction [20.13]. In practice, other fluorine-bearing compounds may form (e.g., HF, CF4, and complex fluorocarbons) by reaction with contaminants in the sample chamber or the sample. It is the fluorocarbons, which pose the greatest problem in sample purification, but these can be generally separated in a timely manner. Considerable purification, sufficient for precise and accurate measurement of 633S and 634S, can be accomplished without gas chromatography by use of a variable temperature trap (Taylor & Beaudoin, 1993; Beaudoin & Taylor, 1994; Taylor & Beaudoin, 2000; Coleman, M. Part 2, Chapter 44, this volume; see Figure 20.8). A variable temperature trap offers greater accuracy and precision in cryogenic separation than is possible using traditional, fixed-point cooling mixtures (e.g., dry ice + acetone), and fewer potential health and safety hazards. Sulfur isotope anomalies have previously been associated with cosmogenic studies (e.g., Hulston & Thode, 1965b; Farquhar & Thiemens, 2000). However, the recent discovery of terrestrial 633S anomalies and their significance for evolution of the Earth's early atmosphere (e.g., Farquhar et al., 2000a) have re-generated interest in sulfide fluorination. Such anomalies can only be detected by isotope ratio measurements made using SF6. Also, the increased the accuracy of the SF6 method (Taylor, et al., 2001, in press), will likely lead to wider spread use of fluorination procedures in sulfur isotope analysis. 20.4.5 Fluorination of selenium and selenides for selenium isotopes

Fluorination of elemental selenium was the original method of choice used to measure the isotopic ratio of 82Se/76Se in various types of samples, from native selenium and from selenides, to rare selenite and selenates, and including trace concentrations of selenium in sulphides, organic materials, and in solution (Krouse & Thode, 1962). Figure 20.8 - (A) High-surface area P-trap ensures efficient cryogenic trapping of even trace amounts of a condensable gas, whether from a high-throughput He-carrier stream or, for example, during evacuation of a sample-chamber filling F2 atmosphere. High surface area is provided by the 60mm filter element (Swagelock| SS-6TF-60) used as a "foot" in the P-trap (see Figure 20.8A). The P-trap design is achieved by arc-welding a 1 / 4" tube to 3/8" OD sleeve, providing for a vacuum fit inside bored-through 3/8" stainless "tee"; (B) Variable temperature trap (VTT; modified after design of DesMarais, 1978b), provides for cryogenic trapping and separation with 1~ accuracy. Temperature of the VTT may be varied from ambient to nearly -197~ by combined heating of Ni-chrome heating wire, with variable transformer connected to low voltage (e.g., 24 V.A.C.) transformer via power feed-through, and simultaneous submersion of VTT in liquid nitrogen. VTT may be dismantled for repair at indium-wire sealed flange if ever necessary. Internal Cajon flexible metal hose, with factory-supplied 1/4" tube ends, is connected to the flange via welded Swagelock @ components and to the vacuum line shown in Figure 9 via 1/4" stainless steel Swagelock | run-tee-components having female pipe thread on one arm. A thermocouple inserted inside flexible hose, and vacuum-sealed in a reducing fitting with Teflon ferrules, provides for monitoring of temperature (slight differences in placement of thermocouple tip may result in slight variation in indicated temperature). One way, 10psirated poppet check valve (e.g., Nupro | B-4C-10) provides for safe release of condensed oxygen should outer envelope of VTT leak while submersed in liquid nitrogen. Modified after Taylor (2003) and Taylor & Beaudoin (1993).


449

450


Krouse & Thode (1962) extracted selenium from a range of sample types using either hydroxylamine hydrochloride for solutions, or a hydrobromic acid-bromine extraction procedure (referenced in their paper) for whole-rock samples and minerals. This extraction procedure provides a way to measure selenium isotope compositions in sulfide and sulfosalt minerals in which selenium may commonly occur as a substitution for sulfur. The fluorination of selenium by F2(g) evidently produces fewer contaminants than when carried out using CoF3(s) (c.f., Krouse & Thode, 1962; Webster & Warren, 1981), yet the uncertainties in measured ~82Se values in these studies was similar (ca. 0.5 %0). Liquid air and liquid oxygen cryogenic traps were particularly useful in removing a fluorocarbon contaminant of some type that caused a mass interference (R. Krouse, pers. commun., 2003). Where F2 is the fluorinating reagent, the reaction is relatively simple, and could be ideally written ignoring products of incomplete fluorination, such as Se2F2 and SeF4, and contaminants (e.g., CF4) as: Se(s) + 3F2 = SeF6

[20.15]

Fractionations among natural samples up to 15%o were discovered using the fluorination process, with the largest isotopic deviations found in organic compounds, presumably due to kinetic isotope effects during selenate reduction (Krouse & Thode, 1962). Krouse & Thode (1962) noted that although not as large in magnitude as known for sulfur isotopes, the association of larger fractionations with biologic processes in was, indeed, analogous to the behavior of sulfur isotopes. In the case of selenium isotopes, preparation of samples via the fluorination technique will likely give way to instrumental methods. Thermal ionization mass spectrometry (TIMS) employing negative ions, or N-TIMS, was shown by Wachsmann & Heumann (1992) to permit determination of the isotopic ratio of 80Se/76Se to a precision of 1-3%o (i.e., poorer than by fluorination) for g80Se. A double-spike technique devised by Johnson et al. (1999) currently yields a precision of 0.2%o. The N-TIMS technique requires the concentration and purification of selenium from the sample, but facilitates isotopic measurements on samples with considerably lower Se contents than previously achieved. Multi-collector ICP-MS (inductively coupled plasma mass spectrometry; see Rehk/imper, M., et al., Chapter 31, this volume) provides yet another step in the direction of permissible lower Se concentrations, and has recently been used to measure 82Se/76Se ratios in a variety of rock types, soil, sediment, and naturally-deposited silica (Rouxel et al., 2002). Although this technique has its own associated potential errors of mass interference and contamination, and requires a pre-analysis acid sample digestion, the precision achieved thus far of ca. 0.25%0 (2~J) for ~82Se surpasses the original results by fluorination with far greater ease and economy of time (Rouxel et al., 2002). Most importantly, the multi-collector ICP-MS instrumentation can determine Se isotopic compositions in samples containing nanogram quantities (gg/g) of Se, with little, if any sample matrix related problems. The multi-collector ICP-MS approach is certain


451

to comprise a principal research tool. The N-TIMS technique is viable, as well, but the current reporting of selenium isotope compositions by N-TIMS as 680Se values rather than traditional 682Se values complicates comparison of results obtained by other techniques. 20.4.6 Fluorination of silicates for silicon isotopes SiF4 is produced during fluorination of quartz, for example, using BrF5 as follows:

SiO2(s) + 2BrF5 = SiF4 + 2BrF3 + 02

[20.16]

The product SiF4 may be cryogenically separated and purified at-80~ (e.g., using dry ice+acetone), or, ca.-70~ or colder, using liquid + frozen ethanol (ethanol-liquid nitrogen mixtures) monitored with a thermocouple. Any trace amounts of active fluorine compounds produced by reaction [20.16] may be removed by reaction with Zn granules at 50-60~ in a Cu tube to form condensable ZnF2 and ZnBr2 in three distillation stages (Ding et al., 1996). These include two successive stages of trapping and vacuum pumping at ca.-80~ followed by reaction with Zn, the trapping of the purified SiF4 at-197~ (liquid nitrogen temperature) and removal, by vacuum pumping, of any non-condensable gases released during the hot Zn treatment. This step constitutes an improvement over previous methods using F2 (Epstein & Taylor, 1971), or BrF5 (e.g., Clayton & Mayeda, 1978; Douthitt, 1982). When silicates or purified silica are fluorinated via a reaction analogous to reaction [20.16] using F2, reaction products do not, of course, include Br and Br fluoride compounds. Rather, these form in the waste line during neutralization of the F2. Ding et al. (1996) note, in particular that purification of the BrF5 reagent by distillation at-70~ may be necessary to attain high precision. 20.4.7 Fluorination of water by halogen fluoride The fluorination method described by O'Neil & Epstein (1966b) was designed to analyze but a few milligrams of water. Although miniaturization of the CO2-H20 equilibration technique (e.g., Kishima & Sakai, 1980; Ohba, 1987; Socki et al., 1999) facilitates indirect analysis of several milligrams of water, the uncertainty in measured 6180 values is, except for the study by Socki et al. (1999) utilizing 10rag of water, larger than by fluorination. Usually at least 1 ml of water is required for analysis by conventional CO2/H20 exchange procedures. Fluorination of water with a 3- to 4fold excess of BrF5 takes place rapidly at 80-100~ according to the reaction:

BrF5 + H20(1) ~ Br + 3/2F2 + 2HF + 1/202

[20.17]

with a yield of 100% (O'Neil & Epstein, 1966b). Although the possibility of fluorination by F2 was acknowledged (O'Neil & Epstein, 1966b), a method for water analysis by this means is not published. Reaction [20.17] proceeds rapidly. Loading of sample and reagent, plus evacuation, occupy the bulk of the analytical time. The principal difficulty is the transfer of water (in this case, small quantities), in a vacuum line containing abundant hygroscopic fluorides on the walls of the reaction vessel. Both of the two methods, capillary and direct cryogenic transfer devised by O'Neil & Epstein (1966b)

452


avoided this potential problem and yielded similar results. This method not only offers direct analysis with excellent precision (5 hours at ~ 100~ and then weighed. The percentage of c~-cellulose in the wood depends on the material (wood or leaves) and the tree species (See section 24.4.1 and Table 24.2). 24.3.7 Remarks 9Prior to each solution step the samples are rinsed with de-ionized water. Between the NaC102 and the NaOH steps, and after the NaOH step, the samples are rinsed to closely match pH ~ 7 (3 - 5 rinsing processes are required). ~ Samples larger than ~ 100 mg may need longer extraction. One has also to be aware that softwood (conifers), as well as leaf material, is less dense than hardwood

515

Sample Homogeneity and Cellulose Extraction from Plant Tissue for Stable Isotope Analyses Table 2 4 . 2 - Cellulose yields. species

10 10' 11 11' 11' 11' 12' 14 14 14

Fagus Quercus Quercus Quercus Fagus Quercus Quercus Quercus Quercus Quercus Quercus Quercus Quercus Quercus Quercus Quercus Quercus Fagus Betula

site

Li Li Li Li Li Sa Sa Sa Sa Sa Sa Sa Sa, Li Sa Li Li Sa Kr WSL

date of extraction

1 / 96 2/96 3 / 96 4 / 96 4 / 96 5 / 96 6 / 96 6 / 96 8 / 96 11 / 96 11 / 96 12 / 96 12/96 12 / 96 12 / 96 4 / 97 1 / 99 1 / 99 1 / 99

mean yield [%]

o

slope

+ slope

intercept

+ intercept

[%]

[%]

[%]

[mg]

[mg]

39.0 41.6 43.9 41.9 39.7 35.5 36.4 29.3 32.4 32.0 35.7 35.7 39.0 36.7 40.1 41.6 32.1 39.1 25.9

0.9 2.4 1.7 1.9 1.2 4.7 3.4 5.1 3.8 6.2 4.2 4.7 4.1 3.8 4.0 3.3 1.1 1.5 2.7

40.3 46.1 49.8 49.5 41.4 39.8 44.9 39.5 32.2 33.5 30.3 43.1 42.8 45.7 40.3 41.8 32.8 40.5 57

0.7 0.6 1.0 3.8 2.2 0.9 1.7 1.2 2.1 7.8 2.4 0.7 1.4 0.7 2.7 1.7 4.1 0.9 12

-1.58 -1.05 -4.99 -8.25 -2.88 -1.84 -3.72 -6.74 -0.02 -4.17 1.20 -2.94 -1.92 -4.95 -0.07 -1.98 -0.72 -0.84 -31.7

0.92 0.19 0.88 4.23 3.99 0.64 0.85 1.43 1.13 4.14 1.23 0.51 0.91 0.50 1.48 0.60 4.35 0.71 12.1

n

r2

26 27 28 8 20 28 28 20 30 30 29 30 30 10 19 19 8 11 6

0.993 0.995 0.991 0.966 0.950 0.986 0.963 0.983 0.937 0.396 0.859 0.993 0.972 0.998 0.931 0.971 0.914 0.995 0.855

Typical cellulose yields obtained with our extraction method for different tree species. Because of losses during the extraction and of the non-quantitative removal of the extracted cellulose from the glass-tube, the real yield cannot be determined directly as the ratio of the w o o d and cellulose masses measured before and after the extraction. In order to take these losses into account, we determined the yield by the use of the linear regression method demonstrated in Figure 24.4 (see text, and especially equation 24.2) The first column of the Table gives the number of the extraction; the second is the tree species; the third is the site name: "Li" is a site in Northern Germany, "Sa" and "Kr" are sites on the Swiss ,,Mittelland", and "WSL" stands for a 1-year old birch grown in a controlled chamber. Column 4 gives the date of extraction (month and year). Columns 5 and 6 are the mean values and the standard deviations (o), respectively, of the ratio of measured cellulose to wood mass. Columns 7 to 12 give the results of the linear regression described above: slope (~ real cellulose yield), error of the slope, intercept (- minus lost mass), error of the intercept, number of samples in the series (n), and correlation coefficient of the regression (r2). Obvious outliers have been excluded prior to the calculations.

(broad-leaved trees), which influences the duration of the extraction (softwood, and leaves, needing longer extraction time, due to the larger volume of the samples). 9In order to improve the extraction, and its rapidity, we suggest to stir the samples approximately once every hour with a glass stick; this is especially necessary for large samples. 9 The drying temperature for organic samples should never exceed 100~ partial combustion which could create large isotopic fractionation.

to avoid

9The whole procedure lasts one to two weeks (excluding milling), depending on the number of NaC102 steps necessary until the samples are white. The total duration

516

Chapter 24 - S. Borella, G. M6not & M.

Leuenberger

will depend on whether one has several Soxhlet extractors to run in parallel.

24.3.8 Other methods Several other procedures for R-cellulose extraction from wood samples have already been described during the last decades. Our method is primarily derived from the technique of Brenninkmeijer (1983), with major practical improvements. Older methods used benzene-methanol instead of toluol-ethanol as organic solvent for the first step. Toluol-ethanol works as well as benzene-methanol mixture, and reduces health risks associated with the use of benzene and methanol. Here we like to briefly mention other techniques for R-cellulose extraction published in the last decades. A review of several older methods can also be found in (Green, 1963). Sheu & Chiu (1995) evaluated cellulose extraction procedures for stable carbon isotope measurements in tree ring research. The method they finally recommend involves two main steps: (1) extraction with benzene-ethanol in a Soxhlet extractor for 12 hours, followed by (2) bleaching and soaking in NaC102-CH3COOH solution. At the end of the procedure they also soak their cellulose in distilled water at 70~ for 6 hours. They recommend this method as being more accurate for 613C measurements than when the NaOH step is added, which means that they work with holocellulose instead of ~-cellulose. However, a closer look at their data shows that the accuracy of their holocellulose ~13C results is not significantly better than their Rc e l l u l o s e ~13C results. Loader et al. (1997) also presented a method which is quite similar to ours. The main differences are: (1) They use organic solvents (2" 1 toluenemethylated spirit azeotrope) only for softwood, (2) instead of milling their samples, they cut them into fine slivers (~ 40 mm), and (3) the different steps take place in an ultra-sonic bath, which enhances the removal of extracts. Recently, Brendel et al. (2000) presented a new rapid and simple method for extracting R-cellulose. With this method they obtained a very constant c~-cellulose content for samples weighing between 10 - 100 mg, indicating that the extraction protocol is highly reproducible. The advantage of this method may be its rapidity of cellulose extraction (single day process), when devoting the whole working time to it. There are also other laboratories which do not extract R-cellulose for their stable isotope analyses, but use holocellulose (Leavitt & Danzer, 1993), or lignin, instead. The method used by Leavitt & Danzer (1993) is mainly the same as ours, except that their organic solvent step involves two sub-steps (one with 2 91 toluene-ethanol and one with pure ethanol), and that they don't extract the hemicellulose with NaOH at the end of the procedure.

24.3.9 R-cellulose extraction from shrubs, herbaceous and non-vascular plants After collection, plants are dried and milled with a modified coffee-mill similar to the one described in section 24.3.1. Shrubs, annual plants and mosses have a much lower cellulose content than the 30 - 50% cellulose typically found in woody plants. For example, cellulose extraction yields are typically around 25% for Calluna vulgaris and Vaccinium uliginosum, around

Sample Homogeneity and Cellulose Extraction from Plant Tissue for Stable Isotope Analyses

517

15% for

Eriophorum vaginatum and around 10% for two non-vascular plant species, Sphagnum magellanicum and S. capillifollium.Therefore, a much larger quantity of plant material is initially required. We started cellulose extraction of sphagnum, for example, with 500 mg of dried material. To handle this larger quantity we used glass tubes of 100 mm height but with a larger outer diameter of 24 mm, compared to 10 mm for the tubes used for cellulose extraction from tree rings (Figure 24:. 3 and Picture 24.A2). The number of samples treatable in one run is then limited by the size of the Soxhlet apparatus. In our case, we can perform 7 sample extraction in parallel. Because the chemical composition differs for vascular plants and for non-vascular plants, they require a slightly different extraction treatment. For shrubs and herbaceous plants, the extraction protocol is quite similar to that described for wood. We use an extraction method adapted from that of Brenninkmeijer (1983). Lipids, resins and waxes are first dissolved by a Soxhlet extraction with a solution of toluol-ethanol (1 91). This step, not really decisive for wood, is much more important and requires more time in the case of non-woody plants. The duration of this step is not completely fixed, we stop when the toluol-ethanol solution, renewed every 15 - 20 min., remains colorless. This may take, for example, up to 14: hours for Andromeda polifolia and about 8 - 10 hours for the grass-like plants Carex pauciflora and Eriophorum vaginatum. Samples are then dried over night at 100~ The lignin fraction is removed next with a solution of NaC102 mixed with acetic acid at 70~ with proportions of 5 g NaC102 and 1.5 ml acetic acid for I 1 water, as for extraction from wood. Again, the only difference between the two protocols is the required time. The NaC102 step may last 8 - 24: hours depending on plant type. It is very important to stir the samples and to renew the solution often (every 1/2 hour and every 2 hours, respectively) during the first hours of extraction, else the solution can not reach the whole sample, due to the large volume of the sample. This "bleaching" step is repeated until the samples turn completely white. Samples are then washed under vacuum filtration and by immersion in de-ionized water until the pH of this rinsing solution reaches neutral values. At this point, samples consist only of holocellulose. The extraction of the hemicellulose fraction is performed in a 4% NaOH solution at 80~ Once again, compared to the wood protocol, this step takes much longer. As a final step, a-cellulose is washed for one day with de-ionized water at 80~ For cellulose extraction from mosses, the protocol used is based on (Wise et al., 1946) and adapted from (Price et al., 1997). The resin and lipid fractions are removed, also by Soxhlet extraction, but using three different organic solvents" first a mixture of chloroform and ethanol (2" 1), then a pure ethanol solution and finally de-ionized water. As previously mentioned, the timing of each step can not be completely fixed, and is dependent on the species. For example, the chloroform / ethanol extraction lasts 12 hours for S.cuspidatum but 20 hours for S.capillifollium. Again, the extraction is complete when the solution is clear. For the 2nd and 3rd Soxhlet extraction steps, the timing is less species dependent. Each of these steps lasts around 10 hours. Between each step samples are dried over night in an oven at 90~ The two next steps (NaC102 and NaOH) are identical to those used for vascular plants. However sphagnae are really difficult to rinse, especially after the NaC102 step. One reason may be that as

518

Chapter 24 - S. Borella, G. M6not & M. Leuenberger

long as plant structures are not completely destroyed, the large hyaline cells which constitute the sphagnum structure fill up with the various solutions, making it very hard to remove the solution within the cellular structure. Each rinsing step lasts at least one day. The de-ionized water is replaced several times and samples are washed by vacuum filtration. The use of an ultra-sonic bath may decrease this "rinsing" time. In both cases, extraction products are dried and the non-vascular plants samples are milled again prior to stable isotope measurements.

24.4 Reproducibility of the c~-cellulose extraction One can think of two independent methods for testing the reproducibility of the c~cellulose extraction. Firstly, by comparing c~-cellulose yields of different extractions of the same type of material (because of different cellulose content for wood and leaves, or for different tree species)(Brendel et al, 2000). Secondly, by comparing the isotope ratios of c~-cellulose obtained from different extractions of the same (well homogenized) wood. The first method is less precise due to sample loss which is difficult to quantify.

24.4.1 (x-cellulose yield The determination of the cellulose yield based on the measured sample weights (before and after the extraction) is only moderately accurate, since material loss is not avoidable and difficult to quantify. One way to take account of this loss is a linear regression analysis between wood and cellulose mass (Figure 24.4). Indeed, the measured cellulose mass of a wood sample is the real cellulose mass minus the lost cellulose mass: Cm = C r - n

[24.3]

where C is the cellulose mass, L the lost cellulose mass, and the subscripts m and r assign the measured and the real value, respectively. The real cellulose mass can be written as the product of the wood mass times the percentage of cellulose in the wood, which gives: Cm = ar 9 W r - L

[24.4]

where W is the wood mass, c~ is the percentage of cellulose in the wood sample, and the other symbols are defined as in equation [24.3]. Hence, a linear regression analysis of the measured cellulose mass as a function of the wood mass gives us the cellulose percentage and the mean loss, under the assumption that the loss is constant, and for each tube the

Figure 24.4 - Cellulose mass as a function of the wood mass. The real cellulose yield and the mean loss can be estimated from the slope (46.1%) and the intercept (1.05mg) of this linear correlation, respectively (see text).

Sample Homogeneityand Cellulose Extraction from Plant Tissue for Stable Isotope Analyses

519

same. This is of course not the case, but since we have no direct determination of this loss, this method gives us at least an estimate of it. To document the usefulness of such a treatment we present results of the regression analyses for fourteen extraction series obtained with our cellulose extraction method (Table 24.2), together with the mean cellulose yield calculated as the ratio of the measured cellulose and wood masses. It would be helpful for other institute that such quality tests are mentioned in publications. The results shown in Table 24.2 suggest that the cellulose yield is generally underestimated by up to 9%, when the material loss (estimated to 1 to 8 mg) is neglected. Nevertheless, one has to be careful in interpreting the results of such correlation analyses, since the loss could be mass and tube dependent, and outliers could strongly falsify the results. Moreover, a sample serie could be inhomogeneous in that earlywood and latewood, or different trees of the same species may have a significantly different cellulose content. This warning is well justified by some extreme results of Table 24.2, like the positive intercept obtained for extraction series number 10', which would correspond to a negative loss! However, this method allows us at least to get an estimate of the loss and to revise the cellulose content accordingly. We found in this way that beech and oak wood contain 40 + 5% R-cellulose. For comparison, we give the values obtained some years ago with the same extraction method for different materials by Borella (1994)" 29.0 + 2.2% for wood of a 1-year old birch; 33.8 + 1.6% for poplar wood; 37- 44% for beech wood; and 6 - 11% for birch leaves.

24.4.2 Stable isotope values In order to test the influence of the cellulose extraction on the isotopic ratio of the obtained R-cellulose, we processed the same wood sample (birch wood) several times and determined the ~13C values of the different subsamples. The results of these tests are shown in Figure 24.5. The extraction series 8 and 9 (Figure 24.5: the three points to the right) cannot be compared with others, since the extraction has been incomplete (series 8) or the material was not well homogenized (series 9). The lower 613C value of series 8 shows that the completeness of the extraction is an important criterion (see also Figure 24.2). After eliminating series 8 and 9 (Figure 24.5), we can see that the inhomogeneity due to different cellulose extractions does not exceed the uncertainty of the ~13C determination (~ 0.1%o for on-line techniques, down to ~ 0.02%0 for off-line techniques, as shown by the error bars to the right of Figure 24.5 for extremely fine - and hence homogeneous - commercial cellulose; see Chapter 23). Even more important to show the reliability of our cellulose extraction method are the results of the samples that have been extracted twice: no significant shift in 613C is observed between the single and the double extractions (compare the triangles and the squares in extractions number 2 and 9, in Figure 24.5). The i513C gradients following the several extraction steps (Figure 24.2) also clearly documents the importance of a complete extraction. Indeed, the difference in 613C between pure a-cellulose and wood is of the same order of magnitude as the year-to-year variations in tree ring cellulose (see for example Borella, 1998). Even the shift caused by skipping one or two NaC102 steps is

520 comparable to the 613C measurement reproducibility. However, the flattening of the trend between the 2nd and 4th NaC102 steps confirms the results of Figure 24.5, that the R-cellulose extraction is complete. It is important to note that these conclusions must not be true for 6180 and 62H. Indeed, Brenninkmeijer (1983) found that a prolonged cellulose extraction reduces the cellulose yield and could produce a shift in the oxygen and hydrogen isotopic composition of the remaining R-cellulose. However, prelimnary results from analogous tests for 6180 show no significant difference between two R-cellulose subsamples extracted from the same well homogenized wood sample.


Figure 24.5 - 613C of different extractions of our home-made standard cellulose (,,WSL", wood from a 1-year old birch), average values for each extraction and measurement technique: off-line (full symbols) or on-line (open symbols) techniques (see chapter from Saurer). The error bars represent the standard deviation (1 o), as determined from multiple measurements; symbols without error bar represent a single measurement. The triangles represent cellulose that has been extracted twice. The circles represent cellulose extracted from coarse milled wood. The error bars at the right handside of the figure represent the standard error of the single measurement, as determined for the off-line (solid line) and on-line (dashed line) method with extremely fine (< 0.01mm) commercial cellulose. The different symbols (circles and triangles, extraction numbers 1, 2, and 9) are shifted to the right to enhance the clearness of the figure. The ,,WSL" cellulose from extractions 7 and 10 has not been analyzed for ~13C.

Furthermore, we like to note here that for certain applications requiring not the highest possible reproducibility it is possible to measure the isotopic composition directly on wood. Year to year variations in lignin and holocellulose are approximately the same (Mazany et al., 1980), although lignin seems to be produced later during the growing season than cellulose by the trees (Wilson & Grinsted, 1977; Tans et al., 1978). At least for broadleaved trees, we could show that it is not necessary to extract cellulose, nor lignin prior to ~13C measurements; wood reflects the same year to year variations within the measurement uncertainties (Borella et al., 1998). However, similar results for ~180 values were not so conclusive. These findings have not to be true for other tree species. 24.5 Conclusions

The method for ~x-cellulose extraction from wood presented in this chapter is highly reproducible in regard to the carbon and oxygen isotopic composition. Moreover, it allows a high sample throughput (up to 120 samples in two weeks are possible, excluding the time necessary for milling). In order to achieve reproducibility of 0.1%o for 613C and 0.2%o for ~i180, it is important to finely mill the samples prior to the extraction, and to perform the NaC102 steps until the samples are white. We found

Sample Homogeneity and Cellulose Extraction from Plant Tissue for Stable Isotope Analyses

521

that a w o o d g r a i n size of a p p r o x i m a t e l y 0 . 1 5 m m h a s to be r e a c h e d , if s u b s a m p l e s of 1 - 1.5 m g are a n a l y z e d . M o r e o v e r , fine milling of the w o o d is n e c e s s a r y to p e r f o r m a c o m p l e t e c~-cellulose extraction, a n d e v e n l o w c o n c e n t r a t i o n of lignin in the c~-cellulose can l o w e r the ~)13C b y u p to 0.2%0.

Appendix

Picture 24.A1 - Coffee-mill used for milling the wood samples. We can see the double blade in the ,,body" of the mill (above), and the additionally mounted deflection plate in the cover (below).

Picture 24.A2- Filter glass tube used for the extraction of cellulose. The ground up glass at the bottom of the tube serves as filter; it lets the solution through, but not the wood (respectively the cellulose). The tube is 10cm high, has an outer diameter of 12mm and an inner diameter of 10mm. (For the dimensions, see also Figure 24.3). The number seen at the top of the glass has been engraved to avoid any hazard of mixing samples. [Picture 24.A2 on next page!]

522


[Picture 24.A2 - see caption on f o r m e r page]

Picture 24.A3 - Teflon holder for the filter tubes (see Picture 24.A2) used for cellulose extraction. The holder has a diameter of 90mm and height of 35mm. There is place for 30 filter tubes. The holes for the tubes are a little smaller at the bottom to avoid the tubes to go through. Only the central hole goes down to the bottom without diameter reduction. The holder with the 30 sample glasses can be hold by putting the metal piece shown at the bottom of the picture into the central hole. Pressing the two long metal bars together permits the teflon holder to sit on the two curved parts of the metal piece, seen at the bottom left of the picture


CHAPTER 25 Analytical Methods for Silicon Isotope Determinations Ding Tiping Institute of Mineral Resources, Chinese Academy of Geological Sciences, Beijing 100037, P. R. China e-mail: [email protected]

25.1 Introduction Four isotopes of silicon exist in the natural environment: 28Si, 29Si, 30Si and 32Si. The first 3 isotopes are stable isotopes and the last one is radiogenic. The relative abundance of 28Si, 29Si and 30Si is 92.23%, 4.67% and 3.10%, respectively (Barnes et al., 1975). By the 1920's, all three stable silicon isotopes had been discovered. Mass-spectrometric studies on silicon isotope variation in the natural environment started in the 1950's (Reynolds & Verhoogen, 1953; Allenby, 1954). In the 1970's, extensive studies on silicon isotope compositions of meteorites and rocks were made (Epstein & Taylor, 1970; Yeh & Epstein, 1978; Clayton et al., 1978, Clayton 1986). Douthitt (1982) reported a number of data on silicon isotope composition of terrestrial samples. Since 1988, the author and his co-workers have carried out a long-term studies on the silicon isotope composition of a variety of natural materials (Ding et al., 1988, 1990; Ding 1991; Ding et al., 1994, 1998; Jiang et al. 1992, 1993; Wu et al. 1997; Song & Ding, 1990). Their results were summarised in a book entitled in" Silicon Isotope Geochemistry"(Ding et al., 1996). As for the isotopes of other elements, silicon isotope compositions of a sample (Sa) are expressed as the 6 values related to a standard (St), i. e.:

629Si (%o)= [(29Si/28Si)sa/(29Si/28Si)st- 1] x 103 630Si (%o)= [(30Si/28Si)sa/(30Si/28Si)st- 1] x 103

[25.1] [25.2]

The standards used for silicon isotopes are described in section 25.2.4. There are two methods presently available for silicon isotope analyses" gas source isotope ratio mass-spectrometry (IRMS), and secondary ion microprobe mass-spectrometry (SIMS). The former is a routine method and suitable for all kinds of samples. It provides excellent precision, but needs relatively large samples. SIMS is able to determine Si isotopes on a small spatial scale but its precision is low, so that it is only applicable to the study of meteorites at present.

524

Chapter 25 - T. Ding

25.2 Gas source mass-spectrometric analysis of silicon isotopes Reynolds & Verhoogen (1953) firstly reported a mass-spectrometric analysis of silicon isotope ratios. Since then many modifications to the method have appeared. The common aspect of these methods is that samples of different types are transformed to SiF4 before their silicon isotope compositions are determined in a gas-source isotope ratio mass-spectrometer. There are two main reasons for using SiF4 as the gas for mass-spectrometric analysis. Firstly, SiF4 is relatively easy to prepare and is stable under laboratory conditions. Secondly, there is no need to make isotope corrections with SiF4, as fluorine has only one stable isotope. Three methods for SiF4 preparation are presently used: a) The BaSiF6 decomposition method (Reynolds & Verhoogen, 1953); b) The direct fluorination method using F2 + HF as fluorination reagent (Taylor & Epstein, 1962); c) The direct fluorination method using BrF5 as a fluorination reagent (Clayton & Mayeda,1963). A brief discussion on the advantages and disadvantages of different methods will be given in the following section. The method developed in Beijing will be described in detail, which is an improvement on the method of Clayton & Mayeda (1963).

25.2.1 A brief introduction of SiF4 preparation methods for mass-spectrometry In order to make reliable isotope ratio measurements, the gas samples produced for isotope determination must satisfy a basic requirement, i. e., the gas sample must reproducibly represent the isotopic composition of the original sample. There are two ways to satisfy this basic requirement. The first is to quantitatively transform the element in the original sample to a gas for mass analysis without any observable contamination. The second is to ensure that there is a fixed relationship between the isotopic composition of the gas sample and that of the original sample. The former is the most reliable of two and has been adopted in the preparation of SiF4 for isotope analysis. In this case, the key points are to guarantee 100% yield and to protect the sample from contamination by foreign silicon and other impurities that might interfere with the isotope analysis. 25.2.1.1 BaSiF6 method This method was firstly developed by Reynolds & Verhoogen (1953) and consists of two basic steps. The first step is to convert the silicon in the sample to BaSiF6, and the second step is to decompose the BaSiF6 to form SiF4 for mass-spectrometric analysis. The complete procedure is described below. The chemical reaction begins with the conversion of silicon in mineral specimen to SiO2, which is done through standard analytical procedures. Then 0.5 g of this silica is mixed with 2.5 g of sodium carbonate. This well-mixed combination is fused into a 20 ml platinum crucible, yielding a clear melt. After cooling, the solidified melt is extracted as a pellet by pressing the crucible gently around the base. This pellet dissolves in 15 ml of water to give a clear solution. To this is added 30 ml of concentrated perchloric acid and the solution is heated on a hot plate (swirling constantly to avoid bumping) until copious dense fumes of perchloric acid have boiled off for ten min-

AnalyticalMethodsfor SiliconIsotopeDeterminations

525

utes. When cooled, the solution solidifies to a gelatinous mass. It is then taken up in 200 ml of water and brought to boil. The precipitate of gelatinous silica is then filtered and transferred to a polyethylene beaker with 60 ml of water. Concentrated HF is added dropwise until the solution just clears reflecting total conversion of the silica gel to fluorosilicic acid, H2SiF6. To the clear solution 20 mL of a solution containing 0.2 g of BaC12.2H20 per mL is added. Under these conditions BaSiF6 precipitates while BaF2 does not. To ensure complete precipitation of the silicon, the beaker is allowed to stand overnight before filtering and drying the BaSiF6 at 110~ A series of analytical tests on the above procedure were done to determine the chemical yield. A typical test consisted of starting with an accurately weighed amount of silica, precipitating BaSiF6 with a calibrated barium solution, weighing the precipitate, and extracting and weighing the unreacted barium in the filtrate as BaSO4. Within the accuracy of these tests (+ 2%) the following observations were made" a. BaSiF6 is the correct formula of the barium chloride precipitate when the above procedure is adhered to strictly, i. e., occluded water and coprecipitated BaF2 are not present in the final precipitate. b. The yield of BaSiF6 always exceeds 95% of the theoretical yield calculated from the amount of silica added at the start of the experiment. This yield was sufficiently satisfactory for the method used in preparing samples for mass spectrometric analysis. The final step is to weigh out 150 mg of the dried precipitate into a short length of 10 mm Pyrex tubing, closed at one end. This tube is sealed into a 75 ml Pyrex sample bulb fitted with a stopcock. The bulb is evacuated with a diffusion pump for 12 hours or more, flaming gently from time to time in order to facilitate the removal of water vapour. Then with the stopcock closed, the BaSiF6 is decomposed by heating the bottom of the 10 mm tube vigorously with an oxygen-gas flame until there is no further change visible in the solid material. The temperature of the glass is brought to such a point in this heating that the glass softens and begins to melt around the solid BaF2 residue. Lastly, the lower half of the 10 mm tube is sealed with a torch from the rest of the sample bulb. This can remove the BaF2 residue and prevent any recombination with the silicon that might be isotopically selective. Control experiments show that the percentage decomposition of the BaSiF6 by this scheme is 90%. The SiF4 obtained from above steps in placed directly into the mass-spectrometer for isotope analysis. This method uses procedures relatively common in wet chemistry. The equipment used in this method is simple and no elaborate techniques are needed. Correct and meticulous operation of this method yields reliable analytical results. However, there are problems with this method that need to be addressed. For example, it needs relatively large amounts of sample (gram level). The procedure is also time consuming and requires great care to prevent silicon isotope fractionation if thermal decomposition of BaSiF6 is not complete. It is also possible that the produced SiF4 may change its Si isotope composition by reacting with the glass during the high temperature decomposition stage. In view of these disadvantages, the BaSiF6 method is rarely used in routine analysis of silicon isotopes except in some special cases. For example, De

526


Bievre and his colleagues have been using this method to calibrate silicon isotope ratios of their silicon isotope reference materials and determine molar mass (atomic weight) of silicon (De Bi~vre et al., 1994, De Bi6vre & Valkiers, 1994 and 1995). In their studies, Cs2SiF6 was chosen as the form of silicon compound to prepare synthetic mixtures used in calibration of absolute isotope ratios. Then the Cs2SiF6 was converted to BaSiF6, which was used in preparation of SiF4 for mass-spectrometry analyses.

25.2.1.2. Direct preparation of SiF4 The main difference between this method and that described above is that no intermediate step of BaSiF6 preparation is needed. Instead, the silicon-bearing rocks and minerals are reacted directly with a fluorinating reagent to produce SiF4. Baertschi & Silverman (1951) used firstly direct fluorination for sample preparation for stable isotope analysis. Allenby (1954) was the first to use the direct flourination method to prepare SiF4 for Si isotope analysis. He used HF as a fluorination reagent. However, due to the inconvenience of using HF and the unsatisfactory results, this reagent is no longer used. The most widely used methods today are those developed by Taylor & Epstein (1962) and Clayton & Mayeda (1963). Taylor & Epstein (1962) used a mixture of F2 and HF as the fluorination reagent in their experiments to extract 02 from silicate rocks for oxygen isotope analysis. Using this method the silicon in silicate was converted quantitatively to SiF4. However, the SiF4 products were initially pumped away as waste as the main aim of their studies was to prepare sample for oxygen isotope analysis. Epstein & Taylor (1970) made some modifications to this method to allow collection of SiF4. This method was then used in silicon isotope analyses of lunar samples and meteorites (Epstein & Taylor, 1970). Clayton & Mayeda (1963) developed an alternative method for preparation of 02 from silicates and oxides for oxygen isotope analyses by using BrF5 as the fluorination reagent. BrF5 has a low vapour pressure (about 33342 Pa at room temperature), so it is easily transferred and handled by evaporation and condensation in a metal vacuum system. Furthermore, BrF5 is an active fluorination compound that reacts with all kinds of silicates and oxides just as effectively as a mixture of F2 and HF. In addition, the wastes and residues of the reaction process are easily removed from the apparatus. For these reasons the BrF5 fluorination method is most popular in preparing samples for oxygen isotope analyses of silicates and oxides. Clayton et al. (1974) made some innovations to their extraction line by adding cryogenic parts for collecting and purifying SiF4. Subsequently, they have largely applied this technique to silicon isotope determination of meteorites and lunar rocks (see Molini-Velsko et al., 1986). Compared with the BaSiF6 method, the direct preparation method of SiF4 reduces the sample size to mg level, simplifies the operation procedure, increases the speed of sample preparation, and improves the precision of the analyses. Hence, this method has replaced the BaSiF6 method and has become the routine method for preparation

527

Analytical Methods for Silicon Isotope Determinations

of SiF4 for silicon isotope analysis. It should be noted that only pure silica and silicate rocks and minerals could be directly fluorinated in the manner described above. For samples containing impurities of C, S and B, pre-treatment is needed before fluorination.

25.2.2 The SiF4 preparation method used in the institute of mineral resources, Beijing 25.2.2.1 The sample preparation line The sample preparation line used in Beijing is shown in Figure 25.1. It is similar to the line of Clayton & Mayeda (1963), but improvements have been made in several key aspects. The whole line is made of metal. The left side of Figure 25.1 shows the section involved in SiF4 preparation and extraction. It consists of reaction vessels, the BrF5 storage vessel, a cold trap for waste, cold traps for separation, pressure gauge, Whity valves, connecting tubes, tubes for input of Ar and output of waste. The right side of Figure 25.1 shows the SiF4 purification and collection section, which includes a copper tube containing Zn particles (0.5 - 1.0 mm size), cold traps, resistance gauge, Whity valves, connecting tubes and a sample tube.

~

v9 LV V1

:~-1V2

Ar

l

V3

~-I V19

[ [

T1 V4

V5

T2

T3

BrF5 V6

V8

l Vll ~~ V12

HV

'7-',

~qV10

T l |

V20 CuT

V13~

l l V14 ~'e V15 "

v

S~F V21

V

'v22 0 T4 R1

R2

R3

R4

R5

Figure 25.1 - SiF4 sample preparation line.

R6

LV

V24

THV

ST

".m

T5

528


The reaction vessels are made of pure nickel. The CuT is made of 2 m long pure copper and is filled with pure Zn particles. The sample tube is made of Pyrex glass. Other tubes, cold traps, storage vessels and the pressure gauge are all made of stainless steel. All the valves are Whity two way metal ball valves. The low vacuum of the system is obtained by using mechanical pumps and the high vacuum is generated using a turbo-molecular pump. The dynamic vacuum of the system is about 6.6 x 10-4 Pa, and the static vacuum of the system can be kept in the range of 3~4 Pa in the 24 hours after pumping has stopped. The major differences between this line with that of Clayton & Mayeda (1963) are that a CuT tube is used for SiF4 purification and a device for waste disposal has been added to this line.

25.2.2. 2 Sample preparation method The sample preparation method adopted in Beijing is the BrF5 method similar to that of Clayton et al. (1963), i. e., the silicon-bearing sample is reacted with BrF5 to produce SiF4. SiO2 + 2BrF5 = SiF4 + 2BrF3 + 0 2 KA1Si308 + 8BrF5 = KF + A1F3 + 3SiF4 + 402 + 8BrF3

[25.3] [25.4]

However, in comparison with the method of Clayton et al. (1963), the method used in Beijing is superior in several aspects; namely in the purification of BrF5 and the SiF4 products and the treatment of waste.

Figure 25.2 - Mass-scanning diagram showing the impurities in the BrF5 agent, which can be frozen by liquid N2 but not by dry ice-acetone.


529

Purification of BrF5 reagent The BrF5 reagent (produced by the Wuhan Institute of Chemical Products, China) contains some impurities of CF4, SiF4 and SF6 (Figure 25.2). These impurities will interfere with the isotope determination of SiF4, so purification of the BrF5 is necessary. The method for BrF5 purification consists of cryogenic vaporization and condensation. At 1 atm (101325 Pa) the melting and sublimating points of SiF4 are -90~ and 95.1~ respectively. The melting and boiling points of SF6 are -80.4~ and -63.7~ respectively; and the melting and the boiling points of CF4 are -184~ and -128~ respectively. In contrast, the melting and boiling points of BrF5 are -61.3~ and 40.5~ respectively. Under vacuum, these temperatures will reduce, but their relative order does not change. Therefore BrF5 can be purified through cryogenic vaporization and condensation. Dry ice-acetone (-80~ and ethanol-liquid nitrogen mixtures (its temperature is adjustable and can be as low as -100~ were tested as cryogenic liquids. At the temperatures of these cryogenic liquids, BrF5 remains in the solid phase, but the impurities, such as SiF4, SF6 and CF4, are present in the vapor phase and can be pumped to waste. In general the ethanol-liquid nitrogen mixture of-70~ is used as the cryogenic liquid and the BrF5 reagent is purified several times to ensure good purification. This is a key step to improve the precision of the Si isotope analyses.

Pre-treatment of samples The SiF4 preparation method can be used for any sample containing silicon for isotopic analysis. However, different pre-treatment procedures are needed for the different types of samples. A. For silica and high purity silicates of containing little or no C, S and B, no pretreatment is needed. The sample can be fluorinated directly after it is grounded. B. For samples containing significant contents of S, C and B (>1%), chemical pretreatment of the sample is necessai:y. a. When the impurities in the sample are carbonates or acid soluble sulphides, the HC1 dissolution method is used. b. When the sample contains only graphite or organic carbon, high temperature oxidation and evaporation are used to remove the contaminants. c. When the sample contains boron compounds, sulfates or sulphides insoluble in HC1, wet chemical pre-treatment is required. In this instance there are several methods for the preparation of SiO2. The procedure adopted at Beijing is as follows; the sample of 0.1--0.5 g (depending on the Si content) is mixed with NaOH and Na202 and is placed in graphite crucible. This mixture is melted in a furnace and the cooled melting products are dissolved with a 1"1 HC1 solution. This solution is dried by low temperature evaporation. Concentrated HC1 of 10 mL is then added to the dried product. The solution is warmed for 10 minutes in a water bath and animal glue is added into the solution to precipitate SiO2. The solution is filtered and the precipitates rinsed several times with 1% HC1 followed by distilled water. The filter paper containing the precipitates is placed in a platinum crucible, and ashed at 700~ then calcined at 1000~ The SiO2 obtained

530


is then fluorinated.

SiF4 preparation and purification A. Sample loading and reaction a. Sample loading: The Ni reaction vessels are filled with Ar after the treatment with BrF5 and left over residues of last reaction. The reaction vessels are opened, the solid products of the previous fluorination are turned over, and the new dried sample entered. Then the reaction vessels are connected back to the vacuum line. The sample loading is normally 5 mg of silicon, but it can be reduced as small as I mg of silicon. b.Adding BrFs" Cool the waste cold trap T1 with liquid nitrogen and evacuate the manifold and reaction vessels by mechanical pump. Waiting for a few minutes, cool the separation traps T2 and T3 with liquid nitrogen, pump the line to high vacuum by a turbo-molecular pump. When the ion gauge near molecular pump show a vacuum of ~ 2 x 10-3 Pa (indicating the reaction vessels are properly connected), the BrF5 is added. Let BrF5 diffuse into the manifold of the line and reach its saturation vapor pressure (33342 Pa at room temperature). Freeze the BrF5 in the manifold into R1 (at liquid nitrogen temperature). Fill the manifolds with BrF5 again. Freeze BrF5 into R2. Repeat above steps to transfer BrF5 into R3, R4, R5 and R6 one by one. Then pump the manifold and reaction vessels with mechanical pump for ten minutes while these vessels are frozen with liquid N2. Then close V3 and open V17, pump the manifold and reaction vessels with turbo-molecular pump. When the ion gauge by the molecular pump indicates a vacuum higher than 2 x 10-3 Pa, close Vii-V16, remove the liquid nitrogen cups from the reaction vessels, and warm in a water bath. In general, the amount of BrF5 added is 4~5 times more than that needed for complete reaction. c.Reaction: A cooling water device is placed around the upper part of R1-R6, just below the valves. These vessels are then heated with electric furnaces. The temperatures of the furnaces are adjusted using adjustable transformers. The reaction temperature varies according to the type of sample. For samples which react easil3r such as" quartz, silica sinter, feldspars, micas and granite rocks, a temperature of 550~600~ is used. For samples that do not react so easily, such as olivine, pyroxene, actinolite and topaz, a temperature of 650~700~ is needed. The reaction time is normally set to be more than 14h (overnight). The above procedure is almost the same as that for oxygen isotope analyses, and hence, sample preparation for Si and O isotope analyses are often done simultaneously. BO Extraction and purification of SiF4

After the reaction is completed, the electric furnaces are switched off, removed from the reaction vessels, and cooled with cold water. Freeze the reaction vessels (R1-R6) and cold traps (T1 -T3) with liquid nitrogen cups. Then open V11-V16, pump the line with the turbo-molecular pump to release 02. In the case of simultaneous preparation of 02 and SiF4 for isotope analysis, the extracting procedure is different. Instead of pumping out the 02, the 02 in each


531

reaction vessel is extracted separately and converted to CO2 before extraction of the SiF4. In the processes of SiF4 extraction, three stages of distillation and condensation at dry ice-acetone and liquid nitrogen temperatures are used for separating SiF4 from 02, N2, BrF5 and BrF3. Through these processes, the SiF4 obtained will contain no impurities of 02 and N2, but may contain trace amounts of BrF5 and other active fluorine compounds still. These impurities will react with the glass and grease of the sample tube to produce contaminants, such as SiF4 and CF4, which will interfere with the isotopic determination. For this reason, a new step for SiF4 purification has been added by using a Cu tube containing pure Zn particles (CUT). Heated Zn particles (50~176 react with BrF5 and other active fluorine compounds (except SiF4) to form ZnF2 and ZnBr2. This step is simple and very efficient. Very pure SiF4 gas can be prepared, as shown in the mass-scanning plot in Figure 25.3. Freeze SiF4 into the sample tube at liquid nitrogen temperature. Then the sample tubes are brought to the mass-spectrometer for isotopic analysis. A set of 6 samples can be prepared each day using this method. CQ Waste treatment After completing preparation and purification of SiF4, the waste left in the reaction vessels and cold traps are treated as follows. Condense the BrF5 and other wastes left in reaction vessels into T1. Fifteen minutes later fill Ar gas into T1 through V4, V5 and V3. Remove the liquid N2 cup from T1 and warm it with a water bath to defrost the BrF5 and other wastes. Open V3 and V2, blow the wastes with Ar into a bucket containing Ca(OH)2 solution. The BrF5 and other active fluorine compounds in the waste gas react with Ca(OH)2 to produce CaF2, CaBr2 and 02. After the reaction in the bucket has stopped, keep Ar blowing for 5 more minutes. Then close the Ar tank, V4 and V2 in order. This waste treatment is simple, secure, and pollution free.

Figure 25.3 - Mass-scanning diagram of purified SiF4gas.

532


25.2.3 Laser probe extraction method for SiF4 preparation The laser probe extraction method for SiF4 preparation was reported by De La Rocha et al. (1996). Their sample preparation line is shown in Figure 25.4:. It consists of two portions" a metal section for fluorination and a glass section for yield determination and sample collection. Purified fluorine for sample reaction is generated inside the vacuum line by heating potassium hexafluonickelate, which decomposes at ~350~ to produce F2. A tank of F2 provides the commercial fluorine for charging the pure fluorine generator. A tank of N2 plumbed into the metal section is used to pressurize the sample chamber for sample loading and melting. Purified silica samples containing 15-100 mmol of Si, are loaded into 0.5 cm deep wells drilled into a cylindrical nickel plate. After sample loading, the reaction chamber and vacuum line are evacuated for several hours. Pieces of pure quartz can be fluorinated directly, but the finely powdered silica that has been purified through precipitation may sputter during fluorination. Sputtering is eliminated by melting the powder under I atm of N2 into lumps of glass that then fluorinate in a controlled fashion. A CO2 laser (22 W) coaxial with a He-Ne sighting laser (3 mW) is mounted on a motorized x-y translation stage and set to fire through a BaF2 window on the top of reaction chamber. The CO2 laser is set for a beam width of 0.8 ms and a pulse period of 0.9 ms. The intensity of the laser beam at this setting is raised from zero to near maximum levels slowly during lasing. Melting the silica with a continuous beam is avoided as partial vaporization of the sample can occur. After the melting of samples the reaction chamber is evacuated and pumped for several hours. Samples are fluorinated under 0.1 atm of purified F2 that has been passed through a liquid nitrogen trap (cold trap 1 on Figure 25.4). Silica samples react with F2 upon being heated by the laser set to fire a continuous beam: SiO2 + 2F2 --* SiF4 + 02

[25.5]

Figure 25.4 - Schematic of SiF4 preparation line with laser microprobe device (After De La Rocha et al., 1996).


533

Lasing of 1-3 mg samples takes approximately 10-20 min, during which the intensity of the continuous laser beam is varied between zero and maximum intensity to maintain controlled, continuous fluorination. When lasing has been completed, the resulting SiF4 is collected in a coil cooled to -195~ with liquid nitrogen (cold trap 2 on Figure 25.4). 02 and F2 are then pumped away through the coil. Any water present either in the silica or in the reaction chamber during fluorination will form HE which may in turn form SiF4 by etching the walls of the glass section of the line. Traps cooled to liquid nitrogen temperature will collect both SiF4 and HE and so will not serve to separate them. For effecting a separation, the SiF4 (and any HF) is transferred from cold trap 2 into the variable-temperature trap, which has cooled to at least -183~ with liquid nitrogen. Remaining noncondensibles are pumped away through the trap. The variable-temperature trap is then heated t o 140~ distilling SiF4 to the glass side of the line, where it is collected at liquid nitrogen temperature in the multitrap. Tests indicate that SiF4 quantitatively distils out of the trap at this temperature but HF remains behind. The liquid nitrogen on the multitrap, where the sample is frozen, is replaced by a dry-ice-2-propanol slush. Distil SiF4 from the multitrap to the capacitance manometer, where the micromoles of sample gas are determined. Samples are distilled into borosilicate tubes, sealed, and are then ready for mass-spectrometer analyses.

25.2.4 Standards for silicon isotopes So far 6 samples have been used or proposed as silicon isotope standards: i. e. NBS-28, Rose Quartz, IRMM 017, IRMM 018, GBW 04421 and GBW04422. NBS-28 is a sample of quartz sand distributed by National Bureau of Standard of the United States (now NIST) as an oxygen isotope reference materials for silicates. It has been used in laboratories of the Chicago University, the Institute of Mineral Resources, CAGS (Beijing) and other institutions as a silicon isotope reference material. Rose Quartz is a quartz sample that has been used at the California Institute of Technology as a reference material for oxygen and silicon isotopes. IRMM 017 (a silicon metal) and IRMM 018 (a sample of silica) are two samples preparedat the Institute for Reference Materials and Measurements (IRMM) (De Bibvre et. al. 1994a). GBW 04421 and GBW 04422 are national reference materials for silicon isotopes in China. The 630SINBS-28 values of these reference materials are listed in Table 25.1. De Bi6vre et al. (1994a) determined the absolute isotopic ratios of silicon in IRMM 017 and IRMM 018. IRMM 017 has 29Si/28Si value of 0.0507715(66) and 30Si/28Si ratio of 0.0334889(78). IRMM 018 has 29Si/28Si value of 0.0508442(48) and 30Si/28Si ratio of 0.0335851(66).

25.2.5 Mass-spectrometry analysis The SiF4 obtained from the above process is analyzed in a gas mass-spectrometer for its isotopic composition. At Beijing a MAT-251 EM mass-spectrometer with multiple collectors is used.

534


Table 25.1 - The ~)30SiNBS-28values of some reference materials for silicon isotopes Sample No.

Description of sample

NBS-28 Rose Quartz IRMM-17 IRMM-18 GBW-04421 GBW-04422

Quartz Quartz Silicon metal Quartz Quartz sand SiO2 chemical agewnt

630SINBS-28 (%o) 0 -0.28 -1.3 0.0 -0.02 -2.68

Reference

Molini-Velsko et al. (1986) This study This study Wan et al. (1997) Wan et al. (1997)

As mentioned above, silicon has 3 stable isotopes" 28Si, 29Si and 30Si. The isotopic compositions are commonly expressed as ~)30Si and ~)29Si values. However, in routine investigations of terrestrial samples, only 630Si value is measured. 629Si determinations are done in special cases, w h e n meteorite Si isotope anomalies or n o n - m a s s d e p e n d e n t isotope fractionation are studied. The most abundant SiF3 + ions are normally used in the isotopic determination. Two collectors are simultaneously used to collect 28SIF3+ and 30SiF3+ at mass numbers 85 and 87. A 10 kV accelerating voltage is used with a magnetic field intensity of 0.5587 T. The ion beam intensity is measured on the 6 V scale. The working standard used is the SiF4 prepared from NBS-28. The results are expressed as 630Si values related to NBS28. As fluorine has only one isotope, there is no need to make an isotopic correction for the 630Si determination of SiF3 +. Six sets of data are collected for each analysis. The precision of the mass-spectrometry analysis is + 0.05 to + 0.10 %o (lo). During mass-spectrometry, the 630Si values will be altered if the SiF4 is not pure or if there are trace amounts of air leaking into the sample tube. If the carbon in the sample is not removed out before fluorination, COF3 + ions will appear in the mass-spectrum and cause unusual low ~)30Si values. In Beijing the present arrangement of collectors in the mass-spectrometer does not allow simultaneous collect ions of SiF3 + of masses 85, 86 and 87, so that 630Si and ~)29Si can not be determined simultaneously. When it is needed, ~)29Si is determined in the form SiF4 +, at masses of 104 and 105. The intensity of SiF4 + is m u c h lower than that of SiF3 +, but reliable data can still be obtained.

25.2.6 Analytical precision Several samples, such as the NBS-28, GBW 04421 and a sample of SiO2 prepared from a diatom (88-43) have been fluorinated and measured repeatedly at Beijing. In addition, a meteorite sample and a tektite sample have been analyzed in duplicate; the results are listed in Table 25.2. From these results the standard deviations of the analyses is estimated to be _+0.03%o (lo) to _+0.10%o (lo).

535

Analytical Methods for Silicon Isotope Determinations Table 25.2 -

~)30SiNBS-28

values and analytic uncertainties for several samples. Average and uncertainty of ~)30SiNBS-28(%o)

Sample No.

Sample type

~)30SiNBS-28(%o)

NBS-28

Quartz sand

GWB-04421

Quartzite

-0.04 + 0.07

88-43

SiO2 prepared from a diatom Thailand tektite

0.03, 0.03, 0.02, 0.05, -0.04,-0.01, 0.00 0.0, 0.0,-0.01, 0.04, -0.13, -0.11, -0.09 -1.18, -0.94, -1.05, -0.06 -0.23, -0.29

-0.26 + 0.03

88-44

0.0 + 0.06

-1.06 + 0.1

Since 1988, all analyses made in Beijing have been monitored by the analyses of NBS-28, and the precisions have always been better than _+0.10%o (lo).

25.3 Ion Microprobe mass-spectrometer analyses The secondary ion microprobe mass-spectrometer (SIMS) developed in the 1970"s differs from traditional instruments, such as gas isotope mass-spectrometers and thermal ionization mass-spectrometers. It measures secondary ions instead of the primary ions. In the ion microprobe, secondary ions are produced by ion bombardment on the surface of the sample to be analyzed. Focusing of the bombarding ions into a fine beam allows the in situ analysis of individual mineral phases. The main advantages of the technique are high sensitivity~ small sample size and measurement of elements that are difficult to analyze by other techniques. However, there are also problems with this technique and these include interference of molecular ions with the atomic ions of interest, large variation of the ionization efficiencies of different elements and a matrix-effect. Several kinds of ion microprobes, such as the Cameca IMS 3f and the SHRIMP (Sensitive High Mass Resolution Ion Microprobe) have been applied to isotopic measurements on extraterrestrial material and on terrestrial rocks. The application includes U-Pb dating of individual zircons and the study of the distribution of Pb, S, H, C, O, Mg, Si, Ca and Ti isotopes in a number of minerals. A schematic diagram of a SIMS is shown in Figure 25.5. The general aspects of isotopic measurement with SIMS are described by McKeegan et al. (1985) and Clement & Compston (1989) and Chapter 30 of this book. Silicon isotopes have been measured as positive (Huneke et al., 1983; Clayton et al., 1991; McKeegan et al., 1985) and negative ions (Zinner et al., 1987). The silicon signal per primary beam current for a given sample is higher than for negative secondary ions. However, since the sputter rate (number of sputtered atoms per incident ion) of Cs + is roughly 40 times that of O-, the ionization efficiency (number of secondary ions

536


Figure 25.5 - Schematic diagram of an ion microprobe mass-spectrometer (After Eldridge et al., 1989). P - pump.

per sputtered atom of a given species) is much higher for positive ions. The instrumental mass fractionation, however, is much larger for positive (~32%0/amu) than for negative ions (~8%o/amu; amu is atomic mass unit). Stone et al. (1991) measured the silicon isotope compositions of SiC grains from carboniferous chondrites and enstatite chondrites with a Panurge IMS-3f ion microprobe mass spectrometer. For the majority of the analyses, silicon was analyzed as positive secondary ions, produced by a 0.2~0.5 nA 160- primary beam rastered over a 20 mm square region surrounding each grain. The low sputter rate, galena > chalcopyrite. Yields were much lower for pyrite and other sulfides. Low yields can result from the mineral being highly reflective of the laser beam, being transparent to the laser beam, and/or having high thermal conductivity so that heat quickly migrates away from the target area. SO2 yields near the edge of the specimens are higher than in more central regions attesting to more energy loss from the target area by thermal diffusion in the latter case. For the same mineral and 02 pressure, SO2 yields decreased with both laser power density and the total energy received by consecutive firings. It was found that above 500 torr 02 pressure, yields of SO2 and ~34S values remained constant with change of pressure. It was also found that with N2-O2 mixtures, total pressures above 500 torr realized highly reproducible data provided that the 02 pressure exceeded a minimum value (50 to 100 torr dependent on mineral). The data were consistent with the requirement that SO2 was produced by gas phase reactions. Hence, the mineral specimen has to reach the boiling point, which is dependent upon the total pressure of the surrounding gas. As observed by others (e.g. Crowe et al., 1990; Kelley & Fallick, 1990) there were mineral dependent shifts of a few per mil towards lower ~34S values than found by traditional analyses. The shifts decreased with increasing 02 pressure but no consistent correlations were found between SO2 yields and ~34S values among different minerals. Li (1996) designed an intriguing experiment to better understand the ~534Sshifts. A piece of thin glass (0.5 mm thickness) was placed 4 mm above a galena surface in the reaction chamber. Enough solid PbS was deposited on the underneath side of the glass for isotopic analysis after 12 laser firings. Whereas the PbS specimen had a conventionally determined ~34S value near 0 %o, laser produced SO2 and the deposit on the glass had ~34S values o f - 1 and +11%0 respectively. From these results, it was concluded that the shifts resulted from kinetic isotope effects favoring lighter isotopes in the vapor phase (probably sulfide molecules, ions, and $2) and preferential removal of heavier isotopes from the vapor phase during sublimation. The ~534Sshifts associated with laser SO2 production appear to be absent or insignificant when SF6 is formed. Another suitable technique for studying ~34S variations over sub-mm distances is secondary ion mass spectrometry (SIMS). The instrumentation and data reduction are somewhat complex and few facilities have attempted S isotope abundance measurements (e.g. Deloule et al., 1986; Eldridge et al., 1988; Eldridge et al., 1987; McKibben and Eldridge, 1995; Riciputi et al., 1996; Ireland, Chapter 30). Two manufacturers of

Procedures for Sulfur Isotope Abundance Studies

549

suitable instruments are Cameca and Australia National University (SHRIMP, Sensitive High Resolution Ion Microprobe). A high energy ion beam is directed at a mineral surface and secondary ions are emitted from a sputtered pit of typically 20 gm diameter and 5 gm depth. For conducting minerals such as pyrite, the specimen must be Aucoated, polished, and a primary beam of positive ions (e.g. 14.5 keV, 113Cs+) is used to generate negative secondary ions, S- (Riciputi et al., 1996). Au-coating is used to reduce static charge build-up on the mineral surface. Because of their insulating nature, sulfates are analyzed using a negative ion beam (e.g. 14.5 keV, 160-) and positive secondary ions (S§ are generated and measured (Riciputi et al., 1996). Since the secondary ions have a wide range of velocities (both direction and magnitude), a double focussing mass spectrometer must be used and only a portion of the ion current reaches the collector. Since the secondary ion currents are low, an electron multiplier is used in counting mode (usually 30%). Most commercial HiVol samplers use 20 x 25 cm filters, which are supported mechanically by resting on a down-stream stainless steel mesh. Air flow rates of 2 m3 min-1 are obtainable without undue sacrifice of collection efficiency of SO2 (Holt, 1975). At these rates, sufficient sample for DI-IRMS (> 1 mg S) can be obtained on the time scale of hours. This makes sampling of specific wind directions or industrial emission events feasible.


559

Dequasie & Grey (1970) were able to obtain milligram quantities of atmospheric sulfur dioxide for DI-IRMS by adsorption of SO2 on a molecular sieve. Since removal of water vapor and separating CO2 and SO2 were not straightforward and quite laborintensive, this technique has never received widespread use. Atmospheric SO3 is rapidly converted to sulfate and hence concentrations are typically too low for isotopic analyses. However, its concentration in flue gas can be quite high. Forest & Newman (1973) sampled flue gas by passage through a quartz wool filter and condensing of SO3 in a coil at a temperature just above the dew point of water. SO2 was not condensed and subsequently oxidized to sulfate in a solution of alkali salts and hydrogen peroxide. If isotopic equilibrium between SO2 and SO3 has been achieved (equation [26.6]), the g34S values of the two gases provide information about the temperature conditions during combustion in oil or coal fired power plants. Where H2S emission from springs or wetlands are significant, sampling of H2S prior to its oxidation in the atmosphere might be feasible. Van Everdingen et al. (1982) used sheets of filter paper soaked in saturated cadmium acetate solution to collect airborne H2S near sulfurous springs. The same authors also used silver acetate treated H2S test paper for the same purpose, reporting 634S values as low a s - 3 0 %o. Other researchers inferred indirectly from low ~34S values in atmospheric SO2 or rainwater sulfate that biogenic S gases released from soils, marshes, and wetlands must have contributed to atmospheric S (Grey & Jensen, 1972; Hitchcock & Black, 1984; Nriagu et al., 1987). Dimethylsulfide (DMS) is a major source of atmospheric S. It is mainly produced via assimilatory sulfate reduction in the oceans and can be transported into the atmosphere through the sea-air interface. DMS can be oxidized to either sulfate or methane sulfonate (MSA). Gold surfaces have been used for selective adsorption of nmol quantities of DMS (e.g. Ammons, 1980). Calhoun (1990) up-scaled such a gold collection system to obtain micromole amounts of DMS for isotopic analysis using TIMS. The collection tube consisted of an ultrapure quartz tube filled with 200 meters of loosely packed gold wire. This setup had a collection capacity of approximately 80 nmoles of DMS. The collected DMS was subsequently converted to As2S3 via thermal desorption of the DMS in an H2 atmosphere to generate H2S, which was trapped as As2S3 in an As3+/NH3 solution. A minimum of 25 collection and purge cycles was used to obtain micromole quantities of DMS for isotopic analysis. Using this system, Calhoun (1990) was able to collect a single seawater-derived DMS sample (10 ~g S) from the remote southeastern Pacific Ocean. TIMS analysis revealed a ~34S value of +17 %0, which is consistent with the sulfur isotope ratios of submicrometer NSS aerosols and MSA reported below. Hence it was concluded that DMS can be the source of NSS in the remote South Pacific (Calhoun, 1990). This is evidence that NSS from fossil fuel combustion is often isotopically distinct from NSS formed by oxidation of biogenic marine S sources such as DMS. This makes sulfur isotope ratio measurements on atmospheric S compounds a useful tool for differentiating between anthropogenic and biogenic sources (e.g. Nriagu et al., 1991).

560

Chapter 26 - B. Mayer & H.R. Krouse

26.3.1.3 High volume atmospheric sampling A typical high volume atmospheric sampler is a motor driven impeller, which draws air through a system of filters at a rate of 1.41 to 1.83 m3 min-1. Filter sheets (8" x 10") are mounted with regular gaskets (rubber or similar material) leaving an effective filtering area of 7" x 9". Motors are usually 115 V AC with 24 V DC versions for mounting on aircrafts. The flow can be maintained reasonably constant as material collects on the filters by using a sensor and a feedback circuit to control the motor's speed. Sometimes the volume flow is kept constant by using manometer sensors. Alternately, mass flow can also be regulated using a combination of velocity and temperature sensors. If the filter paper becomes excessively loaded with particulate matter, the motor may overheat. This is more problematic if the pore size of the paper is too small. For most commercial units, if the flow is < 20 ft3 min-1, there is insufficient air cooling of the motor. Brush life is another limitation, being typically ~ 500 hours for motors run at 115 V. It can be extended to ~ 1500 hours if the motor is run at 90 V. Brushless motors are available from some suppliers. One problem in sampling particulates is that large amounts of material may be collected, which are a nuisance since they load the filter papers but are not the object of the study. Examples are wind blown dust from coal, ore piles or road construction. To some extent, the peaked roof design of high volume sampling chambers reduces the intake of large particles. A better approach is the use of a cyclone pre-separator, which collects particles typically > 5.5 gm. Cascade impactors (discussed below) fail with larger particles because they may bounce or roll through. A cyclone pre-separator has an inlet vent with a vane to align it with the wind direction. Air enters the cyclone body tangentially and sets up a vortex flow pattern. Larger particles migrate to the walls or to the top of the chamber. The Andersen cascade impactor has a different design. There are geometrical patterns of holes in circular metal plates. These differ for adjacent plates but are the same for every other plate.

26.3.1.4 Aerosol sulfate and particulate sulfur High volume samplers have been used to collect both particles and aerosol sulfate. Forrest & Newman (1973) tested the suitability of glass fiber filters and recommended Whatman 81 filters because of their high collection efficiency and little tendency to convert SO2. The glass fiber filter with the S-containing particulates can be subject to Thode reduction (Thode et al., 1961) to convert all S to H2S (see Volume II, Part 3, Chapter 8-3.4), which is trapped as CdS and subsequently converted to Ag2S (Forrest & Newman, 1973; Van Everdingen et al., 1982). Other researchers have collected aerosol and particulate S by passing several thousand m3 air through Whatman 41 cellulose filters (McArdle & Liss, 1995; Nriagu et al., 1991). Sulfate and methanosulfate can be extracted from the filter with deionized water in an ultrasonic bath. The sample should be irradiated with ultraviolet light. Subsequently, dissolved sulfate is converted to BaSO4 by adding BaC12 solution. Using this procedure, McArdle & Liss (1995) were able to generate 1 mg S for DI-IRMS. Other researcher have preferred teflon filters for the collection of aerosol and particulate S (Hitchcock & Black, 1984;


561

Quinn & Bates, 1989). Calhoun (1990) collected particle phase non sea-salt sulfate (NSS) and methanesulfonate (MSA) from marine air using a 1.0 ~m pore size teflon filter. Aerosol sulfate collected on these teflon filters was converted to As2S3 using the method of Paulsen & Kelly (1984), which is described in Volume II, Part 3, Chapter 8-5, but MSA did not convert to As2S3 using this technique (Calhoun et al., 1991). Subsequently, 1.5 ~g S was loaded as As2S3 on a filament and sulfur isotope ratios were determined by TIMS. This technique was chosen because of its low sample volume requirements and the lack of memory effects and isobaric interferences, yet achieving reasonable precision of ~ 2 %o for ~34S measurements including sample collection, handling, and conversion to As2S3 (Calhoun, 1990). To determine the sulfur isotope composition of both aerosol sulfate and methanosulfate (MSA), a subsection of the teflon filter was wetted with I mL spectrophotometric grade methanol in 10 mL distilled water and centrifuged for 30 minutes. The extract was transferred into a Carius tube (Paulsen & Kelly, 1984) and frozen together with 10 mL of 16 M HNO3 and 4 mL of 11 M HC1. The tube was flame-sealed and heated to 240~ for 16 hours to convert MSA to sulfate. Total sulfate (MSA-derived and aerosol sulfate) was subsequently converted to As2S3 and sulfur isotope ratios were determined by TIMS as described above. The isotopic composition of MSA was determined by mass and isotope balances, and was found to be similar to that of NSS (~34S = +15.6 + 3.1%o) from submicrometer aerosol particles in the remote southeastern Pacific Ocean (Calhoun et al., 1991). Aerodynamic sizing of atmospheric particulates can be carried out with a five or six stage cascade impactor fitted to a high volume sampler. One design by Sierra Instruments Inc. consist of stacked plates, which alternately have 9 or 10 parallel slots. Consequently, slots on a given plate occur midway between those of the plates above and below. As the air flow bends going through successive plates, particulates of decreasing size are deposited on slotted glass-fiber filter paper (dimensions 5.625" x 5.375"). The width of the slots decreases with successive plates (stages) and the finest particulates, which pass through the cascade, are trapped on a 8" x 10" "back-up" glass-fiber filter paper mounted above the fan of the high volume sampler. Aerodynamic sizing not only depends on the size of the particle, but also on its shape and density. On a given plate, the physical size of a particle has an inverse dependence on the density. The particle size cut-offs increase with flow; they are chosen on the basis of human inhalation. Those above 7 gm are deposited in the throat. Smaller particles are deposited in the bronchia and the smallest reach the lungs. Sizing of atmospheric particulates for obtaining 634S values seems to have been first reported for a study near two sour gas processing operations (Krouse, 1991). Larger particles had higher 634S values near +20 %o consistent with industrial emissions in the study area. The smallest fraction had ~34S values closer to 0 %0 implying a different source. The collection of size-segregated aerosols for S isotope analyses has also been reported in some recent studies (e.g. Patris et al., 2000b; Turekian et al., 2001). Multiple-stage high volume samplers with glass fiber substrates and backup filters can separate size fractions from more than 5 ~m to less than 0.2 ~m geometric

562


mean radius. With flow rates of approximately I m3 min-1, sufficient S can be sampled for each size fraction in less than 48 hours provided that CF-IRMS is used for the sulfur isotope measurements (Turekian et al., 2001). Using these techniques, it was found that submicron radius NSS over the North Atlantic Ocean is mainly derived from fossil fuel combustion (Patris et al., 2000b). In the above studies, the particulates were not distinguished on the basis of solubility. In a study on Bermuda, Norman & Krouse (1992) separated bulk particulate matter on the basis of solubility in water. They found insoluble particulates to have a v e r a g e ~)34S values near +3 %o, which interestingly is the range reported by Turekian et al. (2001) for the smaller size fraction.

26.3.1.5 Combined sampling of 802 and aerosol~particles High volume samplers can be fitted with several filter papers allowing for the simultaneous collection of SO2 and aerosol/particle S for subsequent isotope analysis. Forrest & Newman (1973) employed a glass fiber prefilter to collect aerosol and particle S followed by two KOH-TEA coated cellulose filters for the collection of SO2. They reported a reproducibility of 634S measurements of _+ 0.2 %o. Newman et al. (1975) used such a setup to sample the plume of an oil-fired power plant with a single engine aircraft. Quinn & Bates (1989) used tandem filters, which separated particle phase non sea-salt sulfate (NSS) and methanesulfonate (MSA) collected using a 1.0 gm pore size teflon filter, from gas phase SO2 collected on K2CO3 coated filters as described above. The reproducibility of sample collection, handling, and conversion to As2S3 for TIMS analysis was reported as better than +_2 %0 (Calhoun, 1990). 26.3.2 Sulfur in the hydrosphere In the hydrosphere, sulfur occurs predominantly in dissolved form either as SO42-, HS- or $2-, with dissolved organic sulfur (DOS) typically being of minor importance. Additionally, gaseous S compounds such as H2S may be present in water samples obtained from reducing environments. This section describes methods to quantitatively recover dissolved and gaseous S compounds from water samples for subsequent isotope ratio mass spectrometry (see Volume IL Part 3, Chapter 15 for additional details). A simplified summary chart for this section is shown in Figure 26.3. The selection of the appropriate procedures to collect water samples for sulfur isotope analyses is critically dependent on knowledge about the existing sulfur species in the sample and their concentrations. It is strongly recommended to fix dissolved or gaseous reduced inorganic sulfur species in the field, since these compounds may otherwise rapidly outgas or oxidize to sulfate. Dissolved sulfate in oxidized water samples is typically fairly stable. Hence preservation of such water samples may not be necessary if they are further processed immediately upon return to the laboratory. If storage of sulfate-containing water samples is necessary, the addition of a bacertiocide (e.g. HgC12) and storage in a refrigerator at 4~ in darkness is recommended. 26.3.2.1 Dissolved sulfate Under oxidizing conditions, SO42- is usually the dominant S species in water samples and reduced inorganic S compounds are typically not present. It is recommended


563

Figure 26.3 - Flow chart summarizing analytical procedures for chemical pretreatment of various S compounds from water samples and their conversion to BaSO4 or Ag2S for subsequent sulfur isotope analysis.

to filter such water samples with 0.45 ~m membranes either in the field or immediately after return to the laboratory. Subsequently, the samples should be acidified to a pH value between 3 and 4 with dilute sulfate-free HC1 to convert dissolved carbonate species into CO2 and hence avoid co-precipitation of BaCO3. Acidification to pH values of less than 2 should be avoided since HSO4-will be the dominant S species and precipitation of BaSO4 will be slow. Note that excess chloride may pose problems if ion exchange techniques are subsequently used. If sulfate concentrations are above 20 mg L-l, BaSO4 can be precipitated directly. After acidification of the water sample, 0.25 M BaC12 solution (typically 10 mL) is added in excess and a white precipitate of BaSO4 will form and settle overnight. Some researchers prefer to heat the water sample prior to adding the BaC12 solution (e.g. Carmody et al., 1998) since this facilitates the outgassing of dissolved CO2 and enhances the rapid formation of a coarse BaSO4 precipitate. However, if oxygen isotope analyses on the sulfate are to be performed it is important not to heat samples with low pH values (< 2) excessively, since oxygen isotope exchange between water and sulfate is promoted under high temperature low pH conditions (Chiba & Sakai, 1985). A 1 L water sample with a sulfate concentration above 20 mg L-1 will yield

564


more than 40 mg BaSO4, which is sufficient for both continuous flow and dual inlet isotope ratio mass spectrometry. Samples with very low concentrations of sulfate (e.g. atmospheric precipitation samples) must be pre-concentrated prior to precipitation of BaSO4 because of its solubility of 2.3 mg L-1. This can be achieved by reducing the volume of the water sample via slow evaporation in a beaker on a hot plate or in a rotary evaporator. Depending on the chemistry of the sample, non sulfate-containing precipitates may form towards the end of the heating process and these must be removed prior to the addition of BaC12 to precipitate BaSO4. Patris et al. (2000a) successfully evaporated up to 2.5 L meltwater from Antarctic firn cores under partial vacuum in a rotary evaporator to a final volume of only 2 mL. Rather than precipitating BaSO4, these authors introduced the remaining sulfate-containing liquid into ChromosorbWM-filled tin cups, where sulfate precipitated upon further heating. Subsequent CF-IRMS analysis yielded precise sulfur isotope ratios for micromolar levels of sulfate from polar ice samples. Another technique of collecting and concentrating sulfate from dilute solutions is the use of anion exchange resins (e.g. Mizutani & Rafter, 1969). Suitable commercially available products include Dowex TM Amberlite IRA-400 or BioRad TM AG 1-X8 and AG 2-X8, which are all in chloride form. Large quantities of water can be passed through such resins in flow-through mode either via gravimetric dripping or by using a peristaltic pump. Since divalent ions have a higher retention affinity than CI-, sulfate will be quantitatively retained at low flow rates (less than 10 mL min-1 recommended). In selecting the size of the ion exchange column it is important to consider the maximum exchange capacity (typically 1.2 meq per mL resin) and the fact that other negatively charged ions such as bicarbonate, nitrate, and some DOC will also be retained hence occupying adsorption sites. High concentrations of chloride may also interfere with the retention process (Carmody et al., 1998) and thus excessive use of HgC12 or HC1 for preserving or acidifying the sample to remove bicarbonate should be avoided. Ion exchange resin columns of any desired size can be made by filling the resin in glass tubes with glass wool on both ends. Preconditioning includes eluting of traces of sulfate with 3M KC1 solution and subsequent rinsing with deionized water to remove excess chloride (Carmody et al., 1998). A convenient alternative are BioRad TM PolyPrep columns pre-filled with 2 mL of anion exchange resin, which need in our experience no pre-conditioning. After passing a sufficiently large sample volume to generate a few mg of BaSO4 through the anion exchange resin, the moist resin can be stored in a refrigerator until further processing. To remove the sulfate quantitatively, a small volume (e.g. 15 mL per 2 mL of resin) of concentrated chloride solution (e.g. 3 M KC1 or HC1) is passed through the resin and the sulfate-containing eluant is collected in a beaker. Subsequently, BaSO4 is precipitated as described above. BaSO4 produced by any of the above techniques is left to settle overnight, filtered off with a pre-weighed 0.45 gm membrane, and subsequently washed thoroughly with deionized water to remove C1- from the precipitate. Thereafter, the precipitate is either air or oven dried. Re-weighing of the BaSO4 containing filter paper allows for an approximate determination of the weight of the precipitate and hence yields an


565

estimated sulfate concentration for the sample. If organic contaminants appear to be present e.g. in DOC-rich solutions, baking of the BaSO4 precipitate at 800~ in a muffle furnace is recommended in case that only sulfur isotope ratio are to be determined. However, this procedure is not suitable if oxygen isotope measurements are planned, because of the risk of oxygen isotope exchange at high temperatures. Direct precipitation of BaSO4 as well as pre-concentration of sulfate e.g. via ion exchange techniques and subsequent isotope ratio mass spectrometry can yield accurate ~34S values for dissolved sulfate with an uncertainty of < +0.2 %o. Incomplete recovery of sulfate from ion exchange resins tends to yield slightly too low 634S values (e.g. Carmody et al., 1998).

26.3.2.2 Dissolved organic sulfur Some solutions (e.g. forest floor seepage water) contain significant amounts of dissolved organic carbon (DOC). It is a common procedure to oxidize the organic carbon either with H202 or with a saturated bromine solution to avoid contamination of BaSO4 precipitates with organics (e.g. Carmody et al., 1998). However, it must be noted that some amino acids of the DOC fraction may be S-containing. This carbonbonded organic sulfur is oxidized to sulfate and will contribute to the sample. One of the co-authors has attempted to determine the isotope composition of dissolved organic sulfur via mass and isotope balances (Mayer, 1993). A DOC-containing water sample was split in half, and dissolved sulfate was precipitated as BaSO4 from one aliquot by addition of BaC12 solution. The second aliquot of the water sample was subject to H202 oxidation under UV light to oxidize DOS to SO42-. Subsequently, total sulfate was precipitated as BaSO4 as described above. Sulfate concentrations in the second aliquot increased only by less than 2% and 634S values of the sulfate from both samples were found to be identical within the uncertainty of the method, hence preventing a conclusive determination of the sulfur isotope composition of the dissolved organic sulfur (DOS). Alternate techniques comprise isolation of a sufficient quantity of DOC via resin techniques (e.g. Fluka XAD 8) and subsequent isotope analysis of total S of the dried DOC sample e.g. by the Eschka method (see section 26.3.3.1). Using this technique, fulvic and humic acids in the Gorleben aquifer (Germany) were found to have 634S values varying between-3 and +17 %0 (Wang et al., 1998).

26.3.2.3 Dissolved sulfide One of the challenges of obtaining reduced inorganic S species for isotope analyses is to prevent their oxidation during sampling. Hence, exposure to atmospheric oxygen must be avoided wherever and whenever possible. Therefore, we do not recommend filtering of sulfide-containing water samples prior to further processing. Generally, there are two different procedures for collecting reduced inorganic S species for isotope analysis" (1) direct precipitation or (2) degassing of H2S from a water sample with a N2 stream. Direct precipitation is particularly suitable for water samples with > 1 mg L-1 dissolved sulfide. The technique involves adjustment of the pH value of the water sample to either strongly acidic conditions under which H2S is thermodynamically stable

566

Chapter 26 - B. Mayer& H.R. Krouse

or strongly basic conditions, under which $2- is prevalent (see Figure 26.3). Sodium or potassium hydroxide have been used to create alkaline solutions, to which either cadmium or zinc acetate is added to precipitate CdS (Van Everdingen et al., 1982) or ZnS (Rye et al., 1981), respectively. Zinc solutions are environmentally more desirable, whereas Cd solutions tend to react faster and more quantitatively (e.g. McKibben & Eldridge, 1989). Direct precipitation of Ag2S by addition of a AgNO3 solution is typically unsatisfactory because of co-precipitation of AgC1. Acidification of the water sample with cadmium acetate and trapping of the evolving H2S as CdS has also been successfully employed (e.g. Fouillac et al., 1990). We recommend this technique for rapid sampling of dissolved sulfide in the field using sampling bottles pre-filled with circa 100 mL of cadmium acetate solution. After pumping the water sample for a few moments and discarding the initial sample, the end of the pump hose should be inserted into the cadmium acetate solution at the bottom of the sampling bottle. This ensures that the water sample is not exposed to atmospheric oxygen prior to mixing with the cadmium acetate solution, in which dissolved sulfides will rapidly precipitate as bright orange-yellow CdS. Upon return to the laboratory, the CdS is filtered off with a 0.45 ~m membrane filter. In case of significant contamination of the CdS with other solids (e.g. silt) it may be necessary to purify the sulfide precipitate by acidification with 6M HC1 and collection of the evolved H2S in a Cd acetate trap (see section 26.3.3). Wet CdS is subsequently converted to Ag2S by titration with 0.1 M AgNO3 solution containing a 3% NH4OH solution to complex C1- and avoid undesirable AgC1 precipitation. The Ag2S precipitate is subsequently filtered, thoroughly washed with deionized water, and air-dried prior to mass spectrometric measurements. As a note of caution we emphasize that Cd acetate is extremely poisonous and advise that utmost caution is necessary during transport and handling of Cd acetate-containing sampling containers. An alternate technique for obtaining dissolved sulfide species for isotope analyses particularly from large water samples (> 20 L) with low sulfide concentrations (< 1 mg L-l) was suggested by Moncaster & Bottrell (1991). The technique involves the transfer of the water sample into a large carboy and subsequent acidification with concentrated HC1 (1 mL L-I) to drive the dissolved sulfide species equilibrium towards H2S. Care must be taken to minimize exposure of the water sample to atmospheric 02 during the filling procedure. The evolving H2S is subsequently purged with a N2 stream (0.1 - 0.3 L min-1) and trapped as Ag2S in a NH4 + containing silver nitrate solution (see Volume II, Part 3, Figure 14.5). To maximize recovery of H2S, the degassing procedure should be performed for several hours. The Ag2S precipitate is subsequently filtered on a pre-weighed 0.45 gm membrane. Re-weighing the filter paper with the dried precipitate enables gravimetric determination of dissolved sulfide contents. Note that Ag2S collected by N2 stripping will be isotopically inhomogeneous with early-formed Ag2S being enriched in 32S relative to the later-formed Ag2S (Carmody et al., 1998) since there is small isotope fractionation between H2S(gas) and H2S(aq) favoring 32S in the former (e.g. Carmody et al., 1998; Szaran, 1996). To avoid erroneous results it is therefore desirable to approach complete recovery and to remove and homogenize the entire Ag2S from the filter paper prior to isotope analysis. Carmody


567

et al. (1998) report that dissolved sulfide collected by N2 stripping yields typically 0.2 to 0.5 %0 lower ~)34Svalues than sulfide obtained by direct precipitation presumably due to incomplete recovery.

26.3.2.4 Simultaneously occurring reduced and oxidized sulfur compounds Many water samples contain both reduced and oxidized forms of sulfur. The challenge is to quantitatively sample the sulfide fraction before processing the sulfate. Since dissolved sulfide may have a 634S value more than 30 %0 lower than that of the associated sulfate in the same water sample, oxidation of even small portions of dissolved sulfide may bias the isotopic composition of dissolved sulfate. The choice of sampling procedures is dependent on both the concentration of dissolved sulfide and the sulfide to sulfate ratio in the sample. For samples with high sulfide concentrations (> 1 mg L-l), the most effective technique to trap dissolved sulfides is using sampling bottles pre-filled with Cd acetate solution in the field, in which sulfide will precipitate as CdS prior to exposure to atmospheric 02 (see section 26.3.2.3). However in samples containing both dissolved sulfide and sulfate, the contribution of co-precipitated pale yellow CdSO4 may be substantial and failure to remove it gives erroneous gravimetric and isotope composition determinations. After recovering the precipitate on a membrane filter, it is therefore recommended to subject the precipitate to reduction with 6 M HC1 in a distillation apparatus as described in section 26.3.3.2 in order to recover S from the sulfide fraction only. To the remaining sulfate-containing water sample, BaC12 solution is added in the laboratory to produce BaSO4, which is subsequently recovered by filtration and subject to isotope ratio mass spectrometry. There are a variety of procedures to deal with samples with low sulfide concentrations and they are somewhat dependent on the sulfide to sulfate ratio in the sample. One option is to conduct the above described N2 stripping technique in the field, but this requires the hauling of heavy equipment (e.g. N2 tank) and several hours of time per sample. By using flow rates between 1.1 and 2.1 L N2 min-1, Carmody et al. (1998) achieved > 80 % sulfide recovery and 634Ssulfide values only 0.25 %o lower than expected for a groundwater sample from the Floridian aquifer. However, if sulfide recovery is less than 50 %, the obtained ~34Ssulfide value may be more than 1%0 lower than expected (Carmody et al., 1998). In some studies only the isotopic composition of dissolved sulfate is of interest. It has been suggested that under such circumstances rapid precipitation of BaSO4 may yield satisfactory results despite the presence of trace amounts of sulfide. At high sulfate concentrations (> 20 mg L-l), BaC12 solution can be added and the rapidly forming BaSO4 precipitate can be filtered off after 5 minutes. Carmody et al. (1998) showed that this technique provides reliable 634S value for sulfate, if the sulfate to sulfide ratio exceeds 40. For sulfate to sulfide ratios below 40, driving off H2S with the N2 stripping technique without capturing the H2S is an alternate option for ensuring accurate isotope analyses on the dissolved sulfate. Carmody et al. (1998) recommend N2 flow

568


rates of 8 L min-1, covering of the sampling container to avoid contact with atmospheric 02, and continuation of the procedure until sulfide concentrations in the water sample decrease below 0.01 mg L-1. In most of their experiments this was achieved in less than one hour and little sulfate was produced via sulfide oxidation during the outgassing process (SO4 2- concentrations < 0.5 mg Lq). However, since sulfate produced via sulfide oxidation has a sulfur isotopic composition similar to that of the sulfide, it may nevertheless modify the isotopic composition of sulfate in the water samples with sulfate to sulfide ratios below 10.

26.3.3 Sulfur in the lithosphere There are numerous forms of sulfur in minerals, rocks, and sediments. They include sulfates (barite, anhydrite, gypsum, etc.), elemental sulfur, mono-sulfides (sphalerite, pyrrhotite, galena etc.), di-sulfides (pyrite etc.), as well as numerous different forms of organic S. Total sulfur refers to the sum of all these individual S compounds in a sample. Total S contents of lithospheric materials may vary from more than 30 % in some sulfide ores to less than 0.005 % in some plutonic rocks such as granites. In this section we summarize techniques suitable for extracting sulfur from magmatic, metamorphic, and sedimentary rocks. Many of the described techniques are also suitable for extracting S from peat deposits or coal (see also section 26.3.4 and 26.3.5). Pretreatment procedures for lithogenic material depend on sample type and the scope of the study. In some cases it is possible to handpick individual sulfur minerals such as pyrite. More typically, rock samples are broken and finely ground in a mill to mesh size 4:0 prior to analysis. In cases where loss of reduced inorganic S compounds such as H2S is anticipated (e.g. fresh lake sediment), sample pretreatment in a glove box under N2 or Ar atmosphere may be necessary (Lasorsa & Casas, 1996).

26.3.3.1 Total sulfur Many studies reported in the literature have determined ~)34Svalues of total sulfur (e.g. Nriagu & Coker, 1983) using a variety of different techniques (see also Volume IL Part 3, Chapters 8 & 13-2). Chemical oxidants may be employed either as high temperature fusion mixtures or as wet acid treatments for extracting total sulfur from geological samples. In both cases all S compounds in the sample are converted to SO4 2- for subsequent precipitation of BaSO4. The amount of raw material required to generate enough BaSO4 for subsequent mass spectrometry is dependent on the S content of the sample and may vary from less than 100 mg to as much as 50 gram. Eschka mixture is suitable for oxidizing S compounds in coal and coke to sulfate (e.g. Smith & Batts, 1974), but has also been used for total S analysis in whole rocks and sediments (e.g. Fry, 1986; Nriagu & Soon, 1985). Eschka mixture consisting of MgO and anhydrous Na2CO3 in a weight ratio of 2:1 is mixed with dry and ground sample material in a crucible and heated to 800 + 25~ After slow cooling, the generated sulfate is rinsed with hot distilled water into a beaker (ASTM, 1993), where it can be precipitated as BaSO4 by adding BaC12 solution (see Volume II, Part 3, Chapter 132.8). Since the amount of Eschka mixture used exceeds the weight of the sample by a

Procedures for SulfurIsotopeAbundanceStudies

569

factor of four, it is important to use an Eschka mixture with the lowest possible S blank. Although commercially available Eschka mixtures (e.g. from MERCK) have S blanks of less than 5 gg g-l, blank corrections may be necessary in determining the true ~34Stotal value of S-poor materials such as granites. Wet chemical techniques for oxidizing lithogenic S compounds to sulfate have the advantage that the S blanks of the utilized chemicals are often negligible, and that there is typically no limit in sample size. This is particularly convenient if mg quantities of BaSO4 must be generated for DI-IRMS from S-poor materials. Digestion of raw materials with a HNO3-Br2 mixture has proven satisfactory (Krouse & Tabatabai, 1986; Zhabina & Volkov, 1978). This technique involves soaking a sample overnight in open beakers in a fumehood with conc. HNO3 and liquid Br2 and subsequent heating to dryness (see Volume II, Part 3, Chapter 13-2.8). The generated sulfate is dissolved in dilute HC1 and after removing undissolved matter by filtration the sulfate-containing solution is transferred into a beaker, where BaSO4 is precipitated by addition of BaC12 solution. Many laboratories prefer Ag2S for conversion to SO2 or SF6 for subsequent isotope ratio mass spectrometry. An alternative to oxidizing all S compounds to sulfate and generating BaSO4 is their reduction to H2S and subsequent formation of Ag2S. This is most commonly achieved by the Kiba technique, which was initially developed for the reduction of sulfate (Kiba et al., 1955). The Kiba solution, which consists of waterfree phosphoric acid mixed with tin(II)-chloride dihydrate (Kiba et al., 1955), reacts in a closed and de-oxygenated reaction flask with geological samples between 120 and 280~ to evolve H2S, which is swept with a N2 carrier gas through a washing solution into a chemical trap to form ZnS or CdS (see Volume II, Part 3, Chapter 8-3.6). These compounds can be subsequently converted to Ag2S by titration with AgNO3 solution. Soon after its invention, it was however found that the Kiba technique also converts organic S compounds (Ohashi, 1955) and sulfide minerals such as pyrite (Kiba et al., 1957) to H2S. Hence, the Kiba technique is often used for total S analysis, although it has been reported that some common sulfide minerals (e.g. arsenopyrite) and elemental S are only partially recovered (Sasaki et al., 1979). If Ag2S is preferred as the final reaction product, it is also possible to convert all S compounds in a geological sample to sulfate, followed by the conversion of the latter to H2S by reduction with Sn(II) strong phosphoric acid (Sasaki et al., 1979) or various hydriodic acid mixtures as described by Johnson & Nishita (1952) and Thode et al. (1961), among others. The produced H2S is trapped as CdS or ZnS in acetate solutions, and these compounds can be subsequently converted to Ag2S by titration with AgNO3 solution. Recently, generation of SO2 from geological materials in an elemental analyzer followed by CF-IRMS has also been successfully attempted (B6ttcher & Schnetger, Chapter 27). This technique is only suitable for samples with sufficient Stotal contents (e.g. > 100 ppm), since commercial autosamplers accommodate only samples of less than 100 mg. It has also been reported that matrix effects may result in erratic results partic-

570


ular with samples containing pyroxene, amphibole or biotite (Studley et al., 2002). Further limitations of this technique have been discussed in section 26.2. If a lithospheric sample is dominated (> 98 %) by one particular S compound, the use of any of the above-described extraction techniques may be justified to asses the 634S value of the predominant form of sulfur. This is because minor S components would presumably not change the result by more than 1%o even if they are isotopically distinct. However, most geological materials contain a mixture of several S compounds with often quite different 634S values. In this case, extraction of individual S compounds is strongly recommended, since total S analysis would provide an often meaningless average 634S value dependent on the quantitative proportions of the respective S compounds and their sulfur isotope ratios.

26.3.3.2 Inorganic sulfur compounds In some cases, it is possible to hand-pick pure sulfur minerals such as pyrite or barite or to use gravimetric techniques (e.g. heavy liquids) for their separation. After isolating these minerals, they may be used for direct conversion to SO2 or SF6 for subsequent sulfur isotope measurement (e.g. Ueda & Krouse, 1986) or for laser ablation (Volume II, Part 3, Chapter 8-1.7). More commonly however, sulfur compounds are finely interspersed in the rock matrix and have microscopic or sub-microscopic size. In these cases, wet chemical extraction procedures are suitable for recovering individual S compounds from geological materials. For this purpose, it has become common practice to describe lithospheric S compounds in broad groupings based on their similarity in chemical properties. It is important to note that these groupings are operationally defined and include terms such as acid volatile sulfur (AVS), chromium reducible sulfur (CRS), and hydriodic acid reducible sulfur (HI-red S), among others. Ideally) methods extracting discrete S compounds should do so without affecting other S forms present, although this is not always the case.

Acid volatile sulfur (AVS) Since sulfide is liable to oxidation once exposed to air, careful sample handling in the field is essential to prevent loss of S compounds. Where possible handling of material under a N2 atmosphere is recommended, and samples should be preferably stored at 4~ or frozen and analyzed within 2 weeks of collection (Lasorsa & Casas, 1996). An alternate solution is the addition of a 20% zinc acetate (ZnAc) solution (w/v) to the sample thereby fixing dissolved sulfide and some acid volatile S by forming more stable ZnS (Duan et al., 1997b). Upon return to the laboratory and after thawing the sample, acid volatile sulfur (AVS) is liberated as H2S by treatment with 6M HC1 (e.g. Tuttle et al., 1986) in a distillation apparatus similar to that shown in Volume II, Part 3, Figure 13-1.10. It is widely assumed that AVS comprises predominantly monosulfide minerals such as sphalerite (ZnS), galena (PbS), pyrrhotite (magnetic Fe0.9S), mackinawite (Fel.IS), greigite (magnetic F3S4), and amorphous monosulfides of other elements. Up to 100 gram of sample can be placed in a 200 mL reaction flask. The flask is attached to the distillation apparatus and the entire system is de-oxygenated for at least 15 minutes by a stream of N2. Thereafter, 40 to 80 mL of 6 M HC1 are introduced with a syringe via the rubber septum and the reaction is allowed to pro-


571

ceed at room temperature. The released H2S is swept by the carrier gas through a washing solution (e.g. pyrogallol) into a trap with Zn or Cd acetate solution, where ZnS or CdS precipitates. After 25 minutes at room temperature, the reaction vessel may be heated to 70~ for circa 15 minutes since minerals such as crystalline pyrrhotite and greigite will not completely react with cold HC1 (e.g. Tuttle et al., 1986). Rinsing the residual sample with hot distilled water yields the operationally defined fraction of acid soluble sulfate (see below). Unfortunately, the selectivity and recovery of this extraction technique is somewhat dependent upon the nature of the sample and maturity of the respective reduced S compounds. One potential problem with AVS extractions may arise if acid soluble Fe(III) is present, which can oxidize generated H2S to elemental S within the reaction flask. Since this oxidation will likely proceed with negligible sulfur isotope selectivity, the 634S values determined for the recovered Ag2S may still be representative for acid volatile S, but gravimetric determinations will underestimate the acid volatile S fraction. A potential solution to this problem is the addition of tin(II) chloride to the sample prior to the reaction. Tin(II) rapidly reduces Fe(III) to Fe(II) and thus prevents it from reacting with H2S (Pruden & Bloomfield, 1968) usually without mobilizing significant amounts of pyrite during treatment with cold HC1 (Chanton & Martens, 1985). However, hot HC1 + SnC12 digestion has been shown to liberate some pyrite S (Cornwell & Morse, 1987; Fossing & Jorgensen, 1989; Rice et al., 1993).

Chromium reducible sulfur (CRS)" Chromium reduction converts sulfur from monosulfide minerals, elemental S, and di-sulfides such as pyrite (cubic FeS2) and marcasite (ortho-rhombic FeS2) to H2S (e.g. Fossing & Jorgensen, 1989; Zhabina & Volkov, 1978) and is hence widely accepted as a means to determine total reduced inorganic sulfur. Chromium reducible sulfur can be extracted from powdered rock samples (< 100 mesh) using the same distillation apparatus as described above and shown in Volume II, Part 3, Figure 13-1.8. After the sample is placed into the reaction apparatus together with 10 mL ethanol, a continuous stream of N2 is established and 40 - 60 mL of 1M CrC12 solution (obtained by reduction of CrC13) and 20 mL 6M HC1 are introduced with a syringe via a rubber septum (Canfield et al., 1986). The sample and the solution are boiled for 1 to 2 hours and the released H2S is trapped as ZnS or CdS and can be subsequently converted to Ag2S for isotope analysis as described above. Canfield et al. (1986) reported recoveries between 92 and 97 % for elemental S, monosulfides, and disulfides using this technique, while sulfate minerals and organic S compounds were shown to be unreactive (exception cystine with 2 % recovery). The method is specific for disulfides (mainly pyrite) if monosulfides and elemental S are removed prior to chromium reduction (see section 26.3.3.4). However, Newton et al. (1995) documented that 634S values are typically shifted by +0.55 %o as a result of incomplete conversion of pyrite (usually 95-96 % recovery) using this technique. Some researchers have attempted to separately extract less mature pyrite (synthetic or recently formed) with cold CrC12 solution (1 hour) followed by extraction of mature pyrite with hot CrC12 solution (1 hour) as described above (Duan et al., 1997b; Fossing & Jorgensen, 1989). Others have reacted rock samples with lithium aluminum hydride (LiA1H4, short LAH) to convert pyrite S into H2S (see Volume IL Part 3, Chapter 8-1.5), which was subsequently trapped as CdS or Ag2S (e.g. Smith et al., 1964; Westgate & Anderson, 1982).

572


We do not recommend this technique since the reagent can be very explosive and yields for LAH reduction of pure pyrite are often low (e.g. Tuttle et al., 1986). Also, extracting pyrite (and sulfate) with 2M nitric acid + bromine (Riley et al., 1990) is not recommended since this treatment is known to also mobilize some organic sulfur (Bottrell et al., 1994; Smith & Batts, 1974).

Elemental S: Elemental (or native) sulfur in rocks and sediments was rarely extracted as a separate fraction prior to the 90's (e.g. Zhabina & Volkov, 1978) but has recently gained increased attention since the significance of bacterial disproportionation of elemental S for the sedimentary sulfur cycle has been detected (e.g. Canfield & Thamdrup, 1994). Measurement of isotope ratios of elemental S usually involves extraction with organic solvents such as benzene, methylene chloride, CC14, dichloromethane (CH2C12), or acetone (Duan et al., 1997b; Hall et al., 1988; Smith & Batts, 1974; Wieder et al., 1985; Zaback & Pratt, 1992). Some researchers have also used carbon disulphide (CS2), but traces of CS2 may decompose upon heating (Fossing & Jorgensen, 1989) and hence we do not recommend this solvent for extraction. Elemental S can be extracted by adding 70 to 150 mL of the organic solvent (e.g. acetone or dichloromethane) to the sediments and shaking for 16 hours (Duan et al., 1997b; Wieder et al., 1985) or by ultrasonic agitation of the sample in a beaker (Hall et al., 1988). Zahbina & Volkov (1978) report that acetone extraction also recovers part of the bituminous organic sulfur from the samples. The elemental S containing solution is filtered through 0.45 ~m PTFE membranes. Subsequently, the solvent with the elemental S is either transferred to the above described distillation apparatus, and reduced to H2S via chromium reduction (Canfield et al., 1986), or evaporated to dryness followed by oxidation to sulfate and precipitation of BaSO4 (e.g. Hall et al., 1988; Smith & Batts, 1974). An alternate technique described by Zaback & Pratt (1992) features extracting approximately 50 gram of rock powder in a Soxhlet apparatus for 48 hours using methylene chloride. Granular copper was added to the collection flask to remove elemental S, which was solubilized during the extraction. The resulting copper sulfide was converted to SO2 for subsequent dual inlet isotope ratio mass spectrometry. Sulfate: Sulfate occurs in rocks as minerals such as barite, anhydrite, gypsum, or less commonly as hydroxy-sulfate minerals (alunite, jarosite, basaluminite etc.). Sulfate is also found in pores in water-soluble form or is occasionally adsorbed on iron and aluminum oxides and hydroxides, particularly under acidic conditions. Carbonate rocks contain sulfate as trace constituents (Staudt & Schoonen, 1994) in what is often referred to as structurally-substituted sulfate. There are numerous techniques described in the literature for removing sulfate from geological samples. Extraction in an open beaker under ultrasonic agitation or in an overhead shaker for several hours at 10 rpm at a solution to sample ratio of 10:1 by weight has been found satisfactory for most procedures described in this section. Extraction with distilled water yields the fraction of water-soluble sulfate (e.g. Hall et al., 1988), which comprises pore water sulfate and some mineral sulfate (e.g. gypsum). Sulfate minerals such as barite, anhydrite, or gypsum can be leached from geological samples by reacting with a 5% Na2CO3 solution (Breit et al., 1985). It is noteworthy that the high pH value of the lat-


573

ter reaction solution may promote conversion of labile organic sulfur to sulfate, but the content of labile organic sulfur in mature geological samples is typically negligible. Reacting a geological raw sample with 6 M HC1 and subsequent rinsing with hot distilled water yields the fraction of acid soluble sulfate. In all cases, dissolved sulfate can be converted to BaSO4 by addition of BaC12 solution. To yield pure BaSO4 it may be necessary to increase the pH of the solution e.g. with ammonia or NaOH and remove Fe(OH)3 by filtration prior to the addition of BaC12 (e.g. Hall et al., 1988). An alternate technique to extract sulfate-S from geological samples is its reduction to H2S. Sulfate-S can be converted to H2S by the Thode reduction mixture consisting of HIH3PO2-HC1 in a ratio of 4:2:7 (Thode et al., 1961), or by the Johnson-Nishita reduction mixture, which comprises HI-HCOOH-H3PO2 in a ratio 4:2:1 (Johnson & Nishita, 1952). The Kiba reduction technique (Kiba et al., 1955) is not recommended for this purpose since it is not specific for sulfates (Sasaki et al., 1979). The selective analysis of structurally substituted sulfate in carbonate specimens requires samples, which are free of other sulfate or sulfide minerals. Pre-treatment with a 5.25 % sodium hypochlorite (NaOC1) solution has been suggested as suitable for removal of organic matter, soluble sulfates, and metastable sulfide minerals associated with the sample (Burdett et al., 1989). Subsequently, the powdered carbonate sample can be digested with 6 M HC1 in a N2 stream as described above, removing sulfide S as H2S and liberating the structurally bound sulfate from the calcite lattice as acid soluble sulfate (e.g. Burdett et al., 1989; Kampschulte et al., 2001; Ohkouchi et al., 1999). Following filtration of the insoluble residue, Hurtgen et al. (2002) raised the pH to values between 3 and 5 by adding NaOH. Subsequently, they added 10 to 15 mL of saturated bromine water to facilitate the precipitation of iron oxyhydroxides. After their removal by filtration, BaSO4 was precipitated by addition of BaC12 solution. The in vacuo Kiba technique constitutes an alternate approach for analyzing trace sulfates in carbonate rocks. This method uses Kiba solution (Kiba et al., 1955, see also section 26.3.3.1) with one tenth (1/10) of the original Sn2+ concentration and extractions are carried out under vacuum (Ueda & Sakai, 1983). Whereas sulfide-S is converted to H2S, sulfate is only reduced to SO2, which is cryogenically trapped and subsequently used for isotope analysis. One problem with this technique is that a small fraction of the sulfate (< 5 %) does convert to H2S accompanied by a slight enrichment of 32S (Krouse & Ueda, 1987). Bottomley et al. (1992) found that this technique is very effective in releasing sulfur from barite as SO2, but is not quantitative in reducing pyrite to H2S.

26.3.3.3 Organic sulfur compounds

In many geological materials, the amount of organic S is minute compared to that of inorganic S. To our best knowledge, there are no reliable techniques, which can extract individual organic sulfur compounds from fresh geological samples for isotope ratio measurements (see also sections 26.2, 26.3.4 and 26.3.5). Hence, the isotopic composition of organic sulfur in geological samples is typically determined by removal of all inorganic S constituents followed by conversion of total S in the sample residue to BaSO4 or Ag2S by any of the previously described techniques (e.g. Eschka method, digestion with HNO3-Br2, etc.). Subsequent isotope ratio mass spectrometry

574


yields a ~34S value which is thought to be representative for organic S. Another option is to obtain the kerogen fraction of sedimentary rocks by treatment with HC1 and HF under N2 atmosphere (e.g. Hitchon, 1974; Zaback & Pratt, 1992) and subject it to a total S isotope ratio determination (section 26.3.3.1).

26.3.3.4 Extraction schemes for isolating individual sulfur compounds In most geological samples, several inorganic and organic sulfur compounds do

occur simultaneously in variable quantities. There are two fundamentally different approaches to extract different S compounds from a powdered rock sample: 1. to extract the S compounds of interest (e.g. acid volatile S, acid soluble sulfate, chromium-reducible S, remainder = organic S) sequentially from one sample, 2. to extract individual S compounds (e.g. total S, acid volatile S, acid soluble sulfate, chromium reducible S) from fresh aliquots of the same sample. The latter approach has often been used for content analyses with some sulfur compounds being calculated by difference (e.g. organic S - total S - 21 inorganic S compounds). However, obtaining sulfur isotope ratios for the calculated S compounds requires isotope and mass balances, which can result in large uncertainties particularly for S compounds representing a small part of the total sample. Hence, sequential extraction schemes are commonly preferred for obtaining ~)34S values for different S fractions of geological samples. It has become common practice to combine a number of the previously described extraction procedures to a sequential extraction scheme (e.g. Bates et al., 1993; Hall et al., 1988; Rice et al., 1993; Zhabina & Volkov, 1978), which recovers the individual S fractions either as Ag2S or BaSO4 to allow for both gravimetric content determination and isotope analyses. Selecting the appropriate extraction scheme for a given sample depends on the relative abundance of organic and inorganic sulfur compounds and on the lability of these compounds. Most rock samples contain sulfur predominantly in form of sulfate, monosulfide and disulfide minerals, and as organic S. For such samples, an extraction scheme was suggested by Tuttle et al. (1986), which is in our view suitable for most magmatic, metamorphic, and mature sedimentary rocks (Figure 26.4). To avoid or minimize the oxidation of reduced inorganic sulfur compounds, it is advisable to extract these compounds at the beginning of a sequential extraction procedure. Monosulfide minerals are initially recovered from the powdered sample by releasing acid volatile sulfur (AVS) as H2S using hot 6 M HC1 in a distillation apparatus described above. Hot HC1 should be used to facilitate complete recovery of minerals such as crystalline pyrrhotite and greigite. Additions of stannous chloride is only recommended if the pyrite fraction is well-crystallized and coarse grained (Chanton & Martens, 1985; Rice et al., 1993), since other pyrite forms may be partially liberated by the vigorous hot HC1 +SnC12 treatment (Cornwell & Morse, 1987; Fossing & Jorgensen, 1989). The evolving H2S is carried with a N2 stream into a cadmium acetate trap, where S precipitates as CdS, which is further converted to Ag2S. Subsequently, the residual rock sample is removed from the reaction vessel and is thoroughly rinsed with hot distilled water to remove acid soluble sulfate, which is precipitated as BaSO4 for subsequent sulfur isotope analysis; some researchers prefer the reduction of sulfate to H2S and subsequent conversion to Ag2S (Zhabina & Volkov, 1978) for sulfur


575

Figure 26.4 - Flow chart summarizing analytical procedures for extracting monosulfide S (acid volatile S), acid-soluble sulfate, elemental S + disulfide S (chromium-reducible S), and organic S from magmatic, metamorphic and sedimentary rocks for subsequent isotope analysis (dw = distilled water).

isotope ratio determinations. The remaining sample is returned to the distillation apparatus and undergoes CrC12 reduction yielding H2S, which is converted to Ag2S. This precipitate is representative of the elemental S and di-sulfide (pyrite and marcasite) fractions of the sample, since monosulfide minerals have been previously removed. In case that all sulfates in the sample were acid soluble, the above-described procedures removed all inorganic S compounds from the sample, and only organic S compounds remain in the sample residue. The sample residue is thoroughly washed with distilled water and subsequently subjected to a total S extraction (e.g. Eschka method, HNO3-Br2 digestion, etc.), which yields either BaSO4 or Ag2S (see section 26.3.3.1), which is typically representative for organic sulfur. Tuttle et al. (1986) tested the above described extraction scheme using oil shales and associated rocks from the Eocene Green River Formation (Colorado, Utah, Wyoming) providing evidence for quantitative recovery of the individual S fractions by 57Fe M6ssbauer spectroscopy. The above-described extraction scheme is not suitable for samples containing acidinsoluble sulfates such as barite. Since acid-insoluble sulfates are non-reactive during HC1 and chromium reduction, they will be part of the residual sample together with organic sulfur. From such residues, Smith & Batts (1974) extracted organic S by adding 40 % NaOH solution followed by repeated additions of 100 % H202. After no more

576


residue dissolves upon boiling, the solution is filtered and sulfur compounds in the filtrate are oxidized to sulfate with aqua regina and precipitated as BaSO4 (Smith & Batts, 1974). The predominantly inorganic residue, which contains the acid-insoluble sulfate minerals such as barite and possibly some recalcitrant organic S, subsequently undergoes a total S extraction (section 26.3.3.1). Whereas Smith & Batts (1974) used the Eschka method, Hall et al. (1988) preferred the Kiba reduction technique because of its excellent yield for barites. Modern sediments usually contain significant amounts of organic S and elemental S and the reduced inorganic S fractions may be considerably more labile than those of ancient rocks. Hence, procedures to sequentially extract S compounds from such samples must be adjusted accordingly. A potential approach is depicted in Figure 26.5. The fresh sediment sample should be immediately treated with a 20% zinc acetate (ZnAc) solution (w/v) to fix dissolved sulfide and some acid volatile S by forming more stable ZnS (Duan et al., 1997b). Upon return to the laboratory and after thawing and potentially freeze drying the sample, elemental S is extracted with acetone or another solvent (e.g. Rice et al., 1993). Duan et al. (1997b) provided evidence that initial extraction of elemental S with dichloromethane does not affect the AVS pool, if AVS was pre-fixed using ZnAc. Subsequently, acid volatile sulfur (AVS) is liberated from the remaining sediment sample as H2S by treatment with cold 6M HC1 (e.g. Rice et al., 1993). This comparatively mild AVS treatment is usually successful in collecting the majority of the S from monosulfides without liberating S from disulfide minerals such as pyrite. The use of stannous chloride in conjunction with AVS extraction from modern sediments is strongly discouraged, since it would almost certainly liberate S from the disulfide (pyrite) fraction. The sample residue is thoroughly washed and filtered with hot distilled water. Acid soluble sulfate is precipitated as BaSO4 from the filtration solution by adding BaC12. The solid sample residue may contain some elemental S, which was generated during AVS extraction via reaction of H2S in the presence of Fe(III). Therefore, it is recommended to conduct an additional solvent extraction (acetone, dichloromethane, or others) on the solid sample residue after AVS liberation. The elemental S containing solvent is evaporated and the Ag2S generated via chromium reduction is representative for part of the acid volatile S fraction (Rice et al., 1993). Subsequently; disulfide sulfur is recovered from the sample residue by chromium reduction (Canfield et al., 1986). After thoroughly washing the sample residue with distilled water, a total S extraction (e.g. Eschka method, HNO3-Br2 digestion, etc.) is performed, which yields either BaSO4 or Ag2S representative for the organic S fraction of the sample, since barite is typically non-existent in modern sediments.

26.3.4 Sulfur in fossil fuels

Sulfur in petroleum and coal specimens occurs as complex assemblies of organic and inorganic compounds. In some cases, organic coatings on inorganic compounds may interfere with their extractions. A potential health hazard is gaseous or dissolved H2S in some samples.


577

Figure 26.5 - Flow chart summarizing analytical procedures for extracting elemental S, monosulfide S (acid volatile S), acid-soluble sulfate, disulfide S (chromium-reducible S), and organic S from modern sediments for subsequent isotope analysis (dw = distilled water).

578


Historically, three techniques have been used to extract total sulfur from liquid and solid fossil fuels for sulfur isotope analyses. 1) In the quartz combustion method (ASTM, 1958), chemically scrubbed air or 02 is passed over the sample in a horizontal porcelain or quartz combustion boat at temperatures above 950~ The combustion products are passed through a H202 solution, which adsorbs sulfur as dissolved sulfate. This technique has never gained widespread popularity for sulfur isotope work. 2) In the bomb combustion method (ASTM, 1964), all sulfur compounds in a sample are converted to sulfate in a closed high temperature reactor under high pressures (2.5 MPa) of 02 (Siegfriedt et al., 1951). Suitable bombs are manufactured by the Parr Instrument Company (Moline, Illinois, USA). In Parr bomb model 1108, up to I gram of material can be placed in a sample cup. One mL distilled water is added to the bomb and oxygen gas is admitted slowly prior to closing. Subsequently, the bomb is placed in a water bath and the sample is ignited by passing a current through a thin Ni alloy wire placed above the cup. Sulfur in the sample is quantitatively converted to SO3, which reacts with the water in the bomb to H2SO4. Circa 15 min after ignition, unexpended 02 and produced CO2 are released and the bomb is opened. The dissolved sulfate is carefully transferred into a beaker through a filter paper (Whatman # 1) and converted to BaSO4 by adding BaC12 solution. A 1 gram sample with a total S content of 0.1% should yield 7.3 mg of BaSO4, sufficient for both dual inlet and continuous flow IRMS. 3) Eschka mixture (ASTM, 1993) has also been successfully used to extract total sulfur from coal and coke (see section 26.3.3.1), but one challenge with this reagent is finding a supply with sufficiently low sulfur blank. Isotope data for total sulfur in oil and bitumen have provided information on their origin, migration, and alteration processes (e.g. Thode, 1981; Thode et al., 1958). Thode & Monster (1970) demonstrated that oil accumulations in the Tertiary and Cretaceous of Northern Iraq had a common origin with extensive vertical migration in contrast to postulating four different epochs of oil formation. Other studies identified changes in composition of oils and condensates with H2S during maturation (Manzano et al., 1997; Orr, 1974). In these and other studies, the isotopic data proved more diagnostic when used in combination with data for other parameters. In a study of the Bolivar Coastal Fields (Venezuela), Manowitz et al. (1990) found an inverse relationship between ~34S values and pristane/phytane ratios of crude oils. They concluded that the five major oil classes were derived from a reasonably uniform source rock and that minor isotopic variability resulted from alteration processes. Despite the successful studies cited above, it is important to realize that isotope data for total S tend to have limited usefulness. This is particularly true for coals with biogenic pyrite having 634S values more variable and quite different from those of organic sulfur (Dai et al., 2002; Lei et al., 1994; Tang et al., 2001).


579

26.3.4.1 Coal and bitumen

Separation of different forms of sulfur from coal and bitumen specimens is not straightforward. Occasionally S-containing minerals such as pyrite or elemental sulfur can be handpicked for subsequent isotope analysis. Various analytical schemes are used for disseminated S compounds, which may involve extraction of most inorganic S first, most organic S first, or alternate extraction of inorganic and organic components. One of the fundamental papers on the isotopic composition of sulfur in coal was published by Smith & Batts (1974). They sequentially extracted elemental sulfur, acid soluble sulfate, pyritic S, organic S, and barite if present. One problem with their analytical scheme, as noted by the authors, is that pyritic S was extracted with bromine and nitric acid, which also mobilizes some organic sulfur. Bottrell et al. (1994) found that the liberated organic S is enriched in 32S compared to the total organic S. Therefore, we recommend chromium reduction for extracting pyrite S from coals, since this procedure does not mobilize significant amounts of sulfate or organic S (Bottrell et al., 1994; Canfield et al., 1986). For a basic characterization of four major S forms in coal (AVS, acid soluble sulfate, CRS, organic S), the extraction scheme shown in Figure 26.4 has proven satisfactory (e.g. Chu et al., 1994; Westgate & Anderson, 1982). A more rigorous procedure for extracting various forms of S from organic-rich sediments was suggested by Zaback & Pratt (1992) and is summarized in Figure 26.6. It features a 48 hour Soxhlet extraction using methylene chloride to divide the sample into a soluble and an insoluble fraction. The soluble fraction comprises elemental sulfur, which is reacted with granular copper added to the collection flask. The copper sulfide is subsequently removed for isotope analysis. The remaining sulfur compounds in the soluble fraction are termed bitumen S, which is converted to BaSO4 using a Parr bomb for subsequent isotope analysis. The insoluble sample residue undergoes extraction of AVS with hot 6 M HC1 + SnC12 followed by extraction of acid-soluble sulfate and chromium-reducible sulfur as described in section 26.3.3.4. For barite-free samples, the residue after chromium reduction is rinsed, dried, and treated with 48% HF in order to isolate kerogen S, which is converted to BaSO4 by Parr bomb combustion (Zaback & Pratt, 1992). If barite minerals are present, extraction of organic sulfur with hydrogen peroxide (Hall et al., 1988; Smith & Batts, 1974) followed by Kiba reduction to convert barite-S to H2S appears to be a preferable procedure (Hall et al., 1988). Oxygen isotope analyses on sulfate in coals have also been attempted, but special precautions (e.g. N2 atmosphere) are necessary to ensure extraction of sulfate with negligible effects on other sulfur compounds in the sample (McCarthy et al., 1998). The organic component remaining after removal of inorganic minerals including carbonates and silicates is defined as kerogen (e.g. Hitchon, 1974), which is converted to crude oil and natural gas during maturation. Comparison of the carbon isotope compositions of different solvent extracts of petroleum and kerogen has been successful in identification of oil source-rock relations (e.g. Stahl, 1977). Kerogen is interesting in that as it matures, the N / C ratio does not alter much but the S/C ratio can decrease by I to 2 orders of magnitude (Durand et al., 1972). Few studies have been carried out on its S isotope composition (e.g. Werne et al., 2003). A study by Hitchon & Krouse

580 Chapter 26 - B. Mayer & H.R. Krouse

Figure 26.6 - Flow chart summarizing analytical procedures for extracting bitumen S, elemental S, monosulfide S (acid volatile S), acid-soluble sulfate, disulfide S (chromium-reducible S), and kerogen S from organic-rich samples such as shales and coals (after Zaback & Pratt, 1992) for subsequent isotope analysis (dw = distilled water).


581

cited in Krouse (1977) revealed the following: whereas pyrite associated with bitumen at Windy Knoll (Derbyshire, England) was found to have ~34S values consistently near-10 %o, those of co-existing kerogen varied between 0 and-30 %o with an inverse dependence on 613C values. Purnell & Doolan (1983) recommended the use of Kiba reducing reagent (Sn2+ in H3PO4) under N2 flow at different temperatures for rapidly determining the concentrations of different forms of inorganic sulfur in coal. Initially, HC1 was added to the sample-containing flask to evolve H2S from non-pyritic sulfides at temperatures up to 120~ Then, Kiba reagent was added. H2S, which evolved between 170 and 210~ and 200 and 270~ were attributed to reaction with sulfate and pyrite, respectively. To test the suitability of this technique for isotope analysis, sulfate minerals and pyrite of widely different isotopic compositions were added to a low sulfur coal in our laboratory. Despite using finely ground mineral mixtures and different temperature-time programming profiles, the isotopic data showed considerable overlap in the H2S evolution peaks due to sulfate and pyrite. It may be possible to collect and analyze many H2S aliquots to properly characterize these evolution peaks, but clearly the simplified approach as published (Purnell & Doolan, 1983) is not recommended for obtaining S isotope data. Purnell & Doolan (1983) also found no evidence of reaction between several aromatic S compounds and Kiba reagent. Whereas their conclusion that organic S in coal does not interfere with the S determinations for inorganic sulfate and sulfide is probably valid, some organic S compounds can be converted to H2S with Kiba reagent (Sasaki et al., 1979). Although it is known that the contents of organic S within individual macerals of coal samples may vary significantly (Demir & Harvey, 1991), few researchers have attempted to analyze the isotopic compositions of different organic S compounds in coal or bitumen. Monster (1972) carried out S isotope analyses on oil sand bitumen from the McMurray Formation (Alberta, Canada). The oil sand bitumen was separated into asphaltenes, saturates plus aromatics, dark oils, and three classes of resins by liquid-solid chromatograhy on activated clay using various organic solvents as depicted in Figure 26.7, followed by converting the individual compounds to BaSO4 with the Parr bomb technique. The S isotope compositions were remarkably uniform among the different S fractions with deviations of less than 1.2 %o from the 634S value of total S in bitumen (Monster, 1972). However, these traditional extraction techniques do not separate different S functional groups. When this is done using the technique described below, the ~34S values may vary by over 20 %o in a single sample (unpublished data, Isotope Science Laboratory, University of Calgary). Attar and coworkers (e.g. Attar, 1979) have developed a method for analyzing different sulfur functional groups in coal and heavy oil. It is based on the premise that all organic S functional groups can be reduced to H2S with strong reducing agents. Their rates of reduction are characterized by unique activation energies. If a sample containing many S groups is reduced during gradual increase of temperature, H2S from an individual functional group evolves as a peak in a characteristic temperature range. Experiments with model compounds containing aliphatic thiols, thiopenes, and aryl

582


Figure 26.7- Flow chart summarizing analytical procedures for extracting asphaltenes, saturates + aromatics, dark oils, and various resins from bitumen samples for subsequent sulfur isotope analysis (after Monster, 1972). More recently, many laboratories have replaced benzene with less carcinogenic toluene

sulfides realized 94 to 99 % S recovery (Attar & Dupuis, 1979). Recovery from pure crystalline pyrite was low, ~ 1%. One complication was considerable overlap of the H2S evolution peak from pyrite with that of aliphatic sulfides. Krouse et al. (1987b) used the findings of Attar and colleagues as an approach to determining the S isotope compositions of different functional groups in bitumen and coal samples. The use of organic reducing agents was not as successful as linear temperature increasing (3~ min-1) pyrolysis with N2 flushing. The results were consistent with those of Attar's group with the following temperature range assignments: 200~ to 325 ~ C: elemental sulfur; 325~ to 375~ thiols and disulfides; 375~ to 500~ saturated sulfides; 500~ to 650~ thiopenes; and > 650~ benzothiopenes. Kinetic analyses of the peaks showed that above 500~ H2S was evolved from a solid phase whereas evolution from a liquid phase occurred below 500~ Since this initial work by Krouse et al. (1987b), pyrolysis has been replaced by hydrogenation, i.e. flushing with a 1:20 He:H2 mixture. In the earlier work, 2 to 5 gram of sample (liquid or solid ground to 40 mesh) to a height of a few centimeters were placed in a tube above a horizontal ceramic frit (Figure 26.8). There was concern that H2S released at lower temperatures might participate in sulfurization reactions before evolution from the sample. With introduction


583

of CF-IRMS, it was possible to reduce the sample size to the order of 100 mg and a height of only few millimeters. The revised apparatus (Figure 26.8) tends to give better resolved H2S peaks and the data are consistent with the earlier work. A number of extractions were conducted with sulfide and sulfate minerals of widely different ~34S values added to low-S coal and bitumen. There was no isotopic evidence of H2S evolved from these minerals. In contrast, Krouse et al. (1987b) showed that addition of elemental S to a bitumen sample produced a distinct H2S evolution peak in the 200~ to 325~ region. Comparison of data by Monster (1972) with the programmed H2S evolution technique is quite revealing. Using whole bitumen from the Fort McMurray oil sands, two resolved H2S evolution peaks were found; the larger in the temperature range identified with saturated sulfides and the smaller at higher temperatures associated with thiopenes. In addition, a minor unresolved peak occurred on the low temperature edge Figure 26.8- Apparatus for extracting sulfur from of the major peak. The 634S values of different functional groups in a coal or heavy oil sample. these S sources were found to be +6, 0, and +7 %o respectively. An asphaltene extract had the same H2S evolution peaks with the one at 0 %o being comparatively slightly more abundant. This is consistent with Monster's finding that asphaltene had a slightly lower bulk sulfur 634S value (Monster, 1972). With a sufficient number of consecutively evolved H2S samples, the temperature programmed reduction procedure is potentially a reliable tool for determining the S isotope composition of different S functional groups. It is limited to solids and high viscosity liquids. Materials with higher vapor pressures are likely to distill out of the reactor before H2S is produced. 26.3.4.2 Oil

Traditional evaluation of petroleum quality was based on separation and measurement of the content of different fractions as defined by a combination of distillation and solvent extraction techniques. Monster (1972) carried out S isotope analyses on fractions prepared from oil samples of the Mission Canyon Formation (Saskatchewan, Canada). The oils were divided into low and high boiling fractions. From the former,

584


Figure 26.9 - Flow chart summarizing analytical procedures for extracting asphaltenes, saturates, aromatics, and polar compounds for subsequent sulfur isotope analysis (after Monster, 1972). More recently, many laboratories have replaced benzene with less carcinogenic toluene. asphaltenes were precipitated by mixing with pentane and subsequent filtration. The de-asphaltened sample was further subdivided into saturates, aromatics, and polar compounds by liquid-solid chromatography on an activated F-20 Alcoa alumina column. After placing the sample on top of the column, saturates were eluted with npentane, aromatics with benzene, and polar compounds with a 1"1 mixture of benzene and methanol as depicted in Figure 26.9 (note that more recently benzene has been replaced in many laboratories by the less carcinogenic toluene). Subsequently, the individual fractions were Parr bombed to produce BaSO4 for sulfur isotope analysis. Despite increasing 634S values with biodegradation, the different solvent extracts from a given sample were markedly uniform, the greatest deviations being found for asphaltene with 634S values ~ 1%o lower than those of total S (Monster, 1972). It would be interesting to ascertain whether the ~)34Svalues of individual S functional groups changed during biodegradation. 26.3.4.3 Natural gas Sulfur in natural gas occurs predominantly in the form of H2S (sour gas). Hydrogen sulfide may be generated by bacterial sulfate reduction (BSR), thermochemical


585

sulfate reduction (TSR) and desulfurization reactions. The highest concentrations of H2S (up to 95%) typically arise during TSR. Sulfur isotope analyses have proven useful in identifying and quantifying these processes (e.g. Krouse, 1977). Natural gas samples are collected from exploratory or recovery wells and refineries using stainless steel cylinders with high pressure valves on each end. Expensive cylinders are available with glass or polymer linings to reduce adsorption of H2S. However, if these linings break down, retention of H2S may be worse than if the interior is not lined. Field sampling of natural gas is frequently done by service companies and most stable isotope laboratories simply receive the filled cylinders. A reliable service company will heat and evacuate these cylinders between samplings. The practice of attaching the cylinder to the field or refinery piping and flushing with the gas is not recommended and highly dangerous. Some gas samples may have pressures of tens of atmospheres and up to 90% H2S. During field sampling, an evacuated cylinder should be attached to the "plumbing". Appropriate valves are opened and left open for a few minutes to ensure that the cylinder contents have the same physical and chemical properties (e.g. temperature) as the system. Because H2S strongly adsorbs to surfaces, two sets of sampling cylinders are desirable for gases with low and high H2S contents respectively. If cylinders arrive in a sealed box, it should be opened in a fume hood in case that H2S is present from leakage. In the laboratory, the natural gas is slowly passed through a cadmium acetate solution, where H2S is quantitatively converted to CdS. This precipitate is converted to Ag2S for subsequent sulfur isotope ratio mass spectrometry as described above. A suitable apparatus for sampling H2S from natural gas and converting it to Ag2S is shown in Volume II, Part 3, Figure 17.13 (Chapter 17.6). Other sulfur compounds present in smaller quantities in natural gas include mercaptan and alkane sulfides. To our knowledge, no sulfur isotope ratios have been reported in the literature for these compounds. In our experience, GC separation of Scontaining gases does not proceed satisfactorily using Poropak Q packed columns. H2S desorbs very slowly producing memory problems from sample to sample and poor reproducibility of 634S values. 26.3.5 Sulfur in peat and soils Sulfur in peat and soils occurs in both organic and inorganic forms. Organic S constitutes often more than 80 % of total S. Organic S is commonly subdivided in two major groups: (1) organic (or ester) sulfates, which constitute a very labile fraction of the organic S pool (Freney, 1986), and (2) carbon-bonded sulfur, which includes S-containing amino acids (e.g. cystine, cysteine, methionine) as well as sulfonates (Biederbeck, 1978). Inorganic S may occur in both oxidized and reduced forms. Inorganic sulfate may exist in water soluble, adsorbed, or insoluble forms (Bohn et al., 1986). In water-logged soils, various reduced inorganic S compounds including elemental sulfur, monosulfide and disulfide minerals may also be present (see section 26.3.3.2). Total S contents in peat and soil vary from more than 0.1% to less than 0.01% depending on peat type or soil horizon.

586


Proper sample handling is essential for obtaining meaningful isotope ratios for peat and soil sulfur compounds. If reduced inorganic S compounds are expected, samples should be taken and transported under N2 or Ar atmosphere to prevent their oxidation. Extraction of these sulfur compounds should occur immediately after return to the laboratory while minimizing the exposure of the field-moist samples to oxygen. In aerated soils, contents of reduced inorganic S compounds are typically negligible. In that case, samples can be transported in plastic or paper bags, while temperatures should be kept < 4~ Upon return to the laboratory, soil material from aerated horizons should be dried at temperatures below 40~ Drying at higher temperatures may result in conversions or loss of some S compounds (Amaral et al., 1989; David et al., 1989; Wieder et al., 1985). Dried soil samples should be sieved (< 2 mm) to remove rocks and roots and ground where appropriate to provide a homogeneous representative sample. In principle, all techniques described in section 26.3.3 are also applicable to isotope analyses on peat and soil S compounds. However, the special characteristics of these materials including their high organic S contents, significant contents of organic sulfates, and occasionally large amounts of adsorbed sulfate, require special considerations in selecting the most appropriate extraction procedures.

26.3.5.1 Total sulfur Several studies have determined isotope ratios of total sulfur in peat and soils (e.g. Bottrell & Novak, 1997; Kusakabe et al., 1976). A variety of techniques are suitable to convert total S to either Ag2S or BaSO4. For example, Ag2S can be generated by converting all S compounds to sulfate via alkaline oxidation (Tabatabai & Bremner, 1970) followed by Johnson-Nishita reduction (Johnson & Nishita, 1952) in a distillation apparatus (see Part 3, Volume-II, Figure 8.1-13) as described by Schoenau & Bettany (1988). Complete conversion particularly during the reduction step is essential to avoid isotope fractionation and thus erroneous results. An alternate method is to convert all peat or soil S compounds to sulfate followed by precipitation of BaSO4. Alkaline oxidation (Tabatabai & Bremner, 1970) results in high DOC contents in the sulfatecontaining solution making precipitation of pure BaSO4 challenging. Therefore, wet chemical oxidation using HNO3/Br2 or other mixtures (Krouse & Tabatabai, 1986; Zhabina & Volkov, 1978) is often preferred (see section 26.3.3.1). Another widely accepted procedure is fusion with sodium carbonate and an oxidizing agent (Tabatabai, 1992). Whereas Kusakabe et al. (1976) relied on a 10:1 mixture of NaHCO3 and Ag20, the commercially available Eschka mixture (see section 26.3.3.1) has been more widely used. All these techniques generate SO4 2-, which is subsequently precipitated as BaSO4. The precipitate is subsequently filtered, weighed, and converted to SO2 or SF6 as described in section 26.2. Although many studies have evaluated isotope ratios of total sulfur in peat and soils (e.g. Kusakabe et al., 1976), more insight can be gained by determining the 634S values of individual S compounds (Chae & Krouse, 1986; Mayer et al., 1995; Schoenau & Bettany, 1989). Selection of the appropriate analytical extraction scheme for peat and soil samples depends upon whether reduced inorganic S compounds are present

Procedures for Sulfur Isotope AbundanceStudies

587

or not. For the former sample type, sequential extraction procedures are often preferred (see section 26.3.5.2). For aerated soils with negligible contents of reduced inorganic sulfur, extraction of individual sulfur compounds from fresh sample aliquots is often more satisfactory (see section 26.3.5.3). Also, a variety of combinations between sequential and individual extractions from fresh sample aliquots are reported in the literature (e.g. Mandernack et al., 2000; Morgan & Mandernack, 1996; Novak et al., 2003a,b; Wieder & Lang, 1988), too numerous to be fully described in this chapter. In the following, we provide a general outline of some of the available analytical options in anticipation that this provides the reader with sufficient information to design a custom-made extraction scheme, which is best suited for the respective sample material of interest.

26.3.5.2 Sequential extraction of individual sulfur compounds Peat and soil samples containing reduced inorganic S compounds should be immediately processed after return to the laboratory. This can be accomplished by using a sequential extraction procedure, which recovers reduced inorganic sulfur compounds prior to further processing of the sample material (Figure 26.10). Fieldmoist samples are transferred to the distillation apparatus shown in Part 3, Volume-IL Figure 8.1-13, while minimizing exposure to atmospheric 02. One option is to extract

Figure 26.10 - Flow chart summarizing analytical procedures for sequential extraction of various sulfur compounds from peat and soil samples (dw = distilled water).

588


the sum of all monosulfides, elemental S, and disulfides by chromium reduction (Canfield et al., 1986) as described in section 26.3.3.2 and shown in Figure 26.10. Alternately, these sulfur compounds can also be extracted individually as described in section 26.3.3.2. In both cases it is difficult to avoid that some organic S compounds are also converted to H2S (Amaral et al., 1993; Brown, 1986). Rinsing the residual sample with hot distilled water yields the fraction of acid soluble sulfate. It is noteworthy that this fraction may contain sulfate generated by hydrolysis of organic sulfates during the prior acid treatment. The residual sample material can be either treated with further sequential extraction steps as shown in Figure 26.10, or as described in section 26.3.5.3. For further sequential extractions it is essential that all inorganic sulfate, including the adsorbed fraction, is completely recovered to avoid erroneous carryover into subsequent fractions. It is therefore recommended to subject the residual sample to an additional sulfate extraction with 16 mM KH2PO4 (see section 26.3.5.3). The remaining sample material after this extraction step contains only organic sulfur compounds. Determination of the 634S values of the total organic sulfur fraction (organic sulfates and carbon-bonded S) can be achieved by processing the remaining sample material with any of the techniques described in section 26.3.5.1. Alternately, organic sulfates, which may constitute between less than 10 to more than 80 % of the organic soil S fraction (e.g. Bettany et al., 1979), can be extracted via reduction with a mixture of hydriodic acid (HI), formic acid (HCOOH), and hypophosphoric acid (H3PO2) in a ratio of 4::2:1 in a distillation apparatus (Part 3, Volume-II, Figure 8-1.13 ). This process is often referred to as Johnson-Nishita reduction (Johnson & Nishita, 1952). The Johnson-Nishita (or HI) reduction mixture converts both inorganic and organic sulfate to H2S (Freney, 1961), a fraction which is often referred to as total sulphate or HIreducible S (SHI-red). Since in the sequential extraction scheme inorganic sulfate has been quantitatively removed from the peat or soil sample (Figure 26.10), H2S generated via HI reduction is solely released from organic sulfates and its 634S value can be directly determined (Spratt & Morgan, 1990; Wieder et al., 1985). The sample residue remaining after quantitative extraction of inorganic S (reduced inorganic S and sulfate) and HI reduction contains only organic sulfur in carbon-bonded form. To determine its 634S value, the remaining sample material can be subject to any total S treatment described in section 26.3.5.1 yielding BaSO4 or Ag2S for sulfur isotope ratio measurement (Figure 26.10). Further subdivision of the carbon-bonded S fraction has been seldom attempted for isotope analysis. Reduction with Raney-Nickel alloy supposedly differentiates between S-containing amino acids, which are converted to H2S, and other carbon-bonded non-reducible organic S compounds (DeLong & Lowe, 1961; Lowe, 1965). However, it has been found that the amount of Raney-Nickel alloy affects the results and that Fe and Mn may interfere with the determination of the carbon-bonded S fractions (Freney et al., 1970). Therefore, this technique has never gained widespread popularity. Schoenau & Bettany (1988) extracted humic and fulvic acid fractions from soils using HC1 and NaOH solutions and converted S in the extracts by alkaline oxidation followed by Johnson-Nishita reduction (Tabatabai &Bremner, 1970) to Ag2S for subsequent isotope analysis.


589

Figure 26.11 - Flow chart summarizing analytical procedures for extracting individual sulfur compounds from fresh aliquots of peat and soil samples.

590

Chapter 26- B. Mayer & H.R. Krouse

26.3.5.3 Extraction of individual sulfur compounds using fresh sample aliquots Extraction of individual sulfur compounds from fresh aliquots of the same sample is an attractive alternative for peat and soil samples, particularly if they have negligible contents of reduced inorganic S compounds. This approach has the advantage that incomplete recovery of inorganic sulfate does not cause erroneous results for subsequently extracted sulfur compounds as is the case during sequential extractions (see section 26.3.5.2). Also, the determination of organic sulfur fractions is presumably more accurate. Some options for extracting sulfur compounds from peat and soil materials are summarized in Figure 26.11. Field-moist samples can be pressure-filtered (e.g. Wieder & Lang, 1988) and, after filtration of the pore water, dissolved sulfate is precipitated as BaSO4 as described in previous sections. Reduced inorganic sulfur compounds including elemental S (Maynard & Addison, 1985), mono-sulfides, and disulfides can be extracted together or individually as described in section 26.3.3.2 on fresh aliquots of compressed peat or soil yielding Ag2S for mass spectrometric determinations. Complete extraction of inorganic sulfate in aerated soils without mobilizing organic S compounds is a major challenge. Dry and sieved samples are typically mixed with an extraction solution and shaken in an overhead shaker for 18 hours (~10 rpm) at a solution to sample ratio of 5:1 by weight for mineral soils and 10"1 for litter horizons. Subsequently, the extraction solution is separated from the soil material by centrifugation and filtration with 0.45 ~m membranes, and sulfate is converted to BaSO4 by adding BaC12 solution. Contents of inorganic sulfate-S in many peat and soil samples vary between 10 and 100 mg kg-1, and therefore up to 100 g of soil must be extracted to yield sufficient quantities of BaSO4 for isotope analysis using dual inlet IRMS. The choice of extraction solution is dependent on soil horizon and soil properties. Different concentrations of phosphate, carbonate, chloride, and acetate solutions, as well as deionized water have been recommended in the literature for concentration analysis of soil sulfate (Tabatabai, 1982), but not all of these are suitable for isotope measurements. An important consideration is that conversion of organic soil S to sulfate should be minimized during the extraction. Since 634S values of organic soil S compounds may differ significantly from those of inorganic sulfate, conversion of even small parts of the often large organic S pools may result in e r r o n e o u s ~34S values for inorganic sulfate. For litter horizons, extraction with distilled water has proven satisfactory. For mineral soils, extraction with distilled water often recovers only a small part of the inorganic sulfate pool. This is because a large fraction of inorganic sulfate may be adsorbed to clay minerals and Fe and A1 oxides or hydroxides, particularly in acid soils. Therefore, slightly alkaline extraction solutions with high charge densities to release sulfate from the sorption sites are preferred for mineral soils. It is widely believed that 16 mM KH2PO4 or NaH2PO4 are the most suitable solutions for quantitative extraction of sulfate from mineral soils (Ensminger, 1954), while minimizing mobilization of organic S. However, one problem with these solutions is that addition of BaC12 results in co-precipitation of BaSO4 with large quantities of Ba3(PO4)2. Hence, the precipitate must be further treated with either Kiba, Thode, or JohnsonNishita reduction to convert sulfate-S into H2S and subsequently Ag2S (Johnson &


591

Nishita, 1952; Kiba et al., 1955; Thode et al., 1961) as described in section 26.3.3. Extraction of peat and soil samples with NaHCO3 (Kilmer & Nearpass, 1960) is not recommended, because the high pH value of this solution may facilitate the conversion of organic S to sulfate. Since there is no significant isotope fractionation during sulfate adsorption and desorption (Van Stempvoort et al., 1990), incomplete recovery of soil sulfate does not prevent the determination of ~34S values representative for the entire inorganic sulfate pool in soils. In our experience, extraction with 0.1 M LiC1 solution yields very reliable 634S values for inorganic soil sulfate, although it is known that this solution only partially removes adsorbed sulfate because of the low charge density of C1- (Krouse et al., 1996). An advantage of this extraction solution is that the Li ion may act as a metabolic inhibitor, minimizing the microbial conversion of organic S to sulfate during the extraction procedure. In carbonate containing soils, sulfate co-precipitated with calcium carbonate can be released by extraction with 1M HC1 (Roberts & Bettany, 1985), but conversion of some organic S to sulfate is difficult to avoid (e.g. Amaral et al., 1993). Reduction with a mixture of hydriodic acid (HI), formic acid (HCOOH), and hypophosphoric acid (H3PO2) in a ratio of 4:2:1 in a distillation apparatus (Part 3, Volume-IL Figure 8.1-13) reduces both inorganic and organic sulfate to H2S (Freney, 1961). The H2S from this so-called Johnson-Nishita reduction is converted to Ag2S. This precipitate represents total sulfate, comprised of both the organic sulfate and inorganic sulfate fractions. Knowledge of the sample masses, which have undergone extraction, and precise weighing of the obtained Ag2S or BaSO4 precipitates provides an approximate estimate of the contents of the individual sulfur fractions. There are also alternate chemical techniques for precise determination of the contents of many of the previously described sulfur compounds (e.g. Tabatabai, 1992). If the concentrations (C) and isotope ratios (634S) of all sulfur compounds shown in Figure 26.11 are precisely determined, it is possible to calculate contents and ~34S values of sulfur fractions, which can not be directly extracted from fresh aliquots of peat and soil samples. For example, the content of organic sulfate is the difference between total (HI-reducible) sulfate and inorganic sulfate: Corg sulfate = CHI-red - Cinorg sulfate

[26.11]

value of organic sulfate can be subsequently determined by mass and isotope balances according to equation [26.12]"

T h e 634S

~34 S

_ CHI - red ~ ~34SHI _ red - C inorgsulfate " ~34S inorg sulfate org s u l f a t e C org sulfate

[26.12]

In case that concentration and ~34S values of total sulfur (see section 26.3.5.1) were accurately determined, it is also possible to calculate contents and isotope ratios of the organic sulfur and the carbon-bonded sulfur fractions. Organic sulfur is the difference

592


between contents of total S and all inorganic S fractions (including reduced species and inorganic sulfate). [26.13]

CSorganic - CStotal - CSinorg

In many aerated soils with negligible amounts of reduced inorganic S, inorganic S is equivalent to the inorganic sulfate fraction. The 634S value of the entire organic sulfur pool can be determined as follows: 34 9 ~)34S Cstotal 9 ~) S s t o t a l - Csinorganic Sinorganic (~34Sorganic = Corganic

[26o14]

Contents of carbon-bonded sulfur can be calculated as the difference between contents of total sulfur and HI-reducible S (total sulfate) and all reduced S species (SCRS): [26.15]

CScarbon-bonded = CStotal - CS-HIred - CS-CRS

In aerated soils with negligible amounts of reduced inorganic sulfur compounds, carbon-bonded S equals: [26.16]

CScarbon-bonded = CStotal - Cs-HIred

In this case, the ~)34Svalue of carbon-bonded S can be calculated as follows:

~)34Scarbon- bonded -

Cstotal

9

34

(~ S s t o t a l - C s - H I r e d

9

634Ss-HIred

[26.17]

Cscarbon- bonded

One disadvantage of this approach is that the uncertainties of the calculated ~34S values become relatively large, if the respective organic sulfur pools are small, as may be the case in mineral horizons of acid soils.

26.3.6 Sulfur in plants Sulfur contents in plants vary from less than 0.05 % to more than 0.2 %. Typically, most of the sulfur in plants is bound organically, e.g. in S containing amino acids of proteins, whereas a smaller fraction of the total S occurs as sulfate (Blanchar, 1986). Interestingly, higher levels of total S in plants seem to be associated with an accumulation of sulfate. In those cases, sulfate may become the dominant S compound in plants. It is widely believed that organic sulfates do not occur in significant quantities in plant tissues. Proper sample handling prior to analysis should include initial storage of freshly sampled plant tissue at < 4~ removal of dust and surface contaminants by washing with deionized water, and subsequent drying at a maximum temperature of 65~ for several days (Jones & Steyn, 1973). Thereafter, samples should be finely ground using


593

a screen size of 40 mesh or finer. It is recommended to re-dry the samples again after grinding at 65~ for 24 hours to remove residual moisture prior to storage in sealed containers.

26.3.6.1 Total sulfur So far, most investigations of the sulfur isotope composition of plants have been based on the determination of 634S values of total sulfur (e.g. Chukhrov et al., 1980; Takala et al., 1991; Trust & Fry, 1992; Wadleigh & Blake, 1999; Yun et al., in press). A variety of techniques are suitable to convert total plant S to either Ag2S or BaSO4 for subsequent isotope ratio mass spectrometry. For instance, Ag2S can be generated by converting all S compounds to sulfate via alkaline oxidation (Tabatabai & Bremner, 1970) followed by Johnson-Nishita reduction (Johnson & Nishita, 1952) as described by Schoenau & Bettany (1988). Alternate techniques comprise the Eschka method (ASTM, 1993), wet chemical oxidation using HNO3/Br2 or other mixtures (Krouse & Tabatabai, 1986; Zhabina & Volkov, 1978), and Parr bomb oxidation (Siegfriedt et al., 1951), which have all been described in earlier sections (26.3.3.1, 26.3.4, and 26.3.5.1). All these techniques generate SO4 2-, which is subsequently precipitated as BaSO4 by adding 0.25M BaC12 solution. The precipitate is subsequently filtered, weighed, and converted to SO2 or SF6 as described in section 26.2. Total S in plant material can also be converted to SO2 in an elemental analyzer followed by CF-IRMS. One problem is that most biological samples produce at least 50 times more water and CO2 than SO2 and the former should be completely removed prior to SO2 entering the mass spectrometer. Monaghan et al. (1999) obtained reliable sulfur isotope ratios for wheat plants with > 1 mg S g-1 dwt. using an on-line continuous flow system. To avoid tailing of the SO2 peak, these authors used removable quartz liners in the combustion tube, PTFE couplings to avoid metal surfaces, and a Nation drying tube instead of the usual magnesium perchlorate water trap. They also prevented CO2 from entering the ion source by using two GC columns. A maximum of 10 mg sample was used together with 20 mg vanadium pentoxide to aid the combustion/oxidation process. Using these modifications, Monaghan et al. (1999) were able to obtain reliable (~34S values for total S in wheat plant samples containing between 10 and 25 ~g S, but samples much below 1 mg S g-1 were not suitable for analysis with their technique. Yun et al. (in press) used a similar technique to determine 634S values of total S in lichen samples. These authors obtained excellent results by weighing up to 15 mg of sample equivalent to ~ 10 ~g S in ultra-light Sn capsules together with V205. However, we reiterate that it is not trivial to normalize the obtained data to the international V-CDT scale (see section 26.2).

26.3.6.2 Individual plant sulfur compounds Few studies have attempted to analyze sulfur isotope ratios for distinct sulfur compounds in plants (e.g. Schoenau & Bettany, 1989). Sulfate-S can be converted to H2S by Johnson-Nishita reduction of plant material without prior oxidation (Johnson & Ulrich, 1959). The generated H2S can be trapped as CdS or ZnS, which is subsequently converted to Ag2S for isotope analysis. Removal of plant sulfate with hot

594


water, CH3OOH, or 0.1 M HC1 solutions has also been attempted, but these extraction solutions do not always achieve sulfate yields comparable to those from JohnsonNishita reduction. Hence, we recommend the latter method as a reliable procedure for accurately determining the content and isotope ratio of plant sulfate. Work in our laboratory has revealed that 634S values of sulfate in pine needles can deviate by more than 3 %o from those of total sulfur. Using the residue of the Johnson-Nishita reduction for subsequent total S analysis is a suitable method to determine the isotopic composition of organic plant sulfur. Alternately, if the contents and ~534Svalues of both the total S and inorganic sulfate fractions are known for a sample, the ~)34Svalue of plant organic S can be calculated by mass and isotope balances according to equations [26.13] and [26.14]. Few investigators have attempted to extract specific organic S compounds from plant material. For example, allylisothiocyanate from mustard plants has been isolated from plant material for sulfur isotope analysis (Remaud et al., 1997c).

26.3.7 Sulfur in humans, animals, and other materials Sulfur isotope analyses can be extremely useful in food web studies, since they often allow the identification of food sources e.g. of marine versus terrestrial origin (e.g. MacAvoy et al., 1998, 2000; Petersen et al., 1986). Sulfur isotope ratio measurements are also increasingly used in tracing origins and migration of wildlife (Hobson, 1999). Many tissues, fluids, and minerals in biological specimens contain sufficient sulfur for isotopic measurements (e.g. Katzenberg & Krouse, 1989). So far, many studies have focused on the isotopic analysis of total sulfur in biological materials including but not limited to muscle (Hesslein et al., 1991; Kwak & Zedler, 1997), liver (Hesslein et al., 1993), blood and skin of fish (MacAvoy et al., 2001), muscle and liver tissues of rats (Hobson et al., 1999), eggs (Hobson et al., 1997), feathers and muscle tissue of birds (Kwak & Zedler, 1997), and hair (Krouse et al., 1987a). Many of the previously described extraction techniques for total sulfur are applicable, including the Eschka method (section 26.3.3.1), HNO3-Br2 oxidation (section 26.3.3.1), and Parr bombing (section 26.3.4). Isotope ratios of total sulfur in tissues can also be analyzed by CF-IRMS techniques (e.g. Hesslein et al., 1991). Recently, ion microprobes have been employed to determine 634S values of total sulfur in fish otoliths with external precisions ranging between I and 5 %o depending on the instrument used (Weber et al., 2002). The reader is referred to previous sections of this chapter for analytical details. Hair is attractive to study because the animal is not sacrificed, the S content is high (~ 4 % mainly in form of keratin), and it is resistive to degradation. Seemingly reliable data have been obtained with 12,000 year old specimens from Siberian mammoth remains and 3,000 year old human hair (Aufderheide et al., 1994). Feathers and claw/ finger nails have similar advantages. All these materials can be easily Parr bombed to generate BaSO4 or converted to H2S and subsequently Ag2S with Kiba reagent (Sasaki et al., 1979). It appears that a number of S-containing amino acids and other compounds in tissues and fluids are capable of conversion to H2S with Kiba reagent (Kiba et al., 1957; Ohashi, 1955).


595

Cystine is relatively insoluble. In some animals, kidney stones may form because of lack of the enzyme to convert cystine to more soluble cysteine. Total sulfur in these stones has been analyzed using the Parr bomb technique (Krouse et al., 1987a). The invacuo Kiba technique (section 26.3.3.2) was applied to phosphorous-containing minerals in kidney stones, bladder stones, and teeth (Krouse et al., 1987a; Krouse & Ueda, 1987). Sulfate in fluids such as blood and urine is difficult to isolate because S-containing organic matter is trapped in BaSO4 precipitates. In our laboratory, one technique used was to dilute the fluids and add BaC12 solution with vigorous stirring. The dried precipitate was further treated by heating at 500~ in an 02 stream. Sulfate in some body fluids is derived from oxidation of organic S. Therefore, dilution should be carried out with deoxygenated H20 and BaSO4 should be precipitated in a N2 atmosphere to minimize organic sulfur oxidation during sample processing. There are many reports in the literature describing chemical extraction of individual organic S compounds or compounds containing the same sulfur moiety (e.g. Mestres et al., 2000). In some cases, the yields are very low. However, if the molecules are large, isotope fractionation associated with the chemical procedures may be acceptably small. Interestingly, very few ~34S values have been obtained for such extracts. One example is methionine-bound S in milk casein (Pichlmayer et al., 1998).

26.4. Summary Traditionally, sulfur isotope ratio determinations comprised three steps" 1) extraction of S from the sample and conversion into BaSO4, Ag2S, or other pure S-containing compounds, 2) preparation of a measurement gas such as SO2 or SF6, and 3) isotope ratio mass spectrometry in dual inlet mode. In the days of off-line gas preparation and dual inlet mass spectrometry, sulfur isotope ratio determinations were cumbersome and labor intensive. Throughout the last 15 years, sulfur isotope abundance studies have benefited from the advent of new technologies. These have resulted in higher sample throughput and greatly reduced sample size requirements. For example, coupling of elemental analyzers to isotope ratio mass spectrometers in continuous flow mode (CF-IRMS) has automated the gas preparation step and amalgamated it with fully computer controlled isotope ratio determinations. Using pure inorganic sulfur compounds such as BaSO4 or Ag2S, it is possible to achieve a reproducibility of better than + 0.2 %o for sulfur isotope measurements within a given laboratory. However, these technical improvements have not eliminated the necessity of extracting the respective sulfur compounds from gaseous, aqueous, or solid samples. Reliable extraction of individual sulfur compounds from often complex sample matrixes remains the key for the successful use of sulfur isotope techniques in many case studies. Complete recovery of the sulfur compound of interest without mobilizing other sulfur fractions is often challenging. Some of the extraction procedures are operationally defined and many of them remain cumbersome and labor-intensive.

596


Repeated extraction of a specific sulfur compound from a given sample often reveals uncertainties in excess of +0.5 %o associated with the analytical procedure. Therefore, it is essential that researchers determine and report the overall reproducibility of (1) sulfur extraction, (2) gas preparation, and (3) mass spectrometric measurements in their publications. In many cases, the overall reproducibility of the entire analytical procedure may be 2 to 3 times larger than that of the mass spectrometer measurement. Fortunately, determinations of 634S values of specific sulfur compounds with an uncertainty of _+ 0.5 %o is more than sufficient for most case studies, since sulfur in nature is characterized by a wide range of sulfur isotope ratios. Unfortunately, there are also uncertainties regarding the values of international reference materials (see Table 26.1). Among different laboratories, discrepancies between measurements on reference materials with very high or very low 634S values have been detected, particularly if different measuring gases such as SO2 or SF6 are used. Therefore, it is essential that researchers include careful descriptions of methodologies used in their publications, and that they report ~34S values obtained for international reference materials to allow for a meaningful comparison of sulfur isotope data reported in the literature. There is no doubt that the number of stable sulfur isotope abundance studies will increase in the future as analytical instrumentation becomes more automated and sophisticated. This expansion will not be restricted to the traditional research themes in Geochemistry and Hydrology since there is a tremendous potential for new applications of sulfur isotope techniques in disciplines such as Ecology and Atmospheric Chemistry, among many others. We anticipate that the key for advancing sulfur isotope studies in the future will not primarily be the development of better instrumentation for sulfur isotope ratio measurements. Improving our ability to reliably extract specific sulfur compounds from complex sample matrixes is more critical for an improved understanding of the sulfur cycle in the atmosphere, biosphere, pedosphere, hydrosphere, and lithosphere. In particular, techniques for extracting different forms of organic sulfur for isotope ratio determinations are still in their infancy. Better chemical resolution of individual sulfur functional groups or advancements in the field of compound specific isotope analyses particularly of organic S compounds coupled with sulfur isotope ratio determinations appear highly desirable.

Handbook of Stable IsotopeAnalyticalTechniques, Volume 1 P.A. de Groot (Editor) 9 2004 ElsevierB.V. All fights reserved.

CHAPTER 27 Direct Measurement of the Content and Isotopic Composition of Sulfur in Black Shales by Means of Combustion-Isotope-Ratio-Monitoring Mass Spectrometry (C-irmMS) Michael E. Bfttcherl & Bernhard Schnetger2 Max-Planck-Institute for Marine Microbiology, Department of Biogeochemistry, Celsiusstr.1, D28359 Bremen, Germany 2 Carl-von-Ossietzky University, Institute for Chemistry and Biology of the Marine Environment (ICBM), P.O. Box 2503, D-26111Oldenburg, Germany e-mail: 1 [email protected], 2 [email protected] 1

Abstract

The content and sulfur isotopic composition of black shales (down to 0.1 wt.% S) were directly measured by means of combustion-isotope-ratio-monitoring mass spectrometry (C-irmMS), and the results are compared to the Kiba reagent method for sulfur isotope preparation and the coulometric method for determination of the concentration. The C-irmMS measurements were not disturbed by the common combustion of sulfur- and carbon-bearing compounds up to 13 wt.% C. The C-irmMS method was successfully applied to a set of different pure synthetic and natural sulfur-bearing compounds and natural shale geostandards. The results show a good agreement, indicating that C-irmMS is a powerful analytical tool both precise and fast in sample preparation, which needs only small amounts of sample material. 27.1 Introduction

The determination of the contents and concentrations and sulfur isotopic composition of sulfur species in modern and ancient sediments is of fundamental interest for the evaluation of biogeochemical reactions in the coupled sedimentary element cycles (e.g., Hartmann & Nielsen, 1969; Goldhaber & Kaplan, 1975; Chanton et al., 1987) and our understanding of the paleo-environment and the evolution of life (e.g., Schidlowski et al., 1983; Ohmoto, 1992; Strauss, 1997). Most analytical schemes developed for the separation of sulfur species from recent and ancient sediments for isotope analysis (e.g., Sasaki et al. 1979; Allen & Parkes, 1995) are based on the early work of Kiba et al. (1955) and Zhabina & Volkov (1978). Due to their different environments of formation, metal sulfides or sulfates are typically occurring as the main respective sulfur-bearing phases in these sediments and, therefore, a number of sulfur isotope studies were based on the analysis of total sulfur (e.g., Brumsack, 1980; Vet6 et al., 1994; Calvert et al., 1996; Bfttcher & Lepland, 2000). The classical off-line scheme for sample preparation is time consuming (e.g., Giesemann et al., 1994; Sasaki et al., 1979) and the

598

Chapter 27 - M.E. B6ttcher & B. Schnetger

preparation may cause some of the previously reported uncertainty in sulfur isotope determination. With the development of modern on-line techniques using a combination of an elemental analyzer directly coupled to a gas isotope mass spectrometer (Pichlmayer & Blochberger, 1988; Giesemann et al., 1994; B6ttcher et al., 1998a) most previous problems related to sulfur isotope measurements (e.g., isotope effects due to chemical conversion of sulfate to sulfide or memory effects in the mass spectrometer) were minimized. Combustion-isotope ratio-monitoring mass spectrometry (C-irmMS) was successfully applied to determine the sulfur isotopic composition of pure barium sulfate (e.g., Giesemann et al., 1994; B6ttcher et al., 1998a, b) and metal sulfides (B6ttcher et al., 1998c, 2001). In the present study we apply C-irmMS to determine the contents and stable sulfur isotopic composition of sulfur in Devonian black shales. The measurements are compared to those obtained with classical off-line preparation methods using the Kibareagent and a coulometric method for stable isotope and content analyses, respectively. The results show a good agreement, indicating that C-irmMS is a powerful analytical tool both precise and fast. 27.2 Materials and methods

34S/32S ratios were measured on-line by means of combustion isotope-ratio-monitoring mass spectrometry (C-irmMS) (Pichlmayer & Blochberger, 1988; Giesemann et al. 1994; B6ttcher et al., 1998a) using a Carlo Erba EA 1108 elemental analyzer connected to a Finnigan MAT 252 mass spectrometer via a Finnigan MAT Conflo II split interface as described by B6ttcher et al. (1998a). Sample amounts equivalent to 20 to 50/~g sulfur were wrapped together with reagent grade V205 as a catalyst in pure tin capsules. All samples were weighed with a high-precision micro balance (Sartorius MC1 Research RC 210P). Natural samples were carefully ground and homogenized prior to further analysis. Except for ZnS, the natural sulfides were drilled by a microdrilling device from polished sections. ZnS corresponds to the NBS-123 reference material, with a sulfur isotopic composition of +17.3 + 0.3%o vs. V-CDT (B6ttcher et al., 1997b). The Sn caps were combusted in a pulse of 02 (grade 4.6, Messer Griesheim) at 1100~ oven temperature, leading to a short increase of temperature in the reaction zone of about 1800~ (Giesemann et al., 1994; B6ttcher et al., 1998a). The liberated sample gas was transported in a continuous stream of He (5.0 grade; Messer Griesheim). Water was removed from the gas stream by a water trap filled with magnesium perchlorate, and SO2 was separated from other gas impurities by a chromatographic column (0.8 m length; PTFE tubing; HekaTech) at 80~ A split of the total gas stream was introduced into the gas mass spectrometer via a fused silica capillary using a Finnigan MAT Conflo II interface, and the ion currents of masses 66 and 64 in the sample gas were compared to the corresponding ion currents of external in-house standards which were combusted every 10 samples. Comparison was done via a commercial SO2 gas (3.8 grade; 666S1 close to CDT composition; Messer Griesheim) which was introduced via the Conflo II into the mass spectrometer. The ion currents were

1. with 666Sis meant the m/z peak 66, including both the sulfur and oxygen isotopes.

Direct Measurement of the Content and Isotopic Composition of Sulfur ...

599

Table 27.1 - Measurements of sulfur recoveries for synthetic and natural sulfur compounds by means of C-irmMS. Absolute range was between 20 and 200 ~g sulfur. Calibration substance to calculate the recovery was synthetic reagent grade sulfanilamide (Carlo Erba). Number of measurements in parenthesis. Composition

Recovery + S.D. (%)

Cometic

BaSO4

104.0 + 10.3 (41) 99.8 + 8.1 (2) 98.9 + 11.9 (16) 103.5 + 1.9 (2) 97.4 + 11.7 (9) 102.8 + 4.3 (3) 94.7 + 9.6 (6) 101.5 + 7.8 (10) 95.4 + 3.0 (6) 100.9 + 19.5 (6) 103.3 + 11.5 (2) 101.5 + 1.7 (2)

synthetic synthetic natural anhydrite natural langbeinite natural celestite synthetic synthetic natural sphalerite (NBS-123) natural galenite natural pyrite natural chalcopyrite natural elemental sulfur

(NH4)2SO4

CaSO4 K2Mg2(SO4)3 SrSO4 CdS Ag2S ZnS PbS FeS2 (CuFe)S2 S~

r e c o r d e d as a f u n c t i o n of time a n d i n t e g r a t e d for m a s s e s 64 a n d 66 u s i n g i n t e g r a t i o n time steps of 0.25 sec w i t h the F i n n i g a n MAT Isodat 5.2 software. The i n t e g r a t e d signal for m a s s 64 w a s c o m p a r e d to a calibration curve d e r i v e d from synthetic s t a n d a r d s w i t h k n o w n sulfur c o n t e n t (CdS, sulfanilamide, BaSO4) to o b t a i n the sulfur contents of black shale samples. Isotope ratios are g i v e n in the 6-notation v e r s u s the C a n y o n Diablo troilite (CDT) s t a n d a r d a c c o r d i n g to: ~34S[%o] - {(34S/32S)sample/(34S/ 3 2 S ) C D T - 1} 103. Replicate a n a l y s e s on p u r e sulfate or sulfide s a m p l e s g e n e r a l l y a g r e e d w i t h i n + 0.2%o. Synthetic b a r i u m sulfate a n d c a d m i u m sulfide s a m p l e s previously m e a s u r e d w i t h the off-line m e t h o d (see below) w e r e u s e d for calibration of the m a s s s p e c t r o m e t e r a g a i n s t the C D T - s t a n d a r d to m a k e b o t h black shale d a t a sets, o b t a i n e d by off-line a n d on-line p r e p a r a t i o n , c o m p a r a b l e . It s h o u l d be noted, h o w ever, that a n e w sulfur isotope s t a n d a r d (V-CDT; V i e n n a - C a n y o n Diablo troilite) h a s b e e n i n t r o d u c e d by the IAEA w h i c h is n o w r e c o m m e n d e d for i n t e r n a t i o n a l calibration of sulfur isotope ratios ( C o p l e n & Krouse, 1998). The a b s o l u t e 32S/34S ratio of VCDT is 22.6436 + 0.0020 (Ding et al., 2001). The isotopic c o m p o s i t i o n s of i n t e r n a t i o n a l rock reference materials g i v e n in Table 27.2 are, therefore, r e p o r t e d v e r s u s V-CDT. Table 27.2 -Comparison of the determination of total sulfur contents by means of C-irmMS with the coulometric method, and sulfur isotope ratios of total sulfur measured by C-irmMS. TW-TUC is an inhouse standard (schist), and SR-1 and Jet-Rock I are international oil shale standards. Number of measurements in parenthesis. Data are taken from B6ttcher et al. (1998b). Note: 634Svalues are given vs VCDT. Standard

S + S.D. (wt.%) Coulometry

S + S.D. (wt.%) C-irmMS

~)34S (%0) C-irmMS

TW-TUC SR-1, 23 Jet-Rock 1, 23

0.37 + 0.02 (32) 1.28 + 0.00 (2) 7.14 + 0.10 (2)

0.34 + 0.03 (2) 1.29 + 0.06 (2) 7.31 + 0.10 (2)

+1.2 + 0.2 (2) +0.9 + 0.1 (2) -17.5 + 0.4 (2)

600


Table 27.3 - Natural black shale samples from the Wismuth location (Germany): Comparison of the determination of the total sulfur contents by means of C-irmMS with the coulometric method, and of sulfur isotope ratios measured by means of C-irmMS with the off-line method after Kiba-reagent-preparation. 634S values are given vs. CDT. nd: not determined, na: not applicable. All C-irmMS and coulometric measurements were run in duplicate. Disseminated total sulfur consists mainly of pyrite sulfur. In the SA samples, sulfidic sulfur was removed by high temperature oxidation (1 h at 1000~ and the remaining sulfur is believed to mainly consist of barite sulfur. Total organic (TOC) and inorganic (TIC) carbon contents were measured by coulometry. Sample

6003-428-S 6003-430-S 6003-433-S 6003-438-S 6175-218-S 6175-249-S 6175-253-S 6175-383-S 6529-952-S 7512-590-S 7515-428-S 6003-430-SA 6003-433-SA 6003-438-SA #1-1 #1-2 #1-3 #1-4 #1-5 #2-1 #2-2 1 in

wt.%;

2

S1

S1

6348 2

Coul.3

MS 4

MS 4

6348 2 off-line

2.1 1.2 3.4 4.5 11.4 3.8 2.8 4.6 4.9 4.0 0.5 nd nd nd nd nd nd nd nd nd nd

2.2 1.2 2.9 3.8 12.5 3.4 2.7 4.4 4.9 3.5 0.5 0.2 0.1 0.1 nd nd nd nd nd nd nd

-18.6 -3.5 -12.2 -7.7 +6.1 -16.4 -9.0 -14.3 -7.7 -6.5 -36.2 -2.1 -3.4 +4.6 -1.7 +0.7 +1.4 +2.1 -0.4 +13.1 +10.8

-19.3 -3.4 -11.8 -8.1 +6.2 -16.9 -8.3 -15.2 -6.8 -5.9 -38.3 -3.2 -3.1 +4.3 nd nd nd nd nd nd nd

in %o; 3 Coul. = Coulometry;

4

TOC (wt.%)

TIC Comment (wt.%)

4.9 3.2 nd nd nd 5.5 12.5 nd 5.7 nd nd na na na na na na na na na na

1.3 7.0 nd nd nd 1.1 0.6 nd 1.0 nd nd na na na na na na na na na na

disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated total sulfur disseminated barite sulfur disseminated barite sulfur disseminated barite sulfur idiomorphic pyrite crystal idiomorphic pyrite crystal idiomorphic pyrite crystal idiomorphic pyrite crystal idiomorphic pyrite crystal massive band of pyrite massive band of pyrite

MS = C-irmMS.

For off-line preparation, black shale samples were prepared by the Kiba reagent method (Kiba et al., 1955; Sasaki et al., 1979) and the isotopic composition of SO2 was measured on a Finnigan MAT 251 mass spectrometer at the Geochemical Institute of GSttingen University. Replicates agreed within about + 0.5%o. Sulfur measurements were carried out by coulometry (Heinrichs & Herrmann, 1990) using a Str6hlein coulomat with a reproducibility of about 10% (Lange & Brumsack, 1977). 27.3 Results and discussion

In the present investigation we demonstrate the applicability of the elemental analyser- isotope mass-spectrometer connection (C-irmMS) to measure directly the contents and sulfur isotope ratios of sulfur in black shales on small samples without any chemical pre-treatment of the samples. The black shale samples contain total sulfur contents between 0.1 and 12% (Tables 27.2 and 27.3).

601


As a first part of the study the dependence of the combustion procedure from the chemical composition of pure sulfur-bearing compounds at a constant temperature of 1100~ (rising up to 1800~ during flash combustion) was investigated. This sample set was build of different anhydrous synthetic and natural sulfates, sulfides, disulfides, and elemental sulfur (Table 27.1). It was found that the signal for mass 64 generally varied linerarily with the sample amount combusted between equivalents of 10 and 200 gg sulfur. The recovery of sulfur as SO2 was always complete and independent from the chemical composition and the oxidation state of sulfur in the samples (Table 27.1). The average recovery was 101%, 99%, 102%, and 101% for sulfates, sulfides, disulfides and elemental sulfur, respectively. Most of the standard deviation observed in the recovery is believed to be due to the weighing procedure. In the second step, both coulometry and C-irmMS where applied to natural samples covering the range between 0.5 and 12 % total sulfur (Tables 27.2 and 27.3) and a good agreement between the two methods was found (Figure 27.1). Again, no influence of the chemical bonding environment of sulfur in the fossil sediments was 15

i

I

I

I

i

I

i

I

i

.

.F

~um

12 ,-" "'~

O

F" F"

9

."

9

F

9

0 6

3

,O

S (Coul.) - 0.92 S (C-irmMS) + 0.39 (r2 = 0.99; n = 14) I

0

3

I

I

6

I

I

i

9

I

12

i

15

S (wt%; C-irmMS) Figure 27.1 - Comparison of the determination of total sulfur of black shale samples by coulometry and C-irmMS (Tables 27.2 and 27.3). Dashed line indicates the 1: I relationship.

602


observed and there was no influence of the absolute sulfur content on the reproducibility and precision of the analytical methods applied. The sulfur isotopic composition of the sulfur fraction in the natural samples was found to range between-36 and +13%o (Table 27.3), and the replicate measurements generally agreed within 0.2%o for samples with sulfur contents exceeding 0.1 wt.% and within 0.3%0 for the two samples with low sulfur contents (Table 27.3). No influence of the formation of carbon dioxide due to the common combustion of sulfur compounds with organic and inorganic carbon (up to 13%; Table 27.3) from the natural samples is observed, indicating a sufficient separation of SO2 from CO2 on the chromatographic column of the elemental analyzer and no later interference in the ion source of the mass spectrometer. Varying primary diagenetic conditions during sediment formation and subsequent overprints are reflected especially by the shift to heavier sulfur isotope values in the idiomorphic grains and massive bands of pyrite when compared to disseminated sulfur. Coexisting sulfate (barite?) sulfur was gener-

10

'

I

'

I

'

I

'

I

'

,"

..."

O

r or,-~

a

-10

O

9

,,O

9

-20

r

-30 m

634S (off-line) - 1.04 634S (C-irmMS) + 0.22 (r 2 - 0.996; n - 14)

,..

,~ a'mdmOn

-40 -40

i -30

I -20

i

I -10

,

I 0

i 10

~)34S (~ Figure 27.2 - Comparison of the determination of the sulfur isotopic composition of total sulfur of black shale samples by the off-line and C-irmMS method (Table 27.3). Dashed line indicates the 1" 1 relationship.


603

ally isotopically heavier than sulfide sulfur as expected from known thermodynamic and genetic relationships (Ohmoto & Goldhaber, 1997). These results are in agreement with other geochemical signatures found in the different black shale samples (Schnetger, unpublished). The agreement between the off-line and the on-line sulfur isotope determinations can be regarded as very good (Figure 27.2) and independent from the sulfur concentrations in the samples used here (Table 27.3). This indicates that the CirmMS method which requires only small amounts of sample material is a precise and fast (with respect to sample preparation) analytical tool for sulfur isotope and content determination of sulfur in natural black shales with sulfur contents down to 0.1% (Table 27.3). Caution has to be taken in more recent sediment with high porosity, where the contribution of residual pore water sulfate may have to be taken into account. It should finally be noted, that for a detailed analysis of microbial reactions leading to an isotope fractionation between different sulfur bearing-species (acid volatile sulfides, pyrite, elemental sulfur, organic sulfur), especially in recent marine sediments or bacterial cultures, the chemical separation of the different sulfur phases may be necessary.

Acknowledgments

The authors wish to thank the German Science Foundation (DFG, Bonn) and Max Planck Society (Munich) for financial support, H. Avak (Finnigan MAT, Bremen) and A. Giesemann (FAL, Braunschweig) for stimulating discussions, and A. Giesemann and V. Reppke for the supply of sample material. J. Hoefs (University of G6ttingen) kindly allowed access to the Finnigan MAT 251. The constructive comments of reviewers M. O. Jedrysek and M. A. Tabatabai helped to improve the manuscript.

Handbook of Stable Isotope AnalyticalTechniques, Volume 1 P.A. de Groot (Editor) 9 2004 Elsevier B.V. All fights reserved.

CHAPTER 28 Summary of Methods for Determining the Stable Isotope Composition of Chlorine and Bromine in Natural Materials H. G. M. Eggenkamp Department of Geochemistry, Faculty of Earth Sciences, Utrecht University, P.O.Box 80021, 3508 TA, Utrecht, The Netherlands e-mail: [email protected]

Abstract

During the past 80 years many attempts have been made to measure natural variations in the stable isotope compositions of C1 and Br. These variations are quite small; variations in the C1 isotope compositions of natural samples were first measured in 1982, and natural variations in Br isotope compositions were first measured in 1997. In this chapter I describe several methods that have been applied over the past 80 years both to separate C1 and Br from natural materials and to measure their isotope ratios by mass spectrometry. It is hoped that this chapter increases interest in the stable isotope geochemistry of halogens. 28.1 Introduction

It was shown by Aston (1919) that C1 consists of two different isotopes with masses 35 and 37. In later years, the C1 isotope ratio was determined quite often (e.g. Curie 1921, Gleditsch & Sandahl 1922, Harkins & Stone 1925, von Kallman & Lasareff 1932, Nier & Hanson 1936, Graham et al. 1951, Shields et al. 1962), but because in nature the differences in C1 isotope ratios are small, no measurable variations were found. Variations in C1 isotope ratios were found in chemical experiments (e.g. Bartholomew et al. 1954, Klemm & Lund6n 1955, Lund6n & Herzog 1956, Herzog & Klemm 1958, Hill & Fry 1958, Howald 1960) and it was found that the diffusion coefficient of 35C1 w a s about 1.0012 to 1.0022 times that of 37C1 (Madorsky & Strauss 1948, Konstantinov & Bakulin 1965). After the development of a new mass spectrometer with double ion collectors (Nier et al. 1947, Nier 1947, 1955, McKinney et al. 1950) it was possible to measure the C1 isotope ratio variation with a precision of +1%o. Hoering & Parker (1961) measured 637C1 values of 81 samples. They found no significant variations from the standard, which was NBS C1 isotope reference standard NBS 105. Two samples of formation water had relatively large (although non significant) deviations from the standard (-0.7 and -0.8%o), but they were not considered to be significantly different relative to the analytical precision. Hoering & Parker (1961) also measured 3 samples of Chilean perchlorate. Although Urey (1947) had predicted that, if hydrogen chloride and perchlorate are in equilibrium the fractionation could be as much as

Methods for Determiningthe StableIsotopeCompositionof Chlorine and Bromine ...

605

92%o, Hoering & Parker (1961) did not find any difference in C1 isotope ratio between their perchlorate and chloride samples and concluded that the perchlorate samples were not formed in equilibrium with the chloride in these deposits. Recently Ader et al. (2001) reported new methods for determining C1 isotope compositions in chlorates and perchlorates, but they also found only small variations, hardly significantly different from the ocean water isotope composition. Morton & Catanzaro (1964) measured the C1 isotope composition of apatites and found no variations larger than their analytical precision of +1%o. Since the early 1980s it has been possible to measure C1 isotope ratio variations sufficient to resolve natural variations. As a result of foresight of Austin Long and his research group in Tucson, Arizona that it was realised that C1 isotopes could probably be measured at these precisions. Kaufmann (1984) published the first thesis in which measurable variation of C1 isotopes in natural materials was shown. The precision of these analyses was better than +0.24%o and improved in later years. In this period, the first results were presented at several congresses (Kaufmann et al. 1983, 1984a, Kaufmann & Long 1984, Campbell & Kaufmann 1984) and published (Kaufmann et al. 1984b). In the years that followed several additional studies on the geochemistry of the stable isotopes of C1 were presented by the Arizona group (Kaufmann et al. 1987, 1988, 1992, 1993 Kaufmann & Arn6rsson 1986, Kaufmann 1989, Desaulniers et al. 1986, Eastoe et al. 1989, Eastoe & Guilbert 1992, Gifford et al. 1985). It was also the group from Arizona that proposed to use the C1 isotope composition of ocean water as an international standard reference for chlorine isotope measurements. This could be justified as they showed that the C1 isotope composition of several ocean water samples did not vary outside the analytical error (Kaufmann, 1984). Considerable progress in C1 isotope geochemistry occurred after 1990. Long et al. (1993) published in detail the method which was developed in Arizona. The following year Eggenkamp (1994) published his thesis, discussing C1 isotope data from a large range of environments. In the same period it was shown that C1 isotope compositions can be measured by positive ion thermal ionisation mass spectrometry (Xiao et al. 1992; Magenheim et al. 1994). This was an important development, following unsuccesful studies attempting to measure C1 isotope compositions using negative ion thermal ionisation mass spectrometry (Vengosh et al. 1989 Gaudette 1990). Besides methods to measure stable C1 isotope compositions in inorganic samples, methods for measuring C1 isotopes in organic compounds were developed and published (Tanaka & Rye 1991; Van Warmerdam et al. 1995; Holt et al. 1997; Jendrezejewski et al. 1997). Only very few studies are known in which the (geo)chemistry of Br isotope variations is described. Although it has long been known already that Br has two stable isotopes (79Br and 81Br, Aston, 1920) no large natural isotope variations were expected due to the small relative mass difference between these two isotopes. Early studies by Cameron & Lippert (1955) showed no Br isotope variations beyond their analytical precision. In later years, however, fractionation due to diffusion was shown in molten

Chapter 28 - H.G.M. Eggenkamp

606

lead bromide (Cameron et al., 1956) and zinc bromide (Lund6n & Lodding, 1960). Willey & Taylor (1978) showed that is was possible to measure the Br isotope composition with a method comparable to the method for C1 isotopes, using bromomethane. Xiao et al. (1993) described a method to measure the Br isotope composition using positive ion thermal ionisation mass spectrometry, which is analogous to the method for measuring C1 isotopes. Eggenkamp & Coleman (2000) described a method to measure Br isotopes in natural samples, which includes a method to separate Br and C1 from samples which contain only (very) small Br concentrations. 28.2 Notation and standards

Chlorine has two stable isotopes, 35C1 and 37C1, with natural abundances of 75.771 and 24.229% respectively (Rosman & Taylor, 1998; abundances still from Shields et al., 1962!). No international standard for C1 isotope ratios has been defined yet officially, but in all recent studies Standard Mean Ocean Chloride (SMOC) is used. Kaufmann (1984) showed that no variations in the C1 isotope composition of ocean water from different locations and depths are found. All laboratories in the world use seawater as their standard. Chlorine isotope data are reported as 637C1 which is defined as: (37C1]

_ (37C1]

1~37C1 _ ~35C1)sample /35C1)standard x 1000 37C1] 35C1)

[28.1]

standard

Bromine also has two stable isotopes, 79Br and 81Br, with abundances of 50.686 and 49.314%, respectively (Rosman & Taylor, 1998; abundances still from Catanzaro et al., 1964!). No formal international standard for Br isotope ratios has been defined, neither have tests been done to check for variations of the Br isotope composition in ocean water. However, as the residence time of Br in the oceans is even larger than that of chlorine, it is assumed that oceanic Br does not show variations, and thus can be used as an isotopic standard reference material. This standard can be called Standard Mean Ocean Bromide (SMOB, Eggenkamp & Coleman, 2000). Bromine isotope data are reported as 681Br which is defined as:

(81BF/ _(81BF/

a81Br _

79Br)sample

79Br)standard

81Br/ 79Br)

standard

x 1000

[28.2]

Methods for Determining the Stable Isotope Composition of Chlorine and Bromine ...

607

28.3 Basic techniques for measuring chlorine isotopes 28.3.1 Chlorine isotope measurements by gas-source ratio mass spectrometry Many studies report attempt to measure C1 isotope compositions in gas source mass spectrometers. Many different gases have been applied in these studies. In this chapter attempts using hydrogen chloride, C12 gas and chloro methane will be described. In early studies, several other gasses were used, including carbonyl dichloride (COC12, phosgene; Aston, 1942), boron trichloride (BC13, Osberghaus, 1950), phosphorus trichloride (PC13, Kush et al., 1937), arsenic trichloride (AsC13, Kush et al., 1937), and antimony trichloride (SbC13, Kush et al., 1937). These studies will not be discussed here. 28.3.1.1 Hydrogen chloride used as m a s s s p e c t r o m e t e r g a s Hydrogen chloride was used in several C1 isotope studies (Nier & Hanson, 1936; Madorsky & Strauss, 1947; Johnston & Arnold, 1953; Hoering & Parker, 1961). Of these, Hoering & Parker (1961) presented the first major study of C1 isotopes. They measured the C1 isotope composition of 81 samples in all types of geological environments. They were able to measure the C1 isotope ratio of HC1 with a precision of 0.8%0. The reason for using HC1 gas was that it can be prepared quantitatively, and because it has a simple cracking pattern. A large disadvantage of HC1 is that it sticks to the walls of the vacuum system, and as a result it has a large memory effect. Hydrogen chloride has a relatively simple mass spectrum in a mass spectrometer, with four ion species formed (Table 28.1). Samples of chloride were precipitated as silver chloride (AgC1). For rock and mineral samples this was produced according to the method described by Kuroda & Sandell (1953). Samples were dissolved in ammonium hydroxide, to which magnesium metal was added in excess. The magnesium displaces the silver, and an ammonium chloride solution was formed from which solids were removed by filtration. The remaining solution was evaporated to dryness (in a vacuum oven). The solution from which the ammonium chloride was precipitated is basic, and care must be taken that it would not react with CO2 from the atmosphere, as this can not be separated from the hydrogen chloride, and as such it interferes during isotope analysis. The ammonium chloride residue was converted to HC1 by reaction with sulphuric acid under vacuum. The gaseous HC1 could then be trapped onto frozen phosphorus pentoxide to remove traces of water, and the dried HC1 was frozen into a sample container for introduction into the mass spectrometer. 28.3.1.2 Chlorine g a s used as m a s s spectromTable 28.1 - Cracking pattern for HC1 (Hoereter g a s Chlorine gas has been used only rarely as ing & Parker, 1961). mass spectrometer gas. Owing to its simple Species m/ z Relativeintensity cracking pattern, however, it could be useful. Unfortunately, it also produces large mem35C1+ 35 17.0 ory effects. H35C1+ 36 100 37C1+ 37 5.4 Chlorine gas has a relatively simple cracking H37C1 + 38 32.5 pattern, which should be approximately as is

608


shown in Table 28.2.

Table 2 8 . 2 - Estimated cracking patters for chlorine gas, based upon the abundances for the isotopes 35C1 and 37C1, and assuming that the total intensity for monoatomic species is 10% of the total intensity for diatomic species.

Bartholomew et al. (1954) used C12 as the mass spectrometer gas in their studies on the isotope effect in reactions of tert-butyl chloride (2-chloro-2-methyl-propane). Hydrogen Species m/z Relative intensity chloride was produced from silver chloride 35C1+ 35 13 in a way comparable to that of Hoering & 37C1+ 37 4 Parker (1961). However, the ammonium 35C12+ 70 100 chloride was heated with concentrated sul72 64 phuric acid, and the product HC1 was 35C137C1+ 37C12+ 74 10 trapped in a bubbler (by means of a stream of nitrogen gas) containing a little cold water. The aqueous hydrogen chloride was then oxidised to C12 gas by persulphate oxidation (Brown et al. 1953) in a stream of helium. The C12 gas was frozen in a cold trap, and the helium pumped away. The C12 then could be transferred to the mass spectrometer. 28.3.1.3 Chloro-methane used as mass spectrometer gas

The most commonly used gas for isotope ratio mass spectrometry of C1 is now chloro-methane (methyl chloride). Several methods have been developed to produce quantitative yields of this gas, which is a lot less reactive than HC1 or C12, and it gives no memory effects in inlet and vacuum systems of mass spectrometers. This gas is generally measured at m / z 50 and 52 at which positions the main peaks are found. It's cracking pattern is relativity complex with many peaks as both chlorine and carbon have a significant minor abundant isotope (Table 28.3). Chlororomethane can be produced from several chloride compounds. In the literature on chlorine stable isotopes three of these have been proposed: chloromethane produced from ammonium chloride Table 28.3 - Cracking pattern from chloromethane (Owen & Schaeffer, 1954; Herzog & D6rnenburg, 1958), chloromethane (Taylor & Grimsrud, 1969). produced from silver chloride (LangSpecies m/z Relative vad, 1954; Hill & Fry, 1962; Taylor & intensity Grimsrud, 1969; Kaufmann, 1984; Long et al., 1993; Eggenkamp, 1994), 35C1+ 35 2.8 and chloromethane produced from H35CI + 36 1.2 copper chloride (Holt et al., 1997). 37C1+ 37 1.0 H37C1+ C35C1+ CH35C1+ CH235C1+C37C1+ CH335 C1+ CH37C1 + 13CH335Cl+CH237C1 + CH337C1+ 13CH337C1+

38 47 48 49 50 51 52 53

0.4 7.7 3.2 9.6 100 3.4 31.4 0.5

Chloromethane produced from ammonium chloride Chloromethane can be produced from ammonium chloride, and this technique was used in studies from the 1950s (Owen & Schaeffer, 1954; Herzog & D6rnenburg, 1958). Silver


609

chloride was precipitated from a sample in dissolved form, which was redissolved in ammonia. Silver then was removed as the sulphide. After drying pure ammonium chloride was present. This could be reacted into chloromethane following a procedure described by Blatt (in Owen & Schaeffer, 1954). 240 ml of concentrated sulphuric acid was diluted with 40 ml of distilled water, to which was added 350 ml of methanol. The temperature was kept below 70~ at all times. Approximately 2 ml of this solution was added to 50 mg of ammonium chloride under vacuum. The following reactions can occur: (CH3)2SO4 +2NH4C1 ~ 2CH3C1 + (NH4)2SO4 CH3OH ~ (CH3)20 (Catalysed by sulphuric acid) 2NH4C1 + H2SO4 --~ 2HC1 + (NH4)2SO4

[28.3] [28.4] [28.5]

Reaction [28.4] can be minimised by keeping the temperature of the mixture below 50~ A weighed sample of ammonium chloride was placed into a bulb and evacuated. Approximately 2 ml of sulphuric acid and methanol were added through a stopcock. The mixture was warmed with an infrared lamp until the reaction ceased. The reaction tube was chilled with dry ice and the chloromethane distilled into an evacuated sample bulb. A potassium hydroxide trap removed the hydrogen chloride formed during this reaction. A disadvantage of this technique is that the yield will not be higher than approxiately 35%. Several tests were done to determine errors caused by this low yield and it was found that data were reliable within an experimental error of 2%0. However, as it is now known that the large majority of 637C1 data are within 2%o of SMOC, an error this large is not acceptable for most natural C1 isotope measurements.

Chloromethane producedfrom silver chloride In most modern studies chloromethane is produced from geological samples via silver chloride. The method was originally described by Langvad (1954), and improved by Hill & Fry (1962) and later by Taylor & Grimsrud (1969). Ultimately a very effective technique was published by Kaufmann (1984). The method described here is taken from Eggenkamp (1994), which is based upon the method from Kaufmann (1984). Long et al. (1993) published a comparable version, also based upon the one described by Kaufmann (1984). Their method was developed to measure C1 isotope compositions by accelerator mass spectrometry for 36C1 measurements, and was designed to remove traces of sulphur from the sample. As such that technique is more complex than the one presented by Eggenkamp (1994). Eggenkamp (1994) used the following three steps to produce unfractionated chloromethane of sufficient purity for isotope measurement: 1) precipitation of silver chloride, 2) reaction of silver chloride with iodomethane, 3) separation by gas chromatography.

610


The procedure to prepare the silver chloride depends slightly on the amount of chloride in the solution. The method aims at precipitating silver chloride from solutions of fixed C1- amount, fixed ionic strength and fixed pH. Kohnen (1988) found that the best results are obtained when the amount of silver chloride formed is about 100 gmole (or 14.3 mg AgC1, corresponding to 3.5 mg C1-). This was confirmed by later studies. Therefore the amount of chloride solution needed is: 3000 p p m chloride

= ml necessary

[28.6]

However, if the amount of sample available is limited, with amounts down to 20 gmole reliable measurements can be made. It is recommended however that samples within a mass spectrometric run all contain comparable amounts of chloromethane. If the amount of solution is less than 10 ml, the following standard procedure is used" 4 ml of a 1 M KNO3 solution and 2 ml of a Na2HPO4-citric acid buffer solution are added to the chloride solution. The purpose of the KNO3 solution is to reach a high ionic strength. Taylor & Grimsrud (1969) found that using a less than 0.4 M KNO3 solution leads to very low chloromethane yields; for instance a 0.2 M KNO3 solution gives only 45% yield. The reason for this effect probably is that smaller crystals form at a high ionic strength. These small crystals can react completely whereas larger crystals form a coating of silver iodide that prevents the inner part of the crystals from reacting. Incomplete reaction inevitably leads to fractionation; Taylor & Grimsrud (1969) found a fractionation of +0.43%o due to this effect. The Na2HPO4-citric acid buffer solution is used to buffer pH at 2.2. This is necessary to remove small amounts of sulphide which otherwise precipitate as Ag2S (Kaufmann 1984), and also to prevent precipitation of other silver salts such as phosphate and carbonate (Vogel 1951). We used a buffer solution after McIlvaine (1921) which contains 0.71 gr (0.004 mole) Na2HPO4.2H20 and 20.6 gr. (0.098 mole) HOC(CH2CO2H)2CO2H. H20 (citric acid) per litre. After adding the KNO3 solution and Na2HPO4-citric acid buffer, the mixture is placed on a boiling ring and heated to about 80~ Then I ml of a 0.2 M AgNO3 solution is added and AgC1 starts precipitating instantaneously. The solution is not stirred because the newly formed AgC1 will coagulate and it is difficult to remove from the stirrer. The suspension then is filtered over a Whatman| glass fibre filter, type GF/F with a retention of 0.7gm and a standardised filter speed of 6 ml/sec. During filtration the suspension is rinsed with a dilute nitric acid solution (1 ml concentrated HNO3 in 500 ml water). When the silver chloride precipitate is rinsed with pure water, it occasionally will become colloidal and pass through the filter. Therefore the rinsing solu-


611

tion must contain an electrolyte; nitric acid is chosen because it has no reaction with the precipitate and leaves no residue upon drying (Vogel 1951). After filtration, the filter with the precipitate is dried at 80~ overnight. Care must be taken to protect the silver chloride from exposure to light. Silver chloride decomposes under the influence of light according to the reaction: 2AgC1 ---, 2Ag + C12

[28.7]

Therefore the filter with silver chloride is covered with aluminium foil. The aluminium foil must not be in contact with the silver chloride otherwise the aluminium will reduce the silver chloride: 3AgC1 + A1 ~ A1C13 + 3Ag

[28.81

which may cause isotope fractionation. The filter is weighed before precipitation and after drying, so that the amount of silver chloride is known. Samples with a chloride content below 300 ppm are treated in a slightly different way because more sample solution is needed. For these samples, KNO3 and the pH buffer are added as dry chemicals, otherwise the amount of solution would become too large. Per 100 ml of sample solution 6.00 g (0.06 mole) KNO3, 2.06 g (0.0098 mole) citric acid and 0.07 g (0.0004 mole) Na2HPO4.2H20 are added. The reaction of AgC1 to CH3C1 takes place in evacuated Pyrex tubes sealed at both ends; the tubes are 8-10 cm long and have an inner diameter of 8 m m and an outer diameter of 12 mm. The filter with AgC1 is loaded in a tube sealed at one end, a capillary drawn at the other end, and the tube is evacuated to a pressure less than 2 x 10-1 mbar. The tube is then filled with nitrogen gas and sealed with a rubber stopper to prevent air coming in. In a fume-hood, 200 ~1 (3.21 mmole) of iodomethane (CH3I) is added. Back on the vacuum line the CH3I is frozen on the AgC1 with liquid nitrogen and the tube is p u m p e d to less than I x 10-1 mbar. The tube is then sealed at the site of the capillary. The sealed tube is placed in an oven at a temperature of 70 to 80~ for 48 hours so that the following reaction takes place: AgC1 + CH3I ~ AgI + CH3C1

[28.9]

This is an equilibrium reaction so the CH3I must be added well in excess to get good CH3C1 yields. If the reaction temperature is too high the CH3I will partly decompose: 2CH3I --+ C2H6 + I2

[28.10]

see Eastoe et al. (1989). When CH3I decomposes, the colourless liquid will become yellow to brown. Samples that have been overheated can give much less accurate 637C1 values. Decomposition of CH3I can also been detected in a background scan that is routinely made after the isotope measurement. Overheated samples then show

612


increased background peaks at m / z 29, 45 and 46. CH3C1 and CH3I are separated by gas chromatography on two 75-cm long, 1/4" OD SS (stainless steel) columns, filled with Porapak-Q 80-100 mesh. Because the columns are easily overloaded with the large amount of excess CH3I, and the CH3C1 must be very pure, the gases are separated in two successive runs. The carrier gas is helium, at a pressure of 3 atm and a gas flow of about 100 ml min-1 (this rate should be adjusted such that a good separation between the three gases is obtained). The column temperature is 140~ The CH3C1 peak is detected by a thermo conductivity detector using a Carle 100 Micro Detector Control. A schematic drawing of the set-up is shown in Figure 28.1. The procedure for processing a series of samples is as follows: the gas chromatograph is first back flushed to remove the excess CH3I from the previous separation, to minimise contamination of the column and detector (3-way valve 3 set to position A, and 3-way valve 2 set to position B, while the openclose valve is closed). The borosilicate glass reaction tube is scratched by a glass cutting knife, and placed in the tube cracker. The tube cracker is evacuated (3-way valve 1 to position A) and liquid N2 is placed around the first coldtrap. 3-way valves 1 and 2 are closed and the reaction tube is broken. At the same time 3-way valve 3 is turned to position B and the open-close valve is opened. After 30 seconds 3-way valves 1 and 2 are turned to position B and the open-close valve is closed. After 3 minutes the liquid N2 around trap 1 is replaced by warm water and liquid N2 is now placed around trap 2. The recorder is started and a CH3C1 peak will be seen after about 2 minutes. As soon as the recorder signal has returned to the base line, and before the CH3I peak would be detected (after about 6 more minutes), 3-way valves 1 and 2 are turned to position A, the open-close valve is Figure 28.1 - Schematic drawing of the gas chromatograph opened and the liquid N2 around (see description in text).


613

the second cold-trap is replaced by warm water. Liquid N2 is now placed around the sample vessel. Valve 5 is closed and just before the expected arrival of CH3C1, 3-way valve 4 is turned to position B. After the pressure in the sample vessel has increased to approximately 3 psi above atmospheric level, valve 5 is turned to position A so that the over-pressure of helium is released. The CH3C1 is trapped in the sample vessel. When all the CH3C1 has been trapped, as indicated by GC, valve 5 is closed and the other valves must be returned to the starting configuration. Helium is pumped out of the sample vessel that should now contain pure CH3C1. The column is then backflushed, so that remaining CH3I is removed. The broken reaction tube can be replaced by the next one to be extracted. CH3CI producedfrom CuCI Holt et al. (1997) produced copper(I) chloride in the process of reducing chlorinated organic material (as summarised below). This copper(I) chloride is reacted in vacuum with iodomethane at 300~ for two hours. This procedure leads to ~100% yields for chloromethane. Holt et al (1997) separated chloromethane from iodomethane cryogenically.

Cryogenic separation of chloromethanefrom iodomethane Holt et al. (1997) developed a method to separate the mixture of chloromethane and iodomethane cryogenically. The mixture of chloromethane and iodomethane, that sits in a scratched tube, is put into a vacuum line that is evacuated. The location of the tube is cooled with dry ice-acetone slush to -79~ The tube where the iodomethane is frozen too is cooled with n-pentane-liquid nitrogen slush (-130~ and the tube where the chloromethane is to be frozen to is cooled with liquid nitrogen (-196~ The tube containing chloromethane and iodomethane is cracked open. As the mixture is cooled with dry ice-acetone slush the chloromethane and iodomethane only evaporate slowly, thus enhancing the cryogenic separation. Ten minutes after breaking the tube the dry ice-acetone slush is removed, and after 10 minutes more the chloromethane is moved to a third cold trap, by moving the liquid nitrogen to this trap. The n-pentaneliquid nitrogen slush is moved to the former chloromethane trap to trap iodomethane that has slipped through the first trap. This whole transfer may take up to 15 minutes. After separation the chloromethane can be frozen into a sample vessel for measurement in the mass spectrometer, while the excess iodomethane can be frozen into a waste tube for subsequent disposal. This is an advantage above the method to separate the two gasses by gas chromatography, where the waste iodomethane is vented to the air.

Measuring chloro methane samples on a stable isotope ratio mass spectrometer Chloromethane can be measured on all normal isotope ratio mass spectrometers. 637C1 is determined from the beams of mass 52 (CH337C1 +) in collector 3 and mass 50 (CH335C1+) in collector 1. The isotope ratio of C1 is much higher than for the light elements for which these mass spectrometers were built. Thus, beam 52 will be off scale at small working pressures. Working with very low pressures gives isotope fractionation in the inlet system. For this reason the ion source should be made less sensitive. This is done by reducing the trap current to 100~A. In this case the minor beam is

614


brought to a value smaller than 10-10A while still maintaining sufficient gas pressure in the inlet system. At these conditions, the results are highly reproducible.

28.3.2 Cl isotopes measurements by P-TIMS Apart from methods to determine C1 isotope ratios using gas-source isotope ratio mass spectrometry they are also being measured successfully by positive ion thermal ionisation mass spectrometry. Xiao & Zhang (1992) first described a method to measure C1 isotope ratios using this technique, followed by Magenheim et al. (1994) who also described several techniques to extract C1 from a variety of geological materials for isotope measurements. Chlorine is used in the form of HC1. Xiao & Zhang (1992) produced this from seawater by cation exchange chromatography, where resin in the Ba2+ form was used. The hydrogen chloride solution was then diluted to a concentration of 3 mg C1/ml. Tantalum filaments were treated with with 3 ~1 graphite slurry (100 ~g graphite, 80 vol% ethanol/20 vol% water), so that it was completely coated. This is almost dried completely, and the sample, I ml of the HC1 solution, neutralised with cesium carbonate is added. This is then dried for two minutes using a current of 1.1 A passing through the filament. After loading the treated filaments into the mass spectrometer isotope analysis begins as the pressure in the instrument is between 2 x 10-7 and 3 x 10-7 Torr (1 Torr = 1,33 mbar). Current on the filament is increased to 1.1 A in ten minutes. The Cs2C1+ ion current is monitored and used to focus the instrument. Its intensity is adjusted to 5 x 10 -12 to 8 x 10 -12 A by adjusting the filament current, which is typically 1.15 to 1.25 A. The data are collected by switching between the masses 301 (133Cs235C1 +) and 303 (133Cs237C1+). The baseline is determined at m / z 300.5. The data are acquired for 1.5 hours in high precision runs. Magenheim et al. (1994) applied a comparable method. They only removed ions such as fluoride and sulphate by ion exchange from the HC1 solution (to diminish interferences in the mass spectrometer), and the CsC1 was produced from the hydrogen chloride solution by ion exchange. This has the advantage that all acidity of the solution is removed.

28.3.3 Comparison of TIMS with IRMS Rosenbaum et al. (2000) compared positive ion thermal ionisation mass spectrometry with gas isotope ratio mass spectrometry and found that samples measured by both techniques gave comparable data. This indicated that data obtained and published from the two methods can be compared to each other.

28.3.4 Cl isotopes measurements by fast atom bombardment (FAB) mass spectrometry Westaway et al. (1998) proposed to measure the C1 isotope composition using a FAB-mass spectrometer. In this type of instrument silver chloride is mounted on a silver plate, heated and bombarded by Xe atoms. Negative C1 ions are then formed, and measured in Faraday cups at masses 35 and 37. The advantage of this method is that conversion of silver chloride to methyl chloride is not necessary. The analytical error

Methods for Determiningthe StableIsotopeCompositionof Chlorine and Bromine ...

615

of this method is 3 to 5 times larger than IRMS of methyl chloride, but within this error results are in good agreement.

28.4 Preparation techniques for different sample types All techniques described above require to get chloride in aqueous solution. Thus, preparation techniques are necessary to convert non-aqueous C1 (such as in organic compounds or rock samples) to aqueous chloride.

28.4.1 Preparation for organic samples Four different techniques have been proposed in the literature to prepare organochlorine compounds for chlorine stable isotope measurements.

28.4.1.1 LiCl-technique In the very first paper on the potential measurement of the C1 stable isotope composition of organochlorine compounds Tanaka & Rye (1991) proposed to react the organic compound with lithium metal. Samples were prepared by reacting the sample with lithium metal in a sealed quartz tube. The lithium chloride was dissolved in water and the chloride reacted to chloromethane according to the method described by Taylor & Grimsrud (1969). Replicate analyses of samples yield a precision of 0.15%o.

28.4.1.2 Parr Bomb technique Van Warmerdam et al. (1995) oxidised the organochlorine compounds in a socalled "Parr Bomb", according to the standard ASTM method D808-91. Five ml of a 5% CaCO3 solution is injected to the bottom of the bomb to absorb combustion products. The sample capsule is then sealed with cellophane tape to prevent loss of sample by volatilisation, after which sample (max. 50 ~1) is introduced by injecting through the tape using a gas tight syringe. The fuse wire loop is then brought in contact with the centre of the tape cover. The bomb is sealed and oxygen is added to a pressure of 20 atm, placed in an ice-water bath and the fuse is ignited. After the combustion, the bomb is depressurised and combustion products are removed using ultra pure distilled water. The yield of combustion products is only 65 to 75%, but no practical effects are reported on the C1 isotope ratio. The combustion products are then converted to chloromethane for isotope measurement by the method described in Long et al. (1993). The analytical reproducibility of this method is fairly low, with a standard deviation of +0.29%o based on 21 analyses of a 1,1,1,-trichloroethane sample. Standard deviations of 10 other chlorinated solvent samples range between +0.10 and +0.68%o.

28.4.1.3 Break seal technique Holt et al. (1997) proposed a method to produce both CO2 and chloromethane from the same chlorinated volatile organic sample. A borosilicate glass tube with a length of 20 cm is prepared with 4 contractions 2.5 cm apart starting from the open end. These tubes contain I g CuO wire, which is preheated for one hour at 550~ This tube is attached to a vacuum line and the sample, containing 10 to 70 ~mol C1, is frozen onto it. The tube is then sealed at the first constriction. Then the tube is heated for 2 hours at 550~ After this period the capsule is put with one end into a Watlow fur-

616


nace, and with the other end into a liquid nitrogen container. When heating the capsule to 750~ for 45 minutes, the formed CuC1 will evaporate and precipitate in the liquid nitrogen part of the tube. Hereafter the remaining CuO is collected at the end of the capsule where the copper chloride is not precipitated, and the capsule is scratched and put into a tube cracker. After evacuation the capsule is broken and the carbon dioxide is transferred to a sample vessel. Water remaining in the capsule is coldtrapped using a dry ice-acetone slush. The CO2 is subsequently stored and can later be measured for the C isotope ratio. Iodomethane is then frozen onto the remaining part of the capsule containing the copper chloride, evacuated and reacted to form chloromethane as described above.

28.4.1.4 Sealed tube technique Jendrezejewski et al. (1997) proposed to measure both the C and C1 isotope ratios from organochlorine compound using a sealed tube method. Copper oxide is added to a quartz tube, to which (via a septum) I to 3 ~1 of the sample is frozen. The tube is then evacuated and sealed. The capsule formed is heated in a furnace between 720 and 820~ for at least one hour. After this period the capsule is allowed to cool slowly. The capsule is then scratched, put in a tube cracker and after evacuation it is broken. The CO2 that was formed is then transferred cryogenically into a sample vessel. All solid residues (glass plus copper oxide and copper chloride) are recovered and transferred to a small glass beaker, with 15 ml high quality water. This is allowed to dissolve the copper chloride for 15 hours. After this period the solution is separated from the solids and AgC1 is precipitated and reacted with CH3I as described above to produce chloromethane for C1 isotope analysis.

28.4.2 Preparation for rock samples

A few techniques, summarized below, have been described in the literature to obtain a solution that contains enough chloride for isotopic analysis.

28.4.2.1 NaOH fusion technique

Eggenkamp (1994) described a method to extract chloride from silicate rocks based upon a technique which was formerly used by Behne (1953). X grams of powdered rock in which X is an amount of rock containing enough chloride to perform one or more ~37C1 measurements, is heated together with 10X grams of NaOH pellets for about 30 minutes in a nickel crucible. The rock dissolves in the molten NaOH and SiO bonds are partly destroyed. The temperature must not be too high since NaOH will evaporate at high temperatures (e.g vapour pressure is I mmHg (1.33 mbar) at 739~ and 10 mmHg (13.3 mbar) at 897~ (Stull, 1947)). After cooling, the sample is dissolved in 35X ml H20. This is achieved by putting the nickel crucible in a beaker with water on a magnetic stirrer. After two hours of stirring the content of the crucible is dissolved or suspended in the water. This suspension must be exposed to air for some time (overnight) to oxidize Fe2+ and other ions. Because the solution has an extremely high pH, which would cause precipitation of Ag20 after addition of Ag + ions, the solution must be acidified. This is

Methods for Determiningthe Stable Isotope Compositionof Chlorine and Bromine ...

617

done by adding 17.5X ml HNO3 65% to produce a colloidal "solution" of silica gel according to" SiO44- + 4HNO3 --+ H2SiO4 ~ + 4NO3-

[28.11]

This colloidal solution can not be filtered and, therefore, 3X ml HF 40% is added. The silica gel reacts with the HF to form a combined silica oxyfluoride. H4SiO4 + mHF --+ H4-mSiO4-mFm + mH20 m_ 5%o (Russ, 1987; Jarvis et al., 1992; Barnes, 1998; Taylor et al., 1998). Somewhat better results (2o of + 1 - 4%o) have been reported for ICP-TOFMS (Vanhaecke et al., 1999a). Isotope ratio data obtained by HR-ICP-MS have even better reproducibilities (Vanhaecke et al., 1996; 1997; Hamester et al., 1999). Such instruments have variable mass resolution. At a mass resolution of 300, reproducibilities (2 o) of as low as + 1%o and + 0.3%o have been achieved for the measurement of 7Li/6Li and 207Pb/206Pb, respectively (Hamester et al., 1999), approaching the precision of TIMS.

The analytical capabilities of such ICP-MS systems are limited by the following shortcomings. In Q-ICP-MS the mass analyzer is unable to transmit more than one mass at a time and HR-ICP-MS instruments have only a single collector. The analytical performance of these instruments is thus restricted by the sequential measurement of the unstable ion beams produced by plasma ionization. The precision of isotope ratio measurements is further limited by the peak shapes produced by the mass analyzer, which are rounded for Q-ICP-MS and triangular for HR-ICP-MS at higher mass resolutions of > 2000. Time of flight ICP-MS systems have very high repetition rates

694

Chapter 31 - M. Rehk~imper,F. Wombacher & J.K. Aggarwal

and thus nearly simultaneous ion detection capabilities. This should aid in the elimination of noise related to plasma flicker, but such systems have the r o u n d e d peak shapes characteristic of quadrupole mass analyzers.

31.2.2 The concept of MC-ICPMS MC-ICPMS instruments were designed specifically to overcome the limitations of other ICP-MS techniques for the m e a s u r e m e n t of isotope ratios. To achieve this goal, they combine an ICP-source with the magnetic sector analyzer and multiple-Faraday cup array of TIMS. The ICP source ensures excellent ionization efficiency, which is > 80% for most elements of the periodic table (Table 31.1). The magnetic sector analyzer (in conjunction with the slits of the ion-optical system) provide the flat-topped peaks necessary for high-precision isotope ratio measurements. The simultaneous collection of the ion beams with multiple Faraday cups cancels out the degradation of analytical precision due to plasma flicker. Interfacing a plasma source, which is at atmospheric pressure, and a magnetic sector mass spectrometer is not straightforward. First, this necessitates a powerful differential p u m p i n g system, to attain a v a c u u m of ~ 10-8 to 10-9 mbar in the analyzer. Such v a c u u m conditions are essential in order to obtain good peak shapes and an abundance sensitivity of less than 5 - 10 ppm. Second, the ions m u s t be accelerated into the mass spectrometer, by establishing a large potential difference (about 5 - 10 kV) between the ICP source and the detectors. Third, the broad circular ion b e a m produced by the ion source must be reshaped to fit the rectangular analyzer entrance slit at minimal loss of transmission. Fourth, the large energy spread of the ions p r o d u c e d in the plasma, generally ~ 20 V, needs to be reduced by about an order of m a g n i t u d e or, alternatively, energy focusing with an electrostatic analyzer m u s t be applied.

31.2.3 Commercial MC-ICPMS instrumentation The first commercial MC-ICPMS instrument built was the model Plasma 54 from VG Elemental (Figure 31.1). The Plasma 54 uses an ion source that is held at ~ 6 kV. A DC q u a d r u p o l e lens reshapes the circular ion beam and an electrostatic analyzer (ESA) serves to match the energy dispersion of the source to that of the magnetic secTable 31.1 - Ionization energies and degree of ionization for singly charged ions in the argon ICP. Ionization energy ( e V )

Degree of ionization

13

~ 60- 90% ~ 30- 75% ~ 15 - 40% ~ 5%

40 amu) Ar is generally used as the collision cell gas, because it provides efficient thermalization. For the measurement of light elements (e.g., Li, B) He is used as the main collision gas, because Ar significantly reduces the

696

Chapter 31 - M. Rehk/imper, F. Wombacher & J.K. Aggarwal

Figure 31.2 - Schematic diagram of the plasma interface and hexapole collision cell of the Micromass IsoProbe (after Rehk~imper et al., 2001b).

Figure 31.3 - Schematic diagram of the Nu Instruments Nu Plasma MC-ICPMS (after Rehk~imper et al., 2001b).

Stable Isotope Analysis by Multiple Collector ICP-MS

697

transmission of the light elements. The IsoProbe collector block can be equipped with up to 9 Faraday cups and 8 channeltrons for the simultaneous measurement of masses with ~ 15% mass difference. The Nu Plasma (Figure 31.3) utilizes a source that is held at ~ 4 kV. An ESA matches the energy dispersion of the ions produced in the plasma source with that of the magnetic sector. The Nu Plasma has a variable dispersion zoom lens system behind the magnetic sector. Because the collector coincidences for static measurements of multiple ion beams are adjusted by changing the dispersion of the zoom lens system only, moveable Faraday cups are no longer required. The instrument can be equipped with 12 Faraday buckets and three discrete dynode electron multipliers for the simultaneous measurement of ions with a mass difference of up to ~ 15%. The Axiom has an ICP-source at ~ 6 kV. The double focusing ion optical system has a Nier-Johnson geometry (the ions pass through the electrostatic sector prior to the magnet) andthe instrument can alternatively be used as a single collector HR-ICP-MS. The collector block is the Plasma 54 design, with nine moveable Faraday cups plus an optional fixed-position off-axis bucket that extends the mass range of the collector array up to ~ 20%, for static measurements of isotope ratios such as 7Li/6Li. The Axiom can be equipped with up to 3 discrete dynode multipliers. The Neptune plasma source is at ground potential whereas the mass analyzer and collector array are floated a t - 10 kV. The instrument provides double focusing with a Nier-Johnson geometry and the capabilities for high mass resolution with flat-topped peaks, albeit at reduced transmission. The collector block features 8 motorized detector carriers that can be equipped either with Faraday cups or channeltron ion counters and permits the measurement of masses with up to 17% mass difference. A dynamic zoom lens system is furthermore positioned between the magnetic sector and the detector block, to support analyses at high mass resolution and to enhance multidynamic measurements by applying variable zooming for the different sequences. At the time of writing, Nu Instruments is in the process of building a second MCICPMS system, the Nu Plasma 1700. This instrument will provide high mass resolution, flat-topped peaks and high sensitivity by using a dispersion much greater than in previous MC-ICPMS systems (Halliday et al., 2000b). High mass resolution and multi-collection is also possible with the Neptune, but only with an order of magnitude lower transmission (Hamester et al., 1999). The Nu Plasma 1700 is calculated to provide a peak flat of ~ 80 ppm at a mass resolution of > 2500 for the axial collector with a transmission similar to that of the conventional Nu Plasma instrument.

31.2.4 Instrument performance In order to obtain wide, flat-topped peaks (with a peak flat width of > 400ppm) for the collection of precise isotope ratio data, MC-ICPMS instruments are generally operated at a mass resolution of about 400 (M/AM, 10% valley definition). Under such conditions, all present MC-ICPMS instruments achieve an abundance sensitivity of < 5 - 10 ppm at mass 237 relative to 238U. Definitions of these (and other) technical terms

698

Chapter 31 - M. Rehk~imper, F. Wombacher & J.K. Aggarwal

can be found in Habfast (1997) and Rehk~imper et al. (2001b). The transmission of MC-ICPMS is high relative to Q-ICP-MS instruments. Sensitivities of > 5 - 10 x 10-11 A / p p m (> 3 - 6 x 108 cps/ppm) are generally achieved for elements such as Nd, HL Pb and U with conventional aspiration systems (this term is used in the following for the combination of a concentric nebulizer with a glass double-pass spray chamber). Assuming a solution uptake rate of 400 gl/min, this is equal to a transmission efficiency of > 0.002% for Pb. The transmission efficiency is defined here as the ratio of the number of ions registered by the collector to the total number of ions available for analysis. Using a low-flow nebulizer with membrane desolvation (Cetac models MCN 6000 or Aridus), sensitivity can be increased considerably, to > 5 x 10-10 A / p p m (> 3 x 109 cps/ppm). At a solution uptake rate of N 75 ~1/min, this translates to a transmission efficiency of > 0.05%. Such sensitivities permit precise isotopic analyses on samples sizes of < 30 - 50 ng. Sensitivities of > 20 x 10-10 A / p p m (> 10 x 109 cps/ppm) have been achieved for Pb with both the IsoProbe and the Nu Plasma using a MCN 6000, at a transmission efficiency of > 0.3% (Rehk~imper & Mezger, 2000; M. Rehk~imper, unpublished results). Low-mass elements, however, have significantly lower sensitivities. With the Nu Plasma, the sensitivity for Zr or Cd is typically about 50% lower than for T1 or Pb. Sensitivities of about 8.5 and 14 x 10 -11 A / p p m were reported for Mg and Ca, respectively, using a Nu Plasma with a desolvating nebulizer (Halicz et al., 1999; Galy et al., 2001). Tomascak et al. (1999a) obtained ~ 6.5 x 10-11A/ppm (4 x 108 cps/ppm) for Li with the Plasma 54 and the MCN 6000. Further reviews and discussions of MC-ICPMS instrumentation and general performance parameters can be found in a number of recent publications (Walder, 1997; Halliday et al., 1998; Halliday et al., 2000a; Rehk~imper et al., 2001b). 31.3 Instrumental mass discrimination and analytical protocols for stable isotope ratio measurements A significant feature of plasma source mass spectrometry is the large instrumental mass bias, which is related to the preferential extraction and transmission of the heavier ions. For the high-mass elements Nd, Hf, Pb, and U, the mass bias is about 0.5 - 1.5% / a m u (Figure 31.4), approximately an order of magnitude larger than the mass fractionation that is observed during TIMS measurements. For light elements such as B, the mass bias is much larger at > 15%/amu (Figure 31.4). The large instrumental mass discrimination of plasma source mass spectrometry is ascribed by many to "space-charge effects" in the plasma interface and the focusing lens region (Jarvis et al., 1992; Douglas & Tanner, 1998). While this model can explain many of the effects seen in Q-ICP-MS, it is presently unclear if space-charge effects are also the primary source of mass bias in MC-ICPMS (Mar6chal et al., 1999; Rehk/imper et al., 2001b). Regardless of the origin, the large mass bias associated with plasma ionization clearly necessitates that mass discrimination is carefully controlled during isotope ratio measurements, if precise and accurate analytical results are to be obtained.

699


100

-

I

I

n

@ -B -A@

'

I

'

I

r~ ~9

10

m

-

A

Ca

Mg A

~D Cd

A~

FeCU

[~

Plasma 54

A

Nu Plasma

Nd i

~A

_

SB

~

-

or3 r

c~

0.1

Hf

,

0

I

Sr Zr

0r3

c~

'

@ IsoProbe

m

I

50

,

I

100

,

I

150

,

T1 Pb ~

m

@: U-

I

200

Atomic Mass Figure 31.4 - Plot of mass bias per amu (in %) vs. atomic mass for various elements measured with different MC-ICPMS instruments. Most data were acquired by the authors and colleagues at Mfinster (IsoProbe) and Zfirich (Nu Plasma); other results are from references summarized in section 31.4.

31.3.1 Mass bias correction by internal normalization Measurements of radiogenic isotope compositions (e.g, 87Sr/86Sr, 143Nd/144Nd) are corrected for the effects of instrumental mass fractionation by normalization of the measured ratio to an invariant isotope ratio of the same element. This procedure, termed internal normalization, most commonly employs either a linear, power, or exponential law correction (Russell et al., 1978; Wasserburg et al., 1981; Hart & Zindler, 1989; Habfast, 1998). All three "laws" were originally derived empirically, for the correction of TIMS data for evaporation-induced fractionation. The lack of a causal law that is able to describe accurately the static mass discrimination effects of plasma source mass spectrometry has led to the use of the same laws in ICP-MS, and their application is detailed below. It is assumed that rA and rB are the values of two isotope ratios measured by mass spectrometry with nuclides of the masses M1 and M2 (for rA) and M3 and M4 (for rB), where the isotopes with the masses M2 and M4 are in the denominator. RA and RB are the "true" isotope ratios, and RB is known because it is not affected by radiogenic ingrowth. It is defined that AMA- M1- M2 and AMB - M3- M4. The linear law states that R A - rA[1 + (xAAMA] where r the mass bias per atomic mass unit, is obtained by:

[31.1]

700


[RB/rB]- 1 ~B =

[31.2]

AM B

The power law states that: R A - rA[1

-

(IA]AMA

[31.3]

with O~B -

[RB/rB] 1/(AMB)- 1

[31.4]

The exponential law differs from the linear and the power law, because the mass bias is assumed be a function of the absolute mass: RA -

[31.5]

rA[M~113A

The mass fractionation coefficient 13is obtained by: ln[RB/rB] [3B = ln[M3/M4 ]

[31.6]

All three "laws" thus assume that the fractionation observed for one isotope ratio (and quantified by the parameters aB, 13B),can be used to correct for the unknown fractionation of the other isotope ratio. This implies that OtA -- (1 B

and

~A-

f3B

[31.7]

for the linear/power and exponential laws, respectively. The choice among these laws will depend on which law best reproduces the true isotopic composition RA following correction for mass discrimination, regardless of the magnitude of c~. All three laws give similar results if ~ is sufficiently small. Detailed studies of the correction schemes for the internal normalization of TIMS data have shown, however, that the power law, and particularly the exponential law, are preferable for the correction of larger mass fractionation effects (Russell et al., 1978; Wasserburg et al., 1981; Hart & Zindler, 1989). Despite the fundamental difference in the physical origin of the mass fractionation in TIMS and ICP-MS, application of the exponential law for the internal normalization of isotopic data acquired by MC-ICPMS has been quite successful. Use of the exponential law is indicated by the fact that plasma ionization is associated with large mass fractionation effects compared to TIMS. Further support for this conclusion is derived from the observation that the mass bias a is clearly a function of mass in


701

plasma source mass spectrometry (Figure 31.4), whereas ~ has been shown to be approximately constant (at ~ 2) over a wide range of masses (Mar6chal et al., 1999). The application of the exponential law for the internal normalization of MC-ICPMS data was thoroughly investigated for a number of radiogenic isotope systems, and has been demonstrated to provide reliable results. Strontium and Nd isotope measurements conducted by MC-ICPMS display reproducibilities similar to TIMS analyses. For elements with high ionization potentials such as Zr, Hf or W, MC-ICPMS provides superior precision. The ability of MC-ICPMS to produce accurate isotope ratios for various elements (e.g., Sr, Nd, Hf) where precise TIMS reference values are available has also been established by a number of studies (Rehk/imper et al., 2001b). 31.3.2 Mass bias correction for stable isotope ratio measurements Internal normalization cannot be used in general for stable isotope ratio measurements, because this eliminates both instrumental and natural mass fractionations. If a double-spike is added to the samples prior to the isotopic measurement this "rule" does not appl~ however, and internal normalization can be used correct for instrumental mass discrimination. Double-spike procedures have been used in the past in conjunction with TIMS for the acquisition of stable isotope data for various elements (e.g., Ca, Fe and Cd) at high precision (Russell et al., 1978; Rosman & De Laeter, 1988; Johnson & Beard, 1999). Siebert et al. (2001) were the first to adopt a double-spike technique for stable isotope analyses by MC-ICPMS, to precisely measure natural Mo isotope fractionations. A particular advantage of double-spike methods is that they afford very precise and robust control of instrumental and laboratory-induced mass fractionation, if the spike is added prior to the chemical separation procedure. Double-spike procedures are, however, labor-intensive and complicated to establish and the target element must have at least four isotopes. A carefully calibrated spike and optimized spike-sample ratios must be used to minimize the propagation of analytical uncertainties. Currentl~ double-spiking is used routinely in only a few laboratories mainly for high-precision Pb and Fe isotopic measurements. With MC-ICPMS, the precise correction of only instrumental mass discrimination is also possible without double-spiking, using measurement techniques that are not applicable in TIMS. These techniques are based on two distinct characteristics of plasma ionization. First, the ICP source operates at steady-state. Thus mass fractionation is not primarily a time-dependent process, as it is in TIMS where the measured isotopic composition changes with time due to the progressive evaporation process. This is beneficial for the correction of instrumental mass bias by external standardization, where the isotope data obtained for a sample are referenced to the value obtained for an isotopic standard. Second, the mass discrimination associated with plasma ionization is, to a first order, a relatively simple function of mass, such that elements with similar or overlapping mass ranges display a nearly identical mass bias (Figure 31.4). Using a solution containing a mixture of two elements with similar masses, the mass discrimination observed for an element of known isotopic composition can be used to determine the unknown isotopic composition of the second element. This procedure of external nor-

702


malization was first suggested by Longerich et al. (1987) to improve the precision of Pb isotopic measurements by Q-ICP-MS (using T1 as the reference element) and has since been applied in a number of Pb isotope studies conducted by MC-ICPMS (Hirata, 1996; Belshaw et al., 1998; Rehk~imper & Halliday, 1998).

31.3.2.1 External standardization

For heavy elements such as U, where the mass bias is ~ 0.5%/amu, the mass discrimination generally varies by < 0.2%0/hr. For light elements below mass 40, where the mass discrimination is > 5 % / a m u (Figure 31.4), the drift of mass bias is significantly larger at up to ~5%o/hr (Tomascak et al., 1999a). Changes in mass discrimination with time thus appear to be more severe for the lighter elements, such that optimized application of external standardization will require different analytical protocols for different elements. At present, the most precise data have been collected by alternating standard and sample measurements, such that each sample is referenced only to the mean of the standards measured immediately before and afterwards. This technique of "standard-sample bracketing" (Figure 31.5) is similar to the standardization method used in gas-source isotope ratio mass spectrometry. For light elements, where the absolute drift in mass bias is particularly severe, the precision of sample measurements can be improved by performing multiple short analytical runs that are each bracketed by standard measurements. Switching between samples and standards can be very rapid, if long washout protocols are not required, and mass spectrometric runs of ~4 min or less have been used to maximize the precision of B and Mg isotopic measurements by MC-ICPMS (Aggarwal et al., 1999; Galy et al., 2001). Longer data acquisition periods are more applicable for heavier elements for which the drift in mass bias is less severe. 31.3.2.2 External normalization

External normalization can be used for stable isotope analyses because it corrects only for instrumental mass discrimination, whereas internal normalization also removes the effects of any natural mass dependent fractionation processes. The procedure involves application of the same mass fractionation "laws" (equations [31.1-6]) originally developed for TIMS and internal normalization. In the case of external normalization, however, RA, rA are ratios of the analyte element and RB, rB are isotope ratios of an admixed element of similar mass that is used for mass discrimination correction. External normalization thus assumes that the c~ or [3 values for the isotope ratios of the two elements are identical (Equation [31.7]). A number of studies have shown that a linear law is least suitable for the external normalization of MC-ICPMS data, but both the power and the exponential law have been used for this purpose with some success (Taylor et al., 1995; Hirata, 1996; Belshaw et al., 1998; Rehk~imper & Halliday, 1998). Mar6chal et al. (1999), however, noted that strict application of the exponential law was unable to provide a precise correction scheme for Cu and Zn stable isotope measurements (using Zn for external normalization and vice versa), such that an empirical procedure was adopted instead. If the mass discrimination of Cu and Zn were to show ideal exponential law behavior, the raw measured 65Cu/63Cu and 68Zn/64Zn ratios should show a linear relationship


703

in a plot of ln(65Cu/63Cu) vs. ln(68Zn/64Zn)(Mar6chal et al., 1999). From equations [31.5,6] it follows that the slope SE of this linear correlation is defined by ]3Cu ln(M65/M63) S E - ]3Zn x ln(M68/M64 )

[31.8]

where M65 represents the atomic weight of the isotope 65Cu, etc. Strict application of the exponential law (in a manner identical to internal normalization procedures) requires that 13Cu= [3Zn, such that ln(M65/M63) S E = ln(M68/M64 ) = 0.5156

[31.9]

Mar6chal et al. (1999) noted that their analytical data indeed defined linear correlations in plots of ln(65Cu/63Cu) vs. ln(68Zn/64Zn). The regression lines, however, varied from day to day with slopes that were always significantly different from the theoretical slope that would be expected for exponential mass fractionation (Figure 31.6a). Based u p o n this observation, it was concluded that the exponential law could not be used for the precise correction of mass discrimination, because ~Cu [3Zn. The observation that a well-defined correlation was obtained for each measurement session, however, showed that the ratio ~Cu/~Zn was constant and this permitted the application of an empirical correction procedure (Figure 31.6b). For each measurement session, multiFigure 31.5 - Schematic plot of 7Li/6Li vs. time that illustrates the appliple standard runs cation of the external standardization technique (standard-sample were performed to bracketing) for the correction of instrumental mass discrimination. The define the slope of plot is based on the Li isotope techniques of Tomascak et al. (1999a). The raw data obtained in each sample run are corrected relative to the mean the ln(65Cu / 63Cu) of the adjacent standard runs only, to optimize analytical precision; each vs. ln(68Zn/64Zn) sample, however, is analyzed several times. correlation for that

704


particular day. Extrapolation of the regression line to the "true" isotopic composition of the admixed Zn was used to obtain the fractionation corrected Cu isotope ratio of the standard (Figure 31.6b). Extra-polation of the sample results with the same slope and to the same Zn reference value then permitted the precise calculation of differences in the Cu isotopic composition between standards and samples (hCu, Figure 31.6b). An important feature of this technique is that it does not yield accurate absolute Cu isotopic compositions, but it accurately resolves small isotopic differences (MarOchal et al., 1999). Note that the calculated isotopic difference hCu does not depend on the choice of the Zn reference value. Rehk/imper & Halliday (1999) and Hirata (1997) applied external normalization for stable isotope ratio measurements of T1 and Ge, respectively, using the power law. The use of a power law for the correction of T1 isotope data was based on the observation that this procedure produced more precise data than an exponential law correction (Rehk/imper & Halliday, 1999), but it is likely that this was fortuitous. The suitability of the power law correction, however, follows directly from a thorough analysis of the analytical data. In a diagram of ln(205T1/203T1) vs. ln(208Pb/206Pb)(Figure 31.7) the raw measured T1-Pb data for the standard (M. Rehk/imper, unpublished results) define a regression line (r2- 0.9998) with a slope S = 1.0001 + 0.0026 (1 o uncertainty), and this is significantly different from the slope that would be expected for exponential law behavior (SE = 1.0147). If mass discrimination follows a power law, the measured T1-Pb isotope data should show a linear correlation in a ln(205T1/203T1) vs. ln(208Pb/206Pb) plot, with a slope given by

Sp -

(M205- M203) (M208 - M206)x

ln(1 + O~T1) ln(1 + ( I p b )

(see equations [31.3, 4]). If the power law is strictly applicable, (M205 - M 203 ) Sp = ( M 2 0 8 _ M 2 0 6 )

= 0.9999

[31.10] RT1 -- C~Pb,

such that [31.11]

Figure 31.6 - Schematic plots of ln(65Cu/63Cu) vs. ln(68Zn/64Zn) that illustrate the empirical external normalization technique used by Mar6chal et al. (1999) for mass bias correction of Cu. Both diagrams are based on results published by these authors. (a) On each of the four separate measurement sessions, the uncorrected isotope data for a mixed Cu-Zn standard solution defines a regression line with a distinct slope (full lines). The slopes of these regression lines are different from the slope that would be obtained if the mass discrimination (m.d.) of both Cu and Zn displayed ideal exponential law behavior (dashed line). Shaded fields encompass the range of raw data obtained in the MC-ICPMS measurements as well as estimates of the true Zn isotopic composition. (b) Application of an exponential law correction for mass discrimination (which assumes f3Cu = [3Zn; dotted lines) would result in poor external precision. Use of the empirical correction scheme (which assumes that [3Cu/[3Zn is constant for one measurement session and identical for standards and sampies; full lines) generates data with superior precision (see text).


705

706


The theoretical slope for power law mass fractionation is thus identical, within error, to the observed slope of the regression line for the analytical data (Figure 31.7). Use of the power-law and 208Pb/206Pb - 2.16701 for NIST SRM-981 Pb (Todt et al., 1996), however, generated a mean corrected 205T1/203T1 ratio of 2.38908 + 11 (2 o) for NIST SRM-997 T1 (Rehk/imper & Halliday, 1999). This is significantly higher than the certified T1 isotopic composition of this reference material which is 2.3871 _+10 (Dunstan et al., 1980). The latter value is considered reliable, within the given errors, because it is based on an absolute measurement of the atomic weight. The discrepancy between the true and the measured isotopic composition is probably related to the fact that the most precise correction procedure does not necessarily generate accurate isotopic data, because the regression line through the measured data does not intersect the true isotopic compositions of the standards at a = 0. Nonetheless, Rehk/imper & Halliday (1999) were able to show that the correction technique was suitable for the accurate and precise measurement of T1 isotopic differences between samples and an isotopic standard. These observations indicate that external normalization using either the approach of Mar6chal et al. (1999) or the exponential/power law do not provide a sufficiently correct description of the isotope fractionation process that is associated with plasma ionization. These procedures therefore are not suitable for the accurate and precise determination of absolute isotopic compositions. Internal normalization of MCICPMS data using an exponential law, however, appears to be able produce isotopic results that are accurate to within < 50 - 100 ppm for a number of elements. This indicates that chemical or physical properties other than mass may cause different elements to display a unique behavior during instrumental isotope fractionation. External normalization, however, can be used to resolve small isotopic differences between samples, even if the mass discrimination of the analyte and normalizing element is (slightly) different. The primary requirement is that the two elements show a highly correlated response to changes in instrumental mass bias but this condition is probably only met by elements with very similar or overlapping mass ranges.

31.3.3 Caveats and potential problems In the following we summarize some of the most important problems that may be encountered in acquisition of precise and accurate stable isotope ratio data by MCICPMS. Further treatments of this subject can be found in a number of publications (Jarvis et al., 1992; Habfast, 1997; Horlick & Montaser, 1998).

31.3.3.1 Spectral interferences Isotopic measurements are particularly vulnerable to artifacts generated by spectral interferences, because they aim to resolve very small differences in isotope composition. Such artifacts are a major concern, particularly below 100 amu, because MCICPMS instruments are generally operated at a mass resolution of N400, to produce peak flats that are sufficiently wide for high-precision isotope ratio measurements. At such operating conditions, the mass spectrometer is unable to resolve interferences between the analyte and other ions with similar mass to charge ratios. Spectral interferences are primarily generated by ions derived from the sample matrix (and the


707

Figure 31.7 - Plot of ln(205T1/203T1) vs. ln(208Pb/206Pb) for the raw data obtained by MCICPMS for a mixed solution of NIST SRM-997 T1 and NIST SRM-981 Pb. The data, obtained over a period of several months with a Plasma 54 (M. Rehk~imper, unpublished results), define a single linear correlation with a slope that is statistically identical (at the 99% confidence level) to the slope expected for power law mass fractionation.

sample solvent) or by molecular ions and refractory oxides that form in the plasma. The collection of accurate isotope ratio data necessitates that spectral interferences are either insignificant, or reduced to tolerable levels, such that an accurate correction can be applied. In most cases this can be achieved by (1) chemical separation of the analyte element from the sample matrix a n d / o r (2) the application of appropriate measurement protocols.

Isobaric interferences from other elements can be corrected for, by monitoring an isotope of the interfering element and application of a suitable interference correction (Halliday et al., 1995). Nonetheless, it is imperative to evaluate the correction scheme, to determine the maximum levels of contamination that can be tolerated for accurate isotope ratio measurements. To avoid the measurement uncertainties that are associated with such interferences, the majority of the published MC-ICPMS isotope data for geological samples have been collected either (1) on "clean" samples that were processed through a chemical separation step prior to the mass spectrometry, or (2) by direct analyses of major elements in a single phase (e.g., Cu in ore samples, Ca in carbonates). Correction for spectral interferences from polyatomic ions such as hydrides, oxides, nitrides, or argides is less straightforward. In many cases, such interferences (e.g., from element-argides) can be avoided by the chemical separation of the analyte prior to the mass spectrometry. The production of hydrides, oxides and nitrides is further-

708


more greatly reduced if a desolvating nebulizer is applied for sample introduction (Montaser et al., 1998). The development of methods for the measurement of Fe isotopic compositions by MC-ICPMS has been particularly difficult, due to interferences from various argides. The published data (section 31.4), however, demonstrate that precise results can nevertheless be obtained for this element, if appropriate analytical protocols are adapted. Interferences from refractory oxides and hydroxides (such as REE oxides) are often not a severe problem in MC-ICPMS because they can be avoided by chemical separation and because oxide formation can be greatly reduced in many cases by the use of a desolvating nebulizer. Collision cells, which are used in some Q-ICP-MS systems (Feldmann et al., 1999b; Bandura et al., 2001) and in the IsoProbe MC-ICPMS, have also been used to reduce polyatomic and isobaric interferences. Experiments conducted with Q-ICP-MS systems, however, have also shown that interfering molecular ions can also be produced in collision cells, for example by reaction of a He-H2 collision gas mixture with residual water from the solvents (Feldmann et al., 1999a).

31.3.3.2 Matrix effects Correction of instrumental mass bias by external standardization requires that the mass discrimination of the analyte element is identical for both the sample and the standard. Similarly, external normalization assumes that the relative mass bias encountered by the analyte and the reference element are indistinguishable for the samples and the isotopic standard. A number of studies conducted by both Q- and MC-ICPMS have shown, however, that instrumental mass discrimination can vary considerably with the analyte matrix and the response of two elements of similar mass to a different matrix may not be sufficiently correlated (Douglas & Tanner, 1998; Horlick & Montaser, 1998; Galy et al., 2001). This indicates that precise measurements of natural stable isotope fractionations are particularly prone to the generation of artifacts and should preferably be conducted on samples that have been chemically isolated from the sample matrix. Because no chemical separation is perfect, a number of MC-ICPMS studies demonstrated that remaining matrix constituents do not have a detrimental effect on data quality, either by influencing the mass bias behavior or through the formation of spectral interferences (Halicz et al., 1999; Zhu et al., 2000a). It is furthermore possible, that the accuracy of stable isotope ratio measurements may be affected by the concentration of the analyte (or the relative concentrations of the analyte and normalizing element) as well as the type and strength of the acid matrix of the sample and standard solutions. The effect of such factors on the accuracy of stable isotope ratio measurements at the 20 - 200 ppm level has yet to be investigated systematically. Future studies should consider such factors, because matrix effects on mass discrimination may vary considerably for different elements a n d / o r instruments.

Stable IsotopeAnalysisby MultipleCollectorICP-MS

709

31.3.3.3 Other factors

A number of MC-ICPMS stable isotope studies have noted that instrumental operating conditions (gas flows, focusing lens settings, acceleration lens potentials) must remain constant during an analytical session (Hirata, 1997; Rehk~imper & Halliday, 1999) because even small changes in these parameters may result in changes in mass bias. Memory effects have been noted as a particular problem for elemental concentration measurements by Q-ICP-MS, particularly for the analyses of samples with highly variable analyte abundances. For (stable) isotope ratio measurements by MC-ICPMS, memory problems are often less severe, because most analyses are conducted with "clean" solutions of separated elements, and large variations in the concentrations of the analyte solutions should be avoided in any case (Rehk~imper & Halliday, 1999). The "instrumental blank" problem is further demagnified, because samples and standards typically display only small differences (generally < 10%o) in isotopic composition. This condition does not apply if a double-spike is used for mass bias correction, such that a more rigorous control of instrumental memory may be required. It has been observed by many that memory effects are typically somewhat larger for desolvating nebulizers in comparison to other sample introduction systems. This disadvantage, however, can generally be overcome with appropriate washout protocols. Particularly severe memory problems for isotope ratio measurements by plasma source mass spectrometry have been encountered for a only a few elements (e.g., B, Os) which are thought to "stick" well to the surfaces of the sample introduction system. Isotopic analyses of these elements are nevertheless possible by ICP-MS, but require the application of designated sample introduction systems such as direct injection nebulizers (Montaser et al., 1998), and/or special cleaning procedures.

31.4 Applications In the last ten years, the technique of MC-ICPMS has found application in a continuously growing number of analytical laboratories wold-wide. The interest of the geochemical community in MC-ICPMS is largely based on the high precision and versatility of this mass spectrometric technique, which has been applied for isotopic measurements of numerous elements (Rehk~imper et al., 2001b). The following section provides a brief review of the application of MC-ICPMS to stable isotope ratio measurements in geo- and cosmochemistry. Stable isotope data for geological samples have already been collected for more than a dozen elements by MC-ICPMS (Table 31.2). In sections 31.4.1 to 31.4.10, we provide a review of the literature concerning those elements for which peer-reviewed papers have been published until August 2001. Section 31.4.11 briefly summarizes results presented in abstract form for additional elements. 31.4.1 Lithium

Lithium has the two stable isotopes 6Li (7.5%) and 7Li (92.5%). Due to the low ionization potential (5.4 eV), the isotopic composition of Li is readily measured by TIMS, but the acquisition of precise data is difficult due to the large instrumental mass fractionation. Significant Li isotopic variations have been detected in natural samples. This has spurred the development of analytical techniques for Li, employing TIMS

710


(Chan, 1987; You &Chan, 1996; Moriguti & Nakamura, 1998a), Q-ICP-MS (Gregoire et al., 1996) and HR-ICP-MS (Hamester et al., 1999). Lithium isotopic measurements by TIMS can achieve precisions of + 0.7- 2.5%0, but the most precise methods require long measurement protocols and either large sample sizes or extremely clean samples for precise control of instrumental mass fractionation. Tomascak et al. (1999a) were the first to develop analytical techniques for the measurement of Li isotopes by MC-ICPMS and these have been applied in a number of geochemical studies (Tomascak & Langmuir, 1999; Tomascak et al., 1999b, c). The Li isotopic data were acquired with a Plasma 54: MC-ICPMS using a Cetac MCN 6000 desolvating nebulizer. Individual analyses required ~ 40 ng of Li to achieve an internal precision of + 0.5%0 (2 o mean). Sample purity was shown to be a less critical factor for MC-ICPMS than for TIMS, thus allowing for a simpler chemical separation procedure. Control of mass discrimination was achieved by external standardization, with sample runs being interspersed between measurements of an isotopic standard. Short analysis times (~ 8 min) were used to minimize sample consumption and to achieve optimal correction for mass bias, because the latter was found to display a significant, but smooth, drift (up to 60%0 over 12 hr). An external reproducibility of + 1.1%o (2 o) was achieved for multiple measurements of samples and isotope standards (Tables 31.2, 3). Accuracy was evaluated by replicate analyses of two well-characterized standard samples: seawater and basalt JB-2 (Table 31.3).

31.4.2 Magnesium Magnesium has the three stable isotopes 24Mg, 25Mg, and 26Mg, with relative abundances of 79.0%, 10.0% and 11.0%, respectively. The first ionization potential of Mg is 7.6 eV. Investigations of mass dependent isotope fractionation of Mg, conducted by TIMS, reported measurement uncertainties (2 ~) of about + 1 - 2%o/amu (Catanzaro & Murphy, 1966; Catanzaro et al., 1966; Wasserburg et al., 1977). Galy et al. (2001) utilized a Nu Plasma MC-ICPMS instrument for Mg isotope analyses. A Cetac MCN 6000 desolvating nebulizer was used for sample introduction. The mass bias was reported to be ~ 7.5%/amu and was found to be sensitive to variations in instrumental operating conditions. The measurements of Mg stable isotope compositions applied a standard-sample bracketing technique for mass discrimination correction, with repeated short (200 s) analyses of both sample and standard solutions. An external reproducibility (2 o) about + 0.06%o/amu was obtained for several samples over a period of 15 months. The addition of Na, A1, and Ca to sample solutions was observed to give rise to higher 26Mg/24Mg isotope ratios, relative to a pure Mg standard. This was thought to result from a mass-dependant matrix effect that was most serious for Ca, followed by A1 and Na. The isobaric interference of 48Ca2+ on 24Mg was found to be significant at Ca/Mg ratios of a 0.5. The Mg isotope compositions of several terrestrial metal, mineral and chlorophyll samples (Figure 31.8) were reported to vary by up to about 2%o/amu (Galy et al., 2001). Results obtained for chondrules from the Allende meteorite were used to infer the conditions at which these objects formed in the early solar system (Galy et al., 2000).

Element Isotope ratio(s)

~

Li B Mg S Ca Fe cu Zn Ge Se Mo Cd Sn Sb Hg T1

Sample Introductiona

~

Mass Bias Correctionb

Mass Bias (70 lamu)

Precision (20) of MC-ICPMS measurements (700 1amu)

~~

Sol Sol Sol LA Sol Sol, LA Sol Sol Sol Sol Sol Sol Sol LA VG Sol

Terrestrial variation measured by MC-ICPMSC (%lamu) ~~

~

ES ES, EN ES EN ES EN, ES EN, ES EN EN ES EN, DS EN EN EN EN EN

-1

- 0.3 - 0.7 0.1 - 0.2 - 0.3

15 - 40 -8 -8 -5 -3 -3 -3 -1.5 -2 -2

- 1.7 - 0.9

- 0.7

0.05 - 0.1 0.03 - 0.15 0.04 - 0.05 0.1 - 0.5 0.15 0.02 - 0.1 0.07 - 0.15 0.05 - 0.1

-4 -5 0.5 (- 1.5") -1 0.3 0.2 0

-

-

-

-

-2

-

0.05 (?) 0.03 - 0.05

-


Table 31.2 -Summary of stable isotope ratio measurements conducted by MC-ICPMS.

-

-

- 0.25 -1

LA = laser ablation, Sol = solution nebulization, VG = vapor generation. b DS = double spike, EN = external normalization using an admixed element, ES = external standardization, including standard-sample bracketing. c No data are given for Li, B, and S because large terrestrial variations have already been determined with other techniques; * variation observed in Fe-meteorites. All data are from references summarized in section 31.4 or unpublished results of the authors. a

711

712


Table 31.3 - Comparison of Li isotope data for standard samples acquired by MC-ICPMS and TIMS. Instrument Type

Ref.

Sample Type

MC-ICPMS

(a)

JB-2 basalt Pacific seawater NIST-688 Li

TIMS

(b) (b) (c) (d)

JB-2 basalt Pacific seawater Atlantic seawater Seawater

Sample Size

67Li + 2o

n

40 ng 40 ng 40 ng

5.1 + 1.1 31.8 + 1.9 2.8 + 1.1

13 15 4

> 250 ng > 250 ng < 100 ng 4 mg

4.9 + 0.7 30.0 + 0.7 31.4 + 1.0 33.3 + 1.2

5 5 6 5

67Li represents the relative deviation of 7Li/6Li for a sample from the L-SVEC Li isotope standard in %. References: (a) Tomascak et al. (1999a); (b) Moriguti & Nakamura (1998a); (c) You &Chan (1996); (d) Chan (1987). n = number of measurements.

31.4.3 Calcium

Calcium has six stable isotopes, 40Ca (96.9%), 42Ca (0.65%), 43Ca (0.14%), 44Ca (2.1%), 46Ca (0.004%), and 48Ca (0.19%). Thermal ionization of Ca is straightforward, due to the low ionization potential of the element (6.1 eV), but the acquisition of precise isotopic data by TIMS is difficult, because of the large instrumental fractionation

Figure 31.8 - Three-isotope diagram in 6-notation for Mg (after Galy et al. (2001). 62XMg is the relative deviation of the 2XMg/24Mg ratio of a sample from the standard in permil. The bold line is the terrestrial mass fractionation line of Galy et al. (2000). Data are shown for two Mg metal samples (open circles), two Mg solutions (filled circles), magnesia (open square), two magnesites (diamonds), and two chlorophyll samples (triangles). Error bars are 2o.

Stable Isotope Analysisby Multiple Collector ICP-MS

713

effects. The most precise TIMS measurements of Ca isotopic compositions (external precision ~ + 0.1%o/amu) were performed by the double-spike technique and small terrestrial variations in Ca isotope ratios were identified (e.g., Russell et al., 1978). Ca isotope ratio measurements have also been conducted by Q-ICP-MS using "cool plasma" conditions (Patterson et al., 1999) and HR-ICP-MS (Sttirup et al., 1997). Halicz et al. (1999) determined Ca isotope compositions in carbonate samples by MC-ICPMS. The measurements were performed with a Nu Plasma instrument and a MCN 6000 desolvating nebulizer. The presence of 40Ar+ prevented the measurement of 40Ca, but backgrounds at higher masses were reported to be sufficiently low to permit accurate data acquisition for 42Ca, 43Ca, and 44Ca. Potential interferences by 84,86,88Sr2+ were corrected, by monitoring 87Sr2+ at mass 43.5. Mass discrimination for Ca was ~ 5%/amu, with a drift of about 0.1 - 0.05%o/amu per hour. A sensitivity of 0.3 x 10-11A/ppm was reported for 44Ca, and analyses were generally performed on Ca solutions with concentrations of ~ 15 - 30 ppm. External standardization by a standard-sample bracketing technique was used to correct for instrumental isotope fractionation during the measurements, and an external precision (2 o) of ~ + 0.05%o/amu was reported. The measured isotope ratios were found to be unaffected by the presence of Mg at concentrations up to twice that of Ca. Several carbonate samples were analyzed, and the data indicate a natural variability of 0.7%o in 42Ca/44Ca isotope ratios. 31.4.4 Iron

Iron has four naturally occurring isotopes, 54Fe (5.8%), 56Fe (91.8%), 57Fe (2.1%) and 58Fe (0.3%). Precise measurements of Fe isotope ratios by TIMS are difficult because of the high ionization potential of the element (7.9 eV), its low mass, and isobaric interferences from 54Cr and 58Ni. With double-spike TIMS procedures the Fe isotopic compositions of geological samples can be measured with an external precision (2 o) of about + 0.1 - 0.3%o/ainu (Beard, 1999; Mandernack et al., 1999). TIMS studies have identified natural variations in Fe isotope ratios for both meteorites (V61kening & Papanastassiou, 1989) and terrestrial materials (Beard, 1999; Beard et al., 1999). Various ICP-MS methods have also been developed for Fe isotopic measurements (Whittaker et al., 1992). The isobaric interferences of ArN + and ArO + on masses 54 and 56, however, are a major obstacle to precise Fe isotopic studies by plasma source mass spectrometry. Anbar et al. (2000) measured 56Fe/54Fe and 57Fe/54Fe ratios with a VG Elemental Plasma 54 using a microconcentric nebulizer coupled to a membrane desolvation system for sample introduction. The use of a desolvating nebulizer (operated without N2) reduced O and N based molecular species, such that the interferences of 40Ar14N and 40Ar160 were typically less than 6%o and 1%o for 54Fe and 56Fe, respectively. Prior to each run, the intensities of 40Ar14N and 40Ar160 were determined, and these data were used to correct the measured Fe isotope ratios. The intensities of sample and standard ion beams were furthermore matched by appropriate dilution. Correction for mass discrimination was achieved by monitoring 63Cu/65Cu of admixed Cu. The

714

Chapter 31 - M. Rehk~imper,F. Wombacher & J.K. Aggarwal

accuracy and precision of the techniques were verified by analyses of a standard solution enriched in 54Fe relative to an isotopically normal standard. An external longterm reproducibility of + 0.1 to 0.15 %o/amu was obtained for the 56,57Fe/54Fe ratios in these measurements, each consuming ~ 1 tog of Fe. Analyses of biogenic FeCO3 samples, produced by in-vitro microbial growth, showed shifts in Fe isotopic composition of ~ 0.4%o/amu relative to the starting material (Anbar et al., 1999). The fractionation of Fe isotopes during elution of Fe from an anion-exchange resin at the 2No/ amu level was also reported (Anbar et al., 2000). Zhu, Belshaw and coworkers reported the measurement of 56,57Fe/54Fe ratios with a Nu Plasma MC-ICPMS instrument for a variety of geological samples and meteorites (Belshaw et al., 2000; Zhu et al., 2000b; Zhu et al., 2001). Sample introduction utilized a MCN 6000 desolvating nebulizer, operated without a flow of N2, to minimize formation of ArN. The isotopic measurements were conducted at typical ion beam intensities of ~ 5 x 10-11A for 54Fe. With such signal levels, the molecular interferences at masses 54 and 57 were typically < 200 ppm; the Cr interference on mass 54 was < 25 ppm. Mass discrimination was ~ 3 % / a m u , and was corrected for during sample measurements by external standardization using a standard-sample bracketing technique. This procedure, however, required that the Fe concentrations of samples and standards were matched to within better than 20 - 50%. An external precision (2 o) of ~ + 0.03%o/amu was achieved for m i n i m u m sample sizes of 20 gg. A time-series analysis of a ferromanganese crust revealed large variations in Fe isotope composition that

Figure 31.9 - Fe and Pb isotope data for a time-series (0-2.5 Ma) of a Fe-Mn crust from the North Atlantic (Zhu et al., 2000b). 657Feis the relative deviation of the 57Fe/54Fe ratio of a sample from the standard in permil. Error bars are 2o. The iron isotope variations are coherent with those of 206Pb/204Pb"

715


were found to be correlated with changes in radiogenic Pb isotope ratios (Figure 31.9). Thus, the variations in 57Fe/54Fe were thought to be caused by changes in the continental input into the oceans (Zhu et al., 2000b). Analyses of various terrestrial minerals and meteorites for 56,57Fe/54Fe revealed differences of up to about 1%o/amu in isotope composition. All samples were found to fall on a single mass-fractionation line when plotted in a three-isotope diagram, indicating the homogenization of Fe isotopes in the early solar system (Zhu et al., 2001). Hirata and Ohno (2001) used a Plasma 54 coupled to a UV laser ablation system for in-situ Fe isotope analysis. An "on-peak" baseline subtraction procedure (similar to that used by Anbar et al., 2000) was applied to correct for interfering signals from ArN +, ArO +, and ArOH + and mass bias correction utilized sample-standard bracketing. External normalization to Cu offered no improvement in measurement precision. The laser beam was rastered over a 250/am2 area during each 50 s-analysis, which consumed about 30 ng of Fe for a total Fe ion current of ~ 6 x 10-11A. With these techniques, internal precisions (2 o mean) of better than 1%o were reported to be achievable for both 54Fe/56Fe and 57Fe/56Fe. The analysis of several natural materials (Feminerals and Fe-meteorites) showed no significant variations in Fe isotope compositions relative to the NIST SRM 665 Fe standard, within the analytical uncertainties of the method.

31.4.5 Copper Copper possesses two naturally occurring isotopes, 63Cu and 65Cu, with relative abundances of 69.2% and 30.8%, respectively. Cu displays a first ionization potential of 7.7 eV. Recent Cu stable isotope data, acquired using low temperature TIMS techniques, display external uncertainties (2 o) of + 0.2 - 0.3%o/amu (Gale et al., 1999). Isotopic analyses of Cu have also been performed by Q-ICP-MS (Lyon & Fell, 1990; Lyon et al., 1996) and HR-ICP-MS (Hamester et al., 1999). Mar6chal et al. (1999) performed the first Cu (and Zn, see below) isotope analyses of geological samples by MC-ICPMS. All measurements were performed with a Plasma 54, using a Glass Expansion nebulizer for sample introduction. Isobaric interferences from polyatomic ions (e.g., 40Ar23Na+, 48Ca160 + and 50,52Cr160+) required chemical separation prior to isotopic analyses of geological samples. The purified Cu fractions were doped with Zn for external normalization using 6 6 Z n / 6 4 Z n and an empirical correction technique (see section 31.3.2). A mass bias of ~ 3.3% / amu was reported for Cu. Repeated measurements of standard solutions and samples were found to display an external

Table 31.4 - Cu and Zn isotopic results obtained by MC-ICPMS for multiple dissolutions of sediment trap material (Mar6chal et al., 1999). Sample No

665Cu %oa

1 2 3 4 5 6 Meanc

0.05 0.06 0.06 0.09b 0.06b 0.16b 0.08 + 0.08

666Zn %oa 0.49 0.48 0.45 0.37 0.42 0.42 0.44 + 0.09

a 6-values expressed relative to NIST SRM-976 Cu and JMC Zn, respectively, b Cu isolated twice by anion-exchange chemistry, c Uncertainties are 2o.

716

Chapter 31 - M. Rehk/imper,F. Wombacher& J.K. Aggarwal

reproducibility of + 0.04%o/amu for 65Cu/63Cu (Table 31.4). As little as 200 ng of Cu were sufficient for one analysis. Data reported for oceanic sediments, biological materials and ore samples indicate the existence of terrestrial Cu isotopic variations with 665Cu values ranging from - 3 to + 5%o, confirming the results of previous TIMS measurements (Shields et al., 1965). 665Cu is the relative deviation of the 65Cu/63Cu isotope ratio of a sample from NIST SRM-976 Cu in %o. Gale et al. (1999) measured the Cu isotopic compositions of ores and archeological artifacts with a Plasma 54. Correction for instrumental mass discrimination of Cu was obtained by external normalization to an admixed Zn solution, using the exponential law. The mass bias of ~ 3.7%/amu was reported to be constant to within 0.8%o (2 o) over 10 h. Repeated standard measurements displayed a 2 o precision of ~ + 0.04%o/ amu. For the analysis of samples, the Cu was isolated from the matrix by a combined anion-exchange/electro-deposition technique. Data obtained for Cu minerals displayed 665Cu values ranging f r o m - 1.6 to + 7.7%o, whilst archeological samples ranged from + 0.2 to + 4.3%o in 665Cu (all data are relative to NIST SRM-976 Cu). Ore samples analyzed by both MC-ICPMS and TIMS displayed identical Cu isotopic compositions, within error. Zhu et al. (2000a) measured Cu isotope ratios of ore samples without chemical separation, to avoid isotope fractionation during ion-exchange chromatography. A Nu Plasma combined with a Cetac MCN 6000 desolvating nebulizer was used. A mass bias of ~ 3% / a m u was reported, and this was observed to vary with changing operating conditions during the course of an analytical session. Two methods of mass discrimination correction were evaluated: (1) external normalization to admixed Zn, and (2) external standardization by standard-sample bracketing. "Zn-doping" was reported to be faster and more precise (2 o external precision ~ + 0.02%o/amu), but required the removal of other elements that could interfere with the mass discrimination correction. The external standardization technique was chosen for analysis, because samples were not subjected to chemical purification. An external reproducibility (2 o) of ~ + 0.03%o/amu was obtained with this method. Isobaric interferences from NaAr +, NaOH+ and MgAr + were found to be negligible for the ore samples and Cu isotope ratios were unaffected by the presence of Fe in the sample solutions at levels up to 15 times higher than the Cu contents. Ore samples that were formed in association with low-temperature aqueous solutions were reported to display somewhat larger differences in 665Cu (- 0.3 to + 1.6%o) than chalcopyrites from igneous rocks (665Cu ~ - 0.6 to + 0.4%o). This was thought to indicate that Cu isotopic variations arise principally through low temperature processes, rather than as a result of source heterogeneities. 31.4.6 Zinc

Zinc has five stable isotopes (64Zn: 48.6%, 66Zn: 27.9%, 67Zn: 4.1%, 68Zn: 18.8%, 70Zn: 0.6%) and displays a very high first ionization potential of 9.4 eV. Stable isotope ratio measurements of Zn using TIMS and a double spike method have achieved an analytical precision of ~1.5%o/amu (Rosman, 1972). Isotopic analyses of Zn have also been conducted by Q-ICP-MS, e.g., for medical and archeological studies (Roehl &

Stable Isotope Analysisby Multiple Collector ICP-MS

717

Gomez, 1995; Budd et al., 1999). The first stable isotope measurements of Zn by MC-ICPMS were conducted by Nishizawa et al. (1998). A Plasma 54 instrument was used to investigate the Zn isotope effects of chemical exchange reactions performed in the laboratory. A precision of better than + 0.5%o was noted for the measurement of the isotopic enrichment factor of 66Zn relative to 68Zn, but no further analytical details were given. Stable isotope measurements for Zn in geological materials were performed by Mar6chal et al. (1999) using analytical techniques identical to those described for Cu. Despite the use of Ni cones on the Plasma 54 MC-ICPMS, only a negligible ( Table 36.1 continued

co2 ~

~

~~~~

~~

Table 36.1 continued > 763

764

> Table 36.1 continued molecule isotope sample condition 1 ratio measurement range

technique

1 laser

setup1

analysis time

10-precision / accuracy2

reference

170, gas phase H20 180,2H -1Opl H20 (1) 6 1 7 0 = - 1 0 0 ~to ~ +30o%~ 6180=-200%0to +1,200%0 62H=-500%0to +12,000%0

Direct absorption, color center laser at 2.7 pm AM-detection

R2

20 min, incl. evac/refill

2 ~m2) portion of the sample, as in the 'Grimm-type' (or 'obstructed') GD ion source (Grimm, 1968), the attainment of mm2-scale sampling capabilities has not yet been demonstrated. Commercial GD instrumentation is not supplied with real-time imaging or sample manipulation capabilities, however such features could be rapidly developed if there was sufficient interest. Some glow discharge analytical applications overlap better developed, and widely accepted geoanalytical techniques such as XRMF, SIMS, and laser ablation ICPMS. Consequently, the utility of GDMS in geochemistry, and particularly in the realm of

790

Chapter 35 - D.M. Wayne

isotope ratio analysis, has not yet been fully explored. Most GDMS instruments are single collector, double-focusingmagnetic sector instruments, such as the VG9000 GDMS (VG Elemental, Winsford, Cheshire, UK). These are designed to provide quantitative elemental analyses by scanning the entire periodic table at high mass resolutions (M/AM = 4:000 to 10,000). The focus of this paper will be on the characteristics of the GD ion source which are relevant to geochemical applications in general, and isotope ratio measurements in particular. 35.2 G l o w discharge ion sources

The history and development of the glow discharge (GD) source is comprehensively reviewed in several other publications (Coburn & Harrison, 1981; Harrison et al., 1986; Harrison & Bentz, 1988; Harrison, 1988). Until the early 1930s, glow discharge (or gas discharge) devices were widely used as ion sources for early mass spectrometric studies (Aston, 1942) but were replaced by spark source sources in the mid1930s (Dempster, 1935). Until the late 1960s and early 1970s, glow discharges were rarely used for analytical purposes. Research activity in the late 1960s and early 1970s (Grimm, 1968; Coburn & Kay) 1971; Harrison & Magee, 1974) culminated in the development of commercial analytical instrumentation based on the GD ion source by several manufacturers in the mid-1980s. Currently, only VG Elemental (Winsford, Cheshire, UK) and LECO (St. Joseph, MI, USA) continue to manufacture GD-based analytical instrumentation. However, GD is widely utilized in commercial and industrial laboratories, and fundamental GD processes and applications are studied in university and government laboratories worldwide. The GD source (Figure 35.1) consists of an electrode couple maintained in a lowpressure (ca. 0.1 to 10 Torr) noble gas (high-purity Ar or Ne) atmosphere. The appeal of noble gases, particularly Ar, in analytical GD applications is the ease with which they attain a long-lived highly energetic, or 'metastable' state. Metastable Ar atoms possess energy levels that exceed the excitation and ionization energies for most elements in the periodic table, with the exception of other noble gases, such as neon (Fang & Marcus, 1993). The sample to be analyzed forms the cathode, while the anode material (usually tantalum, steel, copper, etc.) is of little consequence. When an electric field (about 0.3 to 0.5 kV) is applied to the electrode, the working gas breaks down and electrons, ions and other species are formed in the cathode - anode gap (Figure 35.2). Bombardment of the cathode (sample) surface by incident particles (dominantly working gas ions) erodes atoms from the exposed sample surface. Most of the sputtered atoms remain neutral and redeposit on to exposed surfaces within the ion source, but a small percentage diffuse into a collision-rich zone (the negative glow region) in the cathode - anode gap. The ions in the negative glow can then be sampled via physical transfer into a mass spectrometer.

35.2.1 Glow discharge plasma structure The glow discharge (i.e., the region between the cathode and the anode) is comprised of eight discrete zones, each with different electrical properties, radiation intensities and particle population (Chapman, 1980). The size and importance of each region is dependent on the carrier gas species, pressure, GD potential, sample current,

Glow Discharge Mass Spectrometry: Fundamentals and Potential Applications ...

791

Figure 35.1 - Schematic diagram of the Grimm-type glow discharge ion source used in the Kratos sector GDMS. 1) tantalum ion exit plate, 2) tantalum anode, 3) argon inlet, 4) heating coils, 5) sample (cathode), 6) sample holder, 7) thermocouple, 8) quartz insulator, 9) quartz anode insulator, 10) copper spring, 11) stainless steel cap, 12) boron nitride insulator, 13) Macor insertion probe insulator, 14) stainless steel cell.

GD cell configuration, and cathode-anode gap (Chapman, 1980; Coburn & Harrison, 1981; Harrison, 1988; King & Harrison, 1993; Fang & Marcus, 1993; Bogaerts & Gijbels, 1998). The most important regions are the cathode dark space (also termed the cathode fall), the anode dark space, and the negative glow (Figure 35.2). The cathode dark space (CDS) is a narrow, non-luminescent zone located immediately above the cathode surface. The entire electrical potential difference between the anode and the cathode occurs within this region. Positive working gas ions are accelerated across the CDS and collide with the cathode surface. These collisions cause the continuous release of neutrals and electrons from the cathode surface. The electrons released from the cathode accelerate across the CDS into the negative glow region where they can participate in ionization reactions, and sustain the GD plasma. The

792


Figure 35.2- Schematic depiction of plasma structure (cathode dark space, negative glow and anode dark space) and different ionization mechanisms in a typical analytical glow discharge plasma. Metal and Argon atoms: Mo, Aro; Metal and Argon ions: M +, Ar+; Argon metastables: Ar*; electrons: e-. Diagram shows several possible interactions: 1) collision of Ar + with cathode releasing Mo + e- (metal atom or ion may redeposit on cathode), 2) ionization of Aro and Mo via electron impact, 3) ionization of Mo via interaction with Ar* (Penning ionization).

negative glow (NG) is a broad, light-emitting region adjacent to the CDS. The NG is essentially field-free, charge-neutral and equipotential, as electrons are slowed by numerous collisions with carrier gas and sample atoms. The majority of ionization processes occur in this collision-rich zone, but only ~0.001% of the original sputtered atom population are ionized. The anode dark space (ADS) lies between the NG and the anode (Bogaerts & Gijbels, 1998), and is analogous to the CDS except that electrons are accelerated towards the anode, and cations are repelled.

35.2.2 Cathodic sputtering Sputtering occurs when high energy (e.g. ~100 to 200 eV for a 0.5 kV cathode fall) Ar ions collide with, and penetrate several ~ngstroms into, the cathode surface (Chapman, 1980; Harrison, 1988; King & Harrison, 1993; Fang & Marcus, 1993; Betz & Wein, 1994; Bogaerts & Gijbels, 1998). The kinetic energy of the incident ions is transferred to the cathode material via collision cascades, and if the energy imparted to the lattice atoms exceeds the lattice binding energy, a variety of particles may be ejected from the cathode. The vast majority of these particles are neutral atoms (Fang & Marcus, 1993) which may redeposit on the cathode and other surfaces, or diffuse into the negative glow. Most importantly, cathodic sputtering creates an atom population that

Glow DischargeMass Spectrometry:Fundamentalsand PotentialApplications ...

793

constitutes a representative sampling of the cathode itself. The number of sample atoms produced by a single incident ion is known as the sputter yield, a property dependent on the kinetic energy, mass, angle of incidence, and species of the projectile ion, and the mass, lattice binding energy, surface properties and temperature of the target material (King & Harrison, 1993; Fang & Marcus, 1993; Betz & Wein, 1994; Bogaerts & Gijbels, 1998). Sputter rate, or the number of sample atoms ejected per unit time, is a function of sputter yield and of the operating current of the glow discharge (ca. 0.5 to 3 mA) which is, itself, dependent on carrier gas pressure and GD potential. Factors that alter the sputtering behavior of a material include the presence of oxide layers or other surface contamination, impurities in the carrier gas, phase inhomogeneities at the surface, and heating of the cathode. Compared to volatilization yields, which may vary over several orders of magnitude, sputter yields vary by less than an order of magnitude across the periodic table for a wide range of incident ion energies. Thus, GD processes have little matrix dependence, relative to thermal ionization, and ionization by laser and ion beam bombardment (Betz, 1980; Fang & Marcus, 1993; Bogaerts & Gijbels, 1998). For alloys, mixtures, and compounds, complications may arise due to 'differential sputtering', where elements having higher sputter yields (e.g., Cu, Ag, Au) are sputtered from the surface at a quicker rate than those with lower sputter yields (e.g., C, Si, Nb, Ta). However, a steady state is quickly reached within the glow discharge plasma, and the composition of the sputtered atoms then becomes a predictable function of the surface composition(Fang & Marcus, 1993). 35.2.3 Ionization processes

Ionization in the glow discharge source is spatially and temporally removed from the site of atomization. This separation further contributes to the relative lack of matrix effects observed in analytical glow discharges (Chapman, 1980; Coburn & Harrison, 1981; Harrison et al., 1986; Harrison & Bentz, 1988; Harrison, 1988; King & Harrison, 1990, 1993; Fang & Marcus, 1993; Bogaerts & Gijbels, 1998, 2000; Bogaerts et al., 2003). Working gas atoms, and neutrals derived from the cathode, may be ionized through a variety of collisional interactions in the negative glow region (Chapman, 1980; Harrison, 1988; King & Harrison, 1993; Fang & Marcus, 1993; Bogaerts & Gijbels, 1998). The dominant processes by which carrier gas atoms become ionized in the GD plasma are electron impact, fast Ar ion (or Ar atom) impact, and via interactions with metastable (excited) Ar atoms ('Penning' ionization). Advanced modelling and mathematical simulation of processes within the glow discharge by Bogaerts & Gijbels (1998, 2000) and Bogaerts, et al. (2003) indicate that, in addition to electron impact and fast Ar ion impact, Penning ionization and asymmetric charge transfer are also significant ionization mechanisms for sputtered neutrals.

794

Chapter 35 - D.M. W a y n e

35.3 Glow discharge analysis of insulating materials 35.3.1 Glow discharge spectra As in ICPMS, there are numerous interfering peaks in a typical GDMS spectrum. Most of these are multiply charged Ar and analyte species (absent from most ICPMS spectra), carrier gas clusters (e.g., [ArArl+), and metal - gas clusters (e.g., [FeAr]+ is abundant when sputtering steel samples). For silicate- or oxide-based based ceramics and geological materials, metal oxide clusters are also present. Unlike ICPMS, solvent- and atmospheric-derived peaks are very minor (Feng & Horlick, 1994). However, small amounts of water vapor, carbon, oxygen and various hydrocarbons may be present as contaminants in the GDMS source, and on the sample surface. Such contamination can be minimized prior to analysis by baking the source and sample in vacuo (Wayne et al., 1996; Wayne, 1997), by sputter cleaning the sample, and by cryocooling the GD source (Hall et al., 1989; Ohorodnik & Harrison, 1993). Analysis of insulating materials by dc-GDMS is typically accomplished by combining a powdered sample with a high-purity, electrically conducting metal powder, or binder, such as Ta, Ag, Cu or A1 (Mei & Harrison, 1991; Winchester et al., 1993; Teng et al., 1995; Wayne, 1997). This approach is used routinely for bulk and trace analysis of non-conducting solids, and permits the addition of an internal standard. However, the volume ratio of sample to binder is restricted by the amount of metal binder needed to make good electrical contact, and by the oxygen content of the nonconductor (Teng et al., 1995). Complete homogenization of the analyte in the binder may also be difficult. Compacted powders also have a high net surface area, thus surfaceadsorbed species and occluded atmospheric gases are not routinely sputtered or baked away prior to analysis (Wayne, 1997). There are two alternative methods for the analysis of non-conductors by GDMS. Direct sputtering of non-conducting solids has been demonstrated in radio-frequency (rf) glow discharge sources (Marcus, 1993; Marcus et al., 1994; Becker et al., 1996). Currently, rf sputtering is the preferred method for the analysis of non-conductors via GD atomic emission spectroscopy (GDAES). High-resolution rf-GDMS instrumentation is not available commercially, and the if-powered source must be assembled and interfaced in-house. The use of secondary cathodes (Milton & Hutton, 1993; Schelles et al., 1995; Betti et al.,1996a; Schelles & Van Greiken, 1996; Schelles et a1.,1996; Wayne et al., 1999), while not necessarily superior to the rf-GDMS technique, provides an inexpensive and convenient alternative. In this method, metal atoms sputtered from the secondary cathode redeposit and form an electrically conductive thin film on the exposed surface of the non-conducting sample. Sputtered metal atoms continuously redeposit and replenish the surface film, permitting direct depth profiling and quantitative analysis of the underlying non-conductor.

35.3.2 Quantitative analysis by GDMS One of the primary benefits of GD plasmas for analytical chemistry is the relative lack of matrix effects due to the temporal and spatial separation of atomization and ionization processes. However, GD plasmas are not entirely free from matrix effects, therefore simple calibration procedures are required for quantitative analysis. Relative

Glow DischargeMass Spectrometry:Fundamentalsand Potential Applications ...

795

sensitivity factors (RSFs) are used to relate the peak intensity of an element of unknown concentration to the peak intensity of a reference element of known concentration: RSFx/R = (Ix/Cx) / (IR / CR)

[35.1]

where Ix and Cx are the peak intensity and concentration, respectively, of the unknown element, and IR and CR are the peak intensity and concentration, respectively, of the reference element (Vieth & Huneke, 1990; King & Harrison, 1990, 1993). The RSF values for a series of elements in a given sample are functions of glow discharge processes, and may vary from one sample type to another. Part of this variation is due to differences in the sputtering characteristics and major element chemistry between samples (Vieth & Huneke, 1990; King & Harrison, 1990, 1993). Other studies have shown that changing GD parameters (gas pressure, GD potential, gas composition) can alter RSF values (Smithwick et al., 1993; Saito, 1995). Differences in ionization properties of the sputtered atoms are perhaps even more significant (King & Harrison, 1993; Bogaerts & Gijbels, 1996), and are more difficult to control. Although strict matrix matching is not required for the acquisition of quantitative analytical data by GDMS, calibrations should be performed using standard materials that are chemically similar to the unknown. For example, RSFs determined from steel (or other similar metal) standards are best suited to the quantitation of unknown Fe-CrNi alloys, and a soil sample is best calibrated using a soil, clay, or powdered rock standard (King & Harrison, 1990). Despite the complications summarized above, the relative sensitivities for most elements in typical GD samples vary by less than an order of magnitude, and yield linear calibration curves. Magnetic sector GDMS is a high-sensitivity analytical technique, with a linear dynamic range comparable to that of magnetic sector ICPMS. Numerous studies have shown that magnetic sector GDMS is routinely capable of quantitative analysis down to the nanogram per gram (part per billion) range (e.g., Becker & Dietze, 2003).

35.4 Isotope ratios by GDMS 35.4.1 Previous work

Isotope ratio measurements by GDMS have been the topic of only a handful of publications (Table 35.1) that describe data collected using the VG9000 double-focusing, magnetic sector instrument (Donohue & Petek, 1991; Shimamura et al., 1993; Duckworth et al., 1993; Barshick et al., 1994; Riciputi et al., 1995; Betti et al., 1996b; Itoh & Hasegawa, 1998). Others have reported isotope ratio measurements made using GDMS instrumentation adapted in-house from a conventional magnetic sector TIMS (Ecker & Pritzkow, 1994), and from a double-focusing spark source mass spectrograph (Chartier & Tabarant, 1997). All of these investigations have focused on metallic samples (solid and pressed powders), solutions dried onto metal powders, and powdered insulating materials dispersed in metal powders. In a novel study, Pajo, et al. (2001) applied GDMS to examine oxygen isotope ratios in pressed pellets formed from powdered uranium dioxide. Their results, however, proved to be too imprecise (0.5% to 4% RSD) to reliably detect variations in oxygen isotopes in the samples.

796


35.4.2 M a s s bias in the GD source

Mass bias in mass spectrometry arises when physical phenomena during ionization, sampling or analysis result in the preferential loss of either light or heavy ions from an ion population. Thus, the measured isotope ratio is different from the actual value. As the filament is heated during TIMS analysis, thermal effects favor the evaporation of lighter isotopes, relative to heavier isotopes, causing a bias of ~0.1% per atomic mass unit (amu). However, if the sample is completely consumed, as in total evaporation TIMS (Callis & Abernathey, 1991), mass bias is reduced to negligible level. Isotope ratio measurements made using plasma-source mass spectrometry (e.g., GDMS, ICPMS) are inherently less accurate than TIMS due to greater (>1.0% per amu) mass bias. Mass bias in plasma-source magnetic sector MS appears to be a predictable function of element mass, and is adequately described by a power law function (Russ & Bazan, 1987; Walder & Freedman, 1992): Rtrue = Rmeas (1 + F)n

[35.2]

where Rtrue and Rmeas correspond to the true and measured isotope ratios, (m2/ml) where m~ is the lighter isotope, F is the fractionation factor, and n is the difference between the 2 masses, ml and m2. An exponential fractionation law (Taylor et al., 1995) produces similar results (e.g. Lee & Halliday, 1995). For samples having similar chemistries, the isotopic composition of one element can be normalized to the isotopic composition of an element of similar mass (Longerich et al., 1987; Ketterer et al., 1991; Walder et al., 1993) that can be added directly to the sample solution. For direct solids analysis by GDMS, such mass corrections may be complicated somewhat if the necessary calibrant element (e.g. T1 for Pb) is not present in sufficient quantities to provide a precise determination of its isotope ratio. A calibration on a separate sample may be necessary in some instances, although most isotope systems of geochemical interest have at least one stable pair (e.g. 86Sr/88Sr) which can be used as an internal standard.

For plasma-sourced MS, mass bias is primarily the result of space charge effects that arise following ionization and sampling of the quasi-neutral plasma (Gillson et al., 1988; Li et al., 1995; Niu & Houk, 1996; Chen & Houk, 1996; Heumann et al., 1998; Hang et al., 1999). As positive ions are extracted from the ion source through the skimmer cone, charge separation occurs and lighter ions are repelled farther from the beam axis than heavier ion. The magnitude of the space charge effect in ICPMS is related to the sample matrix, the type of instrumentation (Heumann et al., 1998), and to the total ion current (typically ~ 6 ~A) passing through the skimmer (Gillson et al., 1988; Li et al., 1995; Niu & Houk, 1996; Chen& Houk, 1996; Hang et al., 1999). Experimental studies and theoretical calculations relevant to ionization and sampling processes in GD sources (Hang et al., 1999) indicate that space charge effects may be much less severe in GDMS, due to its lower total ion current (~ 0.02 gA). These calculations further suggest that post-skimmer beam divergence across the GD interface is a function of the square root of the ion mass (Hang et al., 1999). Thus, if space charge is the dominant cause of mass bias in GDMS, theory predicts that isotope ratios so obtained will favor the lighter isotope. However, the improvements in iso-

hlatrix

Type

Pdmetal rl:ld Fe met ecl r i t e -10% soil in Ag pcllder

sollJ.tii:iii

iii

Ag

pc1vder

Re:100% ApJ.. 5 yo

&metal Sr:SrC0, in Ag, Cupowder Re:m e t a1 Ag A$JO, on C u p ~ ~ d ~ r Eb:pwemetal, in Cumetal

B:boric acid in Ag p c d e r sb:in Cumetal 5:in Almetal Umetal

B: i i i r;teel 0s: n,a, U:fiuel rod B: in Zrmetal t a m e t a1


Table 35.1 - Summary of published isotope ratio data acquired using sector GDMS.

797

798


tope ratio accuracy relative to that of the ICP source will be difficult to assess using existing commercial GDMS instrumentation. Commercial GDMS instruments are adapted from existing platforms designed specifically for use with ICPMS, TIMS, or a variety of ionization and atomization techniques used to identify organic molecules. Subsequent evaluation of the isotope ratio capabilities of GDMS would best be carried out using an instrument fitted with an interface optimized for sampling ions produced by GD, rather than by ICP or TI (Hang et al., 1999). Isotope ratio studies to date (Table 35.1) indicate that scanning (i.e. single collector) magnetic sector GDMS provides equal, or slightly inferior, precision relative to that obtained using scanning sector ICPMS (Heumann et al., 1998). The cause of this disparity is not yet known, as the ion signal provided by the GDMS source is as stable as the signal supplied by conventional ICP sources (~2% over 10 min.; Finnigan MAT prom. lit., ca. 1995). Only two studies (Riciputi et al., 1995; Chartier & Tabarant, 1997) have examined mass fractionation trends in GDMS isotope ratio measurements. Riciputi et al. (1995) noted no correlation between GD potential or sample current and the severity of mass fractionation effects, but did observe a systematic increase in 65Cu/63Cu from 0.451 to 0.4:55 with increasing Ar pressure in the ion source region. By contrast, Chartier & Tabarant (1997) found no isotopic effects with changes in ion source pressure. The latter study (Chartier & Tabarant, 1997) was conducted using flat samples in a Grimm-type discharge cell, while Riciputi et al. (1995) used pin-shaped samples in a conventional GD cell. Significantly, Riciputi et al. (1995) speculated that the observed mass fractionation effects could be due to pressure-induced changes in the geometry of the negative glow relative to the ion extraction slits. Although considerable work remains to be done to thoroughly investigate mass bias in GD ion sources, the results of Chartier & Tabarant (1997) suggest that the effects of changing Ar pressure on the shape and position of the negative glow region may not be as severe for flat samples. 35.5 Direct measurement of isotope ratios by GDMS 35.5.1 Experimental

Isotope ratio measurements were performed on two NIST standard reference materials (SRMs) using a Kratos 'Concept' (Kratos Analytical, Manchester, UK) magnetic sector, double focusing, scanning GDMS (Smithwick et al., 1993; Wayne et al., 1996; Wayne, 1997) equipped with Mach 3 software. Standards used in this study include SRM 1264a (modified high-carbon steel), and SRM 611 (trace elements in glass, wafer form). The steel and glass standards were cut to I cm diameter discs, cleaned, affixed to a direct insertion probe using a disposable steel clip, and introduced into the GD source (Figure 35.1). Samples were baked in vacuo at ~200~ for ~30 minutes, cooled, and then pre-sputtered for N20 minutes. As a result, the water content of the ion source and sample surface is minimized, thus reducing the possibility of isobaric interference from metal hydrides. Samples are routinely examined for potential metal hydride (or other) interference at high resolution prior to the collection of isotope data at lower resolution. For all analyses, the distance of the sample surface from the anode was set initially at I mm, and all samples were run using a GD potential of 1.5 to 2.0 kV, with a sample current of 1.0 to 1.5 mA, at an accelerating voltage of 8 kV. Mass resolution, M/

799


AM, (peak width defined at 10% of peak height) was maintained at 1500, in order to obtain a flat-topped peak. All samples and standards were run in 'single ion monitoring' (SIM) mode (Wayne et al., 1996) by scanning either the magnetic field ('magnet scan') or the accelerating voltage ('ESA scan'). Further description of operating conditions for the Kratos GDMS can be found elsewhere (Wayne et al., 1996; Wayne, 1997). 35.5.2 Results NIST SRM 1264a: Isotope ratio analyses of W (0.102%) and Pb (0.024%) in NIST

SRM 1264a are summarized in Table 35.2. All errors and per cent RSDs are based on 2o standard deviations from the mean. Each analysis was performed by continuously sputtering the central region of the sample disc, resulting in the formation of a sputter crater 3 to 4 mm in diameter and several 10s of ~tm deep. Isotopic data was collected in sets of 8 to 15 scans over the mass range for each element, and repeated 8 to 10 Table 35-2 - Thallium, lead and uranium isotope ratios in NIST SRM 611 'Trace Elements in Glass'.

Run Name

205T1 203T1

F (%)

207Pb 206Pb

208Pb 206Pb

235U 238U

03/29/99 Isotop 13 %RSD Isotop 14 %RSD Isotop 16 %RSD Isotop 18 %RSD

2.42 (5) 2.1 2.42 (13) 5.5 2.38 (7) 2.9 2.40 (11) 4.7

-0.7

Trace 51 %RSD Trace 52 %RSD Trace 53 %RSD Trace 54 %RSD Trace 55 %RSD

2.42 (15) 6.1 2.34 (12) 5.1 2.41 (9) 3.6 2.41 (3) 1.4 2.36 (5) 2.1

-0.6

-0.6 0.1 -0.2

0.903 (9) 1.0 0.905 (16) 1.8 0.906 (35) 3.9 0.910 (13) 1.4

2.177 (19) 0.9 2.172 (43) 2.0 2.175 (43) 2.0 2.176 (48) 2.2

0.00220 (21) 9.3 0.00230 (5) 2.3

0.906 (23) 2.6 0.914 (9) 1.0 0.911 (18) 2.0 0.903 (12) 1.4 0.916 (18) 2.0

2.173 (55) 2.5 2.241 (31) 1.4 2.189 (29) 1.4 2.174 (34) 1.6 2.231 (25) 1.1

0.00247 (16) 6.3 0.00236 (14) 5.9 0.00227 (8) 3.7 0.00237 (19) 8.0 0.00236 (19) 7.9

0.9096 (8) 0.9095 (10) 0.9095 (30)

2.1670 (18) 2.170 (2) 2.160 (6)

0.002382

06/30/99 0.9 -0.4 -0.4 0.5

Reference Values MC-LAICPMS MC-TIMS MC-SIMS

2.3871

Pb and U ratios corrected for mass bias using factors (F%) calculated using an exponential law, based on 205T1/203T1 measured in-run. Parenthetical numbers are 2-s standard deviations reported to the last significant figure. Reference value for the T1 isotope ratio is from DeBievre & Barnes (1985). MC-LAICPMS data for Pb from Walder et al. (1993). MC-TIMS and MC-SIMS data for Pb from Belshaw et al. (1994). TIMS data for U in NIST SRM 611 from Barnes et al. (1973).

800


times per analysis. Total elapsed time required for each analysis (i.e., 8 to 10 sets of 8 to 15 scans) was 30 to 40 minutes for magnet scanning, and 4 to 6 minutes for ESA scanning. For the steel SRMs (Table 35.2), blank corrections for Hg (on 204Pb) were found to be unnecessary. Even so, 204Pb peaks were not intense enough to provide statistically useful Pb isotope ratios. Mass fractionation was estimated using the departure of the 186W/183W ratio from an accepted value (Volkening et al., 1991) and an exponential fractionation law (Lee & Halliday~ 1995; Taylor et al., 1995)" F1,2-[ln (Rtrue / Rmeas )] / [ml(ln (m2/ml))] Rtrue - Rmeas (m2/ml)~

[35.3] [35.4]

where 13- F1,2(m2), and ml is the lighter isotope. For magnet scanning, F1,2varies from -0.8% to +0.4% per amu. Mass bias is significantly greater for ESA scanning (2.4% to 6.0% per amu), and the fractionation on 186W changed considerably between runs 7 and 8 (Table 35.2) in response to changing lens settings during refocusing. Although the internal precision of W isotope ratios collected using ESA (N0.5% to 1.3% RSD) and magnet (~0.4% to 1.5% RSD) scanning was similar, the data obtained using ESA scanning was much less accurate (Table 35.2). External precision for raw W isotope ratios was 1.7%. Normalized W isotope ratios were more consistent (0.41% RSD) between runs. Lead isotope ratios were corrected for mass bias using fractionation factors derived from in-run 186W/183W. The 207Pb/206Pb ratios obtained from ESA scans agree reasonably well with those obtained via magnet scanning. However, the 208Pb/206Pb ratios obtained by ESA scanning have a consistent negative bias relative to measurements made by magnet scanning. The precision of the Pb isotope ratios acquired using ESA scanning (~1% to 2.5% RSD) is similar to that observed for magnet scanning (~1.5% to 3.6% RSD). Between-run variation for Pb isotope ratios was 1% to 2% RSD for 207Pb/206Pb, and 2% to 4% RSD for 208Pb/206Pb. NIST SRM 611" A Na-Ca-A1 silicate-based glass-ceramic standard (NIST SRM 611)

contains trace amounts of Pb (426 ppm), T1 (62 ppm), and U (461.5)(Barnes et al., 1973), and was analyzed by dc-GDMS using a tantalum (Ta) secondary cathode. Isotope ratios of Pb (208Pb/206Pb and 207Pb/206Pb) and uranium (235U/238U) in SRM 611 have been characterized using TIMS (Barnes et al., 1973; Belshaw et al., 1991) multicollector ICPMS (Walder et al., 1993c) and multi-collector SIMS (Belshaw et al., 1991). Isobaric interference from (TaNa) + clusters (m/z - 203.9378 amu) prevented the collection of 204Pb at a mass resolution of 1500. All the NIST SRM 611 data (Table 35.3) were collected using magnet scanning, and mass bias on Pb ratios was corrected using fractionation factors calculated from in-run determinations of 205T1/203T1 (Table 35.3), as described by Walder et al. (1993c). Mass bias using ESA scanning was greater by an order of magnitude (Figure 35.3), therefore only data acquired using magnet scanning (mass bias 50 A). Molecular sieves are natural and synthetic crystalline solids that have a microporous crystal structure. The pore dimensions are fixed for a given molecular sieve under specific operating conditions so that only vapor species in a specific size range can pass through. The term molecular sieve includes zeolites (natural and synthetic) and other microporous compounds such as borosilicates, aluminophosphates, and carbon molecular sieves (CMS e.g, carbosphere). Table 36.1 depicts a classification scheme for molecular sieves. 36.1.1 Zeolites

The first known molecular sieves were naturally occurring zeolites such as analcime, chabazite and mordenite. Primarily R. M. Barrer and his group at the Imperial College in London extensively researched the properties of these zeolites in the 1930s. In the 1940s after the first synthetic zeolites were produced, emphasis shifted to syn1. I dedicate this chapter lovingly to the memory of Mrs. Toshiko K. Mayeda, who in the words of Jim O'Neil, was "the first lady of stable isotope geochemistry". Tosh was Harold Urey's laboratory assistant and later Bob Clayton's right hand at the University of Chicago. She advised a number of students and post-docs in the stable isotope laboratory in matters of the lab as well as of life.

806 thetic molecular sieves (Breck, 1974), which can be synthesized in higher purity and designed for specific applications. Currently natural zeolites comprise only a small portion of total molecular sieve production.

Chapter 36- H.R. Karlsson Table 36.1 - Classification of Molecular Sieves (Modified from Szostak, 1989) Silicas Titanosilicates Metallosilicates Zeolites (aluminosilicates) Gallosilicates Chromosilicates Borosilicates Ferrisilicates Metalloaluminates Germanium-aluminates Aluminophospates* A1PO4, SAPO, MeAPO, MeASAPO, E1APO, E1SPO Other Gallogerminates, Gallophosphates, Arsenates Carbon molecular sieves (CMS); molecular sieve carbon

Natural and synthetic zeolites consist of a framework made of silica and alumina tetrahedra (SiO4-4 and A104 -3 units), such that all apices (oxygens) of the tetrahedra are shared with neighboring tetra(MSC) hedra. The packing of oxygen atoms forms a regular pattern * Other than for A1PO4 acronyms are derived from the cation occupying the tetrahedral site. O= oxygen, A = aluminum, of cages and pores. These cages P = phosphorous, S = silicon, Me = metal ion, E1 = some and pores are connected to other element. form a system of channels (Figure 36.1). The size of each pore opening is determined by the number of oxygen atoms surrounding it. In zeolites pore openings consist of 6, 8, 10, and 12 oxygen atoms. Assuming a 2.7 A diameter for the oxygen atom, the openings for a perfect ring are 2.7, 4.4, 6.0, and 7.7/~ for 6, 8, 10, and 12-membered rings, respectively. Substitution of A13+ for Si4+ leads a positive charge deficiency in the crystal structure, which is satisfied by extra-framework cations located in the cages and/or channels (Figure 36.1). In natural zeolites, those cations are typically Ca2+, K + and Na + and less often Ba2+, Cs + or Fe 2+. In synthetic zeolites, charge balance is accommodated by one of these cations, other cations not generally encountered in significant concentrations in natural zeolites (e.g., H +, Ag +, Cu +, rare earth elements) or cation complexes such as NH4 +. Thermally stable and acid-resistant zeolites generally have high Si/A1 ratios (McDaniel & Maher, 1976) and thus require fewer charge-balancing cations. Terms like "zeolite-like" or "zeolitic" refer to crystals whose structures are similar to those of zeolites but lack the Si-A1 framework. The term molecular sieve is therefore more inclusive than zeolite. In natural zeolites, the cations are hydrated by water molecules that are readily removed by heating and/or evacuation. Upon dehydration the framework responds either with little or no distortion or with total collapse (Breck, 1974), the former being preferred in commercial applications. However, with proper selection and careful use, zeolitic molecular sieves can undergo repeated hydration and dehydration up to 1000 times (Hersh, 1961).

The Use of Molecular Sieves in Stable Isotope Analysis

Figure 36.1 - A polyhedral representation of the structures of zeolites A and X (and Y). The truncated cuboctahedron (a) is the basic building block in zeolites A and X (and Y). Silicon or aluminium atoms are located at the corners of the polyhedron. Oxygen atoms are situated approximately midway along the vertices (e.g., O1-O4 in c). In zeolite A, the cuboctahedra are linked to each other along the square faces with rectangles (b). What results is a cubic structure with 4, 6, and 8-membered rings of oxygen atoms. Also shown are the approximate extra-framework cation positions (SI and SII). In zeolite X (and Y) the cuboctahedra are connected at the hexagonal faces. As a result, cages are accessed through 4, 6, 8, and 10-membered rings of oxygen atoms. Approximate extra-framework cation position are also indicated (SI, SI', SII, SII', SIII). Adapted from Wortel (1979).

807

808

Chapter 36- H.R. Karlsson

36.1.2 Other molecular sieves Zeolites and their analogs are probably the best-known and most utilized molecular sieves today. The past four decades have, however, seen a drastic increase in the synthesis and commercial production of non-zeolite molecular sieves such as A1PO4types and their derivatives (Feng et al., 1997), and carbon molecular sieves (CMS or MSC; also see Table 36.1). M a n y of these non-zeolitic molecular sieves are niche applications. For example, CMS is rapidly replacing zeolites in air purification systems (Szostak, 1989; K/irger & Ruthven, 1992). For stable isotope work, however, the author has found few examples of use of CMS (e.g., Carbosphere) and no examples of some other molecular sieves such as a l u m i n o p h o s p h a t e although both have been available for some time. A brief overview of these substances will nevertheless be given as a point of contrast with the zeolites and because these materials m a y well prove suitable in future stable isotope w o r k A l u m i n o p h o s p a t e s are group of zeolite-like substances in which the tetrahedral sites are occupied either by A13+ or p5+. Only A1PO4s with P/A1 - 1 are thermally stable, whereas the Si/A1 ratio in stable zeolites can be varied over a considerable range. On the other hand, the A1PO4 f r a m e w o r k is electrically neutral and there are no charge-balancing cations in the channels. This leads to a decreased preference for polar over non-polar adsorbate molecules. Carbon molecular sieves (CMS) or molecular sieve carbon (MSC) differ from ordinary porous carbon substances in that they have a n a r r o w range of pore widths. CMS pore sizes range from approximately 3 to 5 ~ (see Table 36.2). The CMS pore structure is generated by acid etching of graphite leading to a n e t w o r k of pores or slits running along the basal planes (see e.g., Seaton et al., 1997). Although not well understood, the Table 36.2 - Examples of Synthetic Molecular Sieves Name*

Structure

Linde NaA LTA Linde CaNaA LTA Linde KA LTA Linde NaX FAU Linde NaCaX FAU Linde Y FAU Mordenite MOR ZSM-5 MEL A1PO4-5 AFI VPI-5 VFI CMS graphite

Chemical composition Na12[A112Si12048] 27H20 Na86Ca[A112Si12048] 27H20 (Na, K)86[A112Si12048]27H20 Na86[A186Si1060384] 264H20 Na86Ca[A186Si1060384] 264H20 Na56[A156Si13602841 250H20 Na8[A18Si40080] 24H20 Nan[AlnSi(96-n)O153] 518H20 [Al12P12048] R qH20** [A118018PO72] 42H20 C (anthrasite)

Pore diameter (A) 4.1 4.2-4.4 3.3 7.4 7.8 7.4 6.5 5.3 7.3 12.1 3-5

Product name 4A 5A 3A 13X 10X Y ZSM-5 MSC-4A,-5A

* Linde NaA is also known as simply "A". Similarly Linde NaX is known as "X". Linde of Union Carbide manufactures ZSM-5. Mobil has made a zeolite with the same structure and similar chemistry. It is marketed as silicalite, which is almost A1 free. ** R and q stand for (C3H7)4NOH and number of water molecules, respectively. Pore diameters are for room temperature. Sources: zeolites; Breck (1974) and Int. Zeol. Council Web site (1999): CMS; MacElroy et al. (1999).


809

pore network most likely consists of cavities linked by smaller pores. The size and length of the pores (MacElroy et al., 1999) govern the diffusion and adsorption properties of the material. CMS behave in many respects like small-pore zeolites (e.g. analcime and zeolite A) although they lack discrete pore openings and cages (Chihara et al., 1978). 36.2 Properties of Zeolites and Molecular Sieves

36.2.1 Molecular Sieving The term molecular sieve was coined by McBain (1932). Unlike other porous solids, molecular sieves are characterized by continuous channels with a narrow range of pore sizes (Figure 36.2). The restricted range of pore sizes in molecular sieves allows these materials to separate guest molecules according to size and shape. Currently, about 150 different zeolite framework topologies have been synthesized (International Zeolite Council WEB-site, 19991) with pore sizes ranging from 2.6 ~ to 7.4 ~. Small-pore zeolites such as analcime (2.6 ~) allow only small molecules (e.g., H2, He, H20) to enter whereas large-pore zeolites such as faujasite, zeolite-X, and zeolite-Y

Figure 36.2 - Approximate pore size distribution in molecular sieves and other porous absorbents. Note the logarithmic x-axis. Individual zeolites, aluminophospates and carbon molecular sieves have much more restricted ranges in pore sizes than other absorbents such as silica gel or activated carbon. Modified from Breck (1974).

1. See WEB-site: http: / / www-iza-sc.csb.yale.edu / iza-sc /

810


allow molecules up to 7.4 A to enter. Larger pores occur in aluminophospates (A1PO4), e.g., VPI-5 has 12.1 A wide channels, but nature still holds the record (14.2 A pore diameter in the phosphate mineral cacoxenite; Moore & Shen, 1983). The restricted range in pore size for each individual molecular sieve means that it is possible with careful selection to separate many different sized molecules. Cation identity, temperature, and pre-adsorption all influence the effective pore diameter in zeolites. Consider, for example, the synthetic zeolite Linde A. The Na endmember (Linde 4A) has an effective pore diameter of 4.1 A and will adsorb molecules smaller than that diameter (see Table 36.3). Replacement of Na + by K+ (Linde 3A) narrows the pore opening to 3.3 ~ because K+ is larger than Na +. Both forms will adsorb H20 (is 2.6 A) but the adsorption of 02 (3.5 A) and CO2 (3.2 A) drop off with increased K exchange so that 20% K exchange leads to negligible adsorption of 02. Exchange of Ca2+ for Na + (Ca = 2Na) leading to Linde 5A has the effect of widening the pore diameter by unblocking the pores (fewer crystallographic sites are occupied and "nonblocking" crystallographic sites are populated preferentially at lower site occupancies). Both Linde 3A and 4A exclude N2 (3.6 riO. At liquid nitrogen temperatures it can be sorbed on the Ca zeolite Linde 5A (Breck, 1974). Increasing temperature has two basic effects on the sorption properties of zeolites, effectively opening up the pores allowing larger molecules entrance and enhancing the diffusion rates of adsorbants. The effect of temperature on the properties of molecular sieve Linde 4A is dramatic. At room temperature N2, 0 2 , and Ar are all adsorbed very slowly. As the temperature is decreased the adsorption capacity initially Table 36.3 - Useful properties of some molecules encountered in stable isotope work Species He H2 H20 NH3 Ar N2 NO N20 02 CO CO2 CH4 CF4 SO2 H2S SiF4 SF6 NF3

Kinetic diameter (A) 2.6 2.9 2.6 2.6 3.4 3.6 3.2 3.3 3.5 3.8 3.3 3.8 4.7 3.6 3.6 4.9 5.5

* Sublimation temperature.

Dipole moment (D)

1.854 1.471

0.159 0.161 m 0.110 m

1.633 0.97

0.235

Boiling point (~ -268.9 -252.8 100 -33.35 -185.7 -195.8 -151.8 -88.5 -183.0 -192.0 -78.5* -164 -128 -10 -60.7 -86 -63.8* -128.8


811

Figure 36.3 - Adsorption of 02, N2, and Ar on zeolite 4A as a function of temperature. Note that N2 is most effectively sorbed at around-78~ (dry ice), whereas Ar is most effectively sorbed around-160~ At-195~ (liquid nitrogen) 02 very efficiently sorbed whereas N2 and Ar are not. Adapted from Breck (1974).

increases for each of these species as illustrated in Figure 36.3.02 adsorption increases steadily with further decreases in temperature, whereas N2 and Ar adsorption peak a r o u n d - 100~ a n d - 150~ respectively. Accordingly, the most efficient adsorption of 02 would be at liquid nitrogen temperatures or lower while Ar and N2 are more effectively trapped at higher temperatures. Breck (1974) attributes this temperaturerelated selectivity to narrowing of the effective pore aperture in Linde 4A by 0.1 - 0.2 ~ at low temperatures. Apparently, this narrowing does not inhibit entry of 02 (3.5 ~) but it results in rejection of the slightly larger N2 (3.6 ~). This view is probably overly simplistic since Ar (3.4 ~), which has a smaller pore diameter than 02, has dimin-

812


ished adsorption relative to 02 b e l o w - 150~ The fact that the activation energy for diffusion of N2 (24.3 kJ/mol) is higher than that for Ar (11.3 kJ/mol) in Linde 4A might explain this difference. The adsorption capacity of zeolites can also be influenced by the presence of molecules in the pores. Polar molecules such as H20 and NH4 are particularly effective in reducing the sorption capacity of small-pore zeolites such as Linde A. These polar molecules are strongly adsorbed and not easily displaced by non-polar molecules such as 02. The adsorption capacity of 02 on Linde 4A falls drastically, even with a small amount of preadsorped polar molecules, because cation-adsorbate complexes forming in the zeolite channel act as a blockage (Breck, 1974). In large-pore zeolites (e.g., zeolites X and Y) similar effects are seen when large inorganic cation complexes are formed (e.g. Cu pyridine).

36.2.2 Sorption Depending on grain size, the external surface areas of zeolites range from 1 to 3 m2/g and internal surface areas lie between 700 and 1000 m2/g (e.g., Hersh, 1961). Molecules small enough to enter the pores of the sieve are overwhelmingly adsorbed internally; whereas molecules that are too large will be adsorbed externally. Similarily the diffusion rates of molecules through a bed of molecular sieve will be affected. Molecules migrating through the sieve will diffuse slower than those migrating around the sieve (micro- vs. macro-pore diffusion) because they encounter higher electrostatic barriers. The amount sorbed onto a molecular sieve, or the sorption capacity depends on the maximum available void volume, the size of the sorbate molecule(s), the nature of the gas, pressure and temperature. The void volume is greatest in activated molecular sieves or zeolites, which are generated by either heating, typically in the range 300- 500~ or a combination of heating and evacuation, and varies with sorbed species. For example, in Linde 4A the void space in terms of cm3/g is 0.289, 0.252, and 0.213 for H20, CO2 and 02, respectively. In zeolite Linde 13X, the corresponding values for those molecules are 0.36, 0.33, and 0.31. Provided that the dimensions of an adsorbate molecule are similar to or smaller than the effective pore diameter, it will be adsorbed until an equilibrium state is reached. The equilibrium state varies with temperature and pressure but the sorption capacity is limited by the available void volume as discussed above. Generally, more of a vapor species is adsorbed at higher pressures and lower temperatures, although slow kinetics may make the equilibrium values impractical measures of effective adsorption. The sorption of gas mixtures onto zeolites is still poorly understood and not easily predicted. For example (Figure 36.3), pure 02 is readily adsorbed a t - 183~ on Linde 4A but pure N2 is not. In an N2 - 02 mixture, however, little 02 absorption occurs presumably because N2 hinders absorption of 02. A similar effect is seen at higher temperatures (- 78~ where pure N2 is adsorbed in greater abundance than 02. The overall adsorption of both gas species in a mixture drops significantly.


813

High-silica zeolites (silicalite or ZSM-5), A1PO4s and CMS have molecular sieving properties similar to that of zeolites but differ in their sorption capacity. Since these substances have either few cations (silicalite, ZSM-5) or none (A1PO4, CMS) in the pores, the channels are unobstructed. Furthermore, due to charge-balanced lattice the channels show little preference for polar over non-polar adsorbates. In fact, these substances are hydrophobic compared to other zeolites. For example at 24~ Linde 4A and silicalite adsorb 0.3- 0.35 cm3/g and 0.0- 0.04 cm3/g H20, respectively. CMS are similarly hydrophobic but A1PO4 are intermediate between silicalite and Linde 4A (Szostak, 1989). 36.2.3 Commercial molecular sieves

Commercial zeolites are either natural or synthetic. Natural specimens are usually fine-grained crystals from sedimentary deposits mined primarily for the agriculture and construction industries. Most natural zeolites are not adequately characterized in terms of their chemical composition and purity, and are thus unsuitable for sorption and molecular sieving work. Additionally, large-pore zeolites such as faujasite are extremely rare in nature and therefore not widely available commercially. Synthetic zeolites have well-known chemistries and physical properties and are thus well suited for aiding stable isotope work. Synthetic zeolites are available in two f o r m s - as pure crystals and as aggregates or pellets made from zeolites and an inert binder. The binder is often a clay, silica or silica-alumina mixtures and can make up 10 to 40 wt% of the material (Breck, 1974). Similarly, carbon molecular sieves are made with a binder composed of tar or polymeric material (K~irger & Ruthven, 1992). It is important to know whether or not a binder is present because the binder changes the overall sorption capacity and isotopic exchange properties of the molecular sieve. In this paper I will refer to binder-less materials as crystals and binder-containing materials as pellets. The term granular is non-specific. 36.3 P r o c e s s e s

Various processes can alter the original isotopic compositions of the gas of interest. As the gas diffuses through (e.g., GC) or into (e.g., trap) a molecular sieve, isotopic fractionation will undoubtedly take place. The fractionation can be due to diffusion, chemisorption, and/or isotopic exchange between the molecular sieve and host gas. Kinetic effects due to differences in the diffusivites of isotopic species are well known but should not be a problem when recovery of the gas is complete. Chemisorption can be a problem when the gaseous species reacts with the host. For example, NO sorbed on the zeolites chabazite, A, and X breaks down to form N20 and NO2 or N203. The rate of disproportionation increases with decreasing temperature and is nearly complete in 1 hour a t - 78~ for chabazite (Barrer, 1978). Chemisorption, however, occurs mostly in noble-metal and transition-metal ion-exchanged zeolites. Isotopic exchange could take place between the guest molecule and the host lattice and/or other guest molecules. It is well established that oxygen isotopic exchange takes place between the framework oxygen and molecules such as 02, CO2, and H20

814


at temperatures as low as 0~ (see e.g., Karlsson & Clayton, 1990; Karlsson, 1995, and references therein). However, since the exchange rate increases with temperature and time, it is unlikely that significant isotopic exchange will occur when these gases are desorbed from small traps or cold-fingers in a matter of minutes with gentle heating (< 200~ A worst case scenario will occur when large quantities of sieve material (e.g. sampling traps and GC columns) are heated at high temperature (> 200~ for longer periods of time (hours) to desorb gases. Conceivably H isotopic exchange could take place between adsorbed H2 or H20 and terminal OH groups in zeolites. However, since the concentration of OH is normally small compared to the guest hydrogen-bearing species, it is unlikely to have a significant effect. Exchange between guest molecules is a serious consideration when using molecular sieves in isotopic studies. For example, extensive oxygen isotopic exchange can take place between CO2 and H20 if these coexist in a sieve, either inside the sieve or upon exiting. It is therefore critical that molecular sieves be reactivated between samples. 36.4 Applications of Molecular Sieves to Isotope Analysis - Examples

Table 36.3 lists the properties of gaseous species commonly encountered in stable isotope work. Shown are relevant properties such as atomic or molecular diameter and polarity. The sizes are kinetic diameters. Breck (1974) derived these values from the Lennard-Jones and Stockmayer potentials and by considering the adsorption of species onto zeolites with well-constrained pore openings (e.g., Linde A). Dipole moments (Debye units) were taken from the 75th edition of the CRC Handbook of Chemistry and Physics (Lide, 1994). Breck (1974) evaluated species typically used in light stable isotopic analysis (e.g., H2, 02, CO2, N2, SiF4, SO2, SF6) and some other gases that can interfere with analysis. Simple cryogenic trapping, with the aid of dry ice (- 78~ and liquid nitrogen (- 195~ can be used to transfer and separate gases. However, this procedure does not work in many instances. For example, 02 and N2 cannot be transported quantitatively without the aid of liquid He. Even liquid He cannot transfer H2 and it requires a Toepler pump (or U-chemisorption). Similarly, Ar, N2, 02, and CO cannot easily be separated cryogenically from each other and so are the pairs H2S~SF6, CH4~NO, and CO2~SiF4. Many of these separations can, however, be accomplished with molecular sieves because the sieves allow one to discriminate on the basis of molecular size (kinetic diameter) and polarity (dipole and quadropole moments).

36.4.1 Attainment of good vacuum and trapping of hydrocarbons Molecular sieves such as zeolites can be used to improve vacuum and keep unwanted components from back-diffusing from roughing pumps into the main vacuum line work space. By placing a large quantity of zeolite between the roughing pump(s) and the high-vacuum pump(s) two goals are achieved. First, once activated, the molecular sieve will trap undesirable compounds such as hydrocarbons and prevent them from reaching the main vacuum system. Such fore-traps are common on VG instruments but usually contain activated alumina rather than a molecular sieve


815

(D. Bourne, pers. comm. 1999). Second, by cooling the molecular sieve to liquid nitrogen temperatures it effectively acts like a pump (cryopump) and ~helps to produce high vacuum faster than would otherwise be obtainable. Dr. Irving Friedman's Niertype mass hydrogen isotope mass spectrometer contains such a trap and good vacuum can be obtained within half-an-hour after the machine has been down (I. Friedman, pers. comm. 1998).

36.4.2 Gas sampling Zeolites have been used to collect air samples for carbon isotope analysis. Bol & Harkness (1995) used 13X pellets (7- 8 g, 1/16") packed in quartz tubes for field sampling of CO2 in air. The zeolite cartridges were activated in the laboratory and subsequently filled with clean nitrogen. In the field, the tubes were opened and roughly 8 liters of air passed through each with the aid of a light battery-powered pump. The trapped CO2 was subsequently recovered in the laboratory by heating to 500~ in a vacuum system and separated from other gases by conventional cryogenic trapping. Comparison with CO2 samples obtained by conventional expansion into 2 1 glass bottles showed no appreciable fractionation in carbon isotopes. Bol & Harkness (1995) obtained a precision of 0.2%o or better. Recovery of CO2 from the sampled air was estimated to be > 97%. Neither the pumping rate (200 - 1000 ml air/min) or the amount of air collected (6 - 18 liters) had a significant effect on the carbon isotope values. The technique of Bol & Harkness (1995) may thus prove very useful in collecting gas samples in which virtually the entire isotope of interest resides in one species (e.g., C isotopic composition of CO2 in air). When the sieve retains multiple species, it is likely that isotopes would have been fractionated. If for example, one desired to analyze the oxygen isotope composition of CO2 in air, the procedures of Bol & Harkness (1995) would have been inappropriate because the amount of H20 collected in the trap would have been many times greater than for CO2. Upon heating, oxygen isotopic exchange would have ensued between these two species. Indeed, R Bol found that there was significant oxygen isotopic fractionation in the CO2 (pers. comm. 1999). Schwarz et al. (1996) collected air moisture in the Antarctic using an automatic sampler. Their device featured zeolite molecular sieve 5A (Wolfen Zeosorb A5; G. Schwarz, pers. comm., 1999) to trap the moisture for hydrogen isotope analysis. The device has twelve zeolite cartridges containing 50 g zeolite "balls" (1.5 to 2.0 mm diameter), each of which can be used to collect a sample at a specified time interval. The zeolite was activated by heating to 400~ under vacuum. In the field, air was pumped through the cartridges with the aid of a battery-powered micro-vacuum pump. Sampling times were 6 - 150 hours with airflow rates of 100 - 1201/hour. Thus, at least I g of water was collected onto the molecular sieve. The water was recovered by heating the zeolite cartridge to 400~ under vacuum (Schwarz et al., 1998a) and the evolved H20 reduced to H2 by reaction with hot Cr (Schwarz et al., 1998b). 6D values of moisture collected using the molecular sieves are within 5%o of data obtained by cryogenic methods, which is quite good considering that sampling was done some days apart.

816

Chapter 36 - H.R. Karlsson

36.4.3 Gas separation Cheng & Bremner (1965) report that Linde 5A used in a GC column is effective in separating N2, 02, NO, CO, CH4, N20 and CO2. At room temperature, the sieve retains N20 and CO2 but N2, 02, and NO are separated. N20 and CO2 are released upon heating but no mention was made of possible isotope effects. Meier-Augenstein et al. (1994b) tested the use of zeolite molecular sieves in removing moisture from CO2 produced during breathing. The samples were obtained from human subjects and a 200 ml aliquot passed through columns containing molecular sieves at room temperature. Zeolites 3A, 4A and 5A were tested and compared with results obtained by conventional methods (dry ice trapping). Two 1/4_inch columns were u s e d - the short one was 152 mm and the long one was 994.8 mm. Breath samples were run sequentially through the sieves without reactivation between samples. CO2 passed through the columns containing 4A and 5A experienced large fractionations in both C and O isotopes and the use of these zeolites was therefore not investigated further. More extensive work was carried out with zeolite 3A including tests of memory effects. It appears that C and O are not fractionationed substantially (> 0.5%o) especially in the case of 3A short column. Zeolite 3A may therefore be used as simple way to remove moisture from breath samples without introducing significant isotopic fractionation. The observed isotope effects may readily be explained when the pore sizes of these zeolites are considered. In the case of 3A, little CO2 enters the pores but H20 is readily adsorbed. The CO2 thus flows past the molecular sieve. In the case of 4A and 5A however, both molecules enter the zeolite structure and interact with each other and with the zeolite framework generating significant isotopic effects.

36.4.4 Gas transfer and purification Oxygen (02): Clayton & Mayeda (1983) used zeolite 13X to separate NF3 from 02 prior to isotopic analysis of oxygen isotopes in meteorites. Invaluable information can be obtained by analysis of the three isotopes of oxygen in meteorites and the analysis is best done using 02 rather than CO2. During fluorination of meteoritic samples, NF3 is produced in such large quantities that erroneous 6170 results (NF + has a m / e of 33, the same as 170160) may be obtained. The mixture is, therefore, adsorbed onto zeolite 13X pellets (Anasorb" 30 - 60 mesh) at liquid nitrogen temperatures. Subsequently, the temperature of the sieve is raised t o - 115~ (using a solid/liquid ethanol slurry) releasing the 02 (and Ar) but retaining the NF3 and N2. The purified 02 is then transferred to a mass spectrometer using a liquid He cold-finger. The 13X was regenerated between experiments by heating under vacuum to 250~ for 30 minutes (R.N. Clayton & T. K. Mayeda pers. com. 1999). Miller et al. (1999) used a slight variation of Clayton's & Mayeda's method. In addition to using 13X to purify 02, Miller et al. (1999) used a second 13X cold trap for transporting the 02 to the mass spectrometer inlet system. In order to release 02 into the mass spectrometer, they raised the temperature of the 13X trap to 50~ for 6 minutes. Although, it is certain that oxygen isotopic exchange occurred between the


817

molecular sieve and the 02 gas, the exchange appears to have been negligible due to the low temperature used to heat the sieve and the short duration of heating. The 13X was activated between samples by heating it to 105~ for about 40 minutes (M. F. Miller, pers. com. 1999). MacPherson et al. (1999) used zeolite 5A for storage of 02 used for combustion in their major volatile extraction system. Details of reactivation or trapping on the sieve were not given. Wassenaar & Koehler (1999) employed molecular sieve 5A to purify 0 2 from air, soil gas and water for isotopic analysis in a continuous-flow mass spectrometer (CFIRMS). Sample purification was carried out on an on-line modified Carlo Erba NA1500 elemental analyzer. The sample was contained in a He-carrier stream. After the initial purification steps, that removed H20 and CO2, the sample was passed through a I m GC column containing 5A. The GC column, held at 35~ was used to separate 02 from N2. Finally, the 02 was carried directly into an IRMS utilizing a He carrier gas where it was measured isotopically.

Nitrogen (N2): Nitrogen has been transferred to mass spectrometer inlet systems for isotopic analysis using zeolite molecular sieves. Macko (1981) used the closed-tube combustion technique to produce N2 from organic samples. The reaction also produced CO2, which was retained by freezing the sample tube in liquid nitrogen. The N2 was then transferred onto a zeolite 5A crystals (30 - 60 mesh) frozen at liquid nitrogen temperature for ten minutes (S. A. Macko, pers. comm., 1999) and released into the mass spectrometer by heating the sieve to 300~ for another ten minutes. For shorter periods of either freeze-down or bake-out of the sieve, the N2 experienced isotopic fractionation. During the freezing process, 15N was preferentially condensed onto the molecular sieve relative to 14N and thus too short a freezing period led to increased 615N values. Shorter bake-out period can lead to complex behavior with increasing time. 615N values first decreased, then increased and finally, decreased until the true value was reached. It thus appears that nitrogen is being released from different crystallographic sites within the molecular sieve. The fractionation observed by Macko (1981) resulted in 615N values that were within 0.5%0 of the true value. More recently, Macko (pers. com. 1999) reactivates the sieve by heating at 180~ under vacuum for a couple days. After transferring the N2 to the 5A cold-finger, he now releases the gas into the mass spectrometer at 150~ rather than 300~ This ensures that contaminants that come out of the sieve at higher temperatures are not released. Bebout & Fogel (1990), working with silicate minerals and rocks, compared Macko's molecular sieve method with a gas expansion technique that did not involve a molecular sieve. They found that the sieve trapped CO2 and H20 in addition to N2 and that the N2 blank was higher by 0.1 - 0.2 gmole when the sieve was used. However, heating the sieve to 400~ for 3 hours reduced the blank. Boyd et al. (1988) used a molecular sieve of an unspecified type in a stepped combustion/pyrolysis system designed to produce nmole quantities of N2 for isotopic

818

Chapter 36 - H.R. Karlsson

analysis in a static mass spectrometer. A cold-finger containing pellets of the 5 A pore molecular sieve was cooled in liquid nitrogen (- 196~ to collect gases produced during the combustion or pyrolysis procedure. The sieve was then heated to 200~ in one minute to release all trapped gases (N2, CO2, CO, CH4, N20 etc.) which were then oxidized over hot CuO (850~ to produce H20, CO2 and SO2 (N20 was converted to N2 and 02). Excess oxygen was readsorbed onto the CuO by lowering its temperature (600~ Following the removal of 02, the N2 is purified cryogenically by freezing the other gases onto a liquid nitrogen-cooled cold-finger. The molecular sieve was maintained at 300~ under vacuum between runs (typically overnight). Blanks for the pyrolysis/combustion system were reportedly better than 0.04 nmole suggesting a very low N2 blank for the molecular sieve.

Sulfur (S): Puchelt et al. (1971) employed zeolite 5A to remove Br2 from SF6 prior to sulfur isotope analysis. This was done by passing the gases through a 5 feet by 3 / 4 inch GC column packed with the zeolite and held at 100~ He gas carried the components through the GC and a liquid-nitrogen-cooled trap removed the SF6 from the carrier gas. Tests run on the GC with pure SF6 of known isotopic composition gave quantitative yields and unfractionated sulfur isotopes.

Silicon (Si): Molini-Velsko (1983) determined Si istotope ratios in meteorites using SiF4 as a working gas. The SiF4 was extracted from the samples using the BrF5 method of Clayton & Mayeda (1963). Purification of the SiF4, however, was necessary because contaminants such as SF6, SO2F2 and CF4 were also created. One of the purification methods that Molini-Velsko (1983) examined involved using a molecular sieve column in a gas chromatograph. Four types of molecular sieves were tested in the 3 feet by 0.85 inches chromatographic column: 4A, 5A, 13X (30 - 60 mesh) and CMS (Carbosphere, 60 - 80 mesh). The column was located inside a furnace, which allowed temperature control. Samples were carried through the column using He carrier gas. The 4A sieve failed to separate SiF4 and SF6 since they were too large to enter the zeolite pores. Attempting to open the 4A pores by heating up to 350~ had no effect. SF6 passed easily through 5A and 13X but SiF4 did not emerge even after several hours at elevated temperatures. The CMS (~ 13 ~ pores) eluted SiF4 ahead of SF6 but there was too much overlap between the two gases. According to Breck (1974) SiF4 is to some extent irreversibly sorbed onto 4A and 13X at 0~ suggesting that chemisorption took place. However, if these zeolites are held at 200~ it appears that SiF4 is reversibly sorbed.

Methane (CH4): Jackson et al. (1999a) used molecular sieves in a preparation system designed to handle small quantities of atmospheric CH4 for isotopic analysis in a static mass spectrometer. Air samples were carried through a gas chromatograph containing CMS (Carbosphere, 80/100 mesh) using a high-purity He stream, which had been cleaned by running it through a molecular sieve 13X (IMS-100) to remove H20 and CO2. As CH4 exited the chromatograph it was run through zeolite13X held in liquid nitrogen. Once CH4 collection was complete, non-condensable He was pumped away, CH4 and other trapped gases released from the 13X trap by heating, and the


819

gases transferred to a 13X cold-finger held at liquid nitrogen temperatures in the inlet section of the mass spectrometer. Once transferred, the CH4 was desorbed from the 13X and cleansed of N2 using getters (the eluted gas was only 10% CH4, the balance being N2). The purified CH4 was finally exposed to a trap held a t - 188~ to remove any trace of H20 that might have remained and admitted to the mass spectrometer. Analysis of a CH4 mixed in with the appropriate concentration of N2 yielded a precision of 0.3%o indicating that the 13X had little effect on the isotopic composition of the CH4.

Hydrogen (H2)"The author found no examples of H2 collection or trapping. Because of its low boiling point (Table 36.3), it is unlikely that H2 could be trapped completely on a molecular sieve unless liquid He was used. Preliminary experiments by the author suggest that only 90 - 95% of H2 would condense on zeolites (Linde Corp. ET200 and Oxysiv-5) at liquid nitrogen temperatures. Zeolites are thus similar to activated charcoal (see e.g., Halas & Durakiewiz, 1995) in terms of sorption capacity. According to Breck (1974), D2 is preferentially sorbed over H2 onto zeolite 4A. If this holds true for other zeolites and hydrogen gas adsorption is incomplete even a t 198~ then fractionation of hydrogen isotopes will take place in marked contrast to activated charcoal. 36.5 Conclusions

Molecular sieves have been used to 1) improve vacuum performance, 2) collect gases in the field, 3) separate gases, and 4) transfer gases within vacuum systems. Except for the first of these processes, it is essential that the isotopic composition of the gas of interest remain intact at the completion of the operation. Thus far, only a handful of the available molecular sieve materials have been utilized in isotope work. Rarely have molecular sieve a n d / o r trapping processes been used in combination although such an approach could be fruitful. For example, a small pore sieve could be used to separate small molecules from large ones (molecular sieving). A large pore sieve would then be used to separate the larger gas species according to boiling point (trapping). It is surprising that molecular sieves such as zeolites are not more widely used considering that they are very sturdy and non-toxic substances. Future studies are needed to investigate the use of other molecular sieves such as the high silica zeolites (silicalite, ZSM-types) and non-zeolitic material such as A1PO4s and CMS.

Acknowledgements I am grateful to Mrs. Mayeda and Drs. Clayton, Macko, Miller, Meier-Augenstein and Schwarz for providing further details in their use of molecular sieves. Early versions of the manuscript benefited greatly from reviews by Dr. John R. Beckett. Reviews by an anonymous reviewer and Drs. Bol and Schwarz improved the final version.


CHAPTER 37 Introduction to Isotope Dilution Mass Spectrometry (IDMS) Michael Berglund European Commission- Joint Research Centre, Institute for Reference Materials and Measurements (IRMM), Retieseweg, 2440 Geel, Belgium e-mail: [email protected]

37.1 The beginning The purpose of this chapter is to give the reader a short introduction to IDMS. The basic principles will be outlined and explained. The IDMS equations will be derived and explained. The main benefits of IDMS will be discussed as well as some pitfalls. There will be no detailed discussions on element, or technique specific subjects because the sheer number of topics would be impossible to cover in this chapter without being terribly incomplete. The use of IDMS by Inductively Coupled Plasma Mass Spectrometry (ICP-MS) has been covered to some extent in "Inductively Coupled Plasma Spectrometry and its Applications" (Vanhaecke et al., 1999b). Thermal Ionization Mass Spectrometry (TIMS) and IDMS has been discussed by Fassett & Paulsen (1989). The isotope dilution technique is used in many fields of analytical science, from biochemistry to geology. Despite the very varied applications of isotope dilution in existence they all have one thing in common, the addition of a spike to a sample material. In the most common case the spike consists of a known amount of an isotopically enriched element. The amount of this element in the sample material is then the measurand of the IDMS measurement. From a principal point of view the element in the spike only need to have a different isotopic composition relative to the same element in the sample material. This means that IDMS can only be used on elements with more than one stable isotope, and if these isotopes are measurable on a mass spectrometer, IDMS is in principle possible. So, what does this addition of an isotopically enriched element lead to? To answer this I would like you to make a mental switch and imagine Stanley meeting Dr. Livingstone in the village Ujiji at the Tanganyika lake in Africa in 1871. Little did they know that some 60 years later an entymologist named C. H. N. Jackson were going to study the density of tsetse flies in this very region (Jackson, 1933), and as a tool to count tsetse flies he used fly dilution. This is to my knowledge the first time the principle of what is now named isotope dilution was documented. Etienne Roth (1997) mentions that the dilution method was first used for evaluating populations of rare bird species on islands but no date or references are given. First or not first, let us look at a small scale simplified model of the 1933 African fly dilution.

821

Introduction to Isotope Dilution Mass Spectrometry (IDMS)

In Figure 37.1 below a number of marked flies, coloured black, are added to a normal population of white flies. I am quite sure Jackson did not colour his tsetse flies black but for the sake of argument in this example, normal tsetse flies will be white and marked, or coloured, flies will be black. After waiting for a suitable amount of time to allow for the black flies to blend into the white fly population a sample is taken. This is indicated in Figure 37.1 by the dotted line and seen in Figure 37.2. After Jackson collected the sample blend of flies in Figure 37.2 he calculated the true number of flies using equation [37.1]:

marked flies in sample unmarkerd flies in sample

total marked flies total unmarked flies

[37.1]

which after input of the example data gives: 3

10

total unmarked flies-

6 total unmarked flies

6.10

3

= 20

[37.2]

Is not it beautiful and simple? Do note that in the example above the worries of statistically representative samples and counts have been left out. The criteria for achieving a true count of tsetse flies using fly dilution is given in Table 37.1. Translate these criteria from flies to nuclides and you get the only criteria you need to master IDMS. It is not always easy and there can be numerous difficulties to overcome when applying IDMS. These difficulties can be either in assaying the spike material, i.e. knowing exactly how many black flies you added or achieving a homogenous blend after adding the spike. When you think about it, how did Jackson make sure he got a homogenous mix of marked and unmarked tsetse flies? He does not discuss this, he was more concerned with diffusion of flies into and from his fairly self-contained fly community. Jackson had his difficulties, in IDMS we have ours. The actual count-

. . . . . . . . . . .

......

:~

:.:::~

~ v ~;:~iii!i~i~':

g g ~i!i!!!~i!!~::

~i~

~

i'~i~

~ ~:.v::G::i:

.....

.............

....

..................&::,::::

:~i;::

§

....

Figure 37.1 - Fly dilution experiment. Z ~: ~(s i~i::!:.i!i~:

t

Figure 37.2 - Sample taken from the blend of flies. --)

.... ~=..~.

..... ,,~,N:~:~,:,,,~,~ ~.

.

.

.

~

822

Chapter 37 - M. Berglund

ing of nuclides is not always Table 37.1 - Criteria for an accurate fly dilution experiment straightforward. There can for example be isobaric interfer- 1 you need to know how many flies you have marked and released ences, which can be seen as 2 you need to know that there is a homogenous blend of additional species of flies, non white and black tsetse flies tsetse flies, getting caught in 3 when you start counting flies in your fly mix you must be the sampling process and able to identify and separate black and white tsetse flies wrongfully counted as either 4 you must not loose your count white or black tsetse flies. Another effect that will render an erroneous count of the sampled blend, is the blank contribution. In the fly case this means that you are counting more white flies than you actually sampled in Figure 37.2. They can have entered your little net while you happily walked back to the camp in the African jungle, passing areas with a normal tsetse fly population. Maybe a few landed on your table, drawn there by the light in your tent in the pitch-black African night while you were busy counting.

37.2 At the highest metrological level IDMS has been recognized as a primary method of measurement by the Comit6 Consultatif pour la Quantit6 de Mati6re (CCQM)I or as it is named in English the Consultative Committee for Amount of Substance. In every day life everyone refers to it as CCQM. CCQM is part of the Comit6 International des Poids et Mesures (CIPM) whose main function is to ensure the propagation and improvement of the SI system. The primary method was defined by CCQM in 19952 and slightly redefined in 19983. A primary method of measurement is a method having the highest metrological qualities, whose operation can be completely described and understood, for which a complete uncertainty statement can be written down in terms of SI units. The two terms primary direct method and primary ratio method were introduced in 19983 to better incorporate IDMS since it is not a direct method like coloumetry. In the explanatory notes we read: A measurement traceable to the SI can be made using a primary ratio method in combination with a reference of the same quantity that is itself traceable to the SI. However, a method whose operation cannot be completely described and understood cannot be a primary ratio method. The fact that IDMS has been recognized as a primary method of measurement is not only due to its transparency, which is the qualifying criterion, it has also proven itself in the line of duty. In Figure 37.3 we see CCQM's second key comparison (Papadakis et al., 2001)4 exercise where all participating national metrology institutes (NMI) were using IDMS. In the same figure, measuring identical samples, we see results from the participants to IMEP-9 (Papadakis et al., 1999)5. IMEP is a tool with which field labo1. 2. 3. 4. 5.

http: / / www.bipm.fr / enus / 2_Committees / CCQM.shtml Report of the 1st meeting of CCQM, 1995 Report of the 4th meeting of CCQM, 1995 http://kcdb.bipm.org/BIPM-KCDB / default.asp http://www.irmm.jrc.be / imep


823

Figure 37.3

824


ratories can compare their results against SI traceable values. The participants to IMEP-9 work under normal conditions of their choice, with respect to technique, instrumentation etc. The grey bar in Figure 37.3 is the certified range for IMEP-9. The agreement between the NMI laboratories in CCQM-K2, and other CCQM key comparisons, reflects the state of the art in chemical measurements. IDMS has become the most important technique for accurate measurements of amount content and at the highest metrological level, which the key comparisons represent, it is the method of choice. Even though IDMS has the capacity to yield very small uncertainties it must be pointed out that it does not come automatically. IDMS can be misused as any other technique. It will not surprise you, but very careful planning and correctly applied measurement procedures are needed for optimum performance. For further information on accuracy of IDMS look at De Bi6vre (1990) who has written an article on the accuracy aspect of IDMS. OK, IDMS is transparent and can be very precise, but what makes it such a solid technique is the fact that it works even if we have non-quantitative recoveries from for example digestion or separation steps. It is also independent of instrument sensitivity. This is due to the fact that, with or without sample losses, whether the instrument is perfectly optimized or not, the measured amount ratio of two nuclides of the same element, is always the same. This of course provided that adequate mixing of sample and spike has occurred.

37.3 Deriving the IDMS equation The IDMS equation derived below represents the simplest case. On the other hand it is the building stone for all other variations. We are interested in determining the number of atoms nx(E) of a specific element, E, in a sample, denoted by the subscript x. An isotopically enriched spike with subscript y will be used to prepare a blend of sample and spike. This blend will have a subscript B. The term nuclide will now be dropped in favor off isotope, which is less correct, but is the commonly used term in the mass spectrometric community. So, let us first define some basic relationships"

nx(E ) - ~ nx(iE)

[37.3]

where the superscript i denote an isotope of the element. In the same way the number of atoms in the enriched spike, denoted by the subscript y, will be:

ny(E) - ~ ny(iE)

[37.41

One isotope, usually the most abundant in the sample, is selected as reference isotope and all isotope amount ratios are expressed relative to it. The reference isotope and the most abundant isotope in the spike, are then normally selected to give the master amount ratio. If we now make a blend of sample and spike we can define an isotope


825

Figure 37.4a-c- See text for explanation.

amount ratio of two well-chosen isotopes, aE and bE, as:

RB -

nx(bE) + ny( bE)

nx(aE) + ny(aE)

[37.5]

The reference isotope is thus aE and the master amount ratio to monitor is n(bE)/ n(aE). Let us look at an example. In Figure 37.4ac we see the isotope amount fractions of a hypothetically prepared blend for an IDMS measurement using magnesium as an example. In Figure 37.4a-c we see the sample '4a', with natural isotopic composition, and the spike '4b', enriched in 26Mg, and the blend '4c', which is then a mixture of sample and spike. The most suitable master amount ratio in this example would then be n(26Mg)/n(24Mg). In most cases it is advisable to strive for an RB close to 1 because this can cancel out possible non-linearities in the detector system. However, in many cases this is not possible because of for example, low enrichment in the spike available or a specific choice of master amount ratio is desirable. In these cases a

826


thorough a priori uncertainty analysis should be performed. An investigation of the influence of measurement precision on the choice of blend ratio has been done by De Bi6vre (1994). To find out the exact number of atoms of element E in the sample we will solve for nx(E). We would like to express it in terms of isotope ratios, because those we can measure. Useful as a start are then the following definitions:

Rx Ry

_

nx(bE)

[37.6]

nx(aE) (bE

n

Y

)

[37.7]

ny(aE) ~ny(iE) l

~Riy -

aE

[37.8]

ny( ) ~nx(iE) Rix - ' nx(aE)

[37.9]

where Rx and Ry are the amount ratios of isotopes bE and aE in sample and spike respectively. YRix and YRiy a r e the sums of amount ratios for all isotopes of an element relative to the reference isotope for sample and spike respectively. Now we have all the necessary ingredients. Now it is just a matter of an algebraic rearrangement. Inserting equations [37.6], [37.7], [37.8] and [37.9] into equation [37.5] gives:

RB -

nx(aE) " Rx + ny(aE) " Ry

~nx(iE) t

~ny(iE) +

~iix

aB -

t

~iiy

[tlx(aE) " ax 4- Fly(aE )" ay] " ~ a i x " ~ a i y

~nx(iE)" ~Riy + ~ny(iE)" ~Rix l

RB'I~nx(iE)'~Riy l

l

+~ny(iE)'~aixl

- [nx(aE).Rx+ny( aE ).Ry]. ~ Rix " ~ Riy

l

inserting equations [37.8] and [37.9] into the equation above gives"

827


R B'I~nx(iE)'~Riy +

iE . RixI -

l

I~ Flx(iE) ~Fly(iE) [ ~-a2 "Rx+ aiy

1

_

ix

and if we now insert equations [37.3] and [37.4] in the resulting equation above we get:

nx(E) ny(E) ~ RB'[nx(E)" Riy 4-ny(E).~aix ] - I ~-R~;'Rx+ ~aiy'Ryl" ~Nix" ~aiy now rearrangements and simplifications will give us the IDMS equation

[nx(E) . ~ R i y . R B + ny(E) . ~aix

. aB] - [nx(E) . ~Riy. ax 4-ny(E) . ~aix . ay]

nx(E)'[~Riy'RB-~Riy" Rx] - Fly(E) . [~Rix . Ry- ~Rix. RB] nx(E ) _ ERix. i y - ERix. i B Ry-RB

nx(E)

-

ny(E)-RB_ Rx" ~iiy

[37.10]

In equation [37.10] we have managed to express nx(E) in terms of measurable or given quantities, isotope amount ratios, R, and amount of added spike, ny(E). In the easiest case where the masses of spike and sample are determined gravimetrically, mx and my in (kg), and cy the amount content (mol/kg) of the spike is given, which is the case if a certified spike is used, it will turn equation [37.10] into:

cx

_

cy

.my.[Ry-RB]

mx [RB--Rxx]

~Rix 9

[37.11]

In scientific journals it is common to see equation [37.11] with added input quantities, correcting for various effects, and of course other names for the input quantities. Every now and then you will also see a different looking IDMS equation like in Fassett & Paulsen (1989). This equation can, however, with some simple algebra be turned into equation [37.11]. There is a drawback with the Fassett & Paulsen (1989) equation and that is that amount fractions are used as input quantities. Amount frac-

828


tions of isotopes of the same sample are correlated and this complicates the uncertainty calculation. 37.4 Mass fractionation

correction

In the equation above Cx is the amount content of element E in the sample. Equation [37.10] or [37.11] is what you will find in most articles and textbooks covering IDMS. It may look simple, but there is more to this equation than meets the eye. Following the derived IDMS equation it is clear that the ratios, Ri, given in equation [37.10] and equation [37.11] are absolute amount ratios and not measured ratios. Mass discrimination is a well-known effect in mass spectrometry and is caused by, for example, evaporation, diffusion and electrostatic effects. These effects are mass and time d e p e n d e n t a n d the result is, that in the mass spectrometer, different isotopes do not race on equal terms. Heavier, or lighter, isotopes will, due to the effects mentioned, be favoured which will lead to non-absolute isotope amount ratios. This correction factor, K, is measured using a certified isotopic reference material. The assumption is that the quotient of the measured ratios of sample and reference are equal to the quotient of the true amount ratios of sample and reference. This is a valid assumption if the mass fractionations, fms and fmR, for the measured ratios rs and rR respectively are equal. The subscript R denotes a certified reference material and the subscript S the sample in question. To calibrate a measured ratio an isotopic reference material traceable to the mole must be used. Tables of available reference materials meeting this criterion are listed in Part 2, Chapter 40. For the elements not present in this table the IUPAC Table (Rosman & Taylor, 1998) can be used. This does not mean it is impossible to do IDMS on elements not in the table. We just introduce a small exception. For the IDMS equation to work a mass fractionation correction factor must be applied. An exception can be when the K-factor is the same for two or more measured ratios where some cancellations are possible. fmS" rs fmR rR 9

Rs RR ~ Rs

~

RR

rs ~ Rs - K r s rR ,

,

if

RR K - rR

[37.12]

In equation [37.12], and henceforth, a small r will denote a measured amount ratio and a capital R the absolute isotope amount ratio. This would turn equation [37.11] into Cx _ Cy . m y . [ KRy 9ry - KRB 9rB] 9~ (K i 9rix ) mx [ K R B ' r B - KRx "rx] ~ ( K i ' r i y )

[37.13]

Every measured isotope amount ratio for element E has its own K factor. In equation [37.13] KRB, KRx and KRy are all correction factors for the measured master ratio nx(bE) / nx(aE) but it does not necessarily mean that these mass fractionation correction factors are identical. In ICP-MS mass fractionation can be very time dependent and matrix dependencies of K are seen for both TIMS and ICP-MS. In the sum of ratios, all amount ratios have different K-factors. Since the same reference isotope has to be used

829


for all amount ratios the master ratio, in our case nx(bE)/nx(aE) will enter the sum of ratios as well.

37.5 An IDMS example If we take a look at equation [37.13] using a real element, chromium, as an example we first need to define the reference isotope and the master ratio. The most abundant isotope of chromium is 52Cr so let us use this one as reference isotope. If we use a spike isotopically enriched in 50Cr the master ratio would be n(50Cr)/n(52Cr). This would turn equation [37.13] into equation [37.14]" \

my.

Cx - Cy" mx

IKRy'r50-KRB'rB501

~ (Kx50 9r 50 + 1 + Kx53 9r 53 + Kx54" r 54} x~ x5~ x~ /

IKRB'rn50-KRx'rx501

~ (Ky50 9r 50 + 1 + Ky53. r 53 + Ky54" r 541 Y~-~ Y~-~ Y~-~"

Y5-2

~

[37.14] Here the indexes x,y and B denotes sample, spike and blend respectively. The indexed ratios (i.e. 50/52) indicates the isotope ratio. For the K-factors the indexes indicates the measured ratio they are correcting. Now someone might say: but look, this is a lot of ratio

measurements, not only do I need to measure rx, ry and rB, every K-factor also carries a ratio measurement. Well, normally you never measure all of these because some are given in a certificate, if you use a certified spike material and often IUPAC values (Rosman & Taylor, 1998) are used for isotope amount ratios where the assumption is that the amount ratio in question has a natural isotopic composition. Do not forget that assumptions normally carry an uncertainty. Assuming that a sample has an isotopic composition that is natural and represented by the composition given by IUPAC can be wrong. As a contrast let us look at the simplest case. If we in this example use a certified chromium spike and assumes that the sample material has a natural isotopic composition equation [37.14] would turn into:

cx

Cy my I1 L R y - K R B ' r 50_ . . . . . . ~J ~ mx IKRB " r 50 - Rxl ~

R

(Riy )

where

KRB -- r'~5R

[37.15]

K-~

where only the weighings of mx and my and the measurements of the amount ratios rB and rR need to be performed. The amount content Cy, Ry, and all Riy are taken from the spike certificate and Rx and all isotope ratios for the sample, Rix, are taken from IUPAC. From my experience it is always beneficial to measure rx, rix and relevant Kfactors, even if the isotope amount ratios can be assumed to be close to values given by IUPAC. First of all, mistakes will be avoided and there will also be fewer assumptions to worry about w h e n it comes to uncertainty budgeting. However, there may be cases with severe interference on some masses with the result that all rix cannot be

830


measured without introducing very large uncertainties. In these cases it may be better to use IUPAC values. 37.6 D o u b l e I D M S

Equation [37.10] and the derived equation [37.11], require a well-defined spike. As you may remember it is a must to know how many atoms of the enriched spike you have added, see criterium 1 in Table 37.1. The availability of such certified reference material, enriched in a specific isotope with a certified amount content and isotopic composition is somewhat limited. To overcome this, 'double' IDMS is frequently used. The idea here is to use a material of natural isotopic composition as primary assay standard. Do note that there is no true natural isotopic composition of an element. The isotopic composition of an element varies in nature, unless it is mono-isotopic of course. Sometimes this variation is measurable and sometimes it is not. With every generation of mass spectrometric instrumentation the trend is that isotopic variations are detected for more and more elements. To perform double IDMS we need to make another blend, here called blend B'. For blend B', we use a well-characterised primary assay standard with the amount content Cz and our enriched spike material y. The subscripts x, y and z denotes sample material, spike material and assay material respectively. Let us first set up an equation, equivalent to equation [37.10] but with the assay material (z) instead of the sample (x). c z - cy m ' y . [ R y - R B , ] 9 mz iG;---R-z]" ~ R i y

[37.16]

All quantities in equation [37.16] are equivalent to those in equation [37.11] except that it concerns assay and spike (z, y) instead of sample and a spike (x, y). The mass of spike y for the blend assay and spike is denoted m'y to distinguish it from the mass of spike my used in equation [37.11] If we now divide equation [37.11] with equation [37.16] we get:

CX

Cz

9m y . [ R y - RB] " ~ , R i x Cy m x [ R B - R ] x ~ R i y

Cy"

m ' y . [Ry - RB,]. mz

[R B , - R z ]

~Riy

We can directly see that the amount content of the spike, Cy, and the sum of spike ratios, l~Rip are cancelled out. Reconstructing gives:

C x _ Cz"

m .m

z. m x 9m y [R B - R K ] ' [ R y -

. RB, ] I ~ R i z )

[37.171


831

Very often equation [37.17] is used with the ratio of the sum of ratios equal to unity. This can only be done if the isotopic composition of the sample and assay materials is identical. In most cases when this model is used the sample and assay materials have a natural isotopic composition. The question is can we cancel out the sum of ratios without introducing a bias? To decide that the isotopic compositions are equal, without measuring, requires a good knowledge of natural isotopic variations of the particular element and knowledge of the origin of the sample and assay materials. If in doubt leave it in the equation and measure YRix and ZRiz. In equation [37.17] only calibrated ratios are used and mass fractionation correction factors need to be applied, see equation [37.12] and [37.13].

37.6 Isotope specific IDMS In nuclear mass spectrometry very often the target is not the total amount of an element. Here the total amount of a specific isotope is more interesting. Combining equation [37.5], [37.6] and [37.7] we get equation [37.18] or equation [37.19] depending on preferred master ratio. The master ratio will of course consist of the enriched isotope in the spike and the sought for isotope in the sample.

Ry-RB. 1

Flx(aE) - Fly(bE). G

~

-ay

[37.18]

n x (bE) - rly(aE) 9R G y- R- xR B . Rx

[37.19]

37.7 Some difficulties with IDMS IDMS will work perfectly unless you fail one of the criteria in Table 37.1. Let us look at these criteria in more detail and transfer them to mass spectrometry and see where possible problems may show up.

37.7.1 How many enriched atoms have been added? There are some spikes (Table 41.2, Chapter 41), certified for both isotopic composition and amount content available that are made for IDMS measurements. If you do not have access to any of these, or are measuring on an element where none is available you will have to follow the procedure outlined in Double IDMS.

37.7.2 The necessity of a homogenous blend of sample and spike. This is regarded as one of the most critical steps in IDMS, and this for a good reason. Let us assume we have a digestion procedure and it is not 100% efficient, which is not that unlikely. If the blend is digested, some of the sample atoms of element E may still be bound in a complex of some kind resulting in a measured blend ratio that does not represent the true amount ratio of n(bE)/n(aE) in the blend. There could also be a disproportionate amount of a volatile complex of element E formed due to the initial matrix differences between sample and spike, which would also render an erroneous blend amount ratio. Even if digestion is the most common pre-treatment in inorganic chemistry this problem holds for any pre-treatment of the blend. To achieve an accurate result, losses from blend pro-

832


cessing have to affect element E in the added spike exactly the same way it affects element E in the sample. The best way to identify a digestion problem is to look at the blend to blend reproducibility. If this is alarmingly poor in comparison to the uncertainty of a measurement, it is an indication of a problem in your measurement procedure and the digestion is where you should look first.

37.7.3 Mass spectral interferences and contamination. Mass spectral interferences can be tricky and possible interferences have to be checked for every mass used in the measurements. There are only two ways of handling mass spectral interferences and that is either by estimating the interference, e.g. by measuring, and apply a proper correction, or remove the interfering compound. In IDMS it is also very important to be aware of the contamination risks when working with enriched material. Not only is it easy to contaminate the enriched material but also to change the isotopic composition of sample, blends and standards. It is crucial to realise that contamination on an isotopic composition level is a real danger in IDMS. 37. 7.4 Counting ions. Counting ions is usually done by secondary electron multipliers (SEM) or faraday cups backed up with the appropriate electronic gadgetry. While faraday cups are known to be precise, stable and linear they are much less sensitive than the SEM. The SEM on the other hand, sensitive as it is, is of a more delicate nature. There are three important parameters that need to be controlled when it comes to SEMs and that is dead time, trigger level and the plateau. The dead time is the time the SEM is blind to incoming ions because it is already busy handling a signal pulse. Dead time has been thoroughly investigated by Rameb~ick et al. (2001) and practical ways of correcting for dead time has been handled by Nelms et al. (2001) and Appelblad & Baxter (2000). The trigger level, or discrimination level, is the signal or peak height needed to trigger a count. It is important to set the trigger or discrimination level so that double counting caused by ringing is avoided. For more information look at Hunter & Gray (1993). The plateau is the flat region of the characteristic count rate vs. high voltage curve where a change in the output count rate is small relative to variations of the high voltage of the multiplier. An example of the sensitivity of the plateau can be seen in Figure 37.5. In Figure 37.5 we see change of K-factor versus SEM high voltage (circles) and the plateau curve, signal counts for 187Re versus SEM high voltage (rectangles). The multiplier used for obtaining Figure 37.5 was a nine month old MC-Z-19-TRITON from MasCom, Germany. The instrumental settings were similar to the settings in Rameb/ick et al. (2001). Please note that the data points in Figure 37.5 were obtained in random order. Plateau checks are a convenient tool to determine the optimum high voltage setting for the SEM detector system and it is obvious from Figure 37.5 that the high voltage need to be stable to avoid drifts in the K-factor.


833

Figure 37.5 - SEM high voltage vs. mass fractionation and ion counting (H. Rameb~ick,

pers. com.).

37.8 U n c e r t a i n t y B u d g e t

Because of its transparency IDMS can easily produce traceable results. The way to demonstrate traceability is to use an uncertainty budget, where all parameters influencing the final result are presented together with their uncertainties. Table 37.2 is an example of an uncertainty budget. In this case it is an IDMS measurement of the amount content of C1 in a 37C1 enriched material. In this measurement (M. Ostermann, pers. com.) the reference material IRMM-641, which has a natural isotopic composition, was used as spike. The model equation used is described in equation [37.20]. 9my. [ R y - K . r B ] 1 + K'rx35/37 9 c x - cy mx [ K . r B - K . r x ] 1 + Ry35/37

where

K -

Ry ry

[37.20]

In this example Cy and Ry are taken from the IRMM-641 certificate and rB, rx35/37, ry35/37, mx and my are measured. An uncertainty budget must contain the model equation used in the calculation of the result as well as the input quantities and their uncertainties. The uncertainty is calculated using the concept of error propagation and there are a couple of ways to do this. There is the simple to use spreadsheet approach described by Kragten (1994) but for larger model equations the spreadsheet can get difficult to manage. The help can then be special uncertainty software. There are a few around and it can be the way to go. There is of course the strict mathematical approach, which is used by people who either love partial derivatives and have nothing better to do, or old school know-it-alls. The level of detail in modern IDMS equation modeling, with replicates and blanks, is way beyond paper and pen uncertainty

834


budgeting so save Table 37.2- Uncertainty budget yourself the trouble Value and use one of the Quantity other options. GUM1 was published in 1.13624 1995 and is the referrB 2.50.10-5 mol/g Cy ence work on estiRy 3.1272 mating and reporting rx35/ 37 0.018321 uncertainty. A more ry35/37 3.11291 2.51340 g easy to read guide on mx 1.02090 g uncertainty is the my 4.434-10-6 mol/g EURACHEM guide2 cx which is made for chemists by chemists.

Standard Uncertainty 0.0014 5.50.10-9 mol/g 0.0041 0.00037 0.0031 0.00029 g 0.00012 g 0.012-10-6mol/g

Contribution to total uncertainty 48.1% 0.7% 12.7% 6.5% 31.6% 0.2% 0.2%

Acknowledgements The author would like to acknowledge discussions and exchange of ideas with previous and present colleagues at the isotope measurement unit at IRMM and especially Henrik Rameb~ick, Markus Ostermann, Philip Taylor and Paul De Bi6vre. Helpful comments on the manuscript were made by Bob Loss and Wolfgang Frech.

1. GUM, Guide to the expression of uncertainty in measurement, ISO 1993, ISBN 92-67-10188-9 2. Quantifying Uncertainty in Analytical Measurement, 2nd edition (2000) EURAcHEM: http://www.eurachem.bam.de

PART 2 Calibration and Correction Procedures, Standards, Mass Spectrometers, Experimental Isotope Fractionation Determination and General Information



CHAPTER 38 Mass Spectrometer Hardware for Analyzing Stable Isotope Ratios Willi A. Brand Max-Planck-Institute for Biogeochemistry, PO Box 100164, 07701 Jena, Germany e-mail: [email protected]

Abstract Mass spectrometers and sample preparation techniques for stable isotope ratio measurements, originally developed and used by a small group of scientists, are now used in a wide range of fields. Instruments today are typically acquired from a manufacturer rather than being custom built in the laboratory, as was once the case. In order to consistently generate measurements of high precision and reliability, an extensive knowledge of instrumental effects and their underlying causes is required. This contribution attempts to fill in the gaps that often characterize the instrumental knowledge of relative newcomers to the field.

38.1 Introduction Since the invention of mass spectrometry in 1910 by J.J. Thomson in the Cavendish laboratories in Cambridge('parabola spectrograph'; Thompson, 1910), this technique has provided a wealth of information about the microscopic world of atoms, molecules and ions. One of the first discoveries was the existence of stable isotopes, which were first seen in 1912 in neon (masses 20 and 22, with respective abundances of 91% and 9%; Thomson, 1913). Following this early work, F.W. Aston in the same laboratory set up a new instrument for which he coined the term 'mass spectrograph' which he used for checking almost all of the elements for the existence of isotopes. Aston not only confirmed the neon findings, he also discovered 21Ne which has only a 0.3 atom% abundance. During his scientific career, Aston discovered 212 out of the total 287 naturally occurring isotopes (Aston, 1942). This work brought new order into the periodic table of the elements which had previously been troubled by irregularities between atomic weight and chemical properties of the elements. Aston showed that the isotopic masses are not simple integral masses of a basic nucleon but rather that there is a mass defect that is related to the binding energy of the nuclei. Both J.J. Thomson and F.W. Aston were awarded Nobel Prizes for their achievements (Physics in 1906 and Chemistry in 1922, respectively). In general, a mass spectrometer is used to make a quantitative assessment of the contents of a given sample. The quality of the analysis thus depends on the ability of the mass spectrometer to detect all components of a sample with the same constant

8 36

Chapter 3 8 - W.A. Brand

sensitivity, irrespective of the complexity and chemical nature of the sample.This ideal mass spectrometer does not exist. Instead, the contents of a given sample have to be transformed into something which can be manipulated, separated and detected. In mass spectrometry, ions serve this purpose. The ability to quantitate the contents of a given sample is facilitated if sample complexity is reduced through separation of the individual chemical components prior to the measurement. This principle has led to the extensive use of separation devices (chromatographs) combined with mass spectrometers (as detectors) in chemical analysis. This combination has more recently been used for determination of the stable isotope ratios of the bio-elements (C,N,O,S and H)(Brand, 1996), alongside the more familiar method of isotope ratio measurement by high precision comparison of purified gases in the dual inlet system. A stable isotope ratio mass spectrometer consists of an inlet system, an ion source, an analyzer for ion separation, and a detector for ion registration. The inlet system is designed to handle pure gases, principally CO2, N2, H2, and SO2 but also others such as 02, N20, CO, CH3C1, SF6, CF4, and SiF4. Neutral molecules from the inlet system are introduced into the ion source, where they are ionized via electron impact and accelerated to several kilovolts, and then separated by a magnetic field and detected by Faraday cups positioned along the image plane of the mass spectrometer (Nier, 1940). The principles guiding the design and operation of each of these individual sections of the mass spectrometer are described and discussed in sequence.

38.2 Inlet System Design Inlet systems for gas isotope mass spectrometers are rather simple and clean devices consisting of valves, pipes, capillaries, connectors, and gauges. Home made inlet systems are often made of glass, but commercially available inlet systems are mostly designed from stainless steel components that have no cavities. All components and surfaces are carefully selected for maximum inertness towards the gases to be analyzed. The materials used as components of the valves deserve special attention. The highest quality valves are of 'all-metal' design, with all wetted surfaces made either from stainless steel (the body and membranes) or from gold (the gaskets or seals and the valve seat). The heart of the inlet system is the "Changeover Valve' (Figure 38.1). It was first described in 1947 by B.F. Murphey, who was studying thermal diffusion in gases. The Changeover Valve allows the inlet system to alternately switch within a couple of seconds between two gases which enter in turn into a vacuum chamber (e.g. the mass spectrometer). The gases are fed from reservoirs to the Changeover Valve by capillaries of around 0.1 mm i.d. and about lm in length with crimps for adjusting gas flows at their ends (Honig, 1945; Nier, 1947; Halsted & Nier, 1950). While one gas flows to the vacuum chamber, the other is directed to a vacuum waste pump so that flow through the capillaries is never interrupted. Without capillaries, a flow directly from a reservoir through an orifice into the mass spectrometer would be a direct effusion into vacuum, which would result in a change in isotopic composition over time. The forward flow of gas in the viscous flow regime through the long capillaries prevents the isotopic diffusion profile from penetrating from the crimp back into the sample reser-

Mass Spectrometer Hardware for Analyzing Stable Isotope Ratios 837

Figure 38.1 - Essential components of a 'Dual Inlet System' for gas isotope ratio mass spectrometry. For clarity the pumping infrastructure for the variable volume reservoirs and for the pipework has been omitted from the figure. The 'Changeover Valve' is an arrangement of four valves that allows gas from one reservoir to flow to the mass spectrometer while the other goes to waste. The design of the changeover valve deserves special attention in order to minimize cross talk between the gases.

838

C h a p t e r 38 - W . A . B r a n d

voir (Halsted & Nier, 1950; Habfast, 1997). In 1950, C.R. McKinney et al. applied the Changeover Valve principle to isotope ratio measurements. With their system, McKinney and coworkers were able to measure the stable isotopes of oxygen in 02 and of both carbon and oxygen in CO2 with a precision of about 0.1 per mill (g-notation1). To achieve such high precision, instrumental drifts occurring during measurement need to cancel almost quantitatively. To achieve this goal, the gas reservoirs on either side of the Changeover Valve are normally stainless steel bellows (formerly, mercury pistons) that allow the ion current signals of the two gases to be precisely balanced. Any non-linearity, temperature dependence of electronic components, or changes in sensitivity of the ion source or the magnetic field thus tend to cancel. By comparing the two gases several times within minutes it was possible for McKinney et al. to achieve the reported high precision. The McKinney instrument provided the basis for the "classical" procedure for high precision stable isotope ratio measurements. Its principles have survived for 50 years with little change and they will provide the basis for ultimate precision isotope ratio determination into the foreseeable future. The smallest amount of sample that can be analyzed using the dual inlet system is limited by the requirement to maintain viscous flow conditions. As a rule of thumb, the mean free path of a gas molecule should not exceed 1/10 th of the capillary dimensions. With the capillary dimensions of 0.1 mm i.d., the lower pressure limit for viscous flow and thus accurate measurement is about 15 to 20 mbar. When trying to reduce sample size, it is necessary to concentrate the gas of interest into a small volume in front of the capillary. For practical reasons, such a volume cannot be made much smaller than 250 gl. For condensable gases, it is shaped into a cold finger to be operated as a cryotrap at liquid nitrogen temperature under molecular flow conditions. Using the ideal gas law, the product of pressure and volume yields the smallest sample amount that can be accurately analyzed in a microvolume inlet system to be about 5 bar~l or 220 nmol of clean gas. Because real life samples rarely are the clean gas species used in the dual inlet system, each sample, be it a carbonate, a water sample, a lentil or a piece of tree ring, must be converted into the required simple gaseous form prior to analysis. There is a wide variety of specialized sample conversion and inlet systems including manually operated devices whose output must be manually introduced into the inlet reservoir and automated devices that deliver the final product gas directly to the inlet system under computer control. Other chapters in this book cover the various forms and experimental challenges of sample preparation for high precision isotope ratio determination.

1. ~ [%o1= (Rsa/ Rref - 1) 91000 {for 13C: RSa= 13C/12Cion current ratio of sample gas}

[38.1]

839

Mass Spectrometer Hardware for Analyzing Stable Isotope Ratios

38.3 The Ion Source: Electron Impact Ion Production

Wishful thinking:

If we could only sit and watch the molecules directly distinguishing their different weight through some colorful property, we could calculate isotope ratios just by counting1. This would, however, be a tedious and time consuming task because of the large number of particles required for high precision. 38.3.1 Basic principles

Sample molecules enter the ion source of the mass spectrometer from the inlet system in gaseous form. Here, some of them are ionized by b o m b a r d m e n t with electrons

(Electron Impact, El): M+e-~M+,

+2e -

[38.2]

The efficiency of this process determines the sensitivity of the mass spectrometer. It depends on the ionization cross section, the number of electrons, and the number of molecules presenting themselves to be ionized. Following ionization, the M + ~ molecular radical cation can further fragment into several pieces (e.g. CO2 +o ~ CO + + O.), depending on the internal energy the ion has acquired during the ionization process and the possible reaction pathways. The result of such unimolecular reactions is the mass spectrum of a chemical compound. More specifically, the fragments that form in the ion source within about one microsecond following ionization comprise the mass spectrum. Later reactions give rise to what we refer to as 'metastable' ions. As an example, there is a broad peak at mass 17.8. CO2 +o molecular ions which were accelerated as mass 44 but decayed to CO+ (mass 28) in front of the magnet arrive at the detector plane at mass position m* - m22 / m l . Here, m* is the apparent mass position (17.8), m2 and ml are the mass positions of the daughter (28) and parent ion (44), respectively. 38.3.2 Ion Source Schematics Figure 38.2 is a schematic representation of an electron impact ion source. Electrons are released from a hot filament made from tungsten, rhenium or thoriated iridium and accelerated by electrostatic potentials to an energy between 50 and 150 eV before entering the ionization box. Their velocity, v, can be calculated according to"

v-

,,/2eU/m

[38.3]

where e - elementary charge, U - accelerating potential, m - mass of the particle. The velocity of 100eV electrons is about 6 ~ 108 cm/s. Thus they traverse the ion box in about 2 nanoseconds. The molecules appear virtually motionless because they 1. To avoid confusion: Ion counting is also a special technique using fast secondary electron amplifiers with amplification up to 108 together with time and threshold discrimination techniques. It indeed is a powerful tool e.g. for measuring small abundances of isotopes in thermal ionization mass spectrometry.

840

Chapter 38 - W.A. Brand

Figure 38.2- Schematic layout of an Electron Impact (EI) ion source for gas isotope ratio mass spectrometry. The insulating spacers that also provide an enclosure for the whole source are omitted for clarity.

are moving in the ion source at thermal velocities of only about 3 9 104 cm/s. A homogeneous magnetic field of 100 to 500 Gauss is used to keep the electrons on a spiral path (to increase the ionization probability) through the ionization box effectively confining the ionization region to a diameter of RF, the imbalance signal generated by the bridge is positive and has a defined magnitude. This signal is amplified by the S/H and held at the S/H output, when the negative potential taken from the output of OA turns to positive. The comparator OA produces a positive output voltage if the generator signal is below the voltage held by S/H. In this fraction of the cycle, transistor T conducts. Hence the power supply, V1, is connected in parallel to RF (through T, which plays the role of a switch). No significant current from V1 can pass through the remaining parts of the bridge due to presence of the diode D. Because V1 considerably exceeds Ve, the excitation current cannot pass through the low-resistance legs of the bridge (Rv and RF) when RF is supplied from V1. This is the reason for using a sample-and-hold amplifier instead of a normal operational amplifier. As result of the action of the circuit shown in Figure 39.5; a series of heating pulses is supplied to RF with frequency driven by the generator but their duration is driven by the output voltage of the S/H amplifier. The bridge is always kept close to the balance state, i.e. RF - Rv. If for some reason (e.g. voltage V1 starts to diminish) RF resistance becomes somewhat lower than Rv, then the bridge imbalance signal becomes somewhat higher, which results in a longer duration of the heating pulses. A complete circuit diagram is described by Halas & Durakiewicz (1998b).

869

Techniques of Ion Current Stabilization in Isotope Ratio Mass Spectrometry Figure 39.5 - Conceptual diagram of using the Halas-Kaminski bridge for constant-resistance operation of a filament RF. S / H is a sample-and-hold amplifier; G is a triangular wave generator.

N_

S/H

,," ,,,,-"

o,,,d",.,",,d,,,,,c,,d ,. .,."

R

/'.

;x./

-',.-"%,%

"--,.\.

+ ~ "

I

No

",,% ,,, ",. "%",% ", ,, %., "',,,

"de"

/

I G A/V'~]

% ",,%

Ve

---tl ,.

"%:q

I b,l

I",J

"'.....

,,,,,,".--

D ,z' .,."

F- 4 ~

Vl

lt---~ '1

+

T

The use of a simple voltage stabilizer for temperature control of a filament was described by Halas et al. (2001). Voltage may be stabilized typically by use of the same two wires to feed voltage to the filament and measure the voltage value (2-wire method) or, as it is described below, by use of a separate pair of wires for voltage measurement. In such a 4-wires configuration the voltage drop along the supply lines does not influence the value measured by the stabilizer circuit, hence the fluctuation of the ion beam is significantly reduced. The schematic diagram of the circuit is shown in Figure 39.6. The MA741 operational amplifier with the negative feedback loop constitutes the basis of the stabilizer circuit. The reference voltage supplied to the noninverting input of the amplifier may be set either manually by use of the 10k potentiometer, or digitally by use of the digital-to-analog converter (AD7243), or manually by use of the potentiometer. All the fluctuations of the filament voltage are minimized by the negative feedback loop what allows for the temperature stabilization of the filament. In the 4-wires method the resistance variations of filament power supply wires do not affect the filament voltage measurement. The output signal of the MA741 drives the Darlington circuit comprising the 2N3055 and BD439N transistors. This circuit keeps the filament voltage constant and therefore stabilizes its temperature. Since ions have to be formed into a beam, the filament and the whole supply circuit is fixed on the 2kV potential with respect to the ground. Because of the high potential of the source it was necessary to construct the three channel optical relay between the master computer and slave digital-to-analog converter in the filament power supply. In this way the master computer is protected against high voltage. The remaining components identified in the schematic diagram are used to allow smooth switching between the computer and manual control, and to protect the filament against burning during power-on of the supply.

870

Chapter 39

- S. Halas & T. Durakiewicz

SYNC SDIN LK SYNC ~.x.. ~::'S :,, '~' ~ l SDIN \470K D K"I ~.]N SCLK \.,. ", -1- "~ CLR 470KN\470 i i i rlVref d====++5V\ ;,.X, ./ / ..a,k14v \~-- I MAA I12v _ I c12J, liAD7243{ ZlmF 2.2M~ GND T,, ~ , . , ' T 17812 I i ~ +~5 I I [ +5V

'3

~ .~

1",-.t/ I I FI"L IJ'-It 1

_

I MAA

I

I

_

I I ] 7805 l i i35k "+5V 10mP ~ ~ ~mP,ll~P~ 0 ~10 k Vref

=-15V

Vout

Comp_ Adjust"~ == ~I~F

~--_ MA741 L~~__

----~+2kV

///"N\+ 3.8V "

Il

T+V -3x i2 mF

r

2N3055

91

Figure 39.6 - Schematic diagram of the filament temperature controller based on the stabilizer.

wires voltage

)

9N 4-

The voltmeter indicated on the schematic diagram enables the visual control of the filament temperature. This is possible because the temperature-voltage characteristics of the filament are almost linear, as demonstrated by Halas et al. (2001). The voltmeter was calibrated directly in temperature units what allows the operator immediate information on the filament status. The stabilizer is fully computer controlled via the digital-to-analog converter AD7243 (see Halas et al. 2001 for details). In order to demonstrate the superiority of the 4-wires over the 2-wires approach, we have continuously recorded the ion current of 39K+ for 20 minutes. Potassium was loaded onto a platinum filament, 0.05mm x 0.63mm x 10mm, as 50% K3PO4 solution. After drying in air and degassing under vacuum, the filament was incrementally heated to 900K. Ion emission was measured at a filament temperature about 1000K. Comparison of 39K ion currents obtained using the above-described voltage stabilize with 4-wires and 2-wires configuration is shown in Figure 39.7. 39.8 Integrating of the ion currents The ion detection system of an IRMS consists of a collector assembly containing two or more Faraday cups (as mentioned in section 39.2). Although the ion counting

871

Techniques of Ion Current Stabilization in Isotope Ratio Mass Spectrometry

Figure 39.7 - Comparison of 39K ion currents obtained by the authors for filament voltage stabilization by use of 2-wires and 4-wires methods, respectively. technique is used in specialized instruments for detection of very weak ion beams (like 3He, 230Th), we will not consider these devices here. The ions collected in each cup produce electric currents, which flow through a high-value resistor (i.e. 109 - 1012 f2) to ground. The voltage produced on the resistor is not amplified but rather repeated by a circuit arranged as a voltage follower, or, more frequently, it is reversed by an amplifier with the resistor forming a negative feedback loop. A typical detector system of an IRMS is shown schematically in Figure 39.8. The ion current is converted to a voltage by an operational amplifier with ultra-low input bias current and the high-value resistor. The output voltages are then converted to trains of short pulses by so-called voltage-to-frequency converters (V/F), the repetiN 109Q I

t---"

V/F Faraday cups

,

Counter

N 1011Q I t---

Stop

V/F

Counter

1J Display Figure 39.8 - A typical arrangement for digital measurements of current ratios

872

Chapter 39- S. Halas & T. Durakiewicz

tion rate of which is linearly proportional to the input voltage. These pulses are fed to separate counters which are set to zero at the start of measurement. When the major beam counter reaches a count of 106, the minor beam counter is stopped and its value displayed. The value of the full six-digit display is therefore equal to the current ratio. Unfortunately the detection system described above introduces, its own noise, predominantly from the high-value resistors. This is a fundamental phenomenon which cannot be eliminated by technological improvement of the production of resistors (see Felgett & Usher, 1980, for example). Moreover, the resistors also suffer from fluctuations of their value due to variations of the potential drop along the resistor and temperature changes (Habfast, 1960). The key to the improvement of isotope ratio measurements is to replace each highvalue resistor by a capacitor (Jackson & Young 1973, Halas & Skorzynski 1980, McCord & Taylor 1986). This replacement converts the ion detector from an ion "amplifier" to an "integrator" where the voltage on the capacitor raises in time proportionally to the charge collected: T

V - c f l d t --

I.T

C

[39.10]

0

where C is the value of capacitance, I is the ion current and T is the integration period. The value of the capacitor is selected in such a manner that the final voltages are of order of 10 volts. The integration period may be estimated from the following statistical considerations. Let us assume that the 180/160 ratio is measured using CO2 gas. Typical currents obtained by a Nier type ion source for mass 44 and 46 are 2.5 x 10-9A and 1 x 10-11A, respectively. Hence, after time T the number of electrons collected on each capacitor is: n~

I.T

[39.11]

e

where e is the elementary charge 1.602 x 10-19As. According to the general principles of statistics, the relative uncertainty of An/n is equal n-l/2; hence for a desired uncertainty of about 10-5 (i.e. + 0.01%o) the number of ions collected has to be 1010. From equation [39.11] one obtains for the minor beam: T~ 101~

1.602.10-19As 1.6. 102s 10

-11

A

i.e. the integration time has to be of the order of 100 seconds.

[39.12]

873

Techniques of Ion Current Stabilization in Isotope Ratio Mass Spectrometry

The obvious inconveniences of the integrating system are that the "---. 2 ""--, 1 "-... capacitors have to be discharged periodically, prior to each charging C cycle, and the system cannot be used I I directly for the instrument adjust" I I ments by continuous monitoring of the ion current. Both difficulties, however, can easily be overcame today by use of the computer controlled high quality reed switches for discharging the capacitors and their instant switching to the resistors. Such a solution was described by Halas & Skorzynski (1980). One pair Figure 39.9 - Schematic diagram of the capacitance/ of reed switches is required for each resistance system. amplifier. The switches are connected in parallel to C and R as shown in Figure 39.9. In the integrating mode switch 2 is closed and switch I is used for periodical discharging of the capacitor. For adjustment of the mass spectrometer, recording the mass spectra, etc., switch I is closed whilst switch 2 is open.

i

R

39.9 Final remarks

Good quality isotope ratio results are worth taking the effort. The ideas and general remarks presented in this chapter are certainly not the only solutions that guarantee success. They are, however, tested by many years of operation and maintenance of several instruments in our Mass Spectrometry Laboratory. In the final section we should make our reader aware, that having the ion beams stable enough to produce satisfactory precision of 0.05 permil or better for standard versus standard measurement does not guarantee identical precision for sample versus standard measurement. If this is not the case, one should check the purity of the sample and/or the gas flow conditions through the inlet system, as well as the linearity of the system. Having stable ion beams, pure samples and a high quality inlet system one has a chance for good and long-term performance provided that settings of the IRMS are favorable. The optimum setting of the ion source should assure maximum ion current and best peak shape at minimum electron beam. The gas flow rate through the capillaries should be selected depending on the geometry of the ion source. The pressure of the analyzed gas (as estimated on the basis of major beam current) during gas flow should not exceed significantly the 100-fold background pressure. Too high a pressure leads to ion scattering by the gas molecules and thereby to peak broadening.


C H A P T E R 4O International Stable Isotope R e f e r e n c e Materials Manfred Grtning International Atomic Energy Agency, Agency's Laboratories Seibersdorf, Isotope Hydrology Laboratory, A-1400 Vienna, Austria e-mail: [email protected]

Abstract The availability and the proper use of suitable reference materials is one of the basic preconditions to ensure the comparability of stable isotope ratio measurements as performed by different laboratories. In this contribution the most important international stable isotope reference materials are introduced and described in some detail, which define 6 scales for the elements hydrogen, carbon, nitrogen, oxygen and sulfur. In addition brief comments are given on isotope reference materials for the elements lithium, boron and chlorine. The recent situation, demands and future trends for new isotope reference materials are discussed. Some information is provided on the calibration of measurements using these reference materials. The problem of a proper terminology for reference materials is presented and some terms proposed for general use with the intention to achieve better consistency in discussions, descriptions and publications. 40.1 Stable isotope ratios and reporting scales 40.1.1 Isotopic abundance and conventional scales For reporting an isotope ratio R of isotopes a and b of an element, commonly the ratio is expressed as the abundance [b] of the minor isotope b divided by the abundance [a] of the most abundant isotope a, both abundances being calculated from the same measurement process. For some elements (e.g. B, S), the inverse notation is used to express isotope ratios, in order to start the numerical value with a leading non-zero integer value. For several elements, including boron or lithium, this direct reporting of isotope ratios is commonly used. In most applications of stable isotopes in earth sciences, however, it is of much more of interest to know the differences in isotopic ratios between samples, than to know the "absolute" isotopic ratios of the samples ("absolute" meaning a ratio of amounts of isotopes, traceable to the S.I. system). The improvements of mass spectrometry for gas samples in the 1940s and 1950s (Nier, 1947; McKinney et al., 1950) have facilitated very much the precise measurement of isotope ratios for all those elements that can be converted to a gas for mass spectrometric analysis (e.g. to H2, N2,

International Stable Isotope Reference Materials

875

CO2, CO, CF4, 02, SO2, and SF6). Therefore, the isotope ratios of elements such as hydrogen, carbon, nitrogen, oxygen and sulfur have been commonly measured for decades by dual inlet mass spectrometers (McKinney et al., 1950) and the isotope ratios are reported as deviation from the isotope ratio of an artificially selected reference (~ scale). In this approach the knowledge of the mole fraction of an isotope of a given reference material is not necessary, since these isotope ratio ~ scales are defined completely arbitrarily relative to the isotope ratio of a selected primary reference material. This primary reference material functions as the end of the traceability chain (Chapter 41, this Volume). This concept as described is the realization of a "conventional" scale. It is not traceable back to the S.I. system, since it is not based on fundamental constants, but on arbitrarily selected properties, e.g. the stable isotope ratio in a sample of a chosen primary reference material (the pH scale being another example for a conventional scale). The choice of the primary reference material is completely arbitrary; as we will see, in several cases the selected material is even virtual / hypothetical and does not exist itself physically. In such a case its isotopic composition is just defined in measurable manner in terms of another real existing material and a statement on the isotopic difference of both. Certain basic requirements for the used reference materials have to be fulfilled, e.g. regarding homogeneity and stability (ISO Guide 34, 2000c). Special care should be taken in the isotopic characterization of reference materials; the mole fractions of the isotopes comprising the material should be precisely determined to allow the transition in the future, to a different scale based on another reference material. In this respect, precise determinations of mole fractions of the isotopes have distinct advantages. Firstly, in case a primary reference material is not more available, then the knowledge of mole fractions of its isotopic composition allows to maintain the original scale by a similar mole fraction determination for a suitable new material. Secondly, direct mole fraction determinations using primary methods (traceable back to the S.I. system) will allow to safeguard the linearity of the scale, because such measurements are much less susceptible to non-linearities than for differential measurements. Absolute isotope ratios of selected reference materials - that are discussed later in section 40.3 - are presented in Table 40.1. For the five elements H, C, N, O and S, the abundance of the major and the minor isotopes is different by orders of magnitude, resulting in very small ratio values. Since the natural variations in isotopic composition due to fractionation processes are rather small, the stable isotope ratios of naturally occurring materials on Earth do not differ considerably from each other. For example, the total range of the variation of 180/160 in water and ice on earth is in the order of only nine percent (Coplen et al., 2002; see Table 17 and Fig. 6) and increases to about seventeen percent if all natural oxygen bearing materials are considered (Coplen et al., 2002). Similar statements are true for most other elements except hydrogen (where the variations are exceptionally large). For that reason, an alternative way was introduced of reporting isotope ratios (McKin-

876

Chapter 40- M. Gr6ning

Table 40.1 - Absolute isotope ratios and associated standard uncertainties at 1c-level for selected reference materials discussed in this chapter. Name

Material

Isotope ratio

Isotope ratio value

Reference


877

International Stable Isotope Reference Materials > Table 40.1 continued Name

a

b

Material

Isotope ratio

Isotope ratio value

Reference

Hypothetical carbon isotope mole fraction for VPDB calculated from the carbon isotope ratio of NBS 19. Hypothetical sulfur isotope mole fraction for VCDT calculated from those isotope ratios of IAEA-S-1, IAEA-S-2 and IAEA-S-3

ney et al., 1950; Epstein & Mayeda, 1953) using 6 scales. There, only the deviation of a sample isotope ratio Rsample is reported relative to that of an arbitrarily selected reference material Rreference:

6

-

R

sa

mple- Rreference Rreference

[40.1]

with 6 (e.g. 62H, 813C, 615N, 6180, and ~)34S) being the normalised difference of the isotope concentration ratios R (2H / 1H, 13C / 12C, 15N / 14N, 180 / 160, 34S / 32S) of the sample and the reference (e.g. ~15N with atmospheric nitrogen as reference). ~5-values are therefore unitless numbers, like the isotope ratios itself. As the differences between a sample and reference are normally very small, the ~ values are usually expressed as per mill difference (parts per thousand - per mill - %0 - 10-3). Thus, we have" 6 -

Rsample- Rreferencex 1000 %o Rreference

[40.2]

This modified equation [40.2] results in more convenient numbers being reported. The values can be positive or negative, with e.g. negative numbers indicating a lower abundance of the minor isotope in the sample than in the reference. Note that the per mill sign (%o) is part of the notation and may not be omitted. For a very detailed discussion of all features of the 6 notation as realization of a conventional scale, see Gonfiantini (1981). The reference in equation [40.2] is often chosen to represent the isotope ratio of a major sufficiently homogeneous reservoir of the element (e.g., ocean water for the elements hydrogen and oxygen, atmospheric N for nitrogen or marine carbonate for carbon). This way of reporting isotopic abundances has several advantages (Mook, 2000). First, the mass spectrometers typically used for measuring isotope abundances in naturally occurring materials are not really suitable for obtaining reliable absolute isotope ratios (or mole fractions of isotopes). Secondly, in most applications, the

878


differences of isotope ratios among samples are far more important than the values of the absolute ratios themselves. Moreover, due to the small variations in isotope ratios (most samples showing values close to the reference), all reported R values would need five to six digits after the decimal to be significant. Finall~ for the precision desired in many applications, the mass spectrometric dual-inlet mode or continuous flow mode are most appropriate, where a reference is necessary; because 6 values (differences in isotope ratios) can be determined about an order of magnitude more accurately that absolute isotope ratios or mole fractions of the isotopes of an element in a substance. The necessity for defining an arbitrary material with a well-known isotope ratio to realize the proper calibration of relative isotope data is not a disadvantage per se. That requirement is true for any transfer of properties from instrument to instrument, regardless of the measurement method used. The major disadvantage with this approach is that the property "isotope ratio scale" is defined using a physical material with its limited quantity available. Therefore, the isotope ratio scale, as defined by that material, is at risk to change when an exhausted primary material is replaced by a new one. The comparison of the old with the new material will introduce an extra uncertainty. Comparing isotope ratio values across decades is therefore not obvious. As we will see, such transitions have already taken place several times during the last four decades. It seems, however, that the advantages of using relative isotope ratios, by far exceed the disadvantages, as will be discussed in the next sections. The isotopic abundances of the elements stated above and their natural variations are reported elsewhere (Coplen, 2001b; Coplen et al., 2002).

40.1.2 Historical reporting scales As discussed before, commonly isotope ratios are reported as 6 values using arbitrary reference isotope ratios. The 6 value of the reference is zero by definition (equations [40.1] and [40.2]). Careful investigations were carried out in the early days of stable isotope work in the 1950s and 1960s to define useful reference materials. As a main criterion, the selected reference should represent a major pool of the element under consideration. Therefore, a logical choice for hydrogen and oxygen stable isotope ratios could be ocean water. Indeed, a well-mixed mean ocean water was proposed as the reference in view of the rather small variations in isotopic composition throughout the oceans. This purely hypothetical water was called SMOW (Standard Mean Ocean Water) (Craig, 1961) and its isotopic composition was defined in terms of an isotopic difference from an existing water (NBS 1). This is a good example for a scale based on a virtual material, not realised in nature. SMOWs proposed isotope ratio was compiled by averaging the isotope ratio values of measurements of different sea water compartments available at that time (McKinney et al., 1950; Horibe & Kobayakawa, 1960; Craig, 1961). Unfortunately the same term SMOW was used for a water subsequently prepared by H. Craig, Scripps Institution of Oceanography, La Jolla, USA, and which was isotopically adjusted to match that virtual reference SMOW, as described above. Just to mention, even a third SMOW concept existed, related to a Rose quartz sample. The IAEA then decided to rename the water prepared


879

by H. Craig to VSMOW (Vienna SMOW) in order to remove the confusion. Since that time this material VSMOW is used as primary reference material to define the hydrogen and oxygen isotope ratio scale. Similarly, for stable carbon isotopes, the largest and relatively homogenous carbon reservoirs are marine carbonates. Therefore, the carbon reference was chosen to be the carbonate of the rostrum of a Cretaceous belemnite (Belemnitella Americana) collected in the Peedee formation of South Carolina, USA. This reference was called PDB (Peedee Belemnite) and was used for reporting carbon isotope ratios and additionally as reference for reporting oxygen isotope ratios in carbonates. For nitrogen, atmospheric N2 is the only logical, and nearly ideal, choice, but it needs purification from other gases before being usable. For sulfur, several references were proposed, the most common one was Canyon Diablo Troilite (CDT) iron sulfide (FeS) from an iron meteorite. It was chosen due to the expected primordial isotopic composition in the meteorite sulfur. The major drawback of all these early references (except air N2) was the limited quantity of the selected material physically available to realise a calibration of these isotope-ratio scales and the problem of material homogeneity and long term preservation. Indeed, PDB and CDT have been exhausted for three decades; and even worse, as indicated above, SMOW was just a virtual reference material and never existed physically. The developments associated with the historical isotope-ratio scales are discussed in section 40.3. Also discussed are the consequences and problems of achieving consistency in reporting data using conventional scales, and the problem of transferring the calibration to another successor material and successor scale.

40.1.3 Measurement techniques The main international efforts on standardization of stable isotope ratio measurements so far have concentrated on problems of mass spectrometric methods. Up to now, mass spectrometry has been used almost exclusively for the analysis of stable isotope ratios at environmental levels. Concerning the elements H, C, N, O and S, until the last decade, most measurements on gas samples were performed by using the dual-inlet configuration for mass spectrometers. That involved considerably large amounts of samples being prepared, converted into suitable gases and transferred into steel bellows for inlet into the mass spectrometer. The great advantage was the direct comparison of each sample with one gas sample used as transfer gas or working standard over one full measurement day, which facilitated to achieve a high measurement precision for each individual sample. During the last decade, continuous flow techniques emerged, allowing the preparation of very tiny sample amounts and transfer via a helium carrier gas into the mass spectrometer. There, new on-line sample preparation techniques have been developed

880


to combust samples and to separate effectively the different gases produced during that process via gas chromatographic techniques. New requirements for reference materials emerged on issues like homogeneity and available compounds, due to the smaller sample sizes used and the different standardisation technique applied (injection of reference gas only before or after the samples, no direct simultaneous measurement possible; treatment and behaviour of reference materials should be similar as for samples). Several non mass spectrometric analytical methods are applied for isotopically enriched materials, using often cheaper instruments and robust analytical techniques, but with lower sensitivity (Roth, 1997). These methods are of great advantage, where ultimate precision is not required, and often provide faster and easier measurement possibilities. Even for isotopic compositions at environmental level with their rather small variations in isotope ratios, some other promising techniques are emerging, such as optical techniques using infrared absorption spectra (Kerstel et al., 2001) and the optogalvanic effect (Murnick, 2001). Since the requirements in terms of standardisation and calibration are somewhat comparable and similar to mass spectrometric techniques, no distinction is made in the following considerations on those different analytical techniques. 40.2 Terminology for different kinds of reference materials In the literature different definitions and terms are used for internationally distributed reference materials and for local laboratory standards. Different authors often use the same terms in varying context and meaning, and some expressions are used in an ambiguous manner (e.g. the term 'standard'). So far no clear guidelines exist on the definitions for reference materials to be used in the field of stable isotope ratio measurements. The ISO International Vocabulary of Basic and General Terms in Metrologyl provides a consistent general set of definitions, but unfortunately is not in agreement with well established terms used throughout the stable isotope measurement community, consistently causing misunderstandings.

An effort was made originally to define the various categories of reference materials (Gr6ning et al., 1999) in accordance to the practice in the field of stable isotoperatio measurements (Gonfiantini et al., 1995). The terms below are a slightly modified and updated version of those in Gr6ning et al. (1999). Some examples on real materials are given to illustrate the definition. Those definitions used for the various kinds of materials should be clearly distinguished from each other:

Primary reference material (or international standard 1).

a natural, synthetic or virtual

material, which, by general agreement, serves as the substance against which iso1. International Vocabulary of Basic and General Terms in Metrology (1983), International Organization for Standardization, 2nd ed., Geneva, Switzerland.


881

tope ratios of samples are expressed. In the context of this publication, it is used to define a conventional scale (arbitrarily defined by agreement) for reporting variations of stable isotope ratios. For most of these materials information on their stable isotope abundances as molar fractions are available. Recently, the term 'primary conventional reference material' was suggested to point to the nature of the defined scale, being conventional due to the artefacts used for their definition. Example" 62H: VSMOW, 613C: VPDB. Note: For 613C the primary reference material is a virtual (non existing) material !

Calibration material (Gonfiantini et al., 1995) (or primary standard 1): a natural or synthetic compound, which has been carefully calibrated against the primary reference material. It is used in case the primary reference material is not available to calibrate measurements and instruments (or as substitute if the primary material is not existing at all). Each physical existing primary reference material can be referred to as a calibration material as well. Example" 62H: VSMOW, 613C: NBS 19. Note: VSMOW is both calibration material and primary reference material, NBS 19 is not a primary reference material, as ~13C values are expressed versus VPDBisotopic composition as defined zero-point of the 6-scale.

Reference material (RM): a natural or synthetic compound which has been carefully calibrated against the primary reference material and property values of which are sufficiently homogeneous, well established, and associated with determined uncertainties. It is used to calibrate laboratory equipment and measurement methods for analysis of materials of a composition different from that of the primary reference material. The available reference materials cover a broad spectrum of chemical compositions and a wide range of stable isotope ratios. Most existing stable isotope ratio reference materials were investigated in interlaboratory comparison exercises and were first distributed as intercomparison materials. For the reporting of stable isotope measurements, the isotopic value of the reference material used for their calibration should be stated. Example" 62H and ~13C: several materials, see Tables 40.3, 40.5, 40.6, 40.7, all calibrated against respective primary reference materials.

Certified reference material: see under reference material, with the additional requirement of a detailed description of its calibration and measurement in a certificate and certification report. Example: BCR-IRMM materials in Table 40.3.

Intercomparison material (quality control material)" a homogeneous natural or synthetic compound that provides the means to check the overall quality of measurements performed in comparison with that of other laboratories. It is useful to identify the relative bias of results between laboratories. The second term is used at IAEA for ongoing distribution of materials having been used previously in interlaboratory comparison exercises, since they are used like a quality control material (see next term below).

882


Example: 82H and 813C" various test materials circulated in interlaboratory comparison exercises, since it is used like a quality control material (see below) (e.g. IAEA-OH-1 to IAEA-OH-8 water samples).

Quality control material (or internal standard / internal laboratory standard) (or reference standard1): This term describes a homogeneous material being of similar composition as normal samples, but carefully calibrated versus reference materials and used routinely day by day in a laboratory to transfer the calibration to samples and to check measurements and the measuring instruments. The second term 'internal standard' is commonly used in stable isotope mass spectrometry. Instead of the expression 'internal' similar terms like 'local' are also used. Example at IAEA: 62H: ST7 to ST10 water samples calibrated against VSMOW, 813C: Carrara marble calibrated against VPDB. 1)- The first term is to be preferred, describing an arbitrarily chosen gas used as a reference for analysis of isotope ratios of samples in a dual-inlet mass spectrometer (see section 40.1.3). The second term 'working standard' is well established in stable isotope mass spectrometry, but is somewhat misleading since it is not describing at all a real 'standard'. All dual-inlet measurements of prepared samples and reference materials are made relative to this transfer gas and results are later converted and expressed on an international accepted 8-scale. For continuous flow measurements, these terms should not be used, but rather the expressions as stated directly below.

Transfer gas (or working standard) (or transfer standard

Reference gas (or reference injection gas) (or laboratory reference gas)" a gas used as a reference for analysis in continuous flow isotope ratio mass spectrometry. The gas of known isotopic composition is injected in the carrier gas stream alternate to sample gas.

Standard: The single term 'standard' is a quite ambiguous expression and is not well defined due to its application for many different purposes. Its use should be strictly limited, and used only exceptionally as general expression for a material meeting one of the classifications above. The more specific terms as stated above should be strongly preferred and used whenever possible. According to the Commission on Atomic Weights and Isotopic Abundances of the International Union Standard of Pure and Applied Chemistry (IUPAC) the use of the term "standard" should be reserved in chemistry solely for use in standard states and standard deviation (oral communication).

40.3 International stable isotope reference materials The presentation here focuses on 'international' reference materials being easily accessible and relevant on the international scale. It discusses available reference materials with property values traceable to internationally agreed scales. No attempt 1. International Vocabulary of Basic and General Terms in Metrology (1983), International Organization for Standardization, 2nd ed., Geneva, Switzerland.


883

was made to evaluate comprehensively the market of suppliers and stable isotope reference materials to its whole extend. Many materials are available from providers, where the direct link to the international scales (traceability, see Chapter 42 in this Volume) is not evident, i.e. not fully documented and not accompanied by reliable uncertainty statements. Those materials are not considered here. Please note that so far none of the IAEA stable isotope reference materials is called 'certified' since not all the ISO requirements for certification are fulfilled. 40.3.1 Common characteristics of available stable isotope reference materials The reference materials discussed here can be classified in two categories: first synthetically produced or refined substances, such as carbonates, sulfates, sulfides, nitrates, graphite, sugar or polyethylene, being chemically pure; secondly, natural materials like distilled water, carbonate rock, silicates, refined oil, cellulose and similar compounds, selected and tested for their purity and isotopic homogeneity.

Most of these materials have been prepared with the intention of using them as reference materials. Therefore, much care was taken in the initial purification and homogenisation of the raw materials. In most cases, the recommended values for stable isotope ratios in these materials were determined by interlaboratory comparison exercises. For some of the materials issued twenty or more years ago, the number of participating laboratories was rather limited (in some cases, less than ten), constraining the validity of any statistical evaluation. All of those materials were originally prepared for use with dual-inlet mass spectrometers with off-line sample preparation systems using relative large sample amounts. Therefore, an urgent need exists to assess the homogeneity of those materials at the sub-milligram level, which is the usual sample size for modern continuous flow systems. In this contribution, the recommended isotopic 6 values for available internationally distributed reference materials are reported with associated standard uncertainties as derived during the reference material calibration, in most cases from the reported uncertainties of individual laboratories. All uncertainties are reported at the l o-level. However, limited information is on record on the uncertainty assessment methods used for data reported by laboratories more than a decade ago. The recommended 6 values for the reference materials, as stated in this contribution, are derived from published literature with a few exceptions of data from publications in preparation. All available raw data for existing and currently distributed reference materials have been compiled and re-evaluated using a consistent approach for identifying and excluding outliers. This statistical approach, called 'exclusive sigma-test,' is a variation of the well known 2o-outlier test and was first suggested by H. Meijer, CIO, University of Groningen, Netherlands, during an IAEA experts meeting on stable isotope reference materials in the year 2000 in Vienna. A detailed description of this outlier test method will be given elsewhere together with a discussion on the resulting (slight) changes for some recommended isotope ratio values.

884

Chapter 4 0 - M. Gr6ning

Most of the reference materials discussed in this chapter are distributed by the International Atomic Energy Agency (IAEA), the U.S. National Institute of Standards and Technology (NIST) and the EU Institute for Reference Materials and Measurements (IRMM). The IAEA and NIST have cooperated since the 1960s in the distribution of stable isotope reference materials by sharing the available materials for the elements hydrogen, lithium, carbon, nitrogen, oxygen, sulfur, and silicon. In Table 40.2, the corresponding names for the reference materials, as used by IAEA and by NIST, are presented. 40.3.2 Hydrogen and oxygen stable isotope reference materials

40.3.2.1 Historical development of scales Hydrogen and oxygen are discussed in a common section due to the correspondence for water stable isotope ratio reference materials. First the historical development is presented on the establishment of the scales for reporting hydrogen and oxygen isotope ratios, primarily intended for measurements on water samples and later extended to other hydrogen and oxygen bearing materials. By 1953 "average ocean water" was suggested and used as a reference point for isotope-ratio measurements (Epstein & Mayeda, 1953). Because no 'average ocean water' existed, H. Craig refined this concept in 1961 by defining the hypothetical Standard Mean Ocean Water (SMOW) as zero-point of that conventional scale in terms of Table 40.2 - Synonyms used by IAEA and NIST for jointly distributed stable isotope reference materials. Other materials not listed in the table are not distributed by both organisations. The following reference materials are also known under synonyms: IAEA-S-4 as "Soufre de Lacq", IAEA-CH-6 as "Sucr.Anu', IAEA-CH-7 as "PEF-I", IAEA-S-1 as "NZI", IAEA-S-2 as "NZ2". IAEA-name

NIST-name

compound

IAEA-name

NIST-name

compound

VSMOW GISP SLAP NBS 30 NBS 22 IAEA-CH-7 USGS24 IAEA-CH-6 NBS 18 NBS 19 LSVEC NBS 28 IAEA-N-1 IAEA-N-2 IAEA-NO-3 USGS25 USGS26

RM RM RM RM RM RM RM RM RM RM RM RM RM RM RM RM RM

water water water biotite oil polyethylene graphite sucrose carbonatite limestone Li2CO3 silica sand (NH4)2SO4 (NH4)2SO4 KNO3 (NH4)2SO4 (NH4)2SO4

NSVEC IAEA-S-4 IAEA-S-1 IAEA-S-2 NBS 123 NBS 127 USGS32 NGS1 NGS2 NGS3 USGS34 USGS35 USGS40 USGS41

RM RM RM RM RM RM RM RM RM RM RM RM RM RM

nitrogen gas sulfur Ag2S Ag2S sphalerite BaSO4 KNO3 hydrocarbon gas hydrocarbon gas hydrocarbon gas KNO3 NaNO3 L-glutamic acid L-glutamic acid

8535 8536 8537 8538 8539 8540 8541 8542 8543 8544 8545 8546 8547 8548 8549 8550 8551

8552 8553 8554 8555 8556 8557 8558 8559 8560 8561 8568 8569 8573 8574


885

a real reference water (Craig, 1961). Its isotopic composition was defined as a weighted average of the available measurements of the isotopic composition in the main oceanic water masses. But because SMOW was just the concept of a hypothetical water and never existed as a real water sample, it couldn't be used directly for calibration of laboratory measurements. Thus, the isotopic ratios of SMOW were defined with respect to an existing water standard1 NBS-1 (Mohler, 1960) distributed by the US National Bureau of Standards. This water was used earlier for an water stable isotope interlaboratory comparison. So for the first time a physically existing material was used to calibrate different laboratories to the SMOW scale. NBS-1 water was readily available for worldwide distribution together with a second water standard1 called NBS-1A, which was obtained from melted snow with a lower abundance of the heavier isotopes. During an IAEA interlaboratory comparison in 1965, serious doubts were confirmed concerning the preservation of the NBS-1 water standard and possible changes of its isotopic composition over time. For this reason, it was unsuitable to be used as standard anymore. At an IAEA Panel Meeting in 1966, it was therefore recommended to establish a pair of two new primary water reference materials, the first one being as close as possible to the defined SMOW and the other one with an abundance of the heavier isotopes close to the lowest limits observed in natural water. The new material with an isotopic composition as close as possible to SMOW was prepared by R. Weiss and H. Craig, Scripps Institution of Oceanography, La Jolla, USA. It was obtained by mixing distilled ocean water with small amounts of other waters in order to adjust its isotopic composition as close as possible to that of the defined SMOW. This task was complicated due to the required adjustment of both the isotopic composition of hydrogen and of oxygen. This reference water was ready in 1968 and was called Vienna Standard Mean Ocean Water (VSMOW). According to the control analyses performed by H. Craig, VSMOW has the same 180/16 0 ratio as the defined SMOW, but a slightly lower 2H/1H ratio or respectively a slightly negative 62H value (-0.2%o) relative to SMOW. However, this slight difference is about a factor of four to five lower than the measurement uncertainty of most laboratories. Absolute isotope ratios of VSMOW were determined for 180/160 (Baertschi, 1976), for 170/16 0 (Li et al., 1988) as well as for 2H/1H (Hagemann et al., 1970; De Wit et al., 1980; Tse et al., 1980). The second water reference material was obtained by E. Picciotto, Universit6 Libre de Bruxelles, Begium, from melting a firn sample at Plateau Station, Antarctica. This material was named Standard Light Antarctic Precipitation (SLAP). The absolute isotope ratios of SLAP were determined only for 2H/1H (Hagemann et al., 1970; De Wit et al., 1980; Tse et al., 1980). Absolute 180/160 ratios for SLAP in this report are com1. Here the general and vague term 'standard' is used, since the exact status in the 1960s is not anymore known for the two materials NBS-1 and NBS-1A. This also applies to some other materials mentioned in the text.

Chapter40- M. Gr6ning

886

puted from those of VSMOW and by using its assigned relative 5 value of 5180 = -55.5 %o relative to VSMOW (see below). All four materials, VSMOW, SLAP, NBS-1, and NBS-1A, were then distributed by the IAEA Isotope Hydrology Laboratory (the latter two materials were transferred to the IAEA from the National Bureau of Standards, today the National Institute for Standards and Technology (NIST), Gaithersburg, Maryland, USA). Nowadays, most stable isotope reference materials are distributed both by the IAEA and NIST. In 1976, an IAEA Consultants' Meeting was convened in order to discuss the isotopic results on these reference materials and to advise on future actions on standardization of stable isotope ratio measurements (Gonfiantini, 1978). The recommendation of the experts concerning data of the water reference materials was to express all future results for hydrogen and oxygen isotope ratios as 5 values relative to VSMOW in order to resolve confusion on results expressed in non-corresponding scales. Thus, VSMOW water was recommended to serve as a new primary reference material. It was stated that the coherence between 5 values reported by different laboratories could be improved by adopting a fixed 5 value for a second water reference material. The experts recommended that SLAP be adopted for this purpose and that both the 180/160 and 2H/1H 5 values be normalized relative to VSMOW water, the primary reference material. In Table 40.3 the 5 values of selected reference materials are listed with those of VSMOW being by definition at zero and those of SLAP established by assessing the close agreement of three direct determinations of hydrogen isotope mole fractions as well as the weighted mean of relative isotope ratio measurements of different laboratories. The definition of VSMOW as zero-point for the oxygen and hydrogen 5 scales and the adoption of fixed 5 values for SLAP is therefore a slight modification of the original definition of the 6 scale in equation [40.2]" -

RSAMPLE-RvsMow" RVSMOW

(~3SLAp/RSLAp~RvsMow I 1000%o •

RVSMOW

[40.3]

,

with 5 referring to ~2H and 6180, R being the corresponding 2H/1H and 180/160 ratios and the additional term in the bracket being the normalization of the respective VSMOW-scale in terms of pre-defined hydrogen and oxygen isotope ratios of the two primary reference materials (Table 40.3) (Gonfiantini, 1981). The two scales defined in equation [40.2] for SMOW and in [40.3] for VSMOW coincide only if RSMOW - RVSMOWand if the adopted 5 values for SLAP in Table 40.3 correspond to the true ones as defined by equation [40.2] for both hydrogen and oxygen. From the reported measurements of NBS-1 and VSMOW, a slight offset of the zero-point of the two scales could be concluded (offsets of 0.05%0 and 0.5%0 for 5180 and 62H, respectively), but these offsets are well within the limits of measurement

887


Table 40.3 - Oxygen and hydrogen 6-values versus VSMOW assigned to the existing major water reference materials with associated standard uncertainty at lo-level (where applicable). Available materials are marked in bold. 6-values marked with * are reported versus SMOW. Given references are valid both for 6180 and 62H. CM stands for "Calibration material" and RM for "Reference material". Name

Material

Status

Distribution

6180 [%o]

VSMOW SLAP GISP

water water water

CM CM RM

IAEA, NIST IAEA, NIST IAEA, NIST

0 -55.5 -24.78_+0.08

NBS-1

water

62H [%o] 0 -428 -189.73+0.87

Gonfiantini et al. (1995) discontinued

-7.94*

-47.6*

Craig (1961) -7.89+0.12

-47.1+1.2

Gonfiantini (1978) NBS-1A

water

discontinued

-24.33*

-183.3*

Craig (1961) -24.29+0.25

-183.2+0.7

Gonfiantini (1978)

uncertainty of most laboratories (Gonfiantini, 1977). Both offsets were a bit larger than evaluated before by H.Craig. However, due to the scatter of the individual results, the stated mean offset should not be applied for conversion from one scale to the other. A third water reference material was proposed during the same meeting in 1976 (Gonfiantini, 1977) with an isotopic composition intermediate between that of VSMOW and SLAP. This material was obtained by W. Dansgaard, University of Copenhagen, Denmark, from Greenland firn in 1978 and was called GISP (Greenland Ice Sheet Precipitation). GISP is intended to demonstrate successful calibration performed with VSMOW and SLAP. Results of two interlaboratory comparisons investigating GISP are published in IAEA reports (Gonfiantini, 1977; Gonfiantini et al., 1995). It was noted, that the coherence of results from different laboratories improved by a factor of more than two, when data were normalized using SLAP as second primary reference material. The normalization of oxygen and hydrogen isotope ratio data using the pair of VSMOW and SLAP primary reference materials was recommended by an IAEA panel group (Gonfiantini, 1977) and later by the Commission on Atomic Weights and Isotopic Abundances of the International Union of Pure and Applied Chemistry (IUPAC) (Coplen et al., 1996). 40.3.2.2

Necessary changes of the VSMOW/SLAPscale

The original prepared amount of SMOW was about 70 litres in 1968. Today (in the year 2003), only about 10 litres in total remain at the two distributing institutions, the International Atomic Energy Agency in Vienna, Austria, and the National Institute of Standards and Technology in Gaithersburg, Maryland, USA. In December 1996, a small meeting was convened at IAEA to discuss the possibilities of preparing a successor material before the stocks of VSMOW were totally

888


exhausted. The unanimous recommendation was to try to prepare a successor material isotopically indistinguishable from VSMOW within measurement uncertainty of the best analytical techniques used for routine measurements, in order to minimize any problem of scale conversion. It was decided to prepare the replacement as a mixture of three distilled natural waters, having an initial isotopic composition close to that of VSMOW. The use of water enriched in 180 was rejected due to the aim to reproduce not only 62H and 6180, but also ~170 values. Materials enriched in 180 often show a non mass-dependant fractionation of 170 versus 180. It took three years until sufficient amounts of three suitable water samples were supplied to the IAEA from Lake Bracciano, Italy (G.-M. Zuppi, University of Venice and R. Gonfiantini, Istituto di Geoscienze e Georisorse, Pisa), from Lake Genezareth, Israel (E. Adar, Ben-Gurion University), from a well near Cairo, Egypt (F. Hussein, University of Cairo and I. Nada, Egyptian Atomic Energy Authority, Cairo). All three samples were initially distilled at the IAEA. A careful calibration of the raw waters made directly against VSMOW was performed by five laboratories (T. Coplen, USGS Reston, Virginia, USA; H.A.J. Meijer, CIO Groningen, Netherlands; W. Stichler, GSF Neuherberg, Germany; R Dennis, UEA, Norwich, UK; IAEA IHL Vienna, Austria). The results of all five stable isotope laboratories were in accordance by better than _+0.02%0 for ~180 and by better than +0.3%0 for ~52H(see Figure 40.1). As a result of the calibration, it was decided to produce an amount of 300 litres of this new material. Provisionally, the working name for the new material is set to "NEW VSMOW", subject to change after completion of the production. Special stor-

Figure 40.1 - Calibrated measurements of five laboratories for three raw waters used for preparation of a mixture as a replacement for VSMOW (VSMOW being at the origin of the coordinate cross) and the result of the mixing of the "NEW VSMOW" (big square close to the coordinate origin of the plot), see data in Table 40.4.

889


age containers with a volume of 300 litres were purchased to enable the transfer of water in and out of the container without contact to the atmosphere in order to avoid isotopic fractionation of the water. Due to the particular isotopic compositions of the available three water samples, direct mixing of such a large amount of NEW VSMOW was not possible. Therefore, a deuterium enrichment method was specially designed to enrich several litres of the Egyptian water in its deuterium content without disturbing its natural oxygen isotopic composition. This resulted in the 62H value of this sample water being enriched in 2H by about 40%0, while no significant shift in 6180 was observable (Figure 40.1). Finally appropriate fractions of the three raw waters were mixed gravimetrically. The isotopic composition of the mixture "NEW VSMOW" was assessed by three laboratories (T. Coplen, USGS; H.A.J. Meijer, CIO; IAEA Isotope Hydrology Laboratory) by directly measuring aliquots of NEW VSMOW against aliquots of VSMOW. For 62H determinations, both the equilibration method and the zinc reduction method were used. The results are shown in Table 40.4 and in Figure 40.1. Details of the preparation of this material will be presented elsewhere. The results show no significant deviation within the stated estimated standard error (standard deviation divided by square root of measurements) for all performed measurements in the three laboratories. It is believed that virtually no laboratory with current state-of-the-art techniques will be able to detect this deviation in its routine measurements. Meanwhile, as second project, the preparation of a successor material for SLAP has started. After some efforts, two water samples substantially depleted in 2H and 180 were obtained from the U.S. Antarctic Station at the South Pole (initiated by T.B. Coplen, USGS, Reston, Virginia, USA, with kind support of the U.S. NSF Polar Sciences Program) and from the Antarctic Vostok drilling program (water supplied by J.R. Petit, LGGE-CNRS, Grenoble, France). The two samples are just under calibration at the IAEA. With the experience gathered during the NEW VSMOW production, it is believed that one can reproduce SLAP with only a small difference in isotopic composition from that of SLAP. At least 200 litres of water are planned to be produced from these samples. The provisional working name "NEW SLAP" will be used for this material during its production process. Table 40.4 - Estimate of the deviation of NEW VSMOW from VSMOW for 62H and 6180, expressed in per mill vs. VSMOW for a series of measurements using three mass spectrometers at the IAEA Isotope Hydrology Laboratory. Deviation of NEW VSMOW vs. VSMOW [%0]

6180 62H

0.002 -0.12

Estimated standard error of the mean [%o] for NEW VSMOW (no. of analyses) +0.007 (125) +0.09 (118)

Estimated standard error of the mean [%0] for VSMOW (no. of analyses)

+0.006 (109) +0.08 (1~5)

890


After c o m p l e t i o n of the p r o d u c t i o n a n d bottling of b o t h materials, it is p l a n n e d to p r e s e n t all the results to the scientific community. For the future it is e x p e c t e d to realize any calibration relative to the established V S M O W / S L A P scale by u s i n g the n e w pair " N E W V S M O W " / " N E W SLAP" by assigning precise ~5 values to them. It is h o p e d that the a g r e e m e n t a n d m a t c h i n g of the t w o scales will be convincing, so that no long-lasting debate on the scale issue will e m e r g e as it w a s the case in the p a s t during the transition from S M O W to VSMOW. As an a d d i t i o n a l measure, an absolute calibration of b o t h n e w materials for their h y d r o g e n a n d o x y g e n isotopic c o m p o s i t i o n (isotope mole fractions) s h o u l d be u n d e r taken. Suitable m e t h o d o l o g i e s a n d institutions will h a v e to be identified for this p u r pose. Ideally, this w o u l d p i n d o w n the u n c e r t a i n t y on the isotopic a b u n d a n c e of N E W V S M O W a n d N E W SLAP a n d allow to m a i n t a i n a h i g h degree of consistency of 6180 a n d 62H scales, regardless of the calibration materials used.

40.3.2.3 Other hydrogen and oxygen stable isotope reference materials In Table 40.5 a few m o r e reference materials u s e d for 62H a n d 6180 analyses are listed. Two of these materials, NBS 28 silica s a n d a n d NBS 30 biotite, w e r e p r e p a r e d b y I. F r i e d m a n of the U.S. Geological Survey. They p r o v i d e a link to &180 m e a s u r e m e n t s on silicates. A l t h o u g h no absolute silica isotope a b u n d a n c e m e a s u r e m e n t of NBS 28 has b e e n p e r f o r m e d , NBS 28 has served for several decades as the p r i m a r y reference material for relative 30Si/28Si m e a s u r e m e n t s ; thus, 630SINBS28 = 0%o. H y d r o g e n isotope ratios for NBS 30 biotite s h o u l d be d e t e r m i n e d on the w a t e r fraction (3.5% b y weight). Table 40.5 - Recommended 6-values relative to VSMOW for additional oxygen and hydrogen reference materials with associated standard uncertainties at lo-level. Available materials are marked in bold. For references see text. RM denominates "reference material", CRM stands for "Certified reference material" Name

Material

Status

Distribution

6180 [%o]

62H [%o]

630Si

NBS 28 NBS 30 IAEA-302A

silica sand biotite water

RM RM RM

IAEA,NIST IAEA,NIST IAEA

+9.58 + 0.09 +5.24 + 0.25 -

-65.70+ 0.27 508.4

0 -

IAEA-302B

water

RM

IAEA

-

Parr & Clements (1991)

996

-


IAEA-304A

water

RM

IAEA

251.7

-

-

-

-

-

-

-

-


IAEA-304B

water

RM

IAEA

502.5 Parr & Clements (1991)

BCR-658

water from wine

CRM

BCR(IRMM)

-7.19+0.04 Guillou et al. (2001)

BCR-659

water from wine

CRM

BCR(IRMM)

-7.18+0.02 Guillou et al. (2001)

International Stable Isotope ReferenceMaterials

891

The two reference material sets, IAEA-302 and IAEA-304, consist of two reference materials each, both isotopically enriched at different levels with deuterium and 180, respectively (Parr & Clements, 1991). These materials were prepared by P. Klein, USDA/ARS, Houston, Texas, USA and I.I. Lesk, MSD Canada Ltd., with the aim of providing reference materials for medical and biological applications and for applications needing materials enriched in 180 and 2H. Two certified reference materials, BCR-658 and BCR-659 synthetic wines, were produced by M. Lees, Eurofins Scientific, Nantes, France, in cooperation with C. Guillou and G. Remaud, EC Joint Research Centre, Ispra, Italy, in order to provide the means for proper analysis of wines in the European Community according to officially approved methods (Guillou et al., 2001). Both materials consists of water - ethanol mixtures (7 and 12 %vol.) with added chemicals to mimic the composition of wine. The parameter to be certified was the 6180 composition in the water phase. Recently, two more organic materials were prepared as candidates for ~)180 reference materials. These are benzoic acids of different oxygen isotopic compositions (one of them is enriched in 180), produced within a joint project of A. Schimmelmann, University of Indiana, USA, and W. Brand & R. Wernerl, Max Planck Institute for Biogeochemistry, Jena, Germany. Those materials are prepared with the aim to serve as reference materials for 6180 analysis using online combustion techniques. Initial isotopic measurements were performed using both on-line and off-line combustion techniques. They resulted in provisional isotope values, which have to be verified and further refined by calibration measurements performed in other experienced laboratories. Two nitrate materials were prepared at USGS, Reston, USA by J.K. B6hlke and T. Coplen. One is depleted in 170 and 180 with normal 170/180 ratios. The other is enriched in 170 and 180, with anomalously high 170 content. Those materials could be used for the normalization of oxygen-bearing substances and for calibration of nitrogen-bearing materials, especially atmospheric materials that show 170 anomalies.

40.3.3 Carbon (plus oxygen and hydrogen) stable isotopic reference materials The number of stable isotope analyses of carbon-bearing materials has increased drastically with the development of new mass spectrometric systems and analytical techniques, involving the use of elemental analyzers, gas chromatographic columns and continuous flow systems. Isotope techniques are incorporated in an increasing number of scientific disciplines and new applications emerge continuously. Therefore, suitable isotopic reference materials are requested continuously. The variety of applications has resulted in a change of the priorities for new materials and demanded an improved characterization of existing ones. The assessment of the isotopic homogeneity especially needs further attention due to the smaller and smaller amounts needed for sample analysis in modern on-line mass spectrometric techniques. 1. Now at ETH, Ziirich, Switserland

892


40.3.3.1 Inorganic carbon isotopic reference materials In the early 1950s, 6180 measurements were often expressed against the so-called PDB primary reference material. The PDB material originated from the CaCO3 of the rostrum of Cretaceous belemnites (Belemnitella Americana) collected in the PeeDee formation of South Carolina, USA. The CO2 obtained from the reaction of PDB with 100% H3PO4 (McCrea, 1950) was used for calibration of oxygen isotope paleotemperature measurements (Epstein et al., 1953) and for carbon isotope ratio variations in natural compounds (Craig, 1953). Its 613C and 6180 values were close to the average isotope ratio values of marine limestone. The PDB primary reference material has long been exhausted. The calibration of measurements with respect to PDB in practice was performed by using the reference material NBS 20 (Solenhofen limestone) (Craig, 1957). Doubts on the isotopic stability of NBS 20, especially for 6180, due to potential exchange of this finely ground material with air moisture and CO2 have lead to the discontinuation in the use of NBS 20. As replacement, the carbonate material NBS 19 was produced. NBS 19 is derived from white marble of unknown geological origin (Friedman et al., 1982). It was provided by I. Friedman, U.S. Geological Survey, using a slab of marble representing the form of a toilet seat, and for this reason it is also referred to as TS limestone. The calibration of NBS 19 relative to PDB was performed using NBS 201. It was recommended at two Advisory Group Meetings organised at the IAEA (Gonfiantini, 1984; Hut, 1987) that the use of the PDB scale be discontinued and that it be replaced by a new scale called the VPDB scale (Vienna-PDB), which would be anchored by assigning a fixed 6 value to NBS 19 carbonate, as stated in Table 40.6. VPDB itself never existed as material, but it is defined through NBS 19 in such a way that the VPDB scale corresponds nominally to the previous PDB scale, i.e. the isotopic values of NBS 19 expressed versus PDB are nominally exactly the same as those expressed versus VPDB (Table 40.6). Because no measurement can be performed without an associated uncertainty, the calibration value of NBS 19 versus PDB (via NBS 20) inherently contains some additional uncertainty. Some problems related to the scale transition are discussed in Coplen (1996c). The reason for introduction of the new scale, was to eliminate this (unknown) uncertainty component for all reported data using NBS 19 as a calibration material. Stating a 613CVPDB or 618OVPDB value, therefore, implies the calibration being performed through the existing NBS 19 calibration material and its defined isotopic values. This minimises confusion on data reporting, as long as the supply of NBS 19 is sufficient (at least for the next 20 years). Complementary to NBS 19, an additional calcite material named NBS 18 is used as reference material with slightly more negative (~13C and 6180 values. NBS 18 is a carbonatite from Fen, Norway, and was collected by B. Taylor, University of California, Davis, USA. It was prepared by H. Friedrichsen, University of Ttibingen, Germany and I. Friedman, J.R. O'Neil and G. Cebula, USGS (Friedman et al., 1982). NBS 23, a strontium carbonate, was prepared by I. Friedman, but its distribution was discontinued due to doubts on its isotopic homogeneity. For studies of methane and other 13C depleted materials the LSVEC Lithium carbonate is used. LSVEC was originally pre1. Note: While the NBS 19 isotopic composition was actually measured versus PDB, the values are defined versus VPDB and therefore are used to anchor the VPDB scale.

893


Table 40.6 - Carbon 6-values versus VPDB of the inorganic carbon isotope reference materials with

associated standard uncertainties at lo-level. Available materials are m a r k e d in bold. 6-values m a r k e d with * are reported versus PDB. References given in the 613C column are relevant for both 613C and 6180 values. CM stands for "Calibration material" and RM for "Reference material". Name

Material

PDB NBS 20

carbonate limestone

Status

Distribution

613C [%0]

6180 [%o]

exhausted exhausted

0* - 1.06"

0* Isoto pi cally exchanged

Craig (1957) -1.08+0.06

Hut (1987) NBS 19 IAEA-CO-1

limestone calcite

CM RM

IAEA, NIST IAEA, NIST

carbonatite

RM

IAEA, NIST

+1.95 +2.48 + 0.03

-2.20 -2.44 + 0.07

Gonfiantini et al. (1995) NBS 18

-5.01 + 0.06

-23.00 +_0.07

Gonfiantini et al. (1995) IAEA-CO-8

calcite

RM

IAEA, NIST

-5.75 + 0.06

-22.67 + 0.19


SrCO3

discontinued

-35.32+0.16

-29.98...-30.54

Hut (1987) LSVEC

Li2CO3

RM

IAEA, NIST

-46.48 + 0.15

-26.64 _+0.25

Gonfiantini et al. (1995) IAEA-CO-9

BaCO3

RM

IAEA, NIST

-47.12 + 0.15

-15.28 + 0.09


CO2

exhausted

-41.64+0.17*

-25.75_+0.37*

Gonfiantini (1984) -41.59_+0.06

-25.86_+0.35

Hut (1987) NBS 17

CO2

exhausted

-4.48+0.10"

-8.37+0.25*

Gonfiantini (1984) -4.45_+0.05 RM8562

CO2

-8.51+0.15

Hut (1987) RM

IAEA, NIST

-3.76 + 0.03

-8.45 + 0.11

Verkouteren (1999) RM8563

CO2

RM

IAEA, NIST

-41.56 + 0.06

-23.72 + 0.11

Verkouteren (1999) RM8564

CO2

RM

IAEA, NIST

-10.45 + 0.04

+0.19 + 0.10

Verkouteren (1999) NGS 1

Hut (1987) NGS 2

Hut (1987)

NGS 3

Hut (1987)

Natural hydrocarbon gas Natural hydrocarbon gas Natural hydrocarbon gas

NIST

NIST

NIST

(CH4): -28.6...-29.16 (C2H6): -25.42...-26.4 (C3H8): -19.9...-21.67 (C H4): -43.1...-44.92 (C2H6): -31.1...-32.1 (C3H8): -23.5...-25.9 (CO2): -7.5...-8.9 (CH4): -69.6...-73.1 (C2H6): -51.37...-59.8

-

-

-

894

Chapter 40 - M. Gr6ning

pared as lithium isotope reference material by H. Svec, Iowa State University, USA. Because some reference materials were nearing exhaustion, three additional materials were introduced: IAEA-CO-1 (Carrara marble) and IAEA-CO-8 (calcite from the Kaiserstuhl, Germany), both prepared at the IAEA, and IAEA-CO-9 (BaCO3, prepared by C. Brenninkmeijer, IGNS, Lower Hutt, New Zealand). Due to the recent discovery of additional batches of NBS 18 and LSVEC from the original production time, the supply of those reference materials is secured for at least ten more years. NBS 16 and NBS 17, two pure CO2 materials, were prepared by T.B. Coplen and C. Kendall of the USGS in sealed glass ampoules (Coplen & Kendall, 1982). The two materials are exhausted. Several years ago, three new CO2 gas reference materials were prepared by M. Verkouteren, NIST, Gaithersburg, Maryland, USA (Verkouteren et al., 1998; Verkouteren, 1999), to provide a means for standardizing gas measurements without the necessity of any preparation by the carbonate acid reaction. These three materials span a broad isotopic range, covering values from those of atmospheric carbon dioxide to those of methane. These three gaseous reference materials were flame-sealed in glass tubes to ensure their isotopic stability on long time scales. Special care was taken for the filling procedure to avoid any isotope fractionation between individual tubes (see also Chapter 43 in this Volume). Three natural gases, consisting mainly of CH4, C2H6 and higher hydrocarbons, were collected and named NGS1, NGS2 and NGS3. They were stored in 150-ml pressurized steel cylinders (Hut, 1987). NGS1, being related to coal reservoirs, was collected by G. Hut, University of Groningen, Netherlands; the oil related NGS2 and the biogenic NGS3 were collected by T. Ricchiuto, AGIP, Milano, Italy. The three samples were originally available from IAEA and from M. Schoell, Chevron, La Habra, USA. New 50-ml sample cylinders for all three materials were filled recently at NIST, and those materials are now available from NIST. Isotopic values of all of these carbonate and CO2 reference materials versus VPDB are listed in Table 40.6, and reference materials still available are marked in bold. The ~}18OVPDB values may also be expressed versus VSMOW or VPDB-CO2 through the following two conversion formulae (Hut, 1987)" 618OvsMow = (1.0309 x 518OVPDB) + 30.9 618OVPDB_CO2 = (618OVPDB- 10.25) / 1.01025

[40.4] [40.5]

As a general rule, 6180 data of carbonates and of CO2 gas are reported versus VPDB, whereas for 6180 measurements of all other materials, the data should be reported versus VSMOW (Coplen et al., 1996). The author is of the opinion to avoid, whenever possible, reporting 6180 data relative to VPDB-CO2 due to the risk of confusing readers with data reporting relative to VPDB. Exceptions are only measurements on CO2 gas samples, as in atmospheric air, or on laboratory reference gases, which cases have to be clearly marked.

International StableIsotopeReferenceMaterials

895

No normalization is currently suggested for 613C,but it is good laboratory practice to state the isotopic composition of any reference material, which had been analysed with samples. That, in principle, allows any data user to normalize data later on, in case the need arises. A remark has to be given on the phosphoric acid digestion, which is used for preparation of CO2 gas from carbonate samples, as well as from the carbonate reference materials. The oxygen isotopic fractionation between carbonate and evolved CO2 depends on the acid reaction and associated parameters. It was determined experimentally more than three decades ago and later tabulated (Friedman & O'Neil, 1977) as common reference. It has been suggested that this oxygen isotope fractionation factor be re-determined (Gonfiantini, 1984) using normalised scales for the measurement results. This is not achieved yet. In fact several recipes exist for the preparation of the phosphoric acid, its storage and its use (e.g.McCrea, 1950; Urey et al., 1951; Coplen et al., 1983). With increasing acid concentration, its degree of polymerization and water release increases. Therefore, 100% phosphoric acid does not exist as such, and nominally even higher concentrations can be achieved (see also discussion in Volume II, Part 3, Chapter 4-5.1.5). The acid concentration can be controlled through density measurements. Because concentrated phosphoric acid is quite hygroscopic, each exposure of it to air slightly changes its concentration, and affects the oxygen isotopic fractionation of the acid reaction. Several chemicals added to the acid are thought to remove organic impurities (e.g. Cr203, H202, V205); however, no systematic assessments on possible effects of those additives on the resulting oxygen isotopic fractionation are available (see also discussion in Volume II, Part 3, Chapter 4-5.1.3 on the pretreatment of carbonate samples).

40.3.3.2 Organic carbon isotopic reference materials Quite early in the development of carbon stable isotope measurement methods, organic compounds were analysed to investigate the isotopic abundances and the different kinds of isotopic fractionation effects (Craig, 1953). The obvious principle to try to calibrate isotope ratio measurements of a certain compound by a reference material of similar chemical and physical properties (IT p r i n c i p l e - 'identical treatment') (Coplen et al., 1996; Werner & Brand, 2001) is especially relevant and applicable for organic compounds. There, the combustion techniques used to produce CO2 gas differ from the quite different preparation technique for usual carbonate reference materials (acid digestion). By the 1960s additional combustable reference materials were being produced, including NBS 21 (graphite) (Eckelmann et al., 1962) and NBS 22 (mineral oil, by S. Silverman, Chevron Oil Company, La Habra, USA) (Silverman, 1964). Graphite is included in the category organic reference materials due to its similar preparation as for organic materials, substantially different than for carbonate materials. Refinements of the initial measurements of those two materials were performed by interlaboratory studies (Schoell et al., 1983; Coplen et al., 1983). In the following years, additional materials were produced and calibrated, including polyethylene (PEF-1, now named IAEA-CH-7, by H. Gerstenberger and M. Herrmann, Zentralinstitut fuer Isotopen- und Strahlenforschung, Leipzig, Germany) (Ger-

896


stenberger & Herrmann, 1983), sucrose (Anu.Sucr., now IAEA-CH-6, by H. Polach, ANU, Canberra, Australia) (Hut, 1987) and graphite (USGS24, by T.B. Coplen, USGS, Reston, USA). All of those materials were produced more than two decades ago. Their calibration was performed by off-line combustion methods, using relative large amounts of substance. Newly developed on-line combustion techniques use amounts that are orders of magnitude lower. That imposes questions on the homogeneity of the existing reference materials at such low amount levels. No systematic tests have been reported so far. Only NBS 22 oil is an exception as liquid can be assumed to be homogeneous at all amount levels. Two sets of reference materials were produced 15 years ago consisting each of two materials enriched in 13C for applications in the medical and biological field (Parr & Clements, 1991). These include two 13C-labelled sodium bicarbonates prepared by I.I. Lesk, MSD Canada Ltd. (IAEA-303) and two 13C-labelled UL-D-glucoses, prepared by D. Halliday, Clinical Research Centre, Harrow, UK (IAEA-309). Two certified reference materials, BCR-656 ethanol from wine and BCR-660 hydro alcoholic solution 12% vol., were prepared by M. Lees, Eurofins Scientific, Nantes, France, one certified reference material, BCR-657 glucose powder, was prepared by C. Guillou and G. Remaud, EC Joint Research Centre, Ispra, Italy, in order to provide the means for proper analysis of wines and fruit juices in the European Community according to officially approved methods (Guillou et al., 2001). The parameters to be certified were the ~13C composition of the alcohol or sugar and the site specific deuterium content in the alcohol and/or water phase by deuterium- nuclear magnetic resonance. The isotopic composition of the available organic carbon isotopic reference materials is given in Table 40.7. A steadily spreading number of new applications of stable isotope measurements include more and more different organic compounds and at the same time require standardisation of such measurements using reference materials as similar as possible to the investigated compounds. It is certainly not possible at international level to provide dozens of stable isotopic reference materials for all possible applications, but attempts have been started to produce a limited number of suitable materials. The main criteria in the selection of substances are the stability of the compounds with time, ease of storage, ease of handling and preparation, combustibility, and suitability as reference material for a whole class of substances. Several additional materials are under investigations, including a batch of cellulose that was prepared by IAEA for stable carbon isotope analysis from a larger stock of cellulose by milling it down to a fine powder and homogenizing it. This material

897


Table 40.7 - C a r b o n 6-values v e r s u s VPDB of the available organic carbon isotopic reference materials w i t h associated standard uncertainty at l o-level. References see in text. RM stands for "Reference material", CRM for "Certified reference material". Name

Material

NBS 21

graphite

Status

Distribution exhausted

613C [%o]

62H [%o]

-28.16+0.11

-

Hut (1987 NBS 22 USGS24

oil graphite

RM RM

IAEA, NIST IAEA, NIST

-29.74+0.12

-120+4

Gonfiantini et al. (1995)

Hut (1987)

-15.99+0.11

-

Gonfiantini et al. (1995) IAEA-CH-6

sucrose

RM

IAEA, NIST

-10.43+0.13

-

Gonfiantini et al. (1995) IAEA-CH-7

polyethylene

RM

IAEA, NIST

-31.83+0.11

-100.3+2.1

GonJi"antini et al. (1995) IAEA-303A

NaHCO3

RM

IAEA

93.3

-

Parr & Clements (1991) IAEA-303B

NaHCO3

RM

IAEA

466

-

Parr & Clements (1991) IAEA-309A

UL-D-glucose

RM

IAEA

93.9

-

Parr & Clements (1991) IAEA-309B

UL-D-glucose

RM

IAEA

535.3

-

Parr & Clements (1991) BCR-656

ethanol

CRM

BCR(IRMM)

-26.91+0.07

Guillou et at. (2001) BCR-657

sugar

CRM

BCR(IRMM)

d a t a see in: Guillou et al. (2001)

-10.76+0.04

Guillou et at. (2001) BCR-660

e t h a n o l in w a t e r

CRM

BCR(IRMM)

-26.72+0.09

Guiltou et al. (2001)

d a t a see in: Guillou et al. (2001)

was successfully tested for its 613C isotopic homogeneity (W. Stichler, GSF, Neuherberg, Germany; variability smaller than +0.02%o for samples amounts of about 1-2 mg). A batch of pure caffeine was recently produced by W. Brand & R. Werner, Max Planck Institute for Biogeochemistry, Jena, Germany. It is now at NIST for homogenization and bottling. One non-enriched benzoic acid (natural 6180 level) could potentially also be used as a carbon isotopic reference material. Two L-glutamic acids were prepared by H. Qi and T. Coplen, USGS, Reston, USA, named USGS40 and USGS41. One of them has natural C and N isotopic compositions, and the other is enriched in b o t h 13C and 15N by about 50 per mill. Eventually two batches of oxalic acid at the IAEA could also be milled and homogenized for use as stable carbon isotope reference materials. All of these new materials will need to be carefully calibrated relative to the VPDB scale by direct comparison to NBS 19 derived CO2. At the same time, the isotopic composition of the existing organic reference materials should be re-assessed to try to reduce the uncertainty of calibration values. This can be undertaken only in a carefully planned exercise involving several "high-precision" laboratories measuring the whole suite of organic reference materials and calibrating them versus NBS 19. It is hoped, that after bottling of the candidate materials, such an exercise can be organized

898


in the near future to improve the calibration of laboratories and their internal laboratory standards. While no normalization has been suggested for reporting of stable carbon isotope ratio data, it is recommended that authors report together with their data, also the values for reference material carbon isotope ratios that they did measure or would have measured had they analyzed them along with their samples. This provides a means to perform calculations on possible data normalization in the future.

40.3.4 Nitrogen stable isotopic reference materials The natural choice as the primary reference material for nitrogen relative isotope ratio measurements is atmospheric nitrogen gas, which seems to be isotopically homogeneous (Mariotti, 1983) with respect to the prevailing analytical precision of most laboratories. Several other reference materials with a wide range of isotope compositions exist and are used to calibrate nitrogen isotope measurements of different compounds (Coplen et al., 2002). These are especially useful for solid or liquid samples to test sample preparation methods and to avoid some problems associated with the purification of nitrogen from air. By the 1950s, G. Junk and H.J. Svec, Iowa State University, USA, prepared a nitrogen gas standard (Junk & Svec, 1958), which was later on split into aliquots in sealed glass tubes by C. Kendall, USGS (Kendall & Grim, 1990) and was named NSVEC. An additional nitrogen gas standard, NBS 14, prepared some decades ago has been exhausted since the 1960s. IAEA-N-1 and IAEA-N-2 are ammonium sulfates, prepared by E. Salati, Centro de Energia Nuclear na Agricultura, Brazil. IAEA-NO-3 is a potassium nitrate (formerly called IAEA-N3) and was prepared by A. Mariotti, Universit6 P. and M. Curie, Paris, France. USGS25, USGS26 (ammonium sulfates) and USGS32 (potassium nitrate) were prepared by J.K. B6hlke, USGS, Reston, USA (B6hlke et al., 1995; B6hlke & Coplen, 1995) by mixing commercially available 15Nenriched compounds with compounds of natural terrestrial abundance to achieve the desired 815N composition. The 815N value stated for USGS32 in Table 40.8 is the mean value derived from the whole data set in B6hlke & Coplen (1995). There is strong evidence that some of those USGS32 data were affected by systematic laboratory offsets, and the recommended value will most probably be adjusted to about +180%o in near future. A series of isotopic reference materials substantially enriched in 15N was prepared for the IAEA within an intercomparison study for medical and biological applications (Parr & Clements, 1991). Three 15N-enriched ammonium sulfates were prepared by E. Fern, Vevey, Switzerland (IAEA-305 set and IAEA-311). Two 15N-enriched urea reference materials (IAEA-310 set) were prepared by H. Faust, Leipzig, Germany. The values of the existing nitrogen isotopic reference materials are presented in Table 40.8.

899


Table 40.8 - Nitrogen 6-values versus air-N2 for the existing nitrogen isotope reference materials and their associated standard uncertainties at l o-level. RM stands for "Reference material". Name

Material

Status

Distribution

615N [%o]

exhausted

-1.18 Kendall & Grim (1990) -2.77+0.05 B~hlke & Coplen (1995) +0.43+0.07 B~hlke & Coplen (1995) +20.32+0.09 B~hlke & Coplen (1995) +4.69+0.09 B~hlke & Coplen (1995) -30.25+0.38 B~hlke & Coplen (1995) +53.62+0.25 B~hlke & Coplen (1995) +179.2+1.3 B~hlke & Coplen (1995) + 180( no rmalized ) B~hlke & Coplen (1995) +39.8+0.25[46] Parr & Clements (1991) +375.3+1.2146] Parr & Clements (1991) +47.2+0.7[46] Parr & Clements (1991) +244.6+0.4[46] Parr & Clements (1991) +4693+29[46] Parr & Clements (1991) -1.8+0.1 B~hlke et al. (2003) +2.7+0.1 B~hlke et al. (2003)

NBS 14

nitrogen gas

NSVEC

nitrogen gas

RM

IAEA, NIST

IAEA-N-1

a m m o n i u m sulfate

RM

IAEA, NIST

IAEA-N-2


RM

IAEA, NIST

IAEA-NO-3

potassium nitrate

RM

IAEA, NIST

USGS25


RM

IAEA, NIST

USGS26


RM

IAEA, NIST

USGS32

potassium nitrate

RM

IAEA, NIST

IAEA-305A


RM

IAEA

IAEA-305B


RM

IAEA

IAEA-310A

Urea

RM

IAEA

IAEA-310B

Urea

RM

IAEA

IAEA-311


RM

IAEA

USGS34

potassium nitrate

IAEA, NIST

USGS35

sodium nitrate

IAEA, NIST

6180 [%o]

+25.6+0.2 B~hlkeet al. (2003)

+25.7+0.2 B~htkeet al. (2003)

-14.8+0.2 B~hlke et al. (2003) +51.5+0.3 B~htke et al. (2003)

Two new materials were recently prepared by J.K. B6hlke, T. Coplen, and S. Mroczkowski, USGS, Reston, USA. They include potassium nitrate USGS34 depleted in 170 and 180 (normal 1 7 0 / 1 8 0 ) and sodium nitrate USGS35 enriched in 170 and 180 with anomalous 170 (B6hlke et al., 2003). USGS40 and USGS41, prepared by T. Coplen and H. Qi, USGS, Reston, Virginia, USA, are two L-glutamic acids, one with a natural 615N level and one isotopically enriched in 15N by about 50 per mill. Additional materials are under discussion to provide a means for better calibration of organic nitrogen-bearing samples. They include thiourea, methionine and Nmethyl anthranilic ester (IAEA, 2001; see therein especially Table 2 in its annex and the report of working group B on biogeochemistry, food and ecology. Also found therein are results of discussions on an N20 gas reference material, mainly for atmo-


900 spheric and soil gas studies).

Past discussions on scale contractions observed in mass spectrometric measurements result in the recommendation to authors to report sample data together with the isotopic values of nitrogen isotopic reference materials had they been analyzed with the samples. When reporting values of 15N depleted or enriched reference materials used for the calibration of internal laboratory standards, one can provide information that can be efficiently used for re-assessing those data later on for any kind of normalization. 40.3.5 Sulfur stable isotopic reference materials The early choice of meteorite materials as reference materials for the calibration of sulfur stable isotope ratio measurements was certainly understandable because they represent the primordial sulfur composition as the average of terrestrial sulfur inventory. But this selection caused considerable problems due to chemical impurities, because of isotopic inhomogeneities, and due to the choice of different reference materials in different countries. By 1960, the Canyon Diablo Troilite CDT (FeS phase from a large octahedrite iron meteorite, Meteor Crater, Arizona, USA) had been adopted as the primary reference material (McNamara & Thode, 1950; Jensen & Nakai, 1962); nevertheless, Russian scientists continued to use the Shikote Alin meteorite (Robinson, 1995). The supply of CDT has been exhausted for more than a decade1, but this material was still used to define the internationally accepted stable sulfur isotope ratio scale (634SCDT) until recently.

After problems with the meteorite reference materials were recognized and discussed (Nielsen, 1984) (see also latest confirmation of this by SF6 measurements of CDT: Beaudoin et al. (1995)), a mineral sphalerite was introduced, but it was also found to be inhomogeneous. Several other approaches preparing additional natural sulfur-bearing materials were not quite successful. OGS, a raw precipitated BaSO4 from sea water (by Y. Horibe, University of Tokyo), caused quite early doubts on its homogeneity. The same occurred with Soufre de Lacq, an elemental sulfur that was derived from natural gas (later renamed to IAEA-S-4), provided by E. Roth, CEN, Saclay) France, which also caused doubts on its homogeneity. Distribution of all three reference materials was subsequently terminated by the IAEA. Distribution of NBS 122, a sphalerite supplied by S. Halas, University of Lublin, Poland, which was recognized to be inhomogeneous, was also terminated. Recent investigations on IAEA-S-4 (Soufre de Lacq) by the USGS (Qi & Coplen, 2003; Carmody & Seal, 1999) indicate that this material indeed is homogeneous and interlaboratory discrepancies may rather have been caused by laboratory offsets. Therefore the distribution of IAEA-S-4 by IAEA has started again. Two other materials that were subsequently prepared include NBS 123, another sphalerite, and NBS 127, a BaSO4 (ion exchanged sea water sulfate, prepared by J.R. O'Neil, USGS, Menlo Park, USA). This listing documents the prevailing unsatisfactory situation for sulfur isotope reference materials. 1. I was informed that one solid piece of the original CDT (approx. I kg) is stored at the Riksmuseum Stockholm, Sweden.

901


Finally, the urgent need for pure chemical compounds as reference materials resulted in the subsequent production of three Ag2S reference materials with substantially different isotopic composition (IAEA-S-1, IAEA-S-2, IAEA-S-3) by B.W. Robinson, Institute for Geological and Nuclear Sciences, Lower Hutt, New Zealand. In view of the situation of an inhomogeneous and exhausted CDT reference, the change to a new scale was recommended (Gonfiantini et al., 1995), based on IAEA-S-1 as calibration material. The VCDT scale (Vienna-CDT) was established by agreement of CAWIA (IUPAC-CAWIA, 1997) by adopting a defined ~)34S value for IAEA-S-1 to keep the new scale as close to the CDT scale as possible. Five years ago, two additional BaSO4 materials named IAEA-SO-5 and IAEA-SO6, having different isotopic composition for both (534S and for 6180 values, were produced jointly by H.R. Krouse, University of Calgary, Canada and S. Halas, University of Lublin, Poland. Those two materials were prepared to be used in addition to NBS 127 to calibrate both 634S and 6180 values in sulfate samples. However, an initial interlaboratory comparison conducted in 1998 resulted in large value discrepancies and made it impossible to assign reliable isotope ratio values to the new materials. Only two years ago, results from SF6 measurements in three laboratories and improvements reported for SO2 measurement techniques resolved this dilemma. The isotopic data on the available sulfur reference materials and proposed reference mateTable 40.9- Sulfur 6-values versus VCDT for the available sulfur isotopic reference materials and their associated standard uncertainties at l o-level. 6-values marked with * are reported versus CDT. w provisional data from new unpublished measurements. CM stands for "Calibration material" and RM for "Reference material". Status

Distribution

634S [%o]

Troilite (FeS) BaSO4 S (elemental)

RM

exhausted discontinued IAEA, NIST

sphalerite sphalerite

RM

discontinued IAEA, NIST

NBS 127

BaSO4

RM

IAEA, NIST

IAEA-S-1 (NZ1) IAEA-S-2 (NZ2) IAEA-S-3 IAEA-SO-5

Ag2S AgRS Ag2S BaSO4

CM RM RM RM

IAEA, NIST IAEA, NIST IAEA, NIST IAEA

IAEA-SO-6

BaSO4

RM

IAEA

SF6

-

IRMM

0* +20.42+0.42" +16.90+0.12 Qi & Coplen (2003) +0.18+0.14" + 17.09+0.31" +17.44+0.10w +20.32+0.36* +21.17+0.12 Halas & Szaran (2001) -0.30 +22.66+0.13w -32.30+0.12w +0.49+0.09w +0.15+0.05 Halas & Szaran (2001) -34.18+0.07w -34.04+0.11 Halas & Szaran (2001) +17.33+0.22 e

Name CDT OGS IAEA-S-4 (Soufre de Lacq) NBS 122 NBS 123

IRMM-PIGS-2010

Material

e 634S value calculated from absolute isotope ratio measurements (see Table 40.1).

6180 [%o] +9.00+0.65

+9.34+0.32

902


rials were discussed during the last IAEA experts meeting on reference materials in the year 2000. Details of the discussions will be presented elsewhere. The isotope ratio values of the available sulfur reference materials are provided in Table 40.9. Before that IAEA meeting, an absolute ratio determination of the sulfur isotope abundance in IAEA-S-1 was undertaken using the SF6 technique at the Institute of Mineral Resources, CAGS, Beijing (Ding et al., 1999) and the Institute for Reference Materials and Measurements IRMM, European Commission, Geel, Belgium (T. Ding and S. Valkiers). The results compared quite favourably with relative 6-measurements using the SF6 technique performed by B.E. Taylor in Canada and T. Ding in China. This is a very successful case, which results in a direct link of the conventional relative 634S VCDT-scale to the S.I.-system of units. The differences of results produced by the SF6 and SO2 techniques are discussed elsewhere (Beaudoin & Taylor, 1995). But the establishment of these two sets of three reference materials each (in form of Ag2S and BaSO4), provides the means for the necessary routine normalization of ~534Smeasurements in laboratory operation. In addition, an SF6 gas, distributed by Messer-Griesheim GmbH, Krefeld, Germany, was assessed for its absolute sulfur isotope ratios by the Isotope Measurements Unit at IRMM in Geel and was described by them as Primary Isotopic Gas Standard (PIGS) (Taylor, 1998). Together with the calculated absolute isotopic abundance of the virtual VCDT, a 634S value for this SF6 gas can be calculated, which can be important for a direct calibration of secondary laboratory reference gases. In an advisory group meeting at the IAEA, the need for organic ~34S reference materials was expressed by the participants (IAEA, 2001). As for 615N measurements, the proposed materials included thiourea, methionine and N-methyl anthranilic acid ester. No results on the suitability of these materials are available so far. 40.3.6 Lithium stable isotopic reference materials For reporting of relative lithium stable isotope ratios, the primary reference material is LSVEC, a lithium carbonate prepared by H.J. Svec, Iowa State University, USA. It is assigned a 67Li value of 0%o by definition. For its Table 40.10 - Lithium 67Li-values versus LSVEC for the existing lithabsolute isotope ratios, see ium isotope reference materials. PRM stands for "Primary reference Table 40.1. Note that for material" and RM for "Reference material". lithium, the expression of Material Status Distribution 67Li [%o] relative isotope data as Name 66Li is still common. In LSVEC Li2CO3 PRM IAEA, NIST 0 accordance with IUPAC IRMM-016 Li2CO3 RM IRMM 0 (Coplen, 1996b), the IRMM-015 Li2CO3 RM IRMM -996 reporting as 67Li data is recommended.

International Stable Isotope ReferenceMaterials

903

In addition to LSVEC, two more lithium carbonate materials were prepared by the Institute for Reference Materials and Measurements IRMM in Geel, Belgium. IRMM016 (IRMM, 1997; Qi et al., 1997) is isotopically indistinguishable from LSVEC, and IRMM-015 (IRMM, 1993) is substantially enriched in 6Li. The isotope values of those reference materials are presented in Table 40.10. 40.3.7 Boron stable isotopic reference materials The primary reference material for boron isotope ratio measurements is boric acid SRM 951 (former name NBS 951) as supplied since the 1960s by NIST, Gaithersburg, USA (Catanzaro et al., 1970). Its 11B/10B ratio is given in Table 40.1; see also NIST Certificate of analysis (1999). A few other isotope ratio determinations on this material are discrepant for unknown reasons (Spivack & Edmond, 1986; Leeman et al., 1991). This material is used to define the zero-point of the 611B scale; so, by definition, 611BSRM951 is 0 %0. A boron acid material with a similar boron isotope ratio is provided as IRMM011 by the Institute for Reference Materials and Measurements (IRMM), Geel, Belgium (see Table 40.1) (De Bi6vre & Debus, 1969; IRMM Certificate, 2001). An overview on absolute ratio measurements of SRM 951 can be found in Deyhle (2001). In the past, only few additional materials were available for calibration of boron isotope ratio measurements: first, the JB-2 and JR-2 reference materials (island arc tholeiitic basalt and rhyolite distributed by the Geological Survey of Japan), which are used for boron isotope result normalization in silicates (Nakamura et al., 1992; Tonarini et al., 1997; Deyhle, 2001; Kasemann et al., 2001); second, the two NIST synthetic silicate glasses, SRM 610 and SRM 612, for which their boron isotopic composition was determined recently (Kasemann et al., 2001); third, ocean water, which has a uniform dissolved boron isotopic composition (Spivack & Edmond, 1987). Three other certified boron reference materials were prepared by J. Vogl et al. at Bundesanstalt ffir Materialforschung und-prfifung (BAM), Berlin, Germany with the primary purpose of producing boric acid reference materials for ICP-mass spectrometers used in nuclear reactors. These materials are available from BAM under the names BAM-I001, BAM-I002 and BAM-I003 (Vogl et al., 2002). Recently, the preparation of additional boron stable isotopic reference materials, covering different matrices, was initiated by R. Gonfiantini, Istituto di Geoscienze e Georisorse, Pisa, Italy. Eight different natural materials were prepared at that institute with assistance through an IAEA technical contract (two groundwater and a sea water sample, Elba tourmaline, Etna basalt, Lipari obsidian, Maiella limestone, Montelupo clay) (Tonarini et al., 2003). The provisional mean values were derived from the results of an extended interlaboratory comparison exercise with involvement of 15 laboratories engaged in boron isotopic measurements on natural matrices (Gonfiantini et al., 2003). Significant discrepancies exist in data reported from different laboratories, which can partly be attributed to differences in precision and systematic offsets, but also due to calibration problems (linear shifts of all data between laboratories), disturbing matrix effects and most probably to additional - so far, undetermined - systematic effects. The isotope values of existing boron reference materials are pro-

904


Table 40.11 - Boron 6-values versus SRM 951 for the existing boron isotopic reference materials and quality control materials with their associated standard uncertainty at l o-level. PRM stands for "Primary reference material", RM for "Reference material" and QCM for "Quality control material". Name

Material

Status

Distribution

6UB [%o]

SRM 951 IRMM-011 JB-2

boric acid boric acid basalt

PRM RM RM

NIST IRMM Jap.Geol.Surv.

0 -0.16 +7.09+0.08 Nakamura et al. (1992)

+7.23+0.24 Tonarini et al. (1997)

+7.13+0.34 Vogl et al. (2002)

JR-2

rhyolite

RM

Jap.Geol.Surv.

+2.71+0.8 Vogl et al. (2002)

SRM 610

silicate glass

RM

NIST

-1.05+0.8 Vogl et at. (2002)

SRM 612

silicate glass

RM

NIST

-1.07+0.8 Vogl et al. (2002)

IAEA-B-1

groundwater

QCM

IGG Pisa

+37.7+2.1 Gonfiantini et al. (2003)

IAEA-B-2

groundwater

QCM

IGG Pisa


IAEA-B-3

sea water

QCM

IGG Pisa

-21.3+0.9 Gonfiantini et al. (2003)

IAEA-B-4

tourmaline

QCM

IAEA


IAEA-B-5

basalt

QCM

IAEA


IAEA-B-6

obsidian

QCM

IAEA


IAEA-B-7

limestone

QCM

IAEA


IAEA-B-8

clay

QCM

IAEA


v i d e d in Table 40.11. A d d i t i o n a l efforts will be necessary to i m p r o v e the intercalibration of laboratories active in the field of b o r o n stable isotope ratio m e a s u r e m e n t s .

40.3.8 Chlorine stable isotopic reference materials For chlorine stable isotope m e a s u r e m e n t s , data are c o m m o n l y expressed relative to sea w a t e r chloride, w h i c h w a s t h o u g h t to be isotopically h o m o g e n e o u s for 637C1 w i t h i n a b o u t +0.15%o ( K a u f m a n n et al., 1988; Coplen, 2001b). Therefore, a S t a n d a r d M e a n O c e a n Chloride (SMOC) w a s p r o p o s e d as p r i m a r y reference material for chlorine stable isotope m e a s u r e m e n t s and the scale realized by use of i n d i v i d u a l s e a w a t e r s a m p l e s in laboratories. However, recently Y. Xiao of the Q i n g h a i Institute of Salt Lakes, Xining, China, has p u b l i s h e d sea w a t e r chlorine isotope ratios s h o w i n g sub-

905


stantial variability (Xiao et al., 2002a). To eliminate the difficulty of using sea water as the chlorine isotopic reference material, Y. Xiao has collected sea water and purified about I kg of NaC1 for use as a relative chloride isotope ratio reference material called ISL 354, which can anchor the chlorine isotope ratio scale (Xiao et al., 2002b). Since the early 1960s, NIST SRM 975 sodium chloride (Shields et al., 1962) has been the basis for absolute isotope abundance measurements of chlorine. Since SRM 975 is exhausted, it has been replaced by SRM 975a sodium chloride (NIST Certificate of analysis, 2001). The isotope ratio values of the existing chlorine reference materials are presented in Table 40.12.

40.4 Concluding remarks In view of the preparation of new isotopic reference materials and the recent measurements on existing isotopic reference materials, a re-evaluation of recommended isotopic values for existing reference materials is ongoing and will result in a set of Reference Sheets issued by the IAEA according to the requirements as stated in ISOGuide 31 and its recent revision (ISO/REMCO, 1998). The uncertainties for all recommended isotope ratio values will be recalculated using a consistent approach for the whole dataset available. It has to be kept in mind, however, that many measurement data were provided long ago by laboratories that did not include sufficiently detailed information on their measurement uncertainty. Therefore, often the uncertainty for the reference values cannot be expressed according to international recommendations (ISO/BIPM, 1995; EURACHEM, 2000). The proper use of reference materials as described in ISO-Guide 33 (ISO, 1998) is discussed elsewhere (Gr6ning et al., 1999). The available reference materials are intended to calibrate local laboratory standards that are prepared by the individual laboratories. The reference materials are NOT intended to be used themselves for quality control purposes. For the distribution by IAEA of all stable isotope ratio reference materials, a rather strict rule applies" Each laboratory is entitled to order one unit of any reference material only once in a threeyear period. This limitation was set to preserve the availability of these valuable reference materials for the maximal possible time and therefore to ensure the comparability of results from laboratories for as many decades as possible. It is hoped that with Table 40.12 - Chlorine 6-values versus SMOC for the existing chlorine isotopic reference materials and their associated standard uncertainty at 1a-level. PRM stands for "Primary reference material" and RM for "Reference material" Name

Material

Status

Distribution

637C1 [%o]

SMOC SRM 975

NaC1

PRM RM

Exhausted

SRM 975a

NaC1

RM

NIST

ISL 354

NaC1

RM

IAEA

0 +0.43+0.04 Xiao et al. (2002) +0.2+1.5 Xiao et al. (2002) +0.05+0.02 Xiao et al. (2002)

906

Chapter 40 - M. Gr6ning

the preparation of additional reference materials, this limitation can be eased and more reference materials can be made available to laboratories more frequently for improved calibration of measurements. The following rules should be applied for an optimal calibration using available reference materials (Gr6ning et al., 2004), regardless of the element under consideration: The laboratory should prepare one or more local laboratory standards with chemical characteristics similar to those of the reference material used for the calibration. The amount of local laboratory standards to be prepared should be at least sufficient for all measurements performed within the timeframe of four calibration cycles (one cycle corresponds to three years at this moment for all internationally distributed stable isotopic reference materials supplied by IAEA and NIST). This enables the laboratory to perform several calibrations with the same local laboratory standards within the period of 12 years, therefore improving with time the accuracy of their calibrated isotopic values. The largest possible number of measurements should be performed during one analysis day, including analysis of relevant reference materials and local laboratory standards to minimise the effects of long-term performance drifts. The combination of individually assessed data from several measurement days should improve the reliability of the calculated average isotopic value for local laboratory standards. Additionally, the probability of any systematic offset, due to daily performance fluctuations, can be minimized with this strategy. A long-term average or floating mean should be applied for evaluation of the entire dataset for the internal laboratory standard. In most cases, weighted means should be applied to account for varying performance on individual calibration measurement days and to calculate a reasonable uncertainty for the mean value. This may serve as a proxy for the standard uncertainty achieved in the laboratory for routine measurements during that time.

Acknow legdements I wish to thank Tyler Coplen, Philip Taylor, Tiping Ding and Roberto Gonfiantini for many constructive comments to improve the manuscript. I am especially grateful to the careful review by Tyler Coplen with corrections on language, style and many historical details on reference materials.

Handbook of Stable IsotopeAnalyticalTechniques,Volume 1 P.A. de Groot (Editor) 9 2004 ElsevierB.V. All fights reserved.

CHAPTER 41 The Nature and Role of Primary Certified Isotopic Reference Materials" A Tool to Underpin Isotopic Measurements on a Global Scale P. D. P. Taylor1, P. De Bi6vre & S. Valkiers Institute for Reference Materials and Reference Measurements, JRC-European Commission, B-2440 Geel, Belgium e-mail: [email protected]

Abstract This chapter focuses on isotopic reference materials with certified values that are traceable in the International System of Measurements (the SI) and explains how they can be produced. An overview is given of cases where such materials have already been realised. Although for many end-users the problems of not having such reference materials (and only having materials with 'consensus-value-only' or'assigned' values ) are often hidden, the usefulness of this approach is elucidated in the light of increased pressures to improve comparability of isotopic measurements across the borders of space, time, and scientific disciplines.

41.1 Introduction As amply demonstrated in other chapters of this book, isotopic measurements are used as powerful investigative tools in many areas of science: medicine, biochemistry, geochemistry, climatic and environmental and nuclear chemistry, hydrology and archeology. In all these areas, the isotope ratios as such are either of interest or they are used in the isotope dilution process. The availability of instruments of various kinds has enabled a soaring flight of isotopic measurements. Mass spectrometry is predominantly used, although infrared absorption spectrometry (Roth, 1997) can be applied in particular cases (e.g. isotopes of hydrogen in water, carbon in CO2). In mass spectrometry, instruments with different ion sources (electron impact, thermal ionisation, ICP), various types of spectrometer configurations (quadrupoles, multiple collectors, ...) and interesting ancillary chemical reactors (modules for oxygen or carbon isotope ratio measurements producing CO2, CO, ...). Scientists developed suitable measurement procedures for these purposes, including the chemical sample preparation. Although in many cases measurements can be used primarily inside a 'closed community' (e.g. within a laboratory, a single hospital), in most cases the measurement scientists and even more so their customers want to compare the results 1. Correspondence should be adressed to this author.

908

Chapter 41 - P.D.P. Taylor, P. De Bibvre & S. Valkiers

obtained. How does a particular value compare with results from other laboratories, obtained at a different time? What trends can be observed in the carbon isotope ratios over the past 10 or 30 years ? Can I compare the laboratory test results for this patient coming from another hospital with what is produced in our laboratory? Are the oxygen isotope ratios for a sample of ethanol from wine identical when measured in country A or B? To increase the comparability of measurement results outside the perimeter of a single laboratory, several measures can be taken. Proper training of staff is of the utmost importance and the establishment of quality systems into the laboratory has become widespread. However probably the most important factor, recognised early on, is the use of a common standard or isotopic reference material. How does this help? Most routinely used isotopic measurement instruments today are comparator devices: a comparison is made between a particular measured isotope ratio of the sample with that of a reference sample (see Figure 41.1). The measurand is named delta of the sample versus the reference sample (expressed in %0):

6 x - (R~e f 1) 1000

[41.1]

and the reference sample therefore carries by definition the value (~Ref = 0. The delta was initially used by Urey and his collaborators (Sam Epstein primarily; Epstein & Mayeda, 1953) when developing the 180 thermometer. Being unable to measure a reliable absolute content of 180 of the samples (this was only achieved by Baertschi years

Figure 41.1 - Graphical presentation of possible traceability chains for measurements linked to 'artefact-only' IRMs (see text). The lines on the left (1,2,3) correspond to the 6 values, w h e n hidden uncertainties are taken into account. (It should not be concluded from this figure that delta values can be added linearly, as this only holds to a very rough approximation ; correct equation: 602 = 601 + 612 + 601" 612/1000)

The Nature and Role of Primary Certified Isotopic Reference Materials ...

909

later; Baertschi, 1976), a differential measurement technique was developed to establish the 180/160 value against the ocean value, assumed constant after preliminary investigations. The delta notation was then a very practical way of reporting results. Extension of the use of this notation together with materializing the reference to ocean water for oxygen and deuterium, and using other reference materials for other elements, followed. This procedure in principle enables laboratories, not interested in absolute calibration of their measurements, to compare their data with those of other laboratories. At first sight, it does not seem to be extremely important to know the absolute value Rref. Instead, the reference sample often only has an assigned (consensus) value. This approach is very practical, economical and straightforward and has helped the development of isotope ratio measurement sciences to a great extent. Reference samples such as PDB, CDT, SMOW all have been construed upon this principle (see Table 41.1 and Gonfiantini et al., 1993). Although such materials are regularly called 'primary' materials by practising isotope scientist because they were first used, from a metrological point of view they are not primary. Primary methods of measurements is a concept used (Quinn, 1997) in many areas of scientific measurements (temperature, mass, concentrations) to give 'direct access' to the international system of measurements (the SI). The SI being a fairly 'new' invention which was exactly conceived to improve comparability of measurements via better traceability. In this chapter, the limitations of relying on reference samples with 'consensus values only' will be explained as well as the ways and means to produce reference samples having 'true' ratios, i.e. with isotope ratios having a primary character, which can therefore solidly anchor the delta scales.

41.2 Commonly used reference samples A Reference Material (i.c. Isotopic RM or IRM), according to the International Vocabulary on Metrologyl, is "a material or substance one or more of whose property values are sufficiently homogeneous and well established to be used for the calibration of an apparatus, the assessment of a measurement method, or for assigning values to materials". The definition given for certified reference materials is "a reference material, accompanied by a certificate, one or more of whose property values are certified by a procedure which establishes traceability to an accurate realization of the unit in which the property values are expressed, and for which each certified value is accompanied by an uncertainty at a stated level of confidence". Many commonly used reference samples are of the first type, i.e. comparator samples with an assigned delta value (e.g. 8 - 0) that may be accompanied by an indicative value of the isotope ratio. Table 41.1 summarises most of the commonly used IRMs. Many of them started their life at the 'spur of the moment', in the 'heat of the 1. VIM (Vocabulaire International des Termes Fondamentaux et G6n6raux de M6trologie/ International Vocabulary of Basic and General Terms in Metrology), 2 nd Edition, ISO 1993.

910

Table 41.1 - Overview of 'comparator' IRMs commonly used in the differential isotopic measurement community. Nominal isotopic composition (in parts per 1000)

RM __

V SMOW-water Oa GISP-water -189.8 -428a SLAP-water -66.7 NBS30-biotite -118.5 NBS22-oil PEFI-polyethylene foil -100.3 USGS24-graphite Sucrose ANU-sucrose NBS18-carbonate NBS19-TS limestone LSVEC-lithium carbonate NBS28- silica sand (optical) IAEA-N1-ammonium sulfate IAEA-N2-ammonium sulfate IAEA-N3-potassium nitrate USGS25-ammonium sulfate USGS26-ammonium sulfate NSVEC- gaseous nitrogen Soufre de Lacq-elemental sulfur NZ1-silver sulfide NZ2-silver sulfide NBS123-sphalerite NBS127-barium sulfate USGS32-potassium nitrate

Oa

-24.85 -55.5a +5.1

-

0.0832a

-29.73 -31.77 -15.9 -10.47 -5.04 +1.95a -46.7

+7.16 +28.65 +3.0 +9.58

Oa

+0.4 +20.3 +2 to +4 -30.4 +53.5 -2.81

-

-

+9.4

-+16.0 -0.3 +21.0 -+17.0 +20.32

+179.9

~

Calibrated measurement. In the case of hydrogen, the absolute values of VSMOW and SLAP were determined by R. Hagemann, G. Nief and E. Roth. Tellus XXII, 712-715 (1970). In the case of oxygen, the absolute 1 8 0 / 1 6 0 ratio of VSMOW was determined by P. Baertschi, Earth and Planetary Science Letters 31,341-344 (1976). Table 41.1: IRMs issued by IAEA/NIST based on concencus values. a


8535 8536 8537 8538 8539 8540 8541 8542 8543 8544 8545 8546 8547 8548 8549 8550 8551 8552 8553 8554 8555 8556 8557 8558


911

battle' as it were, in response to a direct and immediate need. Nevertheless, for a material to be a reliable isotopic reference material (IRM), as for other reference materials, there are some basic requirements: 9 the material should be available in a stable form, so that the value of the property it carries will not change over time (i.e. the 6 value must remain zero!). For example, such a material should be made available in a stable chemical form, and stored and distributed in suitable, tight containers to avoid isotopic effects from evaportion, adsorption on the container wall, contamination deriving from the container material, etc. Else the value of the properties it carries (isotope ratios, elemental weight fraction) will change over time. This would imply a hidden shift of its 6 values, which therefore affects in a quantitatively unpredictable manner all measurements calibrated against it. This is graphically illustrated in Figure 41.1 for an unstable material, by the uncertainty given to the baseline. 9 the property to be measured should not vary across subsamples (homogeneity requirement). 9 the material should be available to the user community for a long period. 9 for CIRMs (Certified IRM), the certified value should have a sufficiently small combined uncertainty. Many of these requirements seem obvious, but nevertheless practice has shown that they have often not been met. Not only because the problems associated have been grossly underestimated, but also because guaranteeing these requirements can only be done at considerable expense, requiring resources often only available to specialised laboratories. There are a couple of well know examples, explained at length by Gonfiantini et al. (1993). The reference sample CDT previously used for sulfur isotopic measurements, was found to be inhomogeneous only after being in use for many years (Robinson, 1993). Similarly, SMOW originally used for oxygen and hydrogen isotopic measurements, purportedly corresponded to typical ocean water. In fact SMOW did not exist physically, but instead was only compared to values for NBS-1 a fresh water sample that was later found to be unstable over time. The reference material PDB used for carbon isotope ratios, also ran out of stock. The resulting 'instability' of the consensus reference materials is highly unsatisfactory, as traceability of such measurements is hampered. First of all, there is the considerable Babylonic confusion within the measurement community, which often perpetuates years after a new reference sample is introduced or a new zero point is assigned to the isotopic delta scale (e.g. VPDB instead of PDB, VCDT instead of CDT). Furthermore, when natural materials (especially for the light elements) are taken from the same location to approximate the original reference, it is found that the isotopic composition of the new sample is not identical to the predecessor. As the chance of finding a new reference sample with identical isotopic composition is rather dim, a 6 value relative to its predecessor is usually assigned to the replacement. For example for the sulfur delta scale, the material IAEA-S1 has been assigned a 634S value o f - 0.3 %0 on the new scale VCDT1. Note that a sample VCDT with delta value equal to zero 1. Reporting of relative sulfur isotope-ratiodata, Technical report IUPAC, 1997, Pure & Appl. Chem., 69, 293-295.

912

Chapter 41 - P.D.P. Taylor, P. De Bihvre & S. Valkiers

does not exist, i.e. the zero point of the VCDT scale only refers to a hypothetical or virtual material. This has been done to establish some comparability with previous values on the CDT scale, but it clearly demonstrates the weakness of a traceability scheme which only relies on artefacts and non-metrological values for RRef. A similar situation occurred in the case of PDB, where VPDB is a virtual reference material based on NBS-19, whose values vs. VPDB have been fixed by consensus 613C - +1.95 %o and 6180 = -2.20 %o. VPDB is used in all carbon isotopic studies, and in oxygen isotope studies of carbonates and its case was the first in which a virtual reference material has been adopted. Consequently, when scientists today state their measurements against SMOW, PDB, CDT (instead of VSMOW, VPDB, VCDT), this is probably erroneous. The letter 'V' (which stands for Vienna) refers to the measurement scale used. It implies also that the fractionation values recommended by ad-hoc IAEA working groups (to link references materials and measuring procedures), have been adopted (as for instance the fractionation between water and CO2, or between VPDB and VSMOW, etc.). Omitting the letter V may (unintentionally) imply that some other procedures have been adopted. The process of comparison of the 'old' versus the 'replacement' reference sample is usually performed by a set of selected laboratories (in a particular scientific discipline). The practicalities of this are not to be underestimated. The instruments used as comparators are nearly always all different (so what and how big are their specific biases?), the quality of the remaining predecessor sample available at some of the laboratories involved in this comparison might be questionable and the experience of the operators probably differs, as well as the measurement procedures. The establishment and maintenance of an international system to measure differences in isotope ratios is difficult, because of the difficulties to measure the uncertainty not only in the various comparison processes but also in the original zero line itself. As will be explained later on in this chapter (see e.g. Figure 41.6), this situation leads to floating delta scales, which are not anchored against the SI system. So apparently, there is some value in calibrating isotope ratios. Using a metaphor, the process of performing absolute ratio measurements could be compared to the process of measuring heights. Measuring the height of a house is much more easy to do if you are next to the house and using a ruler (with the zero reference point as the ground level). A more complicated approach would be to calculate this height by the difference of two heights relative to sea level (i.e. height of the roof minus height of the floor). Nevertheless, if you are concerned by global warming and want to know what the risk of flooding by the sea is for your house, its absolute height relative to sea level is probably important! Choosing the right reference system matters.

41.3 Comparability and traceability of the results of isotopic measurements In the past, isotopic measurements have often been confined to highly specialised research laboratories dealing with geo- or cosmochemistry. However, isotopic measurements are now even being incorporated into legislation1 and therefore much less 'non-committing'. Consequently, the comparability and traceability of the results of these measurements is even more important.


913

Traceability1 is not an aim in itself, but if a good traceability system is in place for a particular kind of measurement it will improve comparability. Although often 'willdo' solutions, 'quick fixes' work to a certain extent, at some stage if requirements for comparability increase, such traceability systems collapse. Figure 41.2 represents an 'artist impression' of what happens when isotope scientists continue to have 'artefactonly' based traceability systems. The absence of a good traceability system most often leads to hidden costs. Only in a limited number of cases does this become fully apparent. The 1999 crash of the Mars Pathfinder space craft failed as part of the mission control engineers where still calculating in Imperial Measures instead of using SI units, leading to a multi billion dollar loss.

Figure 41.2 - An 'artist impression' of what can happen if the measurement community does not use the concept of SI traceability (painting of Pieter Breughel the Elder (1520-1596; Museo Nazionale di Capodimonte, Napoli) 'The Blind', where one blind person is leading the others).

1. CEN European pre standards, ENV 12140 (13C/12C in sugars from fruit juice), ENV-xxx (in progress: comparison of 13C/12C in pulp and sugars), ENV 12141 (180/16 0 in water from fruit juice), EC / 822 / 97 (180 / 160 in water from wine). 1. Traceability is the property of the result of a measurement or the value of a standard whereby it can be related to stated references, usually national or international standards, through an unbroken chain of comparisons, all having stated uncertainties (see also Part 2, Chapter 42 on traceability).

914

Chapter 41 - P.D.P. Taylor, P. De Bi~vre & S. Valkiers

Figure 41.3 - Traceability of length measurements. In the past, a unit was realised by some physical artefact, this has been replaced by some measurement process today (e.g. in the SI, one metre is defined as the length of the path travelled by light in vacuum during a time interval of 1/299 792 458 of a second).

If measurement results are not traced back to a same common unit, problems are bound to arise sooner or later. As traceability is a property of the result of a measurement, it helps to clarify the nature of a measurement process: any measurement process consists of a series of comparisons. For a reliable measurement, the nature of these comparisons and their uncertainty should be known. The link thus established can be compared to a chain, whereby the thick lines in Figure 41.3 represent a comparison process with small uncertainty, whereas the opposite holds for the thin lines. Logically, better links (should) exist moving towards the top of the traceability chain. In the case of length measurements, the comparison process is traced back to a common unit. In the past, such units were anchored in some artefact (the King's foot). A 'novel' approach to improve comparability, was the introduction of the Convention of the Metre (1875), whereby several countries agreed to use the same artefact based unit, the realisation of which was based on a bar of Pt/Ir. This was the first truly internationally structured measurement system, which later evolved into what is now known as the International System of units. A remarkable evolution has taken place in the SI 9nearly all of the seven SI base units (except the Kg) have switched from being embedded in artefacts, to being embedded in measurement procedures (De Bi6vre & Taylor, 1997) making use of some fundamental properties of nature (e.g. one metre as a particular number of wavelengths of a 86Kr laser light source1). Thus, a solid anchor is provided, no longer subject to human arbitrariness. 1. strictly speaking, this is the old definition of the metre. In the new definition, the metre is defined as the length of the path travelled by light in vacuum during a time interval of 1/ 299 792 458 of a second.


915

Figure 41.4 - Comparability of delta measurements m a d e in different scientific disciplines suffers, as different communities use different working standards a n d / o r different measurement techniques.

So what is the situation for isotopic measurements? It can safely be stated that artefacts, such as the ones given in Table 41.1, are predominantly used to establish traceability of isotopic measurements. When values are compared over longer time periods, or when they circulate between different scientific disciplines using isotopic data, or between different countries, or have been generated using different measurement techniques (e.g. carbon isotopic measurements via infra red and mass spectrometry) the risk of incomparability increases (Figure 41.4). Traceability can only be realised if uncertainty statements are both reliable and realistic. For isotopic measurements, such statements are often unsatisfactory (or even non-existent). In many cases it is common practice just to quote reproducibility or even merely repeatability. Apart from reasons of prestige which may drive some investigators to assert exceedingly small uncertainties, the clear lack of training in matters of uncertainty also explains the current situation. The concept of evaluation of uncertainty in measurement, where error is not synonymous to uncertainty, is defined in the internationally agreed ISO document1. In this 1. ISO Guide on Uncertainty, Geneva 1993.

916


approach, a pragmatic stance is taken as for different kinds of uncertainty: they are different depending on the way they are evaluated. Uncertainty estimates resulting from repeated measurements are said to be of'type A" and all others are "type B'. Typical type B data are those derived from a certificate (e.g. of a working standard) or from a separate experiment (e.g. measuring the effect of a spectral interferent). For calculating a total uncertainty, both types are simply treated in an identical way. Why bother about uncertainty? Uncertainty is linked to quality. ISO-BIPM1 uncertainty can be seen as a way to make an objective statement about the quality of a measurement. Vague expressions such as 'inaccurate', 'fairly accurate' or 'highly accurate' can then be quantified. The different steps in the uncertainty estimation process are: 1) formulating the mathematical relationship between what is being measured (the measurand) and the experimental factors having an influence (expressed in the measurement equation), 2) assessing the magnitude of the various uncertainty contributions and 3) combining them to produce a combined uncertainty Uc either using the rules of uncertainty propagation or more simple approximations. If needed, an expanded uncertainty U is calculated, by multiplying with a coverage factor k (U = k.uc) depending on the level of confidence (k = 1, 2 ...). Specifically for uncertainties related to certified reference materials, readers are referred to a detailed description by Pauwels et al. (1999) explaining how to evaluate uncertainty contributions from inhomogeneity and instability. Finally, in isotopic measurements there is a strong tradition and tendency to use large numbers of replicate experiments (e.g. 100 replicate measurements on a sample aliquot in a 10 minute period) to reduce the standard uncertainty of the mean (= s/V'n) when carrying out type A uncertainty evaluations. Although it obviously leads to impressively small numerical values, this can be quite misleading. One must be aware of the other (often larger) sources of uncertainty when calculating combined uncertainties. The likelihood of obtaining comparable results is increased if such uncertainties are used. 41.4 Primary IRMs: CIRMs carrying SI traceable isotope ratios with the smallest achievable combined uncertainties The above explains why some laboratories have specialised in producing primary CIRMs. Such materials are also artefacts, but these artefacts only serve to carry the value. These values are far more important, as they are anchored in the SI system. Such Primary CIRMs can play a crucial role in the calibration process for isotopic measurements. In such a process, the output quantity (what one tries to measure, i.e. (6IRM)x, the delta value for the sample X in the IRM-scale) is linked to input quantities (what is actually measured, i.e. (6obs)x) for a particular system (Figure 41.5). By means of working standards and reference samples, the end user establishes the relationship between the measured value 6obs and the value carried by the reference samples with particular ~)IRMvalues. This operation is carried out at various intervals during a measurement session, depending on the required quality (in the example, a single calibra1. BIPM:International Bureau of Weights and Measures.


917

Figure 41.5 - Calibration (by the end user) of a differential isotope measurement using IRMs with certified 'delta' values (top). The calibration is shown for a single point. The bottom of this graph shows how the abscissa can be anchored to SI traceable values by the IRM supplier.. tion point is shown). The uncertainty on the values for 6IRM for the different calibration materials used, should be taken into account when the final value on the sample is calculated (see also Figure 41.1). If the values 6IRM of the used reference sample (or their uncertainties) are questionable or lacking in quality, this implies 'floating' points on the abscissa. The bottom part of Figure 41.5 then clarifies the role of Primary CIRMs when measuring differences in isotope ratios: their values (ratios of amounts of isotope) are 'locked' on the scale. This activity is the responsibility of the IRM supplier and means that the IRM isotope ratio(s) are not based on assumptions (e.g. mass dependent fractionation) or some arbitrary assigned value. Instead, it is possible to reproduce the value by a completely independent experiment for such a material thus reducing the problems arising when the IRMs run out of stock. Table 41.2 is an overview of presently existing CIRMs of this type. Using a real case, Figure 41.6 illustrates the value of 'anchoring' delta scales. After the many problems with the original CDT reference sample (see introduction), the alternative reference samples (e.g. IAEA-S1) carried an assigned value which was now (Ding et al., 1998) proven to be different to the SI-traceable value by about 20%o! So how are the 'absolute' ratios of amounts of isotopes, e.g. n(34S)/n(32S), obtained? A completely independent and reliable method is used to 'synthesise' these

918

Chapter 41 - P.D.P. Taylor, P. De Bi~vre & S. Valkiers

Table 41.2 - CIRMs issued by NIST (previously NBS) and IRMM (previously CBNM) based on calibrated isotope ratio mass spectrometry. The * indicates the sample is a 'spike' (i.e. can be used for isotope dilution). Also consult www.irmm.jrc.be and www.nist.gov for further information Element

Identification

Boron

IRMM 011 NIST SRM 951 NIST SRM 952 NIST SRM 977 IRMM 621-623 NIST SRM 975 IRMM 641 IRMM 642 NIST SRM 979 IRMM 012 IRMM 624 IRMM 625 NIST SRM 994 IRMM 632 IRMM 633 NIST SRM 994 IRMM 014 IRMM 620 IRMM 634 NIST SRM 981 NIST SRM 983 NIST SRM 991 NIST SRM 982 IRMM 016 IRMM 015 NIST SRM 980 IRMM 009 IRMM 637 IRMM 638 IRMM 639 IRMM 640 NIST SRM 986 IRMM 627 IRMM 628 IRMM 010 IRMM 630a IRMM 631 IRMM 290 A-G NIST SRM 985 NIST SRM 989 NIST SRM 984 NIST SRM 990 IRMM 017 IRMM 018 NIST SRM 978a

Boron-10 Bromine *Cadmium Chlorine *Chlorine *Chlorine Chromium Chromium *Chromium-50 *Chromium Copper *Copper-65 *Copper Gallium Iron *Iron-57 *Iron Lead normal Lead radiogen *Lead-206 Lead-206-238* Lithium Lithium-6 Mgnesium *Magnesium *Magnesium *Magnesium *Mercury *Mercury Nickel *Nitrogen-15 (nitrate) *Nitrogen-15 (nitrate) Platinum *Platinum-194 *Platinum Plutonium Potassium Rhenium Rubidium Silicon Silicon Silicon Silver

Year of Issue

Chemical Form

1969 1970 1969 1964 1997 1962 2000 2000 1966 1999 1999 1999 1964 1999 1999 1986 1992 1996 1999 1968 1968 1968 1968 1984 1986 1966 1999 1999 1999 1999 2000 1989 1998 1998 1999 1998 1998 1993 1975 1973 1969 1975 1989 1989 1962

H3BO3 H3BO3 H3BO3 NaBr sol in HNO3 NaC1 Aq. sol Aq. sol Cr(NO3)3 sol in HC1 sol in HC1 sol in HC1 Cu sol in HNO3 sol in HNO3 Ga Fe sol in HC1 sol in HC1 Pb Pb Pb Pb Li2CO3 Li2CO3 Mg sol in HNO3 sol in HNO3 sol in HNO3 sol in HC1 sol in HC1 Ni Aq. sol Aq. sol Metal sol in HC1 sol in HC1 Pu(NO3)4 KC1 Re RbC1 Si Si SiO2 AgNO3 (Table 41.2 continued >)

919

The Nature and Role of Primary Certified Isotopic Reference Materials ... (> Table 41.2 continued) Element

Identification

Strontium *Strontium Thallium Thorium Thorium Uranium Uranium Uranium

NIST SRM 987 IRMM 635 NIST SRM 997 IRMM 035 IRMM 036 NIST SRM U002 NIST SRM U970 IRMM 183-187

Year of Issue

Chemical Form

1982 1997 1980

SrCO3 sol in HC1 Th sol in HC1 sol in HC1 U308 sol in HNO3 UF

1966 1987 1988 / 1993

ratios, leading to a n Rtrue which can then be used to obtain a 'calibration' factor K for a particular ratio: K-

[41.2]

Rtrue

Robs It was A. Nier (Nier, 1950) w h o first reported on the use of such mixtures (for argon) to calibrate isotopic measurements. Later on, the use of such mixtures was further developed, especially in cases where traceability was very important, e.g. for w o r k on fundamental constants as well as for nuclear safeguards (for a review, consult De Bi6vre et al., 1993). For the latter type of measurements, research but especially field laboratories u n d e r s t o o d at a very early stage that the use of a relative measurement scale is not appropriate. As these scales are always based on assumptions or on some agreement, and therefore on m u t u a l trust between all parties, they were not viable in the context of nuclear/fissile material where such measurements are used to IAEA-S-3

IAEA-S-1

i I

.................

! I

23.394 -30 .... t ................

I

IAEA-S-2

...............................

22.644 -20 !

1

-10 . . . . . . . . .

0

I ...............

I

I

I

-30

-20

.............

,~ [n(32S) / n(34S)]

I ...... I

SI-traceable

22.143 +10

+20

+30

1

I

1

I

-10

....

t

0

........

6(34S)v-CDT

I

+10

SI-traceable +20

+30 &(34S)v-CDT Thode-traceable

Figure 41.6 - In the absence of SI-traceable values for the sulfur isotope ratios, the delta scale 'floats' relative to the SI traceable ratio scale.

920


keep track of the inventory of fissile material on a global scale. The smaller the uncertainty on such measurements, the smaller the uncertainty on the inventory, which is linked to the critical amount of U or Pu needed to make nuclear weapons. Questioning the appropriateness of differential measurement scales has just started in the area of chemical amount content measurements and isotope ratio measurements, for similar reasons: global trade, environmental policy issues and the growing interdisciplinary use of data (De Bi~vre, 1993; De Bi6vre et al., 1996). The value assignment process for CIRMs is based on two parts: they are underpinned by reliable and transparent measurement procedures to measure the isotope amount ratio (and not a difference in ratios) and a completely independent method (also called a primary method) is used to obtain the calibration K-factor. Both of these parts will be explained in detail.

41.5 Mixtures with gravimetrically prepared isotope ratios When a high purity substance of known chemical composition (e.g. Fe203 with molar mass M) is weighed on a balance (mass m), the chemical amount n(E) (e.g. n(Fe) is given by: n(E)

-

m(1-~im p )(1-r M

stoi

)

[41.3]

where ~imp and ~stoi are small correction terms for chemical impurity and stoichiometry (e.g. when using Fe203, is the amount ratio iron/oxygen really 2/3, or 2.001/3 ...). As the amount n(E) is not a function of other amounts (but function of molar mass and mass), such a method is called a primary method of measurement. These methods are perhaps not the fastest, most convenient or versatile, cheapest or most user friendly, but they have two important features which make them metrologically superior: the traceability of the results produced with such methods is easily understood and they yield small combined uncertainties. Gravimetry has been used in many cases to produce mixtures of isotopes with isotope ratios having small uncertainties. For this purpose, materials of high isotopic enrichment are used (e.g. typically 99% or more, see Figure 41.7) and the isotope ratio for a mixture X of such materials (A and B) is given by:

n (iE ,X)

n (iE,A) + n (iE,B)

n(JE,x)

n(JE,A) + n(JE,B)

f, A n(E,A) + f i, Bn(E, B) f j.An(E,A) + f j, Bn(E, B)

[41.4]

Where f denotes the isotope abundance (for isotope iE, jE), e.g. fi, A (= Ri,A/~Ri, A) the abundance of isotope i in sample A. The amounts n can be expressed using equation [41.3]. The starting materials A and B must be chemically purified and then transformed into a compound of well known stoichiometry. Their imperfection prior to mixing is then carefully assessed" purity, stoichiometry, degree of isotopic enrichment. The uncer-

ENRICHED Si ISOTOPE MATERIALS (RAW BASE MATERIAL)

PURE TEST MATERIAL (IRMM-018)

GRAVIMETRICALLYBLENDED si CRYSTAL DETERMINATION OF ISOTOPE MIXTURES AVOGADRO CONSTANT GIVEN MATERIALS

dissolve-in NaOH

INTERMEDIATE PRODUCTS

1

with H F and ion exchange convert to

1

con"ert to BaSiF6 then to SiF4 H2SiF6 further purify further Purify in gas phase in gas phase

+

+ convert to BaSiF6, then to SiF4, further purify in gas phase

t

convert to BaSiO6, then to SiF4, further purify in gas phase

with HF and ion exchange convert to H2SiFh convert to BaSiF6, then to SiF further pur&y in gas phase


GENERAL LAY-OUT OF PREPARATION AND MEASUREMENTS OF Si ISOTOPE MIXTURES

TARGET IONS FOR MS MEASUREMENT measure abundances

4

4

K 'higher' 3 K 'lower' 921

Figure 41.7 - Preparation and measurement scheme for gravimetric isotope mixtures of silicon, used in the context of the Avogadro project.

922

Chapter 41 - P.D.P. Taylor,P. De Bihvre & S. Valkiers

tainty of these three factors is measured and then combined with the uncertainty of the mass measurements. Figure 41.8 gives an example of the contribution of these different sources of uncertainty to the total combined uncertainty. Depending on the uncertainty required and on the amount of material available, solutions can either be prepared by starting from the enriched compounds and then mixed, or the solid compounds can be weighed, mixed and then brought into solution. The first approach reduces the amount of material required, but also often limits the combined uncertainty that can be reached. At IRMM, gaseous synthetic isotope mixtures have also been prepared (e.g. for Xe and Kr). Synthetic isotope mixtures have been prepared in the past by researchers at NBS (now NIST) and CBNM (now called IRMM) and also at CEA. Table 41.2 gives an over-

IRMM

Ins~tute for Reference M ~ e d a i s and Measummen~

CERTIFICA~ iliSOTO~C REFERE~E MATERii~ iRMM~72

~ e I ~ ~ i i c Refe~,~n~ M~dia.i ilRMM~72 ~s s u p p i ~ as a set w ~ c e M ~ ,as follows ....

~~r

ratios

Moi~ i ~ t o , ~ A i b : ~ , c , e Ra~o Code Numb~

,n1-3 permil due to uncontrolled variability in the oxygen isotopic composition of the SO2 arising from exchangeable oxygen (e.g., organic oxygen and water) in the sample matrix. Such effects can be overcome through empirical correction factors or by isotopic equilibration of SO2 produced by the elemental analyzer in a 890~ SiO2 buffering furnace (Fry et al., 2002). While the effects of oxy-

974

Chapter 45 - K. Leckrone & M. Ricci

gen isotopic variability in continuous-flow sulfur isotope measurements can be overcome with sufficient experimental controls, an improved isotope correction algorithm capable of correcting for oxygen isotope variability is nevertheless desirable to reduce these experimental constraints and improve the robustness of continuous-flow sulfur isotope methods. 45.1.3 Objectives o f study Given the inherent limitations of the current isotope correction equation, the goal of this project has been to mathematically and experimentally evaluate two alternative methods for oxygen isotope correction which have the potential to correct for sample-to-sample variability in the oxygen isotopic composition of SO2. The first relies on simultaneous measurement of three adjacent ion currents (m/z 64, 65 and 66). This approach is similar to one suggested by Coleman (Coleman, 1980). Here, however, we use numerical techniques to solve exact expressions for S and O ratios assuming a triple collector instrument capable of simultaneous measurement of three distinct ion beams. This approach also assumes mass-dependent fractionations for both S and O isotopes. Accordingly, we first present a review of the sulfur and oxygen isotope literature to determine the most accurate values for these relationships. We then derive an exact expression for S and O isotope ratios in terms of measured mass spectrometric ion current ratios 66/64 and 65/64. Finally, we test the algorithm using a combination of computer-generated datasets and laboratory analyses of natural sulfide mineral samples and reference materials. Additionally, we have evaluated a previously proposed method for calculating 834S based on simultaneous collection of ion currents at m/z 48, 50, 64 and 66 (Holt & Engelkemeier, 1970; Coleman 1980). This approach does not assume mass-dependent fractionation of S or O isotopes. We derive an expression for the sulfur isotope ratio in terms of the measured mass spectrometric ion current ratios 66/64 and 50/48, and test the procedure using a variety of sulfide minerals and reference materials. 45.2 Corrections based on s i m u l t a n e o u s m e a s u r e m e n t of three adjacent ion currents 45.2.1 N o t a t i o n

Throughout this paper, we will use the letter R to indicate the ratio of a particular atomic or molecular isotopic species to the related isotopic species of lowest mass nR _ concentration of isotopic species of mass n concentration of base species related to n

[45.3]

For example, 34R refers to the isotopic abundance ratio 34S/32S of sulfur, while 66R refers to the measured ion current ratio m/z 66/m/z 64 of the sulfur dioxide molecule. The letter F will be used in a similar manner to represent fractional abundances nF = concentration of isotopic species of mass n concentration of all species related to n

[45.4]

For example, 13F refers to 13C/(12C + 13C). For molecular species, the denominator

Oxygen Isotope Corrections in Continuous-Flow Measurements of SO 2

975

refers to the sum of all species of SO2. The symbol cz refers to the fractionation factor, that is, the ratio of isotope ratios produced by a chemical or physical process

n

nR c~ -

products

[45.5]

nRreactants Classical calculations of 6 values for stable isotopes have often proceeded from the position that the mass spectrometer is measuring 6 values. On one level this treatment is correct, but on a deeper level, the mass spectrometer is actually measuring ion currents, which are then turned into ion current ratios, from which delta values are derived. Thus, in our calculations, ion current ratios are related directly to isotopic or molecular abundance ratios, and calculation of delta values is the last step rather than the first.

45.2.2 Mass spectrum of sulfur dioxide For SO2 samples with isotopic abundances near natural levels, ion currents large enough to allow rapid and precise measurements occur only at masses 64, 65 and 66. The fractional abundance of these SO2 molecular species are related to fractional abundance of individual isotopes according to simple probability functions 64F = 32F.16F.16F 65F = 33F.16F.16F + 2.32F.17F.16F 66F = 34F.16F.16 F + 2.32F.18F.16F + 2.33F.17F.16F + 32F.17F.17F

[45.6] [45.7] [45.8]

where factors of two accounts for the fact that there are two sites at which any given oxygen isotopic substitution can occur, hence the probability of oxygen isotopic substitution is twice its atomic abundance. It is convenient to divide equations [45.7] and [45.8] by equation [45.6] in order to express molecular abundance ratios (ion current ratios) in terms of atomic isotope ratios" 65R = 33R + 2.17R 66R = 34R + 2.18R + 2.33R17R + 17R2

[45.9] [45.10]

The molecular abundance ratios 65R and 66R are obtained from a triple collector instrument capable of simultaneous measurement of three distinct ion beams.

45.2.3 Mass-dependent relationships of oxygen and sulfur The fundamental difficulty underlying most methods for oxygen isotope correction of SO2 is that measurement of two (65R and 66R) ion-current ratios does not provide sufficient information to uniquely solve for the four contributing elemental ratios (33R, 34R, 17R and 18R). Functional relationships between 17R and 18R, and between 33R and 34R, are therefore required to provide the necessary additional equations. These relationships are based on the fact that nearly all physical and chemical processes fractionate multiple isotopes of the same element proportionally to differences in mass (Bigeleisen & Mayer, 1958). Consequently, the sulfur or oxygen isotope ratios

976


of two pools (denoted by subscripts I and 2) linked by a single mass-dependent process are related quantitatively by 17R 1 / 17R 2 = (18R1/18R2)a

[45.11]

or 33R1 / 33R2 = (34R 1 / 34R2)a'

[45.12]

where a = In (17(I)/ln(18~) and a' = ln(33~)/ln(34~). However, fractionation factors are not fixed so that a or a' remain constant for all processes. Instead, values of a and a' depend on such variables as the molecular masses and reaction mechanisms of the species involved. In general, for both S and O, a and a' for specific processes take on values in the range of 0.50 to 0.53 (Hulston & Thode, 1965b; Matsuhisa et al., 1978). Furthurmore, equation [45.12] describes only the immediate isotopic relationship between reactant and product involved in a single fractionating process, rather than the overall composition of mixed isotopic pools. Thus, the relationships underlying the standard SO2 isotope correction equation

(18Rsal0"50 17Rst 18Rst)

17Rsa _

[45.13]

and

33asa _ (34Rsal0.50 33Rst 34Rst)

[45.14]

only approximately describe the isotopic relationships produced by the mixing of sulfur or oxygen pools of similar isotopic composition. The choices of a - a' - 0.5 represent minimal values of a and a' rather than the best possible approximations of the quantitative relationship between related isotope ratios in the terrestrial sulfur and oxygen pools. An optimal value of a can be obtained from paired measurements of ~170 and 6180 for a variety of terrestrial materials. The slope of the resulting 6170 vs. ~1180 crossplot is equivalent to a because the 6 values in such plots are identical to the first term in a series expansion of the log terms in

18Rsa In 17 "- a In 1 ~ Rst Rst 17R

sa

[45 15]

which is in turn derived from equation [45.11] (Santrock et al., 1985). An optimum

977


value for a' can be obtained from paired measurements of ~33S and ~34S by analogous logic. A more recent re-assessment (Robert et al., 1992 and references therein) of oxygen isotopes for a wide variety of terrestrial and lunar materials indicates ~170

-

(0.521 + 0.001) 915180

( n - 102)

[45.16]

where the stated uncertaintly is the l o standard deviation of the slope. The resulting value of a is intermediate between the recently reported value of 0.528 (Meijer & Lee, 1998) and the value of 0.516 suggested by Santrock et al. (1985). Paired measurements of ~133S a n d ~34S have been reported for a wide variety of sulfur-bearing terrestrial, lunar and meteorite samples prepared as SF6 (Beaudoin et al., 1994; Gao & Thiemens, 1991, 1993a, b; Heymann et al., 1998; Hulston & Thode, 1965a; Rees & Thode, 1977; Rumble et al., 1993; Thode & Rees, 1971). Our review of the SF6 literature indicates a relationship of 633S -

(0.515 + 0.001)-

634S

( n - 127)

[45.17]

This analysis excludes a small subset of iron meteorites which are known to have been selectively enriched in 33S by nucleogenic processes. In general it is important to note that the above relationships do not apply to materials affected by nucleosynthetic processes, admixture of separated isotopes, or mass-independent isotope effects. The isotope relationships expressed in equations [45.16] and [45.17] convey only the relative differences between the isotopic composition of unknown samples and an isotopic standard. In order to ascertain the absolute values of isotope ratios for use in equations [45.9] and [45.10], relative isotopic relationships need to be recast in the form (Santrock et al., 1985) 17R -

18Ra" K

[45.18]

34Ra'" K'

[45.19]

and 33R -

where K and K' are con- Table 45.1 - Absolute stants characteristic of dards the solar system's oxy- standard 17R gen and sulfur pools. Evaluation of K requires PDB 0.0003790 an absolute isotopic SMOW 0.0003730 abundance measure- Tank 02 0.0003775 ment of both 17R and 18R for at least one isoto- average St. dev. pic standard. As shown rsd in Table 45.1, there have

abundance measurements of oxygen isotope stan-

18R

reference

K

0.0020671 0.0020052 0.0020514

Mook & Grootes (1973) 0.0094915 Hoers (1997) 0.0094904 Santrock et al (1985) 0.0094916 0.0094912 0.0000007 0.007%

978


been at least three independent analyses of the absolute 17R and 18R values in reference materials (Hoefs, 1980; Mook & Grootes, 1973; Santrock et al., 1985 and references therein), which we have used together with the a value of 0.521 to determine an average value of K - 0.0094912. Agreement among the three independent measurements is sufficient to calculate K with a relative standard deviation of 0.007%. Similarly, evaluation of K' requires an absolute isotopic abundance measurement of both 33R and 34R values for a single sulfur isotope standard. To date there have been absolute isotope abundance measurements of five separate sulfur standards (Ding & Zhang, 1989), resulting in an average value of K' - 0.03924 (0.3% relative standard deviation).

45.2.4 Calculation of sulfur and oxygen isotope ratios The exact expressions relating measured 65R and 66R ion current ratios to the isotopic abundances of their constituent elements (equations [45.9] and [45.10]) and expressions for mass-dependent relationships in the terrestrial oxygen and sulfur pools (equations [45.18] and [45.19]) can now be combined to yield a single master equation for 34R in terms of only measured quantities and the constants a, K, a' and K'. This method is analagous to the successful approach of Santrock et al. (1985) for carbon isotopic analysis of CO2, and apart from uncertainties in the values of the constants, the equation is exact. 66R - 3 4 R - 2( K'

65R/~

9

"

34Ra')1/a - K 965R- 34Ra 65R2 34R2a " K'2 ' '+ = 0 4

[45.20]

It is difficult to rearrange equation [45.20] to yield an explicit solution for 34R. Practically, the equation is most easily solved using numerical techniques. The overall procedure for SO2 isotope correction using this approach is as follows. Step 1" Raw ion current ratios are corrected for background currents, amplifier offsets,

mass discrimination effects, etc. by the usual method of differential measurement, in which m65.-.

65Rs a =

lKsa . t65Rs t _ 65 Q. t65 Rst, . Rst

m65

66

Rsa -

66

Q .

t66.-,

lKst

[45.21]

where m indicates a measured ion current ratio and t indicates the true isotope ratio of the standard gas as prepared. Working standards are determined by extension of this process to a primary standard. Step 2" The corrected ion current ratios are inserted into equation [54.20], which is

solved iteratively for 34R using Newton's method. The initial or "seed" value of 34R may be fixed (for example, set to equal 34Rst) or floating (for example, estimated as 34Rseed = 34Rst.66Rsa/66Rst). The former approach is adequate provided sample and standard differ by less than around 25%o, but the latter

Oxygen Isotope Correctionsin Continuous-FlowMeasurementsof SO 2

979

approach leads to more rapid convergence of equation [45.20] and avoids errors in cases of more widely differing samples and standards.

Step 3" Once 34R of the sample is known, 33R, 18R and 17R are calculated from equations [45.19], [45.9] and [45.18], respectively. Step 4" In the last step, 6 values are calculated from atomic isotope ratios according to ~)34S = 1000(34Rsa/34Rst - 1) and

6180 = 1000(18Rsa/18Rst-1)

[45.22]

As an alternative procedure, the four contributing equations can be combined to yield an exact expression for 18R in terms of 66R and 65R 66R 18R_ 1(65R-2KK'18Ra)1/a' 18Ra 3. K2 18R2a 2 2 - K-65R. + ~_ 9 - 0

[45.23]

in which case 17R, 34R and 33R are subsequently calculated by re-insertion in the primary contributing equations. However, we have not found any significant difference in the final isotopic values if 18R rather than 34S is calculted in Step 2. As with the standard SO2 isotope correction (equation [45.1]), corrections based on equation [45.20] require both ~)180 and ~)34Sof the reference gas to be known relative to primary isotopic standards. Unlike the standard correction, this approach does not require ~180 of the sample to be known or matched relative to the standard, and it accounts for all higher-order terms.

45.2.5 Results for model datasets The algorithm for oxygen and sulfur isotope calculations using the above four-step process was initially debugged and tested using two model datasets. In the first dataset, a series of 65R and 66R values were generated for hypothetical samples having 634SCDT ranging from -100 to 100 %o and ~)18OvsMow of 0 %o. In the second dataset, 65R and 66R values were generated for samples having input values of 6180 from-100 to 100%o and ~)34S of 0 %o. In each case, standards were assigned 65R and 66R equivalent to ~)34SCDT - 0%o and ~)18OvsMow of 0 %o. Thus, in the first dataset, samples varied relative to standards in ~)34S but not in 6180, and in the second set, samples and standards shared the same 634S differed in ~180 values. Sulfur isotopic ratios as 634S were calculated based on these spreadsheet-generated 65R and 66R values using three alternative isotope correction procedures. First, (~34S was calculated using the standard SO2 isotope correction equation (equation [45.1]). Second, ~)34S was calculated using an empirical two-point isotope correction of the form ~)34Ssa = m. ~)66Rsa+ b

[45.24]

in which m and b are determined by regression of measured ~)66R against known 634S for a pair of isotopic standards chosen to bracket the expected 634S of the sample (Crowe & Vaughan, 1996). This correction is particularly useful in laser-microprobe

980


Figure 45.1 - Comparison of isotope corrections when oxygen isotopes of sample and standard are matched. IRMS because, although it assumes the oxygen isotopic composition of samples and standards are equivalent, knowledge of the actual value of 6180 is not required. Finally, both 634S and 6180 were calculated using the SO2cal algorithm based on equation [45.20]. Results are expressed as A values, where A34S is the error in the ~34S value calculated by the standard equation (equation [45.1]), empirical correction (equation [45.24]), or SO2cal algorithm, and A180 is the error in 6180 calculated with the SO2cal algorithm. Results for the dataset with constant 6180 but variable ~34S are summarized in Figure 45.1, and results for the dataset with constant 634S but variable 6180 are summarized in Figure 45.2. These modeling results illustrate several points. First, when the oxygen isotopic composition of sample and standard are matched, the standard correction displayed a slight systematic error of 0.0087%0 in the calculated 634S for every 1%o difference between the 634S of the sample and that of an arbitrarily chosen standard. These errors presumably arise due to the exclusion of higher-order

Oxygen Isotope Corrections in Continuous-FlowMeasurements of SO 2

981

Figure 45.2 - Comparison of isotope corrections when oxygen isotopes of sample and standard are mismatched. terms in the derivation of equation [45.1], and are negligible (99% 12C180160). It therefore stands to reason that resolving the underlying sulfur and oxygen contributions to the SO2 mass spectrum based on adjacent ion currents is far more difficult than discerning the underlying carbon and oxygen contributions to the CO2 mass spectrum by a similar treatment. Another correction related to the one explored here could be made from simultaneous measurement of the three adjacent ion currents m/z 48, 49 and 50 of the SO + fragment ion. However, this approach seems to offer no real advantage, since it would not address


987

the fundamental problem related to parallel isotope distribution patterns of S and O, and would involve lower ion currents. If the isotope correction for S02 can in fact be improved, it seems best to turn to an altogether different approach. 45.3 Corrections based on simultaneous collection of parent-daughter fragments 45.3.1 S O + m a s s s p e c t r u m

As an alternative to the SO2cal algorithm, we have made a preliminary evaluation of the method first proposed by Holt & Engelkemier (Holt & Engelkemier, 1970) for isotope corrections based on the simultaneous collection of ion currents from the m/z 64 and 66 molecular ion as well as m/z 48 and 50 ion currents from the SO + daughter ion, which is produced at approximately 70% intensity relative to the parent ion by electron ionization at 70 eV. The parent-daughter correction method is particularly attractive because as outlined below, it does not employ functional relationships b e t w e e n 33R and 34R, o r between 170 and 180, and it therefore makes no assumption of mass-dependent fractionation. However, it implicitly assumes that any isotopic fractionation which may occur in the mass spectrometric measurement during fragmentation of the SO2 parent ion is adequately addressed through the standard isotopic practice of differential measurement. As with SO2, the fractional abundance of the relevant SO + fragment ions should be related to the fractional abundances of individual isotopes according to 48F = 32F16F

[45.25]

49F = 33F16F + 32F17F [45.26] 50F = 34F16F + 32F18F + 33F17F [45.27] In the case of the SO + fragment ion, only one possible site for oxygen isotopic substitution is present, so there are no multiplicative factors in the fractional abundance expressions. Dividing equations [45.26] and [45.27] by equation [45.25] yields expressions for ion current ratios in terms of elemental isotope ratios relationships in terms of ratios: 49R = 33R + 17R

Figure 45.8 - Effect of noise added in the 66/64 ratio on the [45.28] oxygen isotope calculation.

988


50R = 34R + 18R + 33R17R

Table 45.3 - P r o p a g a t i o n of errors in 65R a n d 66R t h r o u g h the

[45.29] SORcala l g o r i t h m Inspection of equations [45.9], RMS noise a d d e d RMS error in RMS error in [45.10], [45.28] and [45.29] to ion c u r r e n t ratio calculated 634S in calculated 6180 show that they can be comA65R A34S A180 bined in various ways to yield ........................................................................................................................................................................................................................... expressions for elemental iso5.7 x 10-6 0.10 1.1 tope ratios in terms of mea5.6 x 10-5 0.95 11. 5.1 x 10 -4 10.0 113. sured ion current ratios. For example, equations [45.28] and A66R A34S A180 [45.9] can be combined to solve f o r 33R

[45.30]

2 . 4 9 R - 65R - 33R

5.3 x 10 -6 5.8 x 10-5 6.1 x 10 -4

0.03 0.31 3.3

0.3 3.1 33.

This approach is attractive in that it solves exactly for 33R, from which 34R can be obtained by equation [45.19]. However, it is based on collection of the least abundant SO2+ molecular ion and the least abundant SO + daughter ion, and so is experimentally undesirable due to its lower signal-to-noise ratio and potential sensitivity to contamination. Focusing on the more abundant ion, combination of equations [45.10] and [45.29] yields an exact expression for 34R in terms of measured ion currents and 17R 2.50R -

66R = 34R - 17R2

[45.31]

If the 17R2 term is ignored because of its low fractional abundance (0.007% of the m/z 50 ion current, compared to 95.7% 34R and 4.3% for 18R), (~34S c a n be determined directly from 50

66

6348 _ (2" ~Rsa - 66Rsa - 1 ) " 1 0 0 0 2 5~ Rst

[45.32]

for any S O 2 sample, regardless of whether or not its oxygen isotope composition matches that of the standard. Combination of equations [45.29] and [45.9], or of [45.28] and [45.10], are also possible, but are less desirable due to their dependence on at least one low intensity signal and the need to eliminate at least one higher order term.

45.3.2 Experimental evaluation of parent-daughter method The parent-daughter method using equation [45.32] was first tested under typical conditions using a set of sulfide and sulfate mineral samples which were converted to SO2 by vacuum-line combustion with V205 combustion. Samples ranged in ~)34S from 21.0%o to -32%o, and were assumed to have constant 6180. Due to limitations of the cup configuration on the mass spectrometer (a Finnigan MAT 252 with 7-cup


989

MEMCO configured for CO2, N2 and 0 2 ) , m/z 48 and 50 ion currents were sequentially measured using a peak-hopping procedure after measurement of the 66R w a s complete. As shown in Figure 45.9, values of ~34S calculated with the equation [45.32] covaried against those calculated with equation [45.1] with a slope of 1.01 + 0.01 and an intercept of 0.02 + 0.1 (R2 = 0.999; ls standard deviations). The external precision of replicate analyses was slightly poorer for the 4mass correction (+ 0.6%0) compared to when the standard correction (+ 0.4%o). However, precision of the parent-daughter method Figure 45.9 - Comparison of parent-daughter and standard may have been limited by the corrections when oxygen isotopes of samples and standards precision of the 50R measure- are matched. ment obtained the use of the peak-hopping procedure rather measurement of a true, simultaneous ion-current ratio. To furthur test equation [45.32], we analyzed a set of SO2 samples in which ~34S the sulfur isotopic composition was constant but the oxygen isotopic composition was varied by equilibrating half of the samples with Table 45.4 - Comparison of isotope corrections when the oxygen isotopes of sample and standard are mismatched water enriched in ~180 by a large evaporative reduc634S calculated by 634S calculated by SO2 sample tion. To promote recovery equation [45.1] equation [45.1] of the SO2, the water was acidified to pH 1 prior to unequilibrated 9.32 9.44 the experiment, and (original 6180) 9.66 9.54 equilibrated SO2 was 9.48 9.23 9.32 9.89 extracted and purified by average + sd (lo) 9.60 + 0.20 9.53 + 0.27 cryogenic vacuum distillation. As in the prior equilibrated 11.83 8.84 experiment, m/z 48 and 50 (altered 6180) 13.13 9.27 ion currents were sequen12.95 8.71 tially measured using a average + sd (lo) 12.64 + 0.70 8.94 + 0.30 peak-hopping procedure

990


after measurement of the 66R for each sample was complete. The sulfur isotopic composition for each sample was then calculated using equation [45.1] or equation [45.32]. Initial results are shown in Table 45.4. For unequilibrated SO2 samples, ~34S calculated using equation [45.32] matched those calculated using equation [45.1], and both methods resulted in similar precision for replicate analyses. For SO2 samples equilibrated with the isotopically enriched water, the 634S values calculated with equation [45.1] were clearly in error, and appeared enriched relative to the unequilibrated control samples by +3.0%o. The 634S calculated using equation [4:5.32] were depleted b y 0.5%0. This difference in ~34S calculated for equilibrated and control samples is significant, but is not necessarily due to failure of equation [4:5.32]. Controlling the oxygen isotopic composition of SO2 by equilibration with water is experimentally difficult. Yields of SO2 for equilibrated samples were relatively low (ca 70%), presumably due the very large Henry's law constant for solubility of SO2 (aq) in water. An unknown amount of fractionation between the extracted SO2 and the residual aqueous SO2 may have occurred. In spite of this ambiguity, the results of this and the previous experiment suggest that equation [4:5.32] yields results which are comparable to standard calculations when oxygen isotopic composition of the samples and standards is matched, and result in more accurate results when the oxygen isotopes of samples vary relative to standards. 45.4 Conclusions and recommendations for future work

While the SO2cal algorithm developed here allows for independent calculation of 634S and 6180 from 65/64 and 66/64 ion current ratios which are routinely collected for SO2, the precision of calculated (~34S is currently limited to ca + 1%o, and the precision of calculated 6180 to over + 10%o. Thus, this approach is not recommended except in cases where large oxygen isotope variations between samples and standards limit the accuracy of other oxygen isotope correction procedures to > 1%o. The precision of ~)34Scalculated by the method described here appears to be limited primarily by the precision with which the 65R ion current ratio is measured. Readily made modifications in mass spectrometric conditions, such as increasing the measurement and idle times and the number of measurement cycles, did not significantly improve the precision of measured 65R and 66R ion current ratios, nor did they improve results from the SO2cal procedure. It seems likely that isotope corrections for SO2 using this approach may be fundamentally limited by the relatively minor contribution of the oxygen isotope ratio terms 17R and 18R to the m/z 65 and 66 ion current ratios respectively. An alternative oxygen isotope correction based on collection of parent ions at m/z 64 and 66, and their daughter ions at m/z 48 and 50, resulted in accurate and precise calculation of ~)34Sfor samples with constant oxygen isotopic compositions, and provided better accuracy and precision than the standard correction when oxygen isotopes of samples were variable. These preliminary results suggest that potential fractionations associated with fragmentation of the parent ion in the ion source of the mass spectrometer can be adequately addressed through the standard isotopic practise of relative measurements. Routine use of the parent-daughter correction in continuous-flow applications would require a mass spectrometer with a collector system capable of true simultaneous monitoring of m/z 48, 50, 64 and 66 ion beams. Although


991

simultaneous measurements were not possible on the Finnigan 252 used for this work, some commerically available instruments are already capable of the measurement.

Acknowledgements This work was supported by NSF-EAR-9811170 to D.C. and K.L., and by an NSF earth sciences postdoctoral fellowship to K.L.


CHAPTER 46 Experimental Measurement of Isotopic Fractionation Factors and Rates and Mechanisms of Reaction Simon M. F. Sheppard Laboratoire de Sciences de la Terre, CNRS UMR5570, Ecole Normale Sup6rieure de Lyon, 46 All6e d'Italie, 69364 Lyon, Cedex 07, France e-mail: [email protected]

46.1 Introduction Laboratory and field experimental studies have played a fundamenal role in the development of both isotope chemistry, since its very beginnings in the 1930's (e.g., Weber et al., 1935) and, subsequently, in isotope geochemistry (e.g., Thode et al., 1949; Epstein et al., 1951). The stable isotope ratios of many light (H, Li, B, C, N, O, Mg, Si, S, C1,...) and some moderately heavy (Ca, Fe, Cu, Zn, Se, Sb,...) elements are measurably variable. Examples for all of these elements are given here with analyses being made by an isotope ratio mass spectrometer (IRMS), an ion micoprobe (SIMS) or a multiple collector inductively coupled plasma source mass spectrometer (MC-ICP-MS). Applications of NMR spectroscopy (170, 27A1, ...) to determine the kinetics of exchange reactions are not considered here (e.g., Casey & Phillips, 2001; Phillips et al., 2003). Interpretation of mass-dependent isotopic variations in natural phases invariably requires knowledge on how the isotopes are partitioned or fractionated among associated phases, and of the mechanisms and rates of exchange. All of these depend on experimental data. More recently, experimental and theoretical studies have also contributed to our understanding of mass-independent isotopic fractionations (Thiemens & Heidenreich, 1983; Thiemens, 1999, 2001; Gao & Marcus, 2001). Fractionation factors are determined by experimental, theoretical and / or empirical methods. All methods have their advantages and limitations. Although the theoretical methods are not discussed here (see Richet et al., 1977; Kieffer, 1982; O'Neil, 1986; Clayton & Kieffer, 1991; Polyakov, 1998 and references therein), it is worth recalling that (1) they require experimental spectroscopic data including, if possible, on the phase strongly enriched with the minor isotope, and (2) experimental fractionation data are essential to confirm or test the statistical mechanical model calculations (e.g., Urey & Rittenberg, 1933; Bigeleisen, 1965; Bigeleisen et al., 1973; Clayton & Kieffer, 1991). Thus, a detailed interplay between the results of theoretical studies and high precision experimental fractionation data is essential.

Experimental Measurementof Isotopic Fractionation Factors ...

993

In this presentation of experimental stable isotope geochemistry, emphasis is placed on the principles, methods and techniques of isotopic experiments, certain aspects of their design and some problems associated with the interpretation of the data. The examples given have been selected using a number of criteria including the pioneering application, and more recent studies to illustrate particular points. Exclusion of a paper in no way reflects on the quality of the work because there are more good studies than can be mentioned here. Such a chapter cannot review the experimentally derived data or end-results of a specific experiment. However, recent reviews of (1) equilibrium isotopic fractionation factors with a listing of systems studied and their calibration(s) has been given by Chacko et al. (2001), and (2) rates and mechanisms of isotopic exchange have been presented by Cole & Chakraborty (2001). Also, details of the equipment and its use, choice of materials for furnaces, thermocouples, etc.., and measurement of temperature, pressure, f02, pH, etc.., are not discussed because they would require a book to themselves. They are available in, for example, Edgar (1973), Ulmer & Barnes (1987), Holloway & Wood (1988) and references therein. Empirical methods of determining fractionation factors are also discussed. In general, these methods can be considered to be field experiments, in contrast to laboratory experiments. Such an approach has been an important source of fractionation data, particularly for biologically mediated reactions (e.g., aragonite or phosphate water systems, or sulphate reduction), and isotope exchange reactions with extremely slow exchange rates on a laboratory time-scale, often, but not invariably, at low temperatures (e.g., clay mineral systems). The experimental determination of fractionation factors is essential to provide constraints on the interpretations of isotopic data on natural systems and for understanding the principles underlying their behaviour. The temperature calibration of the isotopic geothermometers is one facet. Characterizing kinetic isotope effects, and mechanisms and rates of exchange are others. The general principles of isotopic exchange reactions and mechanisms of exchange as well as the notion of equilibrium are briefly discussed because these can influence the design of an experiment. Examples are inevitably dominated by the isotopic ratios of light elements measured by IRMS (H, C, N, O, Si, S, C1,...) because experimental studies have only really just started on some of the "new" elements (Mg, Ca, Fe, Cu, Zn, Se, Sb,...) that are measured by MC-ICP-MS (e.g., Mar6chal et al., 1999; Matthews et al., 2001; Zhu et al., 2001; Beard et al., 2003; Rouxel et al., 2003), or using new microtechniques such as SIMS (e.g., Valley et al., 1998) and laser probes (Rumble & Sharp, 1998; Shanks et al., 1998). 46.2 The isotopic fractionation factor To avoid any misunderstanding, the definition and implications of the isotopic fractionation factor between two substances A and B is: (~A-B = RA/RB

[46.1]

994

C h a p t e r 46 - S.M.F. S h e p p a r d

where R is the isotopic ratio of the less abundant and heavier isotope over the more abundant and lighter isotope of the element of interest (e.g, D / H , 180/160, 65Cu/ 6 3 C u , .... ), with two notable exceptions: 6 L i / 7 L i and 1 1 B / 1 0 B . Note that the definition is both quite general, being applicable to both mass-dependent and mass-independent isotopic fractionations, and practical; it is applicable to any two substances that contain the element and assumes no thermodynamic or other relationship between the two substances. For example, one can talk meaningfully about the c~'s between starting materials of an experiment before any isotopic exchange has occurred. The fractionation between any two substances can be derived directly from the 6-values of A and B that are measured relative to a common standard (O~A-B - [1000 + 6A]/[1000 + 6B]). Clearly, however, c~x-y is more interesting and useful if it is related to a welldefined kinetic or equilibrium process. But more information is required before c~ can be equated with a thermodynamic function such as the equilibrium constant, K. Although the temperature calibration of a fractionation factor is nearly always presented as a function of c~ (e.g. 1000 In (~A-B - f(T) rather that 1000 In K A - B - f(T) with T in Kelvin), it is too often implicitly assumed that c~ and K are interchangeable, without specific or unambiguous evidence for equilibrium, using the relation (~A-B = K1/nA_B where n is the number of exchangeable atoms in the reaction and assuming that the isotopes are randomly distributed over all possible sites. 46.3 Unidirectional and bidirectional reactions

Experiments to measure fractionation factors fall into two major families: unidirectional and bidirectional processes. The former are specifically for measuring kinetic effects since they are time and, therefore, path dependant; equilibrium may be attained but cannot be unambiguously demonstrated. The latter are for deriving equilibrium fractionations and kinetic information such as activation energies and rate constants for the exchange reaction. Basically it is necessary to devise an experimental setup where the uni- or bi-directional reaction can take place under strictly controlled physical, chemical, mineralogical and, if necessary, biological conditions. Although the experiment may appear to be rather straightforward in principle, in practice, a number of more or less surmountable problems may arise. For a unidirectional reaction, the fracfionation factor is strictly kinetic. Being path dependant, its value is not necessarily unique for any chosen overall reaction. Taking the reduction of sulphate to sulphide at 25~ and I bar as an example, experimentally measured fractionation factors vary from 0 to about 50%o, compared with the calculated equilibrium value of about 70%o (e.g., Kemp & Thode, 1968; Goldhaber & Kaplan, 1974; Canfield, 2001a). The magnitude of the kinetic fractionation depends on which step during sulphatereaction is rate limiting, the rate of reduction and whether the reduction is chemical or bacterial. For bacterial reduction, the fractionation varies with the strain of the bacteria, their degree of acclimatization to the various electron donors, nature of food and its supply, etc... Hidden or unidentified factors may play a decisive role during the course of the reaction. The situation can be quite complex. Isotopic exchange reactions that involve chemical reaction, and thus mineral synthesis, are unidirectional reactions. Such systems, which initially are not in chemical equilibrium, have been used to derive fractionation factors and to elucidate equilib-

Experimental Measurementof Isotopic FractionationFactors ...

995

rium fractionations (e.g., Clayton, 1959; O'Neil & Taylor, 1967, 1969; Matthews & Beckinsale, 1979; Bird et al., 1993; Vitali et al., 2001). Further aspects of this critical subject are discussed in section 46.5.

46.4 Equilibrium Measuring equilibrium fractionation factors has played a central role in experimental studies. Thus, the following remarks, concerning bidirectional exchange reactions that are designed to measure equilibrium fractionation factors, discuss criteria for complete exchange experiments and methods of extrapolation to the equilibrium value for partial exchange experiments. To determine equilibrium isotope fractionations, criteria must be built into the experimental setup to demonstrate that equilibrium was attained or, at least, more or less closely approached or bracketed (i.e., the reaction is reversible). Consider an exchange reaction between two substances where only the isotopes of a single element exchange between the two substances, that are neither created nor destroyed during the exchange reaction and experience no change in chemistry or texture" xA1 + yB2- xA2 + yB1

[46.2]

where A and B are the two substances, and I and 2 refer to the molecules totally substituted by the light and heavy isotope respectively. Thus, only the isotopic composition of the substances changes during the reaction. Equation [46.2] can refer to a homogeneous reaction taking place within a single physical phase, such as a fluid, or a heterogeneous reaction where the reaction occurs at the interface between the two phases (no production of a new phase or change in texture). For the latter case, exchange is considered to be purely diffusional and solution reprecipitation or modification of the surfaces is taken to be negligible (see section 46.5). It is assumed, for simplicity, that for substances with more than a single atom of the element of interest (e.g., oxygen in CaCO3), the several atoms are equivalent or at least indistinguishable during the exchange experiment. Although such an exchange reaction can be written between any two constituents, whether it actually works in the laboratory, or in nature, depends on the availability of a suitable mechanism of exchange and its rate. If equilibrium is attained (i.e., 100% exchange), the actual reaction path followed does not affect the value. Here, we assume that an efficient exchange mechanism exists. At any given time, t: CX(A-B)t- (1000 + ~At)/1000 + ~Bt)

[46.3]

but at equilibrium, eq: (Z(A-B)eq = (1000 + 6Aeq)/1000 + 6Beq) - K1/n(A-B)

[46.4]

where K is the equilibrium constant and n is the number of exchangable atoms with K - [(A2)x(B1)Y]/[(A1)x(B2)y] = (A2/A1) x/(B2/B1)y

[46.5]

996

Chapter 46 - S.M.F. Sheppard

Such an exchange reaction (equation [46.2]) implies that the equilibrium fractionation value can be approached from both sides of the value (equilibrium bracketing or reversing) without modifying the substances, except for their isotopic composition. That is, the reaction is perfectly reversible as the rate-controlling factors are independent of whether the isotopic ratio of the substances increases or decreases to arrive at the equilibrium value. The attainment of equilibrium can thus be unambiguously demonstrated. In practice, systems are not usually ideal and criteria for equilibrium are usually more complex, particularly for heterogeneous systems that contain a solid phase. Pattison (1994) gives an excellent discussion of the subtleties of this crucial point. After Pattison (1994), for most experimental isotopic systems between a solid phase and a fluid, what is usually referred to as an equilibrium fractionation factor is actually apparent isotopic compositional bracketing. Examples of the two-direction approach are given in Figure 46.1 for heterogeneous reactions. Figure 46.1a presents a simple, ideal model example from O'Neil (1986) for O-isotope exchange between exactly the same quartz sample (grain size, defect density, initial isotopic composition, etc...) and three isotopically different waters (W1, W2 and W3), that are the same except for their 180/160 ratios. The isotopic exchange rates for the forward and reverse reactions are taken to be identical. Note that isotopic exchange between three isotopically different quartz samples and a single water is not necessarily an equivalent system; the three quartz samples are highly unlikely to have exactly the same physical and chemical properties except for their isotopic compositions. At 500~ (Figure 46.1a), 50% exchange occurs per day in the three parallel experiments W1, W2 and W3 that have exactly the same P-T-t history and therefore identical rate-controlling factors. Thus, the degree of exchange in the three charges is proportional to the distance that each system is from isotopic equilibrium. After more than 5 days, > 97% exchange has taken place; the equilibrium fractionation of 3.0 can be determined with the necessary precision if the initial quartz-water fractionation in charge W2 was chosen to be close to the equilibrium (or final) value. Figure 46.1b, after Graham et al. (1980), presents the measured time dependence of the H-isotope fractionation between zoisite-water for two parallel runs where everything was identical except the D / H ratios of the waters, D and E. After about 10 days, essentially 100% exchange had been achieved. The close similarities between the nature of the curves of Figures 46.1a and 46.1b are evident. The form of the curves in Figure 46.1b 4 6 . 1 - Diagram illustrating the change in the fractionation factor with time for exchange reactions using the two-direction approach to the equilibrium fractionation, a Model O-isotope exchange between quartz-water with starting-material: 618Oquartz at + 10 and three waters at - 5 (A), + 5 (B), and + 15 (C), that proceeds at the rate of 50% exchange per day (after O'Neil, 1986). After about 5 days the fractionations for runs A, B, and C converge and bracket the equilibrium fractionation of 3%o.b - Measured H-isotope fractionations between zoisite and water with time for two companion runs with the same zoisite but isotopically different waters as starting-materials (after Graham, et al., 1980). The initial zoisite-water fractionations are both smaller and larger than the equilibrium value. After about 10 days, the two fractionations converge on the equilibrium value of about- 52%o (i.e.--- 100% exchange). Figure

Experimental Measurement of Isotopic Fractionation Factors ...

997

998


implies that the hypothesis that the rate of exchange, for both the forward and reverse reactions, is proportional to the distance that the run is from equilibrium is quite satisfactory for this system. Also, at the rather similar levels of concentration of the different isotopic species in these experiments, the rate constants for the different isotopic species are indistinguishable. For heterogeneous reactions, the rate of reaction will not necessarily be constant with time because the surface areas of the grains may change with time (see section 46.5). On Figure 46.1b, there are too few data points, particularly during the first few minutes and hours of the experiment, to determine how the rate of exchange changed with time. The process was probably not purely diffusional and without dissolution; it, therefore, probably was not constant. In reality, some exchange reactions do not attain, or even closely attain, equilibrium, because rates of exchange are just too low on the laboratory scale (usually less than a few months and very rarely over a year). Extrapolation techniques have been developed for partial exchange experiments.

46.4.1 Partial exchange: the Northrop-Clayton method Partial exchange techniques were developed by Northrop & Clayton (1966) for two phase systems, where each phase contains only a single site for the element of interest. In such systems, the rates of isotopic exchange for directly associated experiments are assumed to be equal but not necessarily constant. Associated experiments refer to 2 or more runs that are as identical as possible in (1) their preparation- identical solids (grain size and its distribution, isotopic composition, imperfections, mass,...), identical solutions (chemistry, mass,...) except for their isotopic compositions, identical reaction containers (volume, degree of filling, etc...), and (2) their P-T-t paths during their reaction history. Figure 46.1a illustrates the results of a model experiment where the rate of exchange is constant. Usually three runs are performed together. Two of the runs are chosen so that their initial isotopic compositions of the fluids are such that the isotopic differences with the solid, or salt to be dissolved, are both substantial and on both sides of the estimated equilibrium value, i.e. the initial a's are both smaller and larger than the equilibrium value of a. The evolution of these fractionations with time can thus bracket the equilibrium value (Figure 46.1). A third associated run is designed to have an initial fractionation that is as close as possible to the anticipated true equilibrium value. In this way (1) the degree of exchange, and (2) the extrapolated equilibrium fractionation value can be determined with the best precision for the degree of exchange. Northrop & Clayton (1966) derived the following equation [46.6] for the kinetics of partial isotopic exchange reactions in heterogeneous systems between phases A and B for all initial and final values of ~A-B (for simplicity just c~) close to unity (i.e., In a ~ 1 + c0 and assuming that the reaction rates in the forward and reverse directions are the same: In c t i - In CZeq- 1/F(ln ~f- In c~i)

[46.6]

where the subscripts i, f and eq refer to the initial, final and equilibrium fractionations,

999


respectively, and F is the fractional approach to equilibrium. For a set of runs (usually 3), a plot of (In af- In cq) versus In ~i (Figure 46.2) gives a straight line with slope 1/F and an intercept with the y-axis of In aeq, if the assumptions are valid. For a single exchangeable atom, at equilibrium In af = In (~eq = In K and F - 1. The accuracy of the determination of the intercept or equilibrium value increases with increase in the percentage of exchange. Equation [46.6] has also been applied to homogeneous reactions. For hydrogen, the approximation In r ~ (1 + r is often not satisfactory. Therefore Suzuoki & Epstein (1976) used the following modified form of equation [46.6]" (ai

-

1) = (~eq - 1) - 1 / F(cxf - oti)

[46.7]

Equation [46.7] was used by Graham et al. (1980) for the analysis of the data presented on Figure 46.1b. Criss et al. (1987) and Criss (1999) reexamined the derivation of equations [46.6] and [46.7] and proposed a more complex quadratic equation [46.8] for a pair of exchange experiments, labelled by subscripts I and 2, based on intensive

Figure 46.2 - An example of the Northrop & Clayton (1966) partial exchange method where three companion runs, starting with the same mineral (dolomite) and three isotopically different waters, exchange for a given time and temperature (after Northrop & Clayton, 1966 and O'Neil, 1986). The intercept with the Y-axis gives the best estimate of the equilibrium fractionation for the set of runs, and the slope gives the degree of exchange (about 50%). With increase in percent of exchange, the line rotates anticlockwise, increasing the precision on "the best estimate of equilibrium fractionation".

1000

Chapter 46- S.M.F. Sheppard

variables only and ~ values rather than the logarithms of ~ values" ( 1 - H)(o~eq)2 + (Hcqi + HR2f- o~2i- C~lf)(Req + (O~lfC~2i-Ho~2fO~li)- 0

[46.8]

where H = (RBi/RBf)I (RBf/ RBi)2

[46.9]

using the same notation as above plus RB, the isotopic ratio of phase B, often water. Equation [46.8] is used to solve for (~eq. In the Criss derivation, the fractional approach to equilibrium, E is given by: 1- F - (RBf/RBi) [(af- ~eq)/(O~i- ~eq)]

[46.10]

and F is in fact a complex function of the relative mole fractions of A and B (Criss, 1999). All these equations [46.6,7 and 8] assume that the rate of isotopic exchange is proportional to the deviation of the isotopic ratio from the equilibrium ratio. This technique has been widely applied to the geothermometric calibration of heterogeneous single mineral-water (or aqueous solution) systems for H-, C- and O-isotopes (e.g., Northrop & Clayton, 1966; O'Neil & Taylor, 1967, 1969; Clayton et al., 1972; Suzuoki & Epstein, 1976; Matsuhisa et al., 1979; Graham et al., 1980, 1984:, 1987; Stoffregen et al., 1994a,1994b; Fortier et al., 1995; Guo & Qian, 1997; Cole & Ripley, 1998; Saccocia et al., 1998) and homogeneous reactions in water (e.g., Lloyd, 1968; Chiba & Sakai, 1985; L6cuyer et al., 1999; O'Neil et al., 2003). In order to avoid problems such as congruent or incongruent dissolution of the mineral in water, carbonate (e.g., Clayton et al., 1989, Chiba et al., 1989; Fortier et al., 1994), CO2 (e.g., Scheele & Hoefs, 1992; Matthews et al., 1994; Palin et al., 1996) and H2 (Vennemann & O'Neil, 1996) have been used as the exchange media. Based on a substantial number of experiments that have applied this technique and comparisons with other techniques, it has been observed that if the percentage of exchange was less than about 80 - 90% or so, then the intercept value for the incomplete exchange method may be larger than the "correct" equilibrium fractionation (Clayton et al., 1972; Matsuhisa et al., 1978; O'Neil, 1986). The reasons for this difference are not always understood but the reactions usually involve water or carbonate. Additionally, the basic hypothesis that the rates of isotopic exchange are identical for systems that only differ in their isotopic composition of the fluid reactant may not always hold. Recent experiments on exchange between (PO4)aq and H20 suggest that, under certain conditions of pH, the rate of exchange may depend on the distance from equilibrium (O'Neil et al., 2003). Further aspects of this observation need to be explored. Another way round this type of situation in systems with a solid reactant may be provided by application of the ion probe, if the individual crystals are sufficiently large, so that isotopic profiles can be measured (Fortier et al., 1995; Chacko et al., 1999). The Northrop-Clayton method may not be applicable to substances where the element of interest is present in more than a single crystallographic site or there are more


1001

than two phases (e.g., oxygen in micas, hydrogen and oxygen in chlorites, carbon and oxygen in calcite-dolomite-fluid, etc...)(O'Neil & Taylor, 1969; Sheppard & Schwarcz, 1970; Sheppard, 1980; Chacko et al., 1996; Cole & Ripley, 1998). For such systems, the exchange rates among the different sites may be similar or quite different, depending in large part on the mechanism of exchange (see section 46.5). Equations for exchange in three phase systems have been developed by Zheng et al. (1994). In systems with two solids and a fluid, the form of the curves representing the changes in the isotopic compositions of the solids and fluid with time - usually initially rapid followed by slower 8 changes - are much more complex than those given in Figure 46.1 because the kinetics are usually more complicated in systems with two solid phases (Zheng et a1.,1999).

46.4.2 Partial exchange: the three-isotope method The three-isotope method is an important modification of the Northrop-Clayton technique that was introduced by Matsuhisa et al. (1978), with later modifications by Matthews et al. (1983a). The isotope ratios 1 7 0 / 1 6 0 and 180/160 are measured on the same substances so that two fractionations before (or initial) and after (or final) are derived from each experimental capsule. The three-isotope plot, Figure 46.3, illustrates the method, with the assumption that the two isotopic ratios change at the same rate. The natural or synthetic starting mineral (Mo) plots on the primary mass-dependent fractionation line, PF, but the initial isotopic composition of the water sample (Wo) is chosen, by mixing with isotopically enriched water, so that it plots off the line PF in such a way that the 180/160 ratio is close to the presumed equilibrium value whilst the 170/160 ratio is far from the equilibrium value. In this way the change in the 170/160 fractionation sensitively monitors the extent of exchange while the 180/ 160 ratio closely brackets the equilibrium value, enabling accurate determination of its value. For 100% exchange, the final 8170 and 8180 mineral (Me) and water (We) equilibrium values plot on a secondary mass-dependent fractionation line that is parallel to PF and passes through the bulk isotopic composition of the mineral-water system. Isotopic measurements are made on pure oxygen gas. The combination of the three-isotope exchange method with the use of high pressures (e.g., piston cylinder apparatus) to accelerate the rate of exchange (see section 46.6.3) is thus an important technical innovation in experimental calibration. 46.5 Mechanisms and rates of isotopic exchange Many experimental studies designed to measure equilibrium or kinetic fractionation factors also yield information on the factors which affect the exchange reactions (e.g., Graham, 1981; Giletti, 1985; Cole & Ohmoto, 1986; Cole & Chakraborty, 2001). Some experiments were specifically designed to elucidate mechanisms and rates of exchange (e.g., Sakai & Dickson, 1978; Ohmoto & Lasaga, 1982; Giletti, 1985; Cole, 1992, 2000). Knowledge (or anticipation) of the mechanisms and extent or rates of exchange for both heterogeneous and homogeneous isotopic reactions can influence both the design of an experiment and the interpretation of the results. Isotopic exchange between a mineral and a fluid can occur by (1) chemical reaction or mineral synthesis, (2) dissolution-reprecipitation, or (3) diffusion. In reality, more than one

1002


Figure 46.3 - Schematic diagram of the three-isotope exchange method (after Matthews et al., 1983a). Natural samples plot on the primary mass fractionation line (PF) with slope of 0.52 (Matsuhisa et al., 1978). Initial isotopic compositions are mineral (Mo) on PF and water (Wo) whose fractionation with the mineral is well removed from the equilibrium value in 170/160 (by addition of labelled water; see Matthews et al., 1983a for details) but very close to equilibrium in 180/160. Complete isotopic exchange for the exchange reaction is defined by the secondary mass fractionation line (SF) parallel to PF and passing through the bulk isotopic composition of the total mineral plus water system. Isotopic compositions of partially equilibrated samples are Mf and Wf and completely equilibrated samples are Me and We. Values for Me and We can be determined by extrapolation from the measured values of Mo, Mr, Wo and Wf.

mechanism may operate in the system at the same time, but one mechanism often dominates. Chemical reactions involve the production of phases that were not present at the beginning of the experiment. They are unidirectional reactions which may attain equilibrium, but this cannot be unambiguously demonstrated. Although free energy changes associated with chemical reactions are usually one to three orders of magnitude larger than those for an isotope exchange reaction by itself, certain reactions can lead to metastable phases and/or disequilibrium fractionations, that are not necessarily readily detected (e.g., Matsuhisa et al., 1978). Although the subject of chemical versus isotopic equilibrium cannot be discussed here, it is noted that the attainment of chemical equilibrium does not imply that isotopic equilibrium has also been attained. Dissolution-reprecipitation processes group together a number of processes that occur at external surfaces or within the grains, and certain involve chemical reactions.


1003

Experimental and theoretical studies have shown that, in general, the rate constants for exchange are related to surface area, fluid/solid ratio, the fraction of exchange and the time (Cole et al., 1983; Dubinina & Lakshtanov, 1997). A special case of dissolution-reprecipitation is the Ostwald ripening kinetic process (Ostwald, 1900; Baronnet, 1982; Eberl et al., 1990) which has been modelled by Chai (1974) and Stoffregen (1996). Because the starting solids invariably have a range of grain sizes and solubility varies with particle size (smaller grains have a higher surface energy), the larger particles grow at the expense of the smaller ones, without involving nucleation processes. The end products of an experiment can lead to a steady state condition where the grainsize distribution is independent of time and the physicochemical variables of the experiment, if the total number of particles is large enough to give a continuous size distribution. Exchange of O-isotopes between calcite and water during recrystallization of the former is consistent with an Ostwald ripening process (Anderson & Chai, 1974), and similarly for quartz-water (Matthews et al., 1983a), alunite-water (Stoffregen et al., 1994b) and chlorite-water (Cole & Ripley, 1998). Examples of changes in grain-size of run-products compared to their starting materials observed during experimental runs with quartz-water, anorthite-water and calcite-CO2 are given by Matsuhisa et a1.(1978), Matthews et al. (1983a) and Rosenbaum (1994). Note that the cores of the larger, growing grains may retain their initial isotopic composition because they are effectively armoured from isotopic exchange by the newly deposited layers. This leads to mineral-fluid isotopic disequilibrium without necessarily the intervention of a kinetic isotope effect. Chai's (1974) and Stoffregen's (1996) models enable one to calculate the fractional volume of the original material transferred through the fluid medium for a given increment of the average grain-size. For example, doubling the average grain-size implies that 75% of the initial solids have been transported through the solution (see Figure 8 in Chai, 1974). Importantly, isotopic overshooting of the true equilibrium value can occur by such a process. A dissolution-reprecipitation processes was proposed by O'Neil & Taylor (1967, 1969) for experiments where cation exchange took place to promote O-isotope exchange in systems such as albite-KC1-H20, sanidine-NaC1-H20 and muscoviteNaC1-H20. In particular, they showed (see Figures 5 & 6 in O'Neil & Taylor, 1967) that the gross morphology of the crystal was retained as the reaction front of both cation and isotopic exchange swept through the crystal, with the fluid within the crystal remaining in communication with the external fluid reservoir. There was no evidence for armouring of the original grains. Cation exchange dramatically increased the rate of isotopic exchange, but equilibrium cannot be demonstrated in such experiments. Dubinina & Lakshtanov (1997) have modelled these processes. The rate of change of solution-reprecipitation, textures and isotopic exchange in a charge and mineral-volatile fractionations can be influenced by the presence of "impurities". These may have been added during filling of the charge or produced within the capsule during the run and/or pressure-temperature device. For example, production of hydrogen from the dissociation of moisture in the pressure medium of a piston-cylinder apparatus, followed by its infiltration into the charge to react with CO2 to give CO and H20, has been discussed by Rosenbaum & Slagel (1995). Rosen-

1004


baum (1994) documented the role of minor quantities of water on these rates of change in the calcite-CO2 system. Importantly, he also demonstrated that, because the rate of exchange between CO2 and H20 was much faster than between either of these volatiles and the associated calcite, the final measured calcite-CO2 fractionation had been modified from its value attained at the temperature of the run (900~ thereby accounting for the scatter in the results of the runs of Chacko et al. (1991). For minerals like micas and chlorite with more than a single crystallographic site for oxygen, the rates of exchange of the different oxygens may be similar or different. If the mechanism of exchange is dominantly by recrystallization or dissolution-reprecipitation, then the rates may be very similar. On the other hand if pure diffusion is dominant, then the rates of exchange may be quite different. Diffusion, or volume diffusion, implies that each grain retains its dimensions and shape during the transport of isotopes across the surfaces. The magnitude of the diffusion coefficient of the element of interest depends on the nature of the diffusing or carrier species (often not well characterized) that in turn is related, but not necessarily directly) to the nature of the external reservoir (e.g., H20, CO2, 02, H2, CaCO3). In fact the experimental data can be used to make a choice among a number of possible carriers. For example, experiments that compare "dry" conditions with "wet" conditions can define the role, if any, of H-bearing species (e.g., Elphick et al., 1988). In many systems, isotopic exchange commences relatively rapidly with the surface layers and is followed by lower rates as the diffusing species penetrate into the bulk of the grain. Certain of the experimental data aimed at measuring diffusion coefficients have or may have been influenced by solution-reprecipitation processes. Some of the more recent studies measuring diffusion constants use isotopically enriched material and measure the isotopic depth profile in individual grains, often of known crystallographic orientation, using the ion microprobe (e.g., Giletti et al., 1978; Giletti & Yund, 1984; Elphick et al., 1986, 1988; Fortier & Giletti, 1991; Farver & Yund, 1995, 1999). In fact, equilibrium hydrogen isotope fractionation factors have been derived from depth profiles using the ion probe (Chacko et al., 1999). For further discussion of this topic see the review by Cole & Chakraborty (2001) and, for example, Graham (1981), Stolper & Epstein (1991), Zhang Y. et al. (1991a, b), and Palin et al. (1996). Another consequence of pure diffusion controlled exchange processes is that fractionation factors can evolve with depth of penetration, because surface properties are not identical to those of the bulk material (Hamza & Broecker, 1974). Application of suitable self-diffusion coefficient data are necessary to confirm that the length of the experimental run is long enough so that isotopic exchange has penetrated sufficiently deep (> ~ 0.3 mm) into the material to avoid the influence of surface fractionation factors (e.g., Stolper & Epstein, 1991; Matthews et al., 1994). Surface effects may also intervene, but be difficult to detect, if mineral to fluid ratios are very small and essentially complete exchange has only occurred between the surface layers and the fluid. Elphick et al. (1986) avoided grain boundary and pressure solution-reprecipitation processes from affecting their diffusion experiments by using a gel technique that hydrothermally deposited an isotopically labelled overgrowth on the quartz or feld-


1005

spar starting material. Such experimental diffusion data are essential input to model quantitatively the effects of diffusion on stable isotope partitioning among minerals during prograde reaction histories and retrograde exchange processes, with applications to thermometry, open versus closed system behaviour, and cooling rates (Giletti, 1986; Eiler et al., 1992, 1993; Jenkin et al., 1994b; Kohn & Valley, 1998). However, these models can also be used to aid a field experiment aimed at calibrating an isotopic fractionation factor (J. W. Valley, pers. comm., 2001). For example, model results can aid the screening of samples by indicating the relative role of mineralogy, mode and grain size. The isotopic exchange reaction given in equation [46.2] implies that constituents A and B do not change chemically, i.e., they are in chemical equilibrium with each other. However, many experimental exchange studies start with an aqueous solution that initially is not in chemical equilibrium with the solids that are congruently or incongruently soluble in water. A certain amount of chemical reaction will inevitably take place and this can influence the results of the isotopic exchange reaction and quench products (see section 46.6.5). In incomplete exchange experiments which give extrapolated fractionation factors that are different from the true equilibrium value, water and carbonates are usually involved. Such problems appear to be particularly associated with the quartz-water and calcite-water systems (Matthews et al., 1994). Much effort to derive equilibrium fractionation factors has been invested in these important systems (quartz: O'Neil & Clayton (1964), Clayton et al. (1972, 1989), Matsuhisa et al. (1978, 1979), Matthews & Beckinsale (1979), Matthews et al. (1983a); calcite: Clayton (1961), O'Neil et al. (1969), Anderson & Chai (1974), Clayton et al. (1989), Cole (1992), Kim & O'Neil (1997)). On the other hand, dry gas-mineral exchange experiments, some of which may be controlled by pure diffusion, are often free of the above problems (Matthews et al., 1994; Rosenbaum, 1994). Future studies are required to characterize the order of the reactions and the mechanisms of exchange to advance our understanding of such systems.

46.6 Laboratory experimental methods Fractionation factors are so close to 1, usually between 1.05 and 0.95, that free energy changes of isotopic reactions are only a few calories. They are thus too small to be measured calorimetrically. Laboratory experiments therefore have to be designed so that the isotopic composition of the various substances before and after exchange can be measured directly or calculated from other measured parameters of the experiment. Experiments are carried out in a wide variety of containers made out of borosilicate or silica glass, plastic or noble metals, depending in large part on the selected experimental conditions. From an experimental point of view, it is convenient to divide laboratory experiments into two groups: (1) those conducted below, at or near atmospheric pressures, and (2) those contained in capsules that are placed within a high pressure-high temperature device. If the physical parameters cannot be con-

1006


trolled within the container then it is usually placed within an experimental P and/or T controlling device. A given type of device is usually associated with pre- and postexperimental characteristics such as the P-T-t heating up and quenching paths in addition to the designed P-T-t experimental conditions. Many of the P-T machines and techniques exploited by isotope geochemists have been adapted from those developed by the experimental petrology community. Table 46.1 summarises some of the characteristics of the principal devices that have been used for isotopic studies. Except for most of the near atmospheric pressure vessels, details of these devices and, in particular, their construction, calibration, operation, necessary safety precautions, basic principles of the preparation of sample materials in experimental geochemistry and original and key references are given in the books of Edgar (1973), Ulmer & Barnes (1987) and Holloway & Wood (1988). For experiments carried out above room temperature, the inevitable heating and/or quenching time-temperature trajectories can influence the end products, even if certain precautions are taken. However, quantitative or even semi-quantitative data on heating up and cooling histories are usually not given. Additionally; variations among different versions of the same general type of vessel can be quite considerable. Note that temperature-time quench paths follow an exponential curve so that the quench times given in Table 46.1 must only be taken as an extremely rough guide for the high temperature part of the curve.

46.6.1 Containers and capsules Containers can be basically an extremely simple test tube, flask or beaker in borosilicate or silica glass, or plastic, or a plastic box all at near atmospheric pressure, in the presence or absence of air or a controlled atmosphere. Others are much more complicated glass or precious metal capsule systems that require a device to at least tightly control the thermal regime. The principal purpose of the container is to hold the sample material in either a closed or controlled system during the run so that, for example, isotopic exchange can be interpreted in terms of the kinetic or equilibrium fractionation of interest. Under some conditions, usually at high temperatures, Fe loss to the metal capsule, loss of alkalis from a melt, or hydrogen diffusion can be so important that the chemistry of the phases changes substantially during the experiment, thus affecting the end results. From an isotopic point of view, a closed system requirement can readily be met for all elements except hydrogen, because glass and metals can act as semipermeable membranes for hydrogen. Sufficient quantities of hydrogen can diffuse through noble metal and silica glass capsules during a run to affect the interpretation of the end products of an experiment concerning H-isotopes unless special precautions are taken to reduce or eliminate such mass transfer processes (Graham et al., 1980; Vennemann & O'Neil, 1996). Note, as mentioned above in section 46.5, diffusion of hydrogen in piston-cylinder equipment can affect O-isotope fractionations.

46.6.2 Diffusion of hydrogen

The diffusion of hydrogen through metal capsules and its isotopic implications are discussed in some detail in the appendix to Graham et al. (1980). The permeability of precious metals to hydrogen have been measured at 2 kb and elevated temperatures by Chou (1986). Permeabilities increase in the sequence Au, Pt, Ag70Pd30. However,


1007

Table 46.1 - Summary comparison of experimental devices for isotopic studies and their principal characteristics* Vacuum to Atmospheric Pressures#

Cold-seal Vessels

Range of T: Precision, T Accuracy, T


1021

> Table 46.4 continued 10. SO2-H2S

Sulphur and H 2 0 were sealed in a silica glass bulb under vacuum and heated to 500 - 1000~ to exploit reaction: 3S + 2H20 = 2H2S + SO2. After reaction, H20, H2S and SO2 separated by trap to trap vacuum distillation. At each temperature, one of two runs overshot final T by 100~ for 24 hr to check attainment of equilibrium.

11. Kaolinite-water (D/H)

Kaolinite sealed in glass tubes with isotopically different H20 (few % of total hydrogen) and run in furnace at 200 - 352~ for up to 4 months.

12. Silicate glasses-water vapour (D/H)

Rhyolitic and feldspathic glasses and melts were sealed in quartzvycor tubes with large excess of water and run at 530 - 850~ PH20 = 1.4 - 2.8 bars, < 54000 hr. Vapour-dissolved hydrogen fractionation factor determined. Reversal experiments indicate equilibrium was attained.

13. Hydrous mineral-H2 (D/H)

Exchange between molecular hydrogen and mineral (epidote, kaolinite, muscovite, biotite, hornblende) at 150-400~ in silica tubes. Run with either "infinite" reservoir of H2 or "infinite" reservoir of mineral. Small change in 8D of gases during run probably due to exchange with trace D2 present in H2-HD gas (see section 46.6.4.3). Diffusive loss of hydrogen through silica glass became significant for T > 300~

14. CH4-H2 (D/H)

CH4-H2 mixture at 770- 930 m m Hg, circulated in closed system over a Ni-Thoria catalyst at 200 - 500~ After sampling exchanged mixture (in-line sampling volume with no contribution of H2 adsorbed on catalyst) gas-chromatographic separation of H2 from CH4 before IRMS to determine equilibrium fractionation.

15. CH4-H2-H20 (D/H)

Gas mixtures (CD4-H2 and D20-H2) reacted in closed glass tube at 100 - 500~ Most experiments at P < i bar. Rates of reaction monitored from formation of HD after cryogenic separation of gas mixture including CD4-H2.

16. Fischer-Tropsch synthesis

Equimolar mixture of CO and H2 (0.8 1), initially at I atm, heated to 400~ in Vycor flask in presence of Co catalyst. CO2 removed continuously during reaction. Kinetic fractionation measured between CO2 and organics.

17. CO2-CH4 exchange

Exchange and equilibration between CO2-CH4 in pyrex tube along with varying amounts (10 - 50 mg) of catalytic material (iron oxides and commercial transition-metal (Ni, Pd-Pt, Rh) catalysts), at 200 600~

(180/160)

(13C/12C)

(13C/12C)


1022


> Table 46.4 continued 18. 03-02 (170 / 160, 180 / 160)

Synthesis of ozone from molecular oxygen using an electric discharge (Tesla coil) or microwave plasma. Effect of oxygen pressure measured. Ozone trapped cryogenically during its production. 6170 and 6180 measured to unravel mass-dependent and mass-independent fractionations.

# 1. O'Neil & Epstein (1966a); 2. Anderson (1968, 1969); 3. Hamza & Broecker (1974); 4. Rosenbaum (1993); 5. Eiler et al. (2000); 6. Sharp et al. (1991); 7. Matthews et al. (1994), Palin et al. (1996), see also Stolper & Epstein (1991); 8. Derdau et al. (1998), Elphick et al. (1988), Giletti et al. (1978), see also Freer & Dennis (1982), Freer et al. (1997); 9. Feng & Savin (1991, 1993a, 1993b), see also Karlsson & Clayton (1990); 10. Thode et al. (1971); 11. Liu & Epstein (1984); 12. Dobson et al. (1989); 13. Vennemann & O'Neil (1996); 14. Horibe & Craig (1995); 15. L6cluse & Robert (1994); 16. Lancet & Anders (1970); 17. Horita (2001); 18. Heidenreich & Thiemens (1986), Bains-Sahota & Thiemens (1987). w All isotopic measurements made by IRMS except where noted. some quench products may be precipitated as rims on existing crystals and difficult to identify; determination of their mass, which is not necessarily negligible, or their isotopic composition can be very difficult. If the fluid phase contains two constituents such as H 2 0 and CO2, O-isotope exchange can occur between them during and after the quench. In some synthesis experiments, gels and poorly-ordered very fine-grained precipitates may be the initial product. These may undergo isotopic exchange during subsequent aging processes or preparation of the sample for isotopic analysis. Such effects have been documented by Bao & Koch (1999) in ferric oxide-water systems (Table 46.3, N ~ 11).

46.7 The experiment For experiments carried out above room temperature, the working temperature should be approached from below, with no overshooting (the isotopes do not k n o w that they are not meant to start exchanging before arriving at the working temperature!). Note that in most fluid-mineral systems, the surface layer of the solid undergoes isotopic exchange much more rapidly than the bulk of the sample. The physical conditions within the container at any given m o m e n t are intimately related to the nature of the pressure-temperature device and its use. For these reasons, a brief summary of the principle types of equipment and laboratory setups that have been used is given as a complement to the information presented in Table 46.1. The purpose here is to help both the reader to appreciate the advantages and limitations of these techniques from an isotopic point of view, and the potential user to select or design the best system for his situation. This section, which is inevitably selective, will consider (1) experiments close to atmospheric pressures or under vacuum, (2) externally heated vessels - cold-seal and rocking bombs, and (3) internally heated vessels, including solid media apparati.


1023

46.7.1 Experiments near atmospheric pressure or under vacuum

An extremely wide variety of experimental studies have been carried out below, at or slightly above atmospheric pressure in diverse types of containers. Many of these experiments were carried out near room temperatures, but some required the use of very high temperature furnaces (Table 46.1). Isotope ratios of most elements of interest have been studied. This section is therefore rather heterogeneous as there is no specific P-T device associated with these studies. The experiments can be divided into four groups" (1) systems with biological materials, (2) inorganic experiments carried Table 46.5 - Summary of selected open system experiments carried out at low pressures (< 10MPa, < 100 bar) or under vacuum. Ref. #

System

Experimental set-up

1. Silicate-melt or mineral-CO2 or 02 (180/160)

Silicate sample suspended in 1 atmosphere double-ceramic tube quenching furnace through which CO2 or 02 of constant isotopic composition was passed continuously at 1250- 1525~ Calculated diffusion coefficient and equilibrium fractionation factors

2. Calcite-bicarbonate solution (13C/12C)

Inorganic precipitation of calcite or aragonite in temperature controlled (10 - 40~ water-jacketed reaction vessel by bubbling CO2 N2 gas of known chemical and isotopic composition through Na-CaC1 solutions using synthetic calcite or aragonite seeds. Precipitation rates varied from 102.6 to 104.8tcmolm-2h-1.

3. Evaporation of basalt and chondrites (30Si/28Si)

Basalt and two carbonaceous chondrites were evaporated using a solar furnace.

4. Evaporation of forsterite (26Mg / 24Mg)

Single crystals of forsterite were evaporated at 1500 - 1800~ in a vacuum furnace. Measured solid-gas kinetic fractionation of Mg and derived diffusion coefficients.

5. Evaporation of synthetic "solar" material (26Mg/24Mg, 170/160, 180 / 160, 30Si / 28Si)

A FeO-MgO-SiO2-CaO-A1203-TiO2 rare earth element melt system was evaporated in vacuum furnace at 1800 - 2000~ for few seconds to 5 h. Chemical fractionations monitored along with kinetic isotope effects.

6. Evaporation of solid silica (180 / 160, 170 / 160)

Partial evaporation of silica at 1600 - 1700~ under vacuum (10-9 bars) and in H2 or N2 (> 10-5 bars) atmosphere. Degree of 180 and 170 enrichment of solid controlled by kinetic fractionations at ablating surface, rate of sublimation and efficiency of O-self diffusion in solid.

7. Fractionation of CH4 during diffusion (13C/12C)

Methane diffused at 90~ and 9 MPa through shales with varying organic carbon contents in triaxial flow cell. Fractionation during molecular transport (diffusion) measured.

# 1. Muehlenbachs & Kushiro (1974), see also Appora et al. (2003); 2. Romanek et al. (1992), see also McCrea (1950), Turner (1982); 3. Molini-Velsko et al. (1987); 4. Wang et al. (1999); 5. Wang et al. (2001), see also Richter et al. (2002); 6. Young et al. (1998c); 7. Zhang T. & Krooss (2001).

1024


out in glass or plastic flasks where the contents in the flask may or may not strictly be a closed system, (3) experiments run in sealed containers, and (4) open system experiments. The division between (2) and (3) is sometimes rather arbitrary. A selection of such experimental set-ups is summarized in Tables 46.2 to 46.5. Because of their large diversity, details of the individual experiments cannot be presented; the references give the necessary entry to their descriptions. A few general observations and remarks can be made.

46.7.1.1 Biological systems (Table 46.2) Most of these experiments were designed to measure kinetic isotope fractionations. Many of the experiments using biological material require a certain specialized knowledge of the biological elements under study and access to particular equipment that is not brought out directly in Table 46.2. The reader is referred to the references for such details on, for example, the methods of bacterial culture, etc... Similarly, isotopic analysis of certain of the starting materials and end products require special techniques (see references, Table 46.2 and Lyons et al., 2003). Much attention has been directed to growing plants under various controlled conditions. The systems are quite complex because, among several things, (1) the element of interest is usually present in different sites and substances with different degrees of exchangeability, (2) tissues or organs from the same plant may exhibit isotopic variations substantially larger than the analytical precision, and (3) isotopic compositions can be a function of a wide variety of variables including temperature, humidity, water availability, light intensity and wavelength, CO2/02 ratio, nutrient supply, etc... (e. g., Farquhar et al., 1989a; Heaton, 1999; Tang et al., 2000).

46.7.1.2 Inorganic systems in more or less closed flasks (Table 46.3) Such experiments in glass or plastic flasks were designed to measure either equilibrium or kinetic fractionations at temperatures that were usually below 100~ The contents of a flask in certain experiments were essentially a closed system. An extremely wide variety of systems have been explored for most elements including B, C1 and Se. Certain, N ~ 1 and 2 of Table 46.3, are chemical reduction reactions which can be considered as the chemical equivalent of biological reduction reactions (Table 46.2, N ~ 1, 2 and 5). In addition, to measure the thermal diffusion constants for 29N2/ 28N2 in air at temperatures from -60 to O~ Grachev & Severinghaus (2003) developed a two-bulb apparatus that fits into a dual temperature bath. 46.7.1.3 Sealed tubes (Table 46.4) These experiments have yielded kinetic, including mass-independent, and equilibrium fractionation data and diffusion data. Tube and fluid volumes were selected so that internal pressures were less than atmospheric to a few atmospheres at the working temperature, to avoid the use of more complicated pressure controlling devices. Quenching of the tube can usually be carried out in seconds to a few minutes, depending on the mass of the tube system. Alternatively, the substance of interest (e.g., 03; Thiemens & Heidenreich, 1983) can be cryogenically removed during the reaction. Note that silica glass becomes inconveniently permeable to H2 gas at tern-


1025

peratures above 300~ (Vennemann & O'Neil, 1996).

46.7.1.4 Open system reactions (Table 46.5) These systems are relatively limited in number but wide in application and aimed at deriving either kinetic or equilibrium fractionations. They include studies on silicate melts at about atmospheric pressures (N ~ 1) or under high vacuum, the latter directed at fractionations of Si, Mg and O under simulated space conditions (N ~ 3, 4, 5 and 6), precipitation of carbonates at variable rates (N ~ 2), and diffusion of gas through water-saturated shales using a flow cell (N ~ 7). 46.7.2 Externally heated pressure-vessels Externally heated pressure-vessels cover a wide variety of steel and nickel-chromium-cobalt-based alloy vessels, with or without pressure control, that are placed within an electric furnace. Two types of autoclaves have principally been used for isotopic studies: cold-seal vessels and hydrothermal rocking bombs. The fugacity of hydrogen within the capsule or container is controlled by the vessel and its pressure medium unless special precautions are taken (see section 46.6.2). The experimental data are principally for H, C, O and S. 46.7.2.1 Cold-seal vessels These refer to Tuttle-type cold-seal vessels (Tuttle, 1949), or a modified form. Some are fitted with a water cooled jacket near the cold-seal end so that by removing the furnace and turning the bomb, attached to its flexible capillary or pressure tubing, into a vertical position, the capsules inside can drop down into the cold zone for very rapid quenching. Cold-seal vessels of one sort or another have been very widely used for isotopic calibration studies of minerals for all elements since the late 50's (Clayton, 1959). Three capsules are often placed in the same bomb so that P-T-t conditions are identical for the set. The Northrop & Clayton (1966) method (see section 46.4.1) can then be applied to derive the best estimated equilibrium fractionation factor and degree of exchange in incomplete exchange runs. Quite large noble metal capsules have been used (e.g., 5 mm diameter and 40 - 70 mm long by Clayton et al., 1972, and 8 mm diameter and 100 mm long by Czamanske & Rye, 1974). Experimental configurations using big capsules may have large temperature variations along their length (~ 50~ for Clayton et al. system) or apparently relatively small estimated variations (3 - 5~ for Czamanske & Rye, 1974). Note that the presence of a vertical temperature gradient was part of the experimental design by Czamanske & Rye to aid dissolution of reactant sulphides (galena and sphalerite) at the hotter lower end of the tube and, after transport of metals and sulphur in solution, co-precipitation of sulphides at the cooler upper end. Although the temperature may be monitored to + 5~ or better, the actual capsule temperature may not be known to better than + 25~ (Table 46.1). The heating-up times to the chosen temperature may be tens of minutes to an hour and the temperature should be approached from below. Quench rates to 100~ or below are highly variable but can be as long as 10 minutes or more for large conventional coldseal bombs; the rapid quench models, which can quench within a minute or so, have rarely been used for isotopic studies. Some examples are (1) for oxygen in mineralwater systems: Clayton (1959), O'Neil & Taylor (1967, 1969), O'Neil et al. (1969), Clay-

1026


ton et al. (1972), (2) for oxygen and carbon in carbonates: Northrop & Clayton (1966), Sheppard & Schwarcz (1970), (3) for sulphur in sulphides: Czamanski & Rye (1974), and (4) for hydrogen in mineral-water systems: Suzuoki & Epstein (1976), Graham et al. (1980, 1984).

46.7.2.2 Hydrothermal rocking bombs These autoclaves were developed by Barnes (1963) and Dickson et al. (1963), with modifications by Seyfried et al. (1987), to study aqueous sulphur systems, hydrothermal alteration of rocks, mineral solubilities, speciation and kinetic studies where substantial sample volumes were necessary. The working volume is very large by other autoclave standards. The rocking action increases reaction rates. Parameters such as pH can be varied. Note that aliquots of the fluid phase (e.g., 10 ml of solution), liquid and/or gas phase, can be sampled under working conditions. Although the temperature quenching of such a fluid takes place in a few seconds, the fluid may be such a highly reactive, multicomponent solution that chemical separation must be carried out immediately on sampling (e.g., removing sulphide from the solution to reduce back reaction between H2S and SO4 2- or loss of S2- as H2S which takes a few minutes; Sakai & Dickson, 1978). Any solids in the run can only be sampled after quenching the autoclave which can take an hour or so. This is an important technique for sulphur isotope studies. Some examples for sulphur are Robinson (1973), Sakai & Dickson (1978), and see Ohmoto & Lasaga (1982) for evaluation of the data, and for oxygen, Chiba & Sakai (1985). For sulphur isotopes in the dissolved sulphide-sulphate system, fractionation and rate data as a function of pH were obtained by Igumnov (1976) and Igumnov et al. (1977) using large volume, but non-rocking, titanium alloy autoclaves. 46.7.3 Internally heated vessels Internally heated vessels have the furnace within the pressure device. Their main advantage is that they can operate up to both higher pressures and temperatures than externally heated autoclaves (Table 46.1). The experimental data are principally for H, C and O. There are two principal types of apparati used by experimental geochemists: gas vessels and piston-cylinders. 46.7.3.1 Internally heated gas vessel Such a vessel is relatively complicated because both the furnace and thermocouples are inside the thick-walled pressurized vessel whose outer walls are water-cooled to retain its strength. The internal working volume at constant temperature is relatively large as the furnace is usually double or triple so that internal tempertaure gradients can be reduced. The sample, sealed in a metal capsule, can be as large as 150 250 mg. Gases such as CO2 are introduced as oxalic acid (anhydrous, dihydrate) or silver oxalate (Holloway & Wood, 1988) which decompose to CO2, CO and H20. Because of redox reactions in the charge and diffusion of hydrogen through the Pt capsule walls, the CO is largely converted to COa. The pressure medium is argon up to about 12 kbar (argon solidifies at about 13 kbar or 1.3 GPa), or nitrogen. Because the volume of argon or nitrogen under high pressure and high temperature can be substantial and the compressibility of argon or nitrogen is relatively high, an operating vessel is literally a big bomb, so extremely strict safety precautions must be in force


1027

during the operation of such autoclaves. Large vessels can take several hours to heat up, but smaller versions need only 15 minutes or so. Quench time is often several minutes (Table 46.1) but very rapid quench versions exist (~ 100~ Such devices have been used for mineral-water H- and O-isotope systems (Graham et al., 1980; Karlsson & Clayton, 1990), and albite melt-H20, CO2-basaltic and H20-basaltic magma fractionations (Javoy et al., 1978; Richet et al., 1986; Pineau et al., 1998).

46.7.3.2 Piston-cylinder apparatus A solid medium piston-cylinder apparatus consists of a piston that presses into a cylinder containing the solid materials, capsule and furnace (often graphite) assembly (Boyd & England, 1960). Filling the small capsule with microquantities of substances, including fluids, followed by sealing (arc welding, etc.) requires a certain expertise. Pressure is transmitted to the charge by the solid material(s) such as salt, talc or barium carbonate and it may not be perfectly uniform. Considerable care has to be exercised to derive the accurate value of the pressure from the geometry of the apparatus and corrections made for friction and other effects. Although the size of the capsule is small (Table 46.1), temperature gradients can be important. Corrections have to be made for the effect of P on thermocouples. Since operation up to 30 kbar and 1700~ is relatively straightforward, and Clayton et al. (1975) have demonstrated that increasing pressure increases the rate of O-isotope exchange and recrystallization without modifying the equilibrium fractionation factor, piston-cylinder apparati are now often used for stable isotope studies. Because the capsule volumes are small, two phase mixtures are usually chosen to contain roughly equal proportions of the element of interest and all end products are analyzed isotopically. There may only be sufficient material for a single analysis of each phase with the volatile phase being analyzed by puncturing the capsule with a steel needle under vacuum. Despite the short length of the capsule (few mm), CO2 is added either by freezing in a known mass of CO2 or as oxalic acid dihydrate (giving CO2 and H20 at low fH2 or silver oxalate (giving CO2 only). Since the sample material is contained in a sealed Au or Pt tube that is permeable to hydrogen under the run conditions, the nature of the packing materials around the capsule can influence the end products. Rosenbaum & Slagel (1995) have examined this problem in detail emphasizing the importance of eliminating all traces of water and creating an environment with low hydrogen activity around the capsule to improve experimental reproducibility. Some examples are: (1) for oxygen in mineralH20 systems: Matsuhisa et al. (1978, 1979), Matthews et al. (1983b), (2) for mineralcalcite or BaCO3: Clayton et al. (1989), Chiba et al. (1989), Rosenbaum et al. (1994), Rosenbaum & Mattey (1995), (3) for carbon and oxygen between CO2 and calcite: Chacko et al. (1991), Rosenbaum (1994), (4) for carbon between CO2 and basaltic melt: Mattey (1991), and (5) for hydrogen in mineral-H20 systems" Horita et al. (1999).

46.8 Field experiments Field experiments refer to those conducted under natural or field environmental conditions that are more or less well defined. Such experiments have and will continue to yield crucial fractionation data that often cannot be obtained during laboratory experiments. Although these experiments are not ideal, comparison of such empirical calibrations of fractionation factors with laboratory calibrations of equilib-

1028


rium fractionations can, for example, give much insight into certain biological processes. The limitations of this approach are usually rather different from those associated with laboratory experiments and they can be difficult to access. The major weakness is the general lack of unambiguous criteria for equilibrium. A number of field calibrations of fractionation factors give such highly systematic temperature dependent fractionations that they have been interpreted as equilibrium values (see section 46.8.2). It is also possible that some so-called non-equilibrium fractionation factors represent equilibrium for a reaction that has not been clearly identified. For example, certain biogenic carbonate fractionations only consider the fractionation between the biogenic mineral and open sea or fresh water rather than the body fluids that were actively involved in the precipitation process. If the pH of the body fluids is not identical to that of the external water reservoir, then different fractionations may be found because carbonate-water or-DIC fractionations are pH sensitive (speciation is pH sensitive). Important aspects of the role of changing speciation in the solution on fractionation factors for elements such as B, C, O, and S are discussed in Fry et al. (1986), Gessler & v. Gehlen (1986), Usdowski et al. (1991), Hemming & Hanson (1992), Usdowski & Hoefs (1993), Zeebe (1999) and O'Neil et al. (2003). Table 46.6 summarizes a number of experimental field calibrations of isotopic fractionations for H-, C-, N-, O-, and S-isotopes as a function of environmental parameters as well as a few examples for Li, B, Si, C1 and Ca. A large number of systems have been selected to emphasize the wide variety of field experiments that have been carried out. Four principal types of systems dominate Table 46.6: (1) biominerals and plants, (2) weathering environments, (3) hydrothermal and magmatic systems, and (4) between major earth reservoirs. Such a table can be neither exhaustive nor comment on the quality of the results. Some of the temperatures are not known with any precision (e.g., surface temperatures for weathering reactions could cover the range 0~ to 50~ or so). Alternatively, the temperature today, measured with precision, in a geothermal system is not necessarily the same as that during the formation of the mineral. Such things must be borne in mind when using Table 46.6. Also, the remarks column can only include one or two key words about the system. Four specific systems are discussed in a little more detail.

46.8.1 Carbonate-water systems One of the very first field calibrations was of the carbonate (CaCO3)-water system by Epstein et al. (1951, 1953) where the O-isotope fractionation was directly measured between a living mollusk and its environmental seawater at a measured temperature or small range of temperatures. Importantly, at least two series of experiments were conducted in controlled temperature baths of seawater by analysing the regenerated carbonate infilling a notch or drill hole that had been made in the shell at the beginning of the experiment (i.e., Table 46.2 type laboratory experiments). The results were identical and also agreed satisfactorily with one of the relationships for inorganically precipitated CaCO3 determined experimentally in the same laboratory by McCrea (1950). The similarity of the organic and inorganic temperature scales is highly suggestive, but not proof, that equilibrium was attained. Although this interpretation is supported by subsequent experimental studies (O'Neil et al., 1969; Kim & O'Neil,

System ~

Element

Temperature

Remarks

References ~

~~~~~~

_

_

_

_

A. Direct temperature measurement

1. Aragonite-H20DIC

2-22°C

Living and modern corals, forams, mollusks.

Grossman & Ku (1986), Patterson et al. (1993), Leder et al. (1996), Swart et al. (1996), McConnaughey et al. (1997), Bohm et al. (2000)

2. Calcite-HzO

7-30°C

Marine mollusks and associated seawater (sampling strategy, metabolic effects, salinity, growth rate).

Epstein et al. (1951, 1953); Klein et al. (1996)

3. Carbonate-sea water

8-11°C

Living foraminifera and coccolithophorids and sea water.

De La Rocha & DePaolo (2000)

4. Epidote-H20

90-320°C

Geothermal wells: Reykjanes, Salton Sea, Wairakei.

Graham & Sheppard (1980), Bird et al. (1988)

5. Kaolinite-H20

80°C 65-280°C

Geothermal mudpot, Yellowstone Geothermal wells: Valles Caldera; 6D water inferred.

Sheppard et al. (1969), Lambert & Epstein (1980), Gilg & Sheppard (1996)

6. Magnetite-H20

9°C Chiton "teeth" and associated seawater 112". 175°C Geothermal well: Wairakei.

O'Neil & Clayton (1964) Blattner et al. (1983)

7. Opaline-H20

11-18°C

Opaline phytoliths from stems and leaves of wheat relative to stem and leaf water.

Shahack-Gross et al. (1996), see also Webb & Longstaffe (2000; 2002)

8. Phosphate-H20

3-40°C

Biogenic shell, bone or tooth apatite and associated or inferred water; T measured or calculated from carbonate using Epstein et al. (1953)equation.

Longinelli & Nuti (1973a), Kolodny et al. (1983); D'Angela & Longinelli (1990), Lecuyer et al. (1996a);Iacumin et al. (1996) 1029



Table 46.6 - Some experimental field calibrations of isotopic fractionations

1030

Table 46.6 continued > 9. Phosphatecarbonate

0

37°C

Biogenic apatite (bone, teeth) of modern mammals. Carbonate of phosphate.

Iacumin et al. (1996), Bryant et al. (1996)

10. Quartz-adulariacalcite

0 0

275 k2O"C 275 +20"C

Geothermal well: Broadlands. Geothermal well: Broadlands.

Blattner (1975) Blattner (1975)

11. Quartz-illite

0

160-270°C

Geothermal well: Broadlands.

Eslinger & Savin (1973)

12. Silica-H20

0

-1.5-29°C

Living marine and freshwater diatoms and sponges, and associated waters.

Labeyrie (1974), Knauth & Epstein (1975); Schmidt et al. (1997), Shemesh et al. (1992a); Clayton (1992), Shemesh (1992)

13. Sphalerite-galena

S

273°C

Geothermal well: Broadlands; fluid inclusion T.

Czamanske & Rye (1974)

14. C02-CH4

C

20-600°C

Geothermal systems: kinetic and equilibrium fractionations.

Sheppard (1981), Giggenbach (1982,1997)

15. Plant-CO2

C

surface

Photosynthesis: C3 and C4 plants.

Park & Epstein (1960), Smith & Epstein (1971)

16. Cellulose-H20

H, 0

15-30°C

Plant and animal cellulose: terrestrial, marine and freshwater.

DeNiro & Epstein (l979,1981a), Yapp & Epstein (1982), White et al. (1994) Tang et al. (2000), Aucour et al. (2002)

17. Calcite-graphite

C

600-800°C 400-680°C

Kfeld-plagio and magnetite-ilmenite. Calcite-dolomite solvus thermometer.

Valley & O'Neil (1981) Wada & Suzuki (1983)

18. Dolomite-calcite

C, 0

200-650°C

Mg-calcite-dolomite solvus thermometer.

Sheppard & Schwarcz (1970) Table 46.6 continued >

Chapter 4 6 - S.M.F. Sheppard

B. Chemical and isotopic thermometers

Experimental

Table 46.6 continued > ~~

~~

~

~~~~~~

~~

~

~

~

~

_

_

_ ~~

_ ~~

_

~~~~~~~~~~~

~

200-300°C

Internal mineral fractionation. [OH] oxygen of illite measured by vacuum dehydration and/or low-temperature fluorination.

Bechtel & Hoernes (1990)

20. Silicates and magnetite

0

500-800°C

Various selected isotopic thermometers and coexisting mineral systematics.

Bottinga & Javoy (1973,1975)

21. Pyroxene-olivine

0

855-1300°C

Various chemical thermometers.

Kyser et al., (1981), Gregory & Taylor (1986)

22. Calcite-diopside

0

800°C

Diffusion rates of oxygen in diopside estimated using both infinite and finite reservoir model, mineral modal abundance and 8180 variations.

Sharp & Jenkin (1994)

23. Quartz-calcite

0

80-730°C

Various chemical, mineralogic and isotopic geothermometers.

Sharp & Kirschner ( 3 994)

24. Serpentinemagnetite

0

50-500°C

Quartz-muscovite 1 8 0 thermometer, chlorite-Fe-Ti oxide fractionations.

Wenner and Taylor (1971)

25. Tourmaline-H20quartz

H

200-600°C

Quartz, muscovite, illite, chlorite, biotite isotopic thermometers.

Kotzer et al. (1993)

26. Zircon-titanite Quartz-titanite

0

650°C

Various experimental and empirical isotopic thermometers. Comparison with semi-empirical increment method. Closure temperature taken into consideration.

King et al. (2001)

surface

Farm, zoo and wild eggs. Fractionation between calcite, total organic C or individual acids (9) and diet, local water, climate.

Johnson et al. (1998)

0

o f I s o t o p i c F r a c t i o n a t i o n F a c t o r s ...

0

Measurement

19. Illite (whole)illite [OH]

C. Estimated or inferred temperatures 27. Calcite (Ostrich eggshell)-diet

C, N, 0

1031


1032

Table 46.6 continued > seawater surface

H- and 0-isotope systematics in cherts. T calibration from oceanic chert.

Knauth & Epstein (1975,1976), Kolodny & Epstein (1976)

29. Gibbsite-Hz0

surface

Weathering zones, bauxites.

Lawrence & Taylor (1971), Bernard et al. (1976), Chen et al. (1988, 1990), Bird et al. (1989, 1990), Vitali et al. (2001)

30. Goethite-H20

surface

Weathering and supergene zones, isotopic comp. water inferred.

Yapp & Pedley (1985),Yapp (1987,1990,2000), Bao et al. (2000a)

31. Kaolinite-H20

surface

Weathering zones.

Savin & Epstein (1970), Gilg & Sheppard (1996), Sheppard & Gilg (1996)

32. Opaline-seawater

seawater

Fractionation with dissolved silicic acid.

De La Rocha et al. (2000)

33. Smectite-H20

surface surface

Weathering and oceanic clays. Weathering; function of octahedral Fe3+/(Fe3f + Mg + Al). Interlayer water-water.

Savin & Epstein (1970),Lawrence & Taylor (1971), Gilg & Sheppard (1995), Sheppard & Gilg (1996) France-Lanord & Sheppard (1992)

surface

Hydrated obsidians and perlites; H20 >-2wt.%.

Friedman & Smith (1958), Taylor (1968), Cerling et al. (1985),Gilg & Sheppard (1999)

35. Zeolite-Hz0

surface

Analcime, natrolite: channel water (chabazite, clinoptilolite, laumontite: rapid exchange with ambient water vapour).

Karlsson & Clayton (1990)

36. Bone collagen and phosphate-water

surface

Deer bone from across N. America related to environment (meteoric water, humidity).

Cormie et al. (1994)

37. Plant leafwater transpiration

surface

Leaves of astomatal plants sampled under measured air T and humidity conditions.

Cooper et al. (1991)

seafloor 34. Volcanic glass-

Hz0



28. Chert-water

38. C3 plantsenvironment

C

surface

Spatial, species and temporal variations in leaves, needles and wood.

Heaton & Crossley (1998), Heaton (1999)

39. Tree ringenvironment

H, c , 0

surface

Isotopic variations of cellulose nitrate of tree rings with climatic variables (temperature, humidity, precipitation, topography, soil, within site and between species,. . .).

Edwars & Fritz (1986), Feng & Epstein (1995a, 1995b), White et al. (1994), Tang et al. (2000), Aucour et al. (2002).

40. Nitrification

N

surface

Fractionation during nitrification and algal assimilation of NH4+ estimated from measurement of dissolved NH4+ and N03in river water.

Cifuentes et al. (1989), see also Mariotti et al. (1981)

41. Bacterial sulphate S reduction in organicsediments

surface

Fractionation during sulphate reduction in cyano-bacterial microbial mats and sediments at known rates of sulfate reduction. Part of experiment carried out in laboratory

Habicht & Canfield (1997)

42. Microbial sulphate reduction

0, S

sea floor

S- and 0-isotope fractionations at oil and gas seeps. Rates of sulphate reduction estimated. Seep sediments compared with reference sediments.

Aharon & Fu (2000)

43. Magma (granitic)-

H

700°C

Rayleigh distillation model with several assumptions.

Nabelek et al., (1983)

44. Magma (rhyo1itic)H20

H

magmatic

From obsidians in tephra flows, Rayleigh model with assumptions.

Taylor et al. (1983)

45. Altered oceanic crust-seawater

Li

crustal

Chan et al. (1992) Fractionation estimated between altered basalt and seawater from 66Li measurements of fresh and altered oceanic basalts.

H20

1033




1034

Table 46.6 continued > 46. Continental crustseawater

B

crustal

Model fractionation derived from 611B values of continental tourmalines with time.

Chaussidon & AlbarPde (1992)

47. Evaporite mineralbrine

B

surface

Fractionations associated with precipitation of borates and adsorption onto clays in natural evaporitic environments.

Xiao et al. (1992b), Vengosh et al. (1995)

48. Diffusion coefficient in water

C1

surface

Fractionation associated with diffusion of C1-derived from S37C1 in groundwatersaline water mixing zone, in absence of advection.

Desaulniers et al. (1986)

49. Surface-mantle reservoirs

c1

crust to mantle

Fractionation (37C1/35C1) estimated from fresh MORB glass, high and low T altered oceanic crust.

Magenheim et al. (1995), see also Eggenkamp & Koster van Groos (1997)

Chapter 4 6 - S.M.F. Sheppard


1035

1997), it is well documented that the equation does not apply to all low temperature carbonate-water systems. Certain biogenic and inorganic carbonates are not precipitated in equilibrium with their associated fresh or saline waters (see sections 46.6.4.2 and 46.8) and the nature of the biogenic carbonate (aragonite, low- or high-Mg calcite) has to be considered. 46.8.2 Phosphate-water systems The phosphate-water fractionation expression was initially derived by Longinelli and his coworkers (Longinelli, 1965, 1966; Longinelli & Nuti, 1973a) from the analysis of biologically precipitated (biogenic) invertebrate phosphate and its associated biogenic calcium carbonate (high- and low-Mg calcite, aragonite) from the same shell, plus application of the Epstein et al. (1953) equation for calcium carbonate-water to determine the average growth temperature. Subsequently, Kolodny et al. (1983) measured directly the biogenic fish phosphate-water fractionation, in both sea and fresh waters, at measured temperatures. They also demonstrated that the isotopic composition of fish bone phosphate is only influenced by the isotopic composition and temperature of the water, and is not inherited from either the food or dissolved phosphate oxygen. L6cuyer et al. (1996a) combined measured isotopic compositions of living inarticulate brachiopods (lingulides) with inferred isotopic composition, salinity and temperature of seawater data to confirm that a single equation describes the phosphate-water fractionation of many marine vertebrate and invertebrate organisms, except mammals. All these temperature dependent empirical fractionations have usually been considered to represent equilibrium rather than kinetic effects, although direct evidence for equilibrium was lacking. Recent laboratory experimental determination of the equilibrium isotopic fractionation between the dissolved inorganic orthophosphate ion, H2PO4-, and water by L6cuyer et al. (1999) indicates that biogenic apatites are systematically 1sO-depleted (~ 8%o at 20~ However, this ~ 8%o difference represents the equilibrium fractionation between phosphate solutions with seawater type pHs, where H P O 4 2- is the dominant ion, and lower pH solutions, where H2PO4- is the dominant ion (O'Neil et al., 2003). This example emphasizes the importance of speciation on mineral-solution fractionations. 46.8.3 Plants Isotopic compositions and their variations of plants, their organs (e.g., leaf, wood, roots) or constituents (e.g., cuticles, phytoliths,...) are increasingly being used as proxies for environmental conditions in palaeoclimatic studies. Although a certain number of controlled laboratory experiments have been conducted (see section 46.7.1.1 and Table 46.2), field experiments have also played their part (N ~ 7, 15, 16, 36-39, Table 46.6). In particular, such studies have demonstrated that variations substantially larger than the analytical precision commonly occur in the same substance from different parts of the same plant as well as between plants of the same species from the same site, not to mention differences between different species from the same site (for references see above examples from Table 46.6). Because the isotopic composition is usually a function of several environmental variables- temperature, soil water, humidity, insolation, etc...- unscrambling the isotopic signal(s) in terms of a particular parameter becomes quite complex. Additionally, because the analytical techniques are

1036


often quite labour intensive, isotopic analyses of, for example tree rings, are rarely taken from more than a single tree per site (Tang et al., 2000; Aucour et al., 2002). Moreover, an isotopic calibration carried out on a currently living species may not be applicable, in detail, to fossil material that is similar, but not identical, to the so-called calibrated reference living material.

46.8.4 Between reservoir fractionations Examples N ~ 1 to 38 (Table 46.6) measured fractionations between substances minerals, plants, magma, liquid or gaseous species. In contrast to these more classical examples, N ~ 3 and 44 to 48 are for the more recently studied isotopic ratios of Li, B, Si, C1 and Ca. The empirical experiments were principally designed to measure fractionations among major earth reservoirs - seawater, oceanic crust, mantle, etc...- in order to understand the large-scale geochemical cycle of these elements. 46.9 Conclusions

Stable isotope geochemistry has come a long way over the past 50 years or so. The results of carefully designed laboratory and field experimental isotopic studies, many of which have been extremely time-consuming, have contributed fundamentally to this progress. On quick inspection, many experimental laboratory studies seem to be quite straightforward, but, as partially summarized in Tables 46.1 to 46.5, much time, ingenuity and a wealth of ideas have been invested behind the results. The development of new analytical techniques has played its role, of which the ion microprobe and multiple collector inductively coupled plasma source mass spectrometer are two examples. Future laboratory and experimental field studies will be equally challenging. Interactions among the three approaches to characterize isotopic fractionation fact o r s - experimental, empirical and theoretical- are fundamental to improving our understanding of isotope effects, because each method suffers from certain difficulties and limitations. Laboratory experiments, however, hold a key position in this trinity because they can be repeated or redesigned to increase their credibility in a way that is often less evident for the others. It is recalled that the theoretical modelling of isotopic fractionations requires experimental spectroscopic data as basic input. Surprisingly few minerals, however, have been synthesised experimentally with strong enrichment in the usually minor occurring isotope(s) so that the effects of isotopic substitution on the vibrational spectra (infrared, Raman,...) can be measured (e.g., Sato & McMillan, 1987; Gillet et al., 1996; B6ttcher et al., 1997c) rather than calculated using some model. Comparison of experimental fractionation data with theoretical models is essential to assess whether a harmonic approximation is satisfactory or not. Although emphasis has been placed on how isotopic fractionation factors have been measured experimentally both in the laboratory and in the field, much detailed information and understanding have also been derived on mechanisms and rates of reactions. These have many applications. In fact there is a chicken and egg situation because understanding the mechanism of the reaction can aid both the design of the experiment and the interpretation of the results.


Acknowledgments

1037

Since my introduction to experimental isotope geochemistry at McMaster University, Canada with H. P. Schwarcz, D. M. Shaw and H. G. Thode, I have particularly benefited from my collaborations and/or interactions with S. Epstein, C. M. Graham, J. J. Hemley, M. Pichavant and H. R Taylor, Jr., all of whom are gratefully acknowledged. This chapter has benefited from the helpful comments of A. Matthews and J. W. Valley.

Handbook of Stable Isotope AnalyticalTechniques, Volume 1 P.A. de Groot (Editor) 9 2004 Elsevier B.V. All fights reserved.

CHAPTER 47 Laboratory Set-Up for GC-MS and Continuous-Flow IRMS Wolfram Meier-Augenstein Queen's University Belfast, Environmental Engineering Research Centre, David Keir Building, Belfast, BT9 5AG, UK e-mail: [email protected]

47.1 The Ideal MS / IRMS Laboratory It is probably fair to say that every scientist dreams about access to unlimited resources to design and build the ideal laboratory for their work. It seems to be equally true that those scientists involved in mass spectrometry (MS) and isotope ratio mass spectrometry (IRMS) are more likely than others to have their wish granted. In reality though, this is mainly due to administrative considerations concerning the protection of a major cash investment in the instruments rather than assuring high quality analytical work. However, if you should find yourself in the fortunate position of being asked to submit a wish list for your IRMS laboratory, take advantage of this situation as best as you can. The performance of your instruments and the quality of the results they produce not only depend on the amount of time and effort you put into sampling, sample preparation, sample separation and instrument maintenance but also on the quality of the environment your instruments operate in. In the following paragraphs, consideration is given to all the points that affect instrument installation and its performance. This listing is the result of foresight and hindsight in conjunction with experiences made or rather suffered when setting up two MS / IRMS facilities at two universities. Since individual circumstances and precise requirements vary, a broad and general outline of bare necessities is given that can easily be adapted to and incorporated into the situation at hand. 47.2 Location While planning a new or completely refurbished laboratory, usually the first step is to decide where this new facility should be located. The choice of locale should be dictated by four considerations" (1) Access; (2) Floor space; (3) Floor stability and (4) Climate. Of course, there are other things to consider as well such as noise level, the lab's location within the exisiting power grid (section 47.4), gas supply (section 47.5) and connections to telephone and IT services. These and similar points are secondary to identifying the premises that will accomodate the future MS / IRMS laboratory and

Laboratory Set-Up for GC-MS and Continuous-Flow IRMS

1039

will be discussed separately. The first three points almost always lead to a choice of ground floor or basement level accommodation. The new lab should have level access or access via a gently sloped ramp. The doors must be wide enough to manoeuvre the instrument(s) through with ease. The floor must be rock solid, i.e. free of vibrations, and the floor space generous enough to allow easy access to the instrument or instruments from all sides, e.g. instrument floor space plus 1 metre on each side. In terms of overall floor space, one must also consider the spatial requirements of ancillary instruments, future equipment, distance of computers and monitors from the IRMS magnets and the installation of a fume cupboard. In line with local safety regulations, think about space required for gas cylinder storage (see section 47.5). Lastly, the largest outer wall area should face North (Northern Hemisphere locations) or South (Southern Hemisphere locations) since it's easier and cheaper to keep a place warm than transport excess heat out.

47.3 Temperature Speaking of heat, all your analytical instruments will benefit from an environment with constant ambient temperature, preferably around 23 ~ with a temperature stability of + 1 ~ i.e. drift in temperature should not exceed I ~ per hour. So, air conditioning is a must on the check-list for the new laboratory. If money is no object, one should chose an air conditioning system that provides humidity control as well. The size and capacity of the air conditioner depends on the number of instruments in your laboratory and their respective heat output at peak. This includes GCs, vacuum pumps, drying ovens, personnel and spare capacity for future equipment.

47.4 Power Supply When planning your new laboratory, make sure it will have its own dedicated power supply line from the nearest local sub-station. In other words, insist that your laboratory will not be the last in line of an already several times over extended campus grid serving several other major consumers. All analytical instruments are sensitive to fluctuation in the power supply and IRMS instruments are no exception. Apart from maintaining a constant voltage, your power supply should be free of surges, spikes and brown-outs. If that cannot be guaranteed, make sure one or more UPS units (uninterrupted power supply) are installed and connected between wall socket and instrument. An UPS will enable you to safe your data and shut down your IRMS in a controlled fashion should there be a power cut. In addition, modern UPS units act as surge filters and provide a constant voltage output even if your mains power supply does not. An UPS unit providing up to 2 hours backup supply is probably the best compromise in terms of expense and space requirement in a situation where you are not served by an emergency power generator. Should your mains power supply be backed up by a emergency power generator, a smaller UPS unit will suffice, e.g. 15 to 30 minutes power backup, but is still essential since its short reaction time will bridge the time before the emergency power generator kicks in.

1040

Chapter 47 - W. M e i e r - A u g e n s t e i n

Another point worth mentioning is the number of sockets in the new laboratory. In a variation of Murphy's Law there never seem to be enough sockets in a lab and if there are, they are never where you need them. This inevitably leads to the intricate and intertwined system of tripwires otherwise known as extension leads and extension sockets. As a general rule, for an IRMS lab you will require at least one dedicated three-phase, five-connector power line that must be fused separately. In addition, each wall should contain one dedicated, separately fused circuit feeding four blocks of four sockets each, two at floor level and two at bench-top level. In addition, consider the possibility of one or two socket cubes suspended from the ceiling in the centre of the lab.

47.5 Gas Supply In the context of an IRMS (and / or MS) facility, gas supply means the supply of gases required to run your instruments, i.e. gases such as Helium, Oxygen, Hydrogen, Nitrogen, Synthetic Air, Carbon Dioxide and Carbon Monoxide as well as a supply of oil-free compressed air. The overriding rule for gas supply in instrumental analysis in general and IRMS in particular is "the cleaner the better". This maxim can only be achieved if every step in the gas train is set up in such a way as to exclude leaks, atmospheric break-in and contamination. To this end, all materials used for gas supply should be made of stainless steel. This includes pressure regulators that should contain a stainless steel diaphragm, tubing, connectors, valves and ferrules. Regulator diaphragms made from other materials will deteriorate over time and start leaking. Please note, using stainless steel tubing and then connecting it with brass unions and brass ferrules is false economy. Brass ferrules are softer than stainless steel and will therefore not cut into the stainless steel tubing and, hence, not afford a leak tight seal. By the way, the stainless steel tubing must be clean on the inside, i.e. free of any contaminants associated with its manufacture. Gases should be of highest purity, i.e. 99.998% (= 4.8 in European notation) or 99.999% (= 5.0). Ideally, the carrier gas should be of even higher quality, i.e. 99.9999% (= 6.0). However, Helium of this quality is rather expensive and the same quality can be achieved by using 4.8 or 5.0 Helium in conjunction with a high-capacity gas purifier. The best gas purifier on the market is Supelco's thermo-chemical absorption system, the "High Capacity Gas Purifier" that consists of a purifier tube and a small oven. To monitor the performance of this system as well as providing a back-up once the purifier tube is exhausted, it is recommended to install an additional self-indicating purifier tube (e.g. Supelco's OMI-2) in line behind the thermo-chemical trap. The two major reasons for the high purity demand on the carrier gas He are the capillary columns in the GC and the IRMS itself. Stationary phases of medium to high polarity are extremely sensitive to oxygen being present in the carrier gas, which leads to column deterioration and excessive column bleed. To achieve the best possible accuracy and precision of IRMS measurements, the carrier gas must also be free of organic contaminants and moisture.

Laboratory Set-Up for GC-MS and Continuous-Flow IRMS

1041

With the possible exception of the smaller reference gas cylinders, ideally all the gases should be stored outside the lab in two safety cabinets, accommodating four cylinders each. A typical set-up would be N2 and H2 in one cabinet, and He, 02 and synthetic air in the other, thus keeping combustable and flamable gases separate. To safeguard against running out of He when one can afford it least, two He cylinders should be connected simultaneously to a change-over regulator with a pressure dependent toggle switch. Unfortunately, these change-over regulators come with a high price tag and, hence, in most cases must be regarded as a luxury albeit a desirable one. From the central gas storage point, 1/4" stainless steel tubing should be run into the lab and along the walls at head height. From these 1/4" supply lines 1/8" stainless steel tubing tee off at or near the instruments. A pressure gauge and a shut-off valve should be placed between the T-union and the instrument supply tubing. Prior to its first use, such a gas supply line set-up should be flushed with clean, dry nitrogen to remove air and moisture. Similarly, compressed air should be supplied from an external compressor and brought into the lab via a dedicated line that incorporates an oil mist filter. Also, consider a natural gas supply line for your Bunsen burners or, alternatively, use gas cartridge operated burners. Finally, an exhaust line for the vacuum pumps should be installed, either plumbed into the exhaust of the fume cupboard or leading directly outside. A brief note on gas safety. Even with gas cylinders stored outside the laboratory, it is a good idea to have a H2 monitor with alarm in the laboratory. In addition, due to the increased interest in on-line IRMS measurement of 1 8 0 / 1 6 0 isotope ratios of organic compounds, consideration should be given to the installation of a CO monitor (with alarm), because 5180 measurement by on-line thermal conversion (pyrolysis) of organic compounds requires the use of CO as reference gas. For this application provision must be also made for an extraction line positioned at the conversion interface of the GC-IRMS, either branched into a fume hood or connected to a separate extractor fan, to remove CO emanating from the reference gas inlet and potential leaks in the interface.

47.6 Finishing Touches When choosing the floor cover, practicality should overrule appearances. A carpetlike floor cover might look nice but is prone to induce static electricity discharges and is difficult to clean. A chemically inert surface is better suited, especially if the lab floor has been fitted with a drainage point at the lowest point of the floor. Make provisions for ample work surface / bench-top area with underneath storage facilities. Ideal are mobile cupboard, drawer and shelf units on castors with a small work surface of their own. Ask for a solvent storage cabinet that is earthed and connected to the exhaust of your fume cupboard.

1042


Good, i.e. sufficient and evenly spaced lighting is also important. Soft fluorescent light tubes should be suspended from the ceiling at a height of approx. 2.5 metres. Equally important are telephone points and sockets to connect your computer to your local server and / or local intranet. From a practical point of view, a sink with running hot and cold water is desirable as long as it is not directly adjacent to the mass spectrometer. Similarly~ a supply line of de-ionised water is a valuable thing to have. Mass spectrometer laboratories are notoriously noisy, chiefly due to the high vacuum systems. Accommodating the rotary pumps in ventilated yet noise insulated cupboards will go a long way to reduce the noise level in the lab. However, this still leaves to high-pitched noise of the turbo-molecular pumps. Again, if money is no object a sound engineer should be consulted to suggest appropriate sound proofing measures tailor-made to the situation at hand. Lastly, if you want to see all your planning and requests come to fruition, do not sit back and wait for the hand-over day of your new lab. Be involved and get involved during the entire process of your new lab being designed, built and fitted. Be prepared for specifications not being met, the drainage point not being at the lowest point of the floor and doors not being wide enough despite clear instructions and specifications. By the same token, be prepared to discover that the architect, the engineers, the builders, plumbers, electricians etc. will all work on the assumption that you, the MS / IRMS specialist and end-user of this lab will have no idea of what you are talking about and what you really need. For that reason they will make changes to plans, dimensions and specifications without telling you about it. So, get involved armed with a tape measure, a calibre, a volt meter (or even better a voltage monitor) and the pre-installation requirement booklet of your IRMS and fight your corner.

Isogeochem List

1043

APPENDIX A I s o g e o c h e m List ISOGEOCHEM is an e-mail discussion list in stable isotope geochemistry. The list currently counts 1900 subscribers but is still increasing in size. As a result, there are active discussions on all different topics involving stable isotopes. The list represents a useful tool for obtaining information on stable isotope analysis and a range of related subjects. The list-owner, Andrea Lini, is acknowledged for his allowance to present details of the list, including procedures to subscribe or unsubscribe, in this book. Although there always is a risk that this information will become invalid because of changes and improvements in the fast developing internet technology, it is considered highly valuable to include current information here for those not yet aware of the list. The objectives of the ISOGEOCHEM list are to promote the exchange of news and information among those with an interest in stable isotope geochemistry, and to provide new contacts and enhance collaboration among researchers from different disciplines (e.g. geology, biology, chemistry). The list is intended not only as a discussion forum for isotope geochemists but also as a source of information and help for researchers from other fields interested in applying stable isotopes as an additional tool in their own studies.

Procedures for (un)subscribing to the ISOGEOCHEM list In order to gain the benefit of this discussion list, you must subscribe to it. Simply send an electronic mail message to this address: [email protected] Leave the subject field blank. Construct a subscribe command as follows and place it as the first line of your message: sub ISOGEOCHEM fname lname where fname is your first name and lname is your last name. To signoff from the list, email to [email protected] with the following request: signoff ISOGEOcHEM or unsubscribe ISOGEOcHEM

1044

Appendix A

If you have any questions regarding the ISOGEOCHEM discussion list, feel free to contact the "list-owner", Dr. Andrea Lini ([email protected]). I S O G E O C H E M and related WEB-Sites ISOGEOCHEM also has a web-site:

http: / / www.isogeochem.com This site provides information about conferences, vacant positions, IRMS manufacturers and suppliers of isotopic materials. It also contains links to numerous stable isotope laboratories worldwide. Archive All postings to the ISOGEOCHEM list are recorded in ISOGEOCHEM's Archive. The archive has a full search engine for quick retrieval of anything ever posted on the list. This archive can be found on WEB-Site:

http" / / list.uvm.edu / archives / isogeochem.html or by connecting to the ISOGEOCHEM WEB-Site given above, and selecting the link to the Archive. Address list and E-mail address list An address list of subscribers to the list can be found on WEB-Site:

http://geology.uvm.edu / address.html This site can also be accessed from the ISOGEOCHEM WEB-Site. Note on the addresslist: adresses of only about 10% of the subscribers is only given; only those who provided a complete address are included in that list.

The Web Stable Isotope Fractionation Calculator

1045

APPENDIX B The Web Stable Isotope Fractionation Calculator Georges Beaudoinl & Pierre Therrien D6partement de G6ologie et de G6nie G6ologique, Universit6 Laval, Qu6bec, Canada GIK 7P4 e-mail: i [email protected]

The fractionation of stable isotopes between two isotopic species is an inverse function of temperature. It is convinient to transform the isotope fractionation (a) into a polynominal inverse function of temperature such as: ( 10181 ( 10121 ( 109/ (~26) 1 )+F 10001nc~ - A T6 ) + B T4 ) + C - ~ j +D +E( 3T

[B.1]

where A, B, C, D, E, and F are variables and T is the temperature in K. The advantage of the 1000 In (Z(a-b)logarithmic transformation is that its numerical value is similar to the difference (Aa-b) in g-values of each isotopic species a and b (Hoefs, 1997) that can be determined experimentally in the laboratory" [B.2]

Aa-b = 6a-~lb

where the h-value is the per mil difference in isotope ratios relative to a standard: 180/16Oi 6i -

180/16OvsMow

-1) 91000

[B.3I

where VSMOW is a standard, here Vienna-Standard Mean Ocean Water. The fractionation factor O~a-bis related to the 5-value by the following relation: 1000 4- ~}a

[B.4]

~ a - b = 1000 + 5 b

Considering the uncertainty of 5-values, which is commonly up to 0.2%0 for 513C, 6180, ~34S, and of 1-5%o for 5D, propagation of errors yields uncertainties on Aa-bvalues of about 0.3%0 and up to 7%o, respectively. These experimental uncertainties are larger than the difference between the Aa-b and 1000 In Or(a-b) for Aa-b values below

1046

Appendix B - G. Beaudoin & P. Therrien

20%o and 6-val- Table B1 - Evaluation of the 1000 In a(a-b) logarithmic transformation. ues below about 6a 6b O~(a-b) Aa-b 1000 In O~(a-b) 20%o for (~13C, 6180, (~34S (Table 0 1.0100 10 9.95 B1) For hydro10 10 1.0099 10 9.85 9 20 gen isotope data, 20 0 1.0200 20 19.80 higher 6Da and 30 10 1.0198 20 19.61 30 1.0194 20 19.23 Aa-b have higher 50 -30 1.0619 60 60.02 numerical values 30 -20 0.9388 -60 -63.18 and uncertainties -80 such that use of the 1000 In O~(a-b)logarithmic transformation remains a good approximation in most situations. A large number of experimentally and theoretically derived 1000 In C~(a-b)equations exist in the literature and they have been compiled periodically (Friedman & O'Neil, 1977; O'Neil, 1986; Kyser, 1987b). Traditional hardcopy compilation become

Figure B1 - Main w i n d o w of the WEB S t a b l e I s o t o p e Fractionation Calculator that permits input of the temperature and selection of the equation of interest.

The Web Stable Isotope Fractionation Calculator

1047

Figure B2 - Result of the computation of the stable isotope fractionation at the temperature and for the equation selected. The coefficients of the equation are given with the reference (s) for the equation. quickly obsolete, however, and may contain typographic errors that are not easily corrected (Savin & Lee, 1988). The Internet offers an alternative application that is widely accessible and which can be conveniently updated to account for new data or to correct errors. The Internet Stable Isotope Fractionation Calculator:

http" / / www.ggl.ulaval.ca / cgi-bin / isotope / generisotope.cgi is based on a regularly updated database listing a large number (currently more than 450) of fractionation equations between isotopic species of hydrogen, carbon, oxygen and sulfur. The user inputs a temperature (in C) and selects an equation between two isotope species (Figure B1). The program returns the 1000 In ~(a-b) value for the input temperature, the values of coefficients A to F used for the computation, and the reference(s) for the equation (Figure B2). It thus offers a simple yet powerful tool to obtain a list of fractionation equations and their coefficients to interpret stable isotope data.

1048

Appendix C

APPENDIX C S u p p l i e r s of R e f e r e n c e Materials

C1 Adresses for reference and stable isotope materials International Atomic Energy Agency (IAEA) Analytical Quality Control Services Agency's Laboratories Seibersdorf P.O.Box 100 A-1400 Vienna Austria Stable isotope reference materials at environmental level: M. Gr6ning Phone: +43-1-2600-21740 Fax: +43-1-26007 E-mail: [email protected] General Phone: Fax: E-mail:

information on all reference materials: Mr. Radecki +43-1-2600-28226 +43-1-2600-28222 [email protected]

Institute for Reference Materials and Measurements Joint Research Centre of the European Commission B-2440 Geel Belgium For any information or for ordering: E-mail: [email protected] Andr6 Verbruggen Coordinator Isotope Reference Materials Phone: +32 14 571 617 Fax: +32 14 591 978 E-mail: [email protected] Information also can be obtained from the IRMM Web site: http"//www.irmm.jrc.be / mrm.html

Suppliers of Reference Materials

1049

National Institute of Standards and Technology (NIST) Standard Reference Materials Program Room 204, Building 202 Gaithersburg, Maryland 20899-0001 USA

Phone: +1-301-975-6776 Fax: +1-301-948-3730 E-mail: [email protected] Orders for NIST must be accompanied by an order number of a selected reference material (see: Coplen, 1996, or contact NIST for order numbers). CEA: distributes stable isotopes through its daughter company" EUROISO-TOP Parc des Algorithmes (Bat. Hombre) F-91194 St. Aubin France http://www.eurisotop.fr / Messer Griesheim (MG) Futingsweg 34 47805 Krefeld Germany Phone: +49-203-6002-388 Fax: +49-203-6002-460

A list of suppliers of stable isotope compounds and stable isotope labelled compounds can be found on the WEB-site: http: / / www.uvm.edu / --geology / geowww / suppliers.html

1050

Appendix C

C2 Indiana Zinc

The Biogeochemical Laboratories at Indiana University use several methods to reduce water to elemental hydrogen for subsequent determination of D / H ratios. For quantitative production of hydrogen from microliter-quantities of water by use of a simple and inexpensive batch process, we recommend the use of zinc metal (Coleman et al., 1982). Pure zinc is unsuitable because it will not react with water quantitatively. Zinc with suitable characteristics had been available from BDH. As BDH zinc is not produced specifically for hydrogen isotopic analysis its characteristics in the application are unreliable. During the 1980's, privately funded research at Indiana University resulted in a proprietary method to prepare zinc turnings for rapid quantitative reduction of water to H2. Our "optimally contaminated" zinc, which we are calling "Indiana Zinc", is used in dozens of laboratories worldwide. It has a long-standing track record of providing reproducible hydrogen stable isotope ratios (for example, see Schimmelmann & DeNiro, 1993). Since 1999, two types of Indiana Zinc are available. The first type is the traditional formula with its long track record of proven reliability for laboratories that can keep a constant zinc" water ratio. A newly developed zinc formula is available upon request that has a reduced hydrogen blank ("amount effect"; see: Dem6ny, 1995) and is especially suited for analyses where the zinc" water ratio cannot be controlled within a narrow range. For more information on how to obtain "Indiana Zinc" and how to convert water to elemental hydrogen by using our zinc, send e-mail to: [email protected], or fax (812) 855-7961, or send letter to: Arndt Schimmelmann, Indiana University, Department of Geological Sciences, Biogeochemical Laboratories, 1001 East Tenth Street, Bloomington, IN 47405-1405, USA. Indiana Zinc is a research material offered for use without guarantees and without acceptance of any responsibilities for damages arising from its use or possible failure in any application. It is distributed as a service to those engaged in stable isotope research. For additional information, see: http" / / www.indiana.edu / ~geosci / research / biogeochem / biogeochem.html

Suppliers of ReferenceMaterials

1051

C3 H and C stable isotope standards for organic compound-specific investigations Compound-specific hydrogen isotope ratios for organic hydrogen are now analytically accessible by a combination of GC, high-temperature pyrolysis to elemental hydrogen, and subsequent on-line irm-MS (Sessions et al., 1999; Scrimgeour et al., 1999; Hilkert et al., 1999; Tobias & Brenna, 1997). Complete irmGCMS-systems with interfaces are commercially available, but there has been no isotopically defined set of organic compounds for routine calibration and D / H quality control. Standards that reliably establish or confirm isotopic calibrations must be in the same form as the unknown analytes. In contrast, the use of intermittent spikes of introduced elemental "standard" hydrogen gas is fraught with potential problems because it does not take into account D / H fractionations that may occur in the analytical train between injection in the GC and the exit of the pyrolysis reactor. The Biogeochemical Laboratories at Indiana University, in collaboration with Woods Hole Oceanographic Institution (A. Sessions, J.M. Hayes), first established the purity of n-alkanes (range from C-12 to C-50) and n-alkanoic acid methyl esters (C-10, C-20, C-30 FAMEs) by GC-MS, followed by the measurement of D / H and 12C/13C ratios for each compound. Up to five replicate analyses for each compound were performed off-line, via conventional combustion of milligram-amounts of individual compounds in quartz ampules and cryogenic purification of combustion gases in a vacuum line. Water was converted to elemental hydrogen in contact with uranium, followed by collection of hydrogen gas using a Toepler pump. Gas yields and elemental H / C ratios were routinely monitored manometrically for quality control. Hydrogen and carbon isotopic ratios were determined using MAT252 mass-spectrometers at Indiana University. The hydrogen isotopic calibration employed the conventional normalization to VSMOW (zero per mil) and SLAP (-428 per mil), according to Coplen (1996). Our standards are typically requested by the research community in form of solutions in hexane that are sealed under argon in glass ampules. We offer individual solutions of n-alkanes (C-12 to C-50) and fatty acid methyl esters (C-10, C-20, C-30). They are useful for co-injection as internal isotopic standards. One particular'mixture B' should be mentioned here because it is designed to test the accuracy of H3 + correction in hydrogen-isotope-IRMS (see Sessions et al., 1999). Mixture B can be used to measure the H3 + factor under conditions closely matching those experienced by analytes. It contains fifteen n-alkanes (C-16 to C-30) containing a 5-fold range of concentrations (arranged in three pentads with rising concentrations), from 20 nmol H2 to 100 nmol H2 per compound per microliter of solution (see Figure C1). Custom mixtures may be available upon request.

1052

Appendix C

Figure C1 - irm-GCMS chromatogram showing ion currents (m/z 2 and 3) of n-alkane mixture "B" (courtesy of Alex Sessions, Woods Hole Oceanographic Institution).

Although intended primarily for D / H research, our standards are equally useful for irmGCMS work on carbon isotope ratios. For more information, please send email to: [email protected], fax (812) 855-7961,or send letter to: Arndt Schimmelmann, Indiana University, Department of Geological Sciences, Biogeochemical Laboratories, 1001 East Tenth Street, Bloomington, IN 47405-1405, USA. Our isotope standards are research materials offered for use without guarantees and without acceptance of any responsibilities for damages arising from its use or possible failure in any application. As a public university, we supply this reference standard material as a service to those engaged in stable isotope research, rather than as a commercial product. For additional information, see: http: / / www.indiana.edu / ~geosci / research / biogeochem / biogeochem.html

REFERENCES


References

1053

--

(1983) International vocabulary of basic and general terms in metrology. International Organization for Standardization, Geneva, Switzerland, 2nd edition. [40]1 -(1990)Commission regulation (EEC) No 2676/90 of 17 September 1990 determining community methods for the analysis of wines. Off. J. Eur. Commun., L 272:33 p. [5] -(1995) Reference and intercomparison materials for stable isotopes of light elements. IAEATECDOC-825, IAEA, Vienna: 165 p. [5] -(1997) Element sector ICPMS (1997) and the VG-MicroTrace 9000 sector GDMS (ca. 1995). Promotional publication, Finnigan-MAT [35] -(2000) Commission regulation (EC) No 2729/2000 of 14 December 2000 laying down detailed implementing rules on controls in the wine sector. Off. J. Eur. Commun., L 316:16 p. [5] Abbas M. M., Guo J., Carli B., Mencaraglia F., Carlotti M. & Nolt I. G. (1987) Heavy ozone distribution in the stratosphere from far-infrared observations. J. Geophys. Res., 92: 13231-13239. [18] Abelson P. H. & Hoering T. C. (1961) Carbon isotope fractionation in formation of amino acids by photosynthetic organisms. Proc. Natl. Acad. Sci. U.S.A., 47: 623-632. [9] Abragam A. (1961) The principle of nuclear magnetism. Clarendon Press, Oxford. [5] Abrajano T. A. Jr., Murphy D. E., Fang J., Comet P. & Brooks J. M. (1994) 13C/12C ratios in individual fatty acids of marine mytilids with and without bacterial symbionts. Org. Geochem., 21: 611-617. [9] Adams L. H. (1949-1950) Ratio of oxygen isotopes in quartz of contrasted origin. In: Carnegie Inst. Of Washington Year Book, Vol. 49: 36-37. [20] Adams N. G., Smith D. & Henchman M. J. (1982) Isotope exchange in the reactions H30 + + D20, NH4 + + ND3, CH5 + + CD4 and their mirror isotope reactions at thermal energies. Int. J. Mass Spectrom. Ion Phys., 42: 11-23. [4] Adams R. M., Hamilton S. A. & McCarl B. A. (1985) An assessment of the economic effects of ozone on U.S. agriculture. J. Air Pollut. Cotrol Assoc., 35: 938-943. [18] Adamsen E J. & Reeder J. D. (1983) A comparison of diffusion and distillation methods for preparing samples for 15N analyses. Soil Sci. Soc. Am. J., 47: 612-613. [15] Ader M., Boudou J.-P., Javoy M., Goffe B. & Daniels E. (1998) Isotope study on organic nitrogen of Westphgalian anthracites from the Western Middle field of Pennsylvania (U.S.A.) and from Bramsche Massif (Germany). Org. Geochem., 29: 315-323. [16] Ader M., Coleman M. L., Doyle S. P., Stroud M. & Wakelin D. (2001) Methods for the stable isotopic analysis of chlorine in chlorate and perchlorate compounds. Anal. Chem., 73: 4946-4950. [28] Aggarwal J. K. & Palmer M. R. (1995) Boron isotope studies of natural materials: A review of analytical methods. Analyst, 120: 1301-1307. [7] Aggarwal J. K., Rehktimper M. & Mezger K. (1999) Boron isotope measurements by multiple-collector ICP-MS. EOS Trans. AGU, 80: Fl118. [31] Aggarwal P. K. & Dillon M. A. (1998) Stable isotope composition of molecular oxygen in soil gas and groundwater: A potentially robust tracer for diffusion and oxygen consumption processes. Geochim. Cosmochim. Acta, 62(4): 577-584. [15] Aggett J., Burton C. A., Lewis T. A., Llewellyn D. R., O'Connor C. & Odell A. L. (1965) The isotopic analysis of oxygen in organic compounds and in co-ordination compounds containing organic ligands. Int. J. Appl. Radiat. Isot., 16: 165-170. [23] Aguilera R., Becchi M., Casabianca H., Hatton, C. K., Catlin D. H., Starcevic B. & Pope H. G. (1996) Improved method of detection of testosterone abuse by gas-chromatography combustion isotope ratio mass-spectrometry analysis of urinary steroids. J. Mass Spectrom., 31: 169-176. [8] Agyei E. K. & McMullen C. C. (1968) A study of the isotopic abundance of boron from various sources. Can. J. Earth Sci., 5: 921-927. [7] Ahoron P. & Fu B. (2000) Microbial sulfate reduction rates and sulfur and oxygen isotope fractionations at oil and gas seeps in deepwater Gulf of Mexico. Geochim. Cosmochim. Acta, 64: 233-246. [46] Aiken G., Thurman E. M., Malcolm R. L. & Walton H. E (1979) Comparison of XAD macroporous resins for the concentration of fulvic acid from aqueous solution. Anal. Chem., 51: 335-366. [15] 1. The number(s) between the square hooks at the end of each reference relate to the chapter(s) where that reference is cited.

1054

References

Akagi T., Franchi I. A. & Pillinger T. (1993) Oxygen isotope analysis of quartz by neodymium-yttrium aluminium garnet laser-fluorination. Analyst, 118: 1507-1510. [20] Akagi T., Franchi I. A. & Pillinger T. (1995) Isotope analysis of oxygen in minerals using Nd:YAG laserfluorination: the use of stainless-steel wool trap as a fluorine getter. Geochem. J., 29: 115-122. [20] A1-Aasm I. S., Taylor B. E. & South B. (1990) Stable isotope analysis of multiple carbonate samples using selective acid extraction. Chem. Geol. (Isot. Geosci. Sect.), 80, 119-125. [20] Alderton D. H. M. & Harmon R. S. (1991) Fluid inclusion and stable isotope evidence for the origin of mineralizing fluids in south-west England. Mineral. Magazine, 55: 605-611. [3] Aleu J., Fronza G., Fuganti C., Serra S., Fauhl C., Guillou C. & Reniero F. (2002) Differentiation of natural and synthetic phenylacetic acids by 2H NMR of the derived benzoic acids. Eur. Food Res. Technol., 214: 63-66. [5] Alewell C., Michell M. J. & Likens G. E. (1999) Sources of stream sulfate at the Hubbard Brook experimental forest: Long-term analyses using stable isotopes. Biogeochemistry, 44(3): 281-300. [45] Alibert C., Vicet A., Yarekha D. A., P6rona A., Rouillard Y., Chevrier F., Glastre G. & Baranov A. N. (2001) Gas analysis using room temperature diode lasers in the 2 - 2.5 micron range. Proceedings of the 3rd International Conference on Tunable Diode Laser Spectroscopy, Zermatt, Switserland, July 8-12, 2001. [34] Alisuaskas R. T., Klaas E. E., Hobson K. A. & Ankney C. D. (1998) Stable-carbon isotopes support use of adventitious color to discern winter origins of lesser snow geese. J. Field Ornith., 69: 262-268. [9] Allen R. E. & Parkes R. J. (1995) Digestion procedures for determining reduced sulfur species in bacterial cultures and in ancient and recent sediments. Am. Chem. Soc. Symp. Ser., 612: 243-257. [27] Allenby R. J. (1954) Determination of the isotopic ratios of silicon in rocks. Geochim. Cosmochim. Acta, 5: 40-48. [20, 25] Allison C. E. & Francey R. J. (1995) High precision stable isotope measurements of atmospheric trace gases. In: Reference and intercomparison materials for stable isotopes of light elements. IAEATECDOC-825, IAEA, Vienna: 131-153. [43] Allison C. E., Francey R. J. & Meijer H. A. J. (1995) Recommendations for the reporting of stable isotope measurements of carbon and oxygen in CO2 gas. In: Reference and intercomparison materials for stable isotopes of light elements. IAEA-TECDOC-825, IAEA, Vienna: 155-162. [14, 43] Amaral J. A., Hesslein R. H., Rudd J. W. M. & Fox D. E. (1989) Loss of total sulfur and changes in sulfur isotopic ratios due to drying of lacustrine sediments. Limnol. Oceanogr., 34: 1351-1358. [26] Amaral J. A., Kelly C. A. & Flett R. J. (1993) Measurement of organic-35S and organic-S in lake sediments: Methodological considerations. Biogeochemistry, 23: 61-78. [26] Amberger A. & Schmidt H.-L. (1987) Nati~rliche Isotopengehalte van Nitrat als Indicatoren f~ir dessen Herkunft. Geochim. Cosmochim. Acta, 51: 2699-2705. [15] Ambrose S. H. (1990) Preparation and characterization of bone and tooth collagen for isotopic analysis. J. Archaeol. Sci., 17: 431-451. [9] Ambrose S. H. (1991) Effects of diet, climate and physiology on nitrogen isotope abundances in terrestrial foodwebs. J. Archaeol. Sci., 18: 293-317. [9] Ammons J. M. (1980) Preconcentration methods for the determination of gaseous sulfur compounds in air. PhD Thesis, University of South Florida. [26] An Y. H. & Friedman R. J. (2000) Handbook of Bacterial Adhesion. Humana Press, Totowa, NJ. [46] Anbar A. D., Papanastassiou D. A., Creaser R. A. & Wasserburg G. J. (1997) The determination of iridium in natural waters by clean chemical preconcentration and negative thermal ionization mass spectrometry. Anal. Chem., 69: 2444-2450. [29] Anbar A. D., Zhang C., Barling J. & Roe J. E. (1999) Iron isotope fractionation studies using MC-ICPMS. Ninth Annual V.M. Goldschmidt Conference, Abstract #7612. [31] Anbar A. D., Roe J. E., Barling J. & Nealson K. (2000) Nonbiological fractionation of iron isotopes. Science, 288: 126-128. [29, 31] Anbar A. D., Knab K. A. & Barling J. (2001) Precise determination of mass dependant variations in the isotopic composition of molybdenum using MC-ICPMS. Anal. Chem., 73: 1425-1431. [31] Anderson D. Z., Frisch J. C. & Masser C. S. (1984) Mirror reflectometer based on optical cavity decay time. Appl. Opt., 23: 1238-1245. [34]

References

1055

Anderson S. & Mauersberger K. (1981) Calibration of a mass spectrometer experiment for ozone. Rev. Sci. Instrum., 52: 1025-1028. [18] Anderson S. M. & Kaye J. A. (1987) Comment on "Heavy ozone in the stratosphere". Geophys. Res. Lett., 14: 91-93. [18] Anderson S. M., Morton J. & Mauersberger K. (1989) Laboratory measurements of ozone isotopomers by tunable diode laser absorption spectroscopy. Chem. Phys. Lett., 156: 175-180. [18, 34] Anderson S. M., Mauersberger K., Morton J. & Schueler B. (1992) Heavy ozone anomaly, evidence for a mysterious mechanism. In: Isotope effects in gas-phase chemistry. J. A. Kaye (ed.). Am. Chem. Soc., Washington DC: 155-166. [18] Anderson S. M., Hupalo P. & Mauersberger K. (1993) Rotational structure in the near-infrared absorption spectrum of ozone. J. Chem. Phys., 99: 737-739. [18] Anderson S. M. & Mauersberger K. (1995) Ozone absorption spectroscopy in search of low-lying electric states. J. Geophys. Res., 100: 3033-3048. [18] Anderson S. M., Mauersberger K. & Morton J. (1995) The ozone molecule: Isotope effects and electronic structure. In: Progress and problems in atmospheric chemistry. J. R. Barker (ed.). World Scientific, Signapure: 473-497. [18] Anderson S. M., Hiilsebusch D. & Mauersberger K. (1997) Surprising rate coefficients for four isotopic variants of O + 02 + M. J. Chem. Phys., 107: 5385-5392. [18] Andersen T., Burke E. A. J. & Austrheim H. (1989) Nitrogen-bearing, aqueous fluid inclusions in some eclogites from Western Gneiss Region of the Norwegian Caledonides. Contrib. Mineral. Petrol., 103: 153-165.[16] Andersen T., Burke E. A. J. & Neumann E.-R. (1995) Nitrogen-rich fluid in the upper mantle: fluid inclusions in spinel dunite from Lanzarote, Canary Islands. Contrib. Mineral. Petrol., 120: 20-28. [16] Anderson T. F. (1968) Surface area measurement in calcite grains by isotopic exchange with C14-1abeled carbon dioxide. Geochim. Cosmochim. Acta, 32: 1177-1186. [46] Anderson T. F. (1969) Self-diffusion of carbon and oxygen in calcite by isotope exchange with carbon dioxide. J. Geophys. Res., 74: 3918-3932. [46] Anderson T. F. & Chai B., H.-T. (1974) Oxygen isotope exchange between calcite and water under hydrothermal conditions. In: Geochemical transport and kinetics. A. W. Hofmann, B. J. Giletti, H. S. Jr. Yoder & R. A. Yund (eds.). Carnegie Inst. Washington Publ., Vol. 634:219-227. [46] Anderson T. F., Bopp B. N., Williams A. C., Ho L. Z. & Hudson J. D. (1994) The stable isotopic record of fossils from the Peterborough member, Oxford clay formation (Jurassic), U.K.: Palaeoenvironmental implications. J. Geol. Soc. London, 151: 125-138. [22] Anderson W. B. & Polls G. A. (1998) Marine subsides of island communities in the Gulf of California: evidence from stable carbon and nitrogen isotopes. Oikos, 81: 75-80. [9] Andrawes F. F. & Gibson E. K. Jr. (1979) Release and analysis of gases from geological samples. Am. Mineralogist, 64: 453-463. [3] Andreae M. O. & Raemdonck H. (1983) Dimethylsulfide in the surface ocean and the marine atmosphere: A gobal view. Science, 221: 744-747. [26] Angerosa F., Camera L., Cumitini S., Gleixner G. & Reniero F. (1997) Carbon stable isotopes and olive oil adulteration with pomace oil. J. Agric. Food Chem., 45: 3044-3048. [8] Angert A. & Luz B. (2001) Fractionation of oxygen isotopes by root respiration: implications for the isotopic composition of atmospheric 02. Geochim. Cosmochim. Acta, 65: 1695-1701. [46] Annyas J., Bicanic D. & Schouten F. (1999) Novel instrumental approach to the measurement of total body water: optothermal detection of heavy water in the blood serum. Appl. Spectrosc., 53: 339-343. [34] Anonymous (2000) Porewater extraction from argillaceous rocks for geochemical characterization: methods and interpretation. OECD Publications, Paris: 188 p. [2] Anovitz L. M. & Grew E. S. (1996) Mineralogy, petrology and geochemistry of boron: an introduction. In: Boron: Mineralogy, petrology and geochemistry. E. S. Grew & L. M. Anaovitz (eds.). Mineralogical Society of America, Reviews in Mineralogy, Vol. 33: 1-40. [7] Appelblad P. K. & Baxter D. C. (2000) A model for calculating dead time and mass discrimination correction factors from inductively coupled plasma mass spectrometry calibration curves. J. Anal. At. Spectrom., 15: 557-560. [37].

1056

References

Appora I., Eiler J. M., Matthews A. & Stolper E. M. (2003) Experimental determination of oxygen isotope fractionation between CO2 vapor and soda-melilite melt. Geochim. Cosmochim. Acta, 67: 459471. [46] Aragulis-Aragu~s L., Rozanski K., Gonfiantini R. & Louvat D. (1995) Isotope effects accompanying vacuum extraction of soil water for stable isotope analyses. J. Hydrol., 168: 159-171. [1, 2] Aravena R., Schiff S. L., Trumbore S. E., Dillon P. J. & Elgood R. (1992) Evaluating dissolved inorganic carbon cycling in a forested lake watershed using carbon isotopes. Radiocarbon, 34: 636-645. [10] Aravena R., Evans M. L. & Cherry J. A. (1993) Stable isotopes of oxygen and nitrogen in source identification of nitrate from septic systems. Ground Water, 31: 180-186. [15] Arden J. W. (1983a) Distribution of lead and thallium in the matrix of the Allende meteorite and the extent of terrestrial lead contamination in chrondites. Earth Planet. Sci. Lett., 62: 395-406. [31] Arden J. W. (1983b) Electrochemical separation and isotopic determination of thallium at the nanogram level by surface ionization mass spectrometry. Anal. Chim. Acta, 148: 211-223. [31] Aregbe Y., Valkiers S., Mayer K., De Bi6vre P., Wessel R. M. & Alink A. (1998) Measuring amount ratios of gas isotopes by two primary methods. Metrologia, 35: 7-16. [43] Aregbe Y., Valkiers S., Poths J., Norgaard J., Kipphardt H., De Bi6vre P. & Taylor P. D. P. (2001) A primary isotopic gas standard for krypton with values for isotopic composition and molar mass traceable to the Syst6me International d'Unit6s. Int. J. Mass Spectrom., 206: 129-136. [42] Arlinghaus H. F., Spaar M. T., Tanigaki T., McMahon A. W. & Holloway P. H. (1994) Sputter-initiated Resonance Ionization Spectroscopy - An analytical technique for quantitative and sensitive measurement of impurities and ultra-shallow doping profiles in semiconductors. J. Vac. Sci. Technol., B 12: 263-268. [33] Arlinghaus H. F., Guo X. Q., Whitaker T. J. & Kwoka M. N. (1996) Fish & Chips: Analytical applications of resonance ionization mass spectrometry. Resonance Ionization Spectroscopy 1996, Eight International Symposium. N. Winograd & J.E. Parks (eds.). AIP Press, Woodbury, NY, AIP Conference Proceedings 388: 139-144. [33] Ash R. D., Arden J. W., Grady M. M., Wright I. P. & Pillinger C. T. (1990) Recondite interstellar carbon components in the Allende meteorite revealed by preparative precombustion. Geochim. Cosmochim. Acta, 54: 455-468. [13] Asghar R. B., Diskin A. M., Spanel P., Smith D. & Davies S. J. (2003) Measuring transport of water across the peritoneal membrane. Kidney Int., in press. [4] Asprey L. B. (1976) The preparation of very pure fluorine gas. J. Fluorine Chem., 7: 359-361. [20] ASTM (1958) Tentative method for sulfur in petroleum oil (quartz-tube method). In: Annual Book of ASTM Standards. American Society for Testing and Materials, Philadelphia: 828-832. [26] ASTM (1964) Standard method of test for sulfur in petroleum products by the bomb method. In: Annual Book of ASTM Standards. American Society for Testing and Materials,, Philadelphia. Vol. D129: 76-80. [26] ASTM (1993) Standard test methods for total sulfur in the analysis sample of coal and coke. In: Annual Book of ASTM Standards. American Society for Testing and Materials, Philadelphia. Vol. 05.05, D317789: 333-336. [26] Aston F. W. (1919) The constitution of the elements. Nature, 104: 393. [28] Aston F. W. (1920) The mass-spectra of chemical elements (Part 2). Phil. Magazine, 40: 628-634. [28] Aston F. W. (1942) Mass-spectra and isotopes. Longmans, Green & Co, NY, 2nd ed.: 276 p. [28, 35, 38] Assonov S. S. & Brenninkmeijer C. A. M. (2003) A redetermination of absolute values for 17RVPDB-CO2 and 17RVSMOW.Rapid Commun. Mass Spectrom., 17: 1017-~029. [40] Atekwana A. E. & Krishnamurthy R. V. (1998) Seasonal variations of dissolved inorganic carbon and 613C of surface waters: Application of a modified gas evolution technique. J. Hydrol., 205: 265-278. [10] Atkins P. W. (1988) Molecular quantum mechanics. Oxford University Press, Oxford: 199 p. [33] Atkins P. W. (1990) Physical chemistry. Oxford University Press, Oxford. [4] Attar A. (1979) Sulfur groups in coal and their determination. In: Analytical Methods for Coal and Coal Products. C. Karr (ed.). Academic Press, New York NY, USA, Vol III: 585-624. [26] Attar A. & Dupuis F. (1979) Data on the distribution of organic functional groups in coal. Amer. Chem. Soc. Div., Fuel Chem., 24: 166-177. [26]

References

1057

Aucour A.-M., Tao F.-X., Sheppard S. M. F., Huang N.-W. & Liu C. Q. (2002) Climatic and monsoon isotopic signals (6D, 613C) of northeastern China tree rings. J. Geophys. Res., 107(D9: Atmospheres): 1.1-1.9. [46] Aufderheide A. C., Kelley M. A., Rivera M., Gray L., Tieszen L. L., Iversen E., Krouse H. R. & Carevic A. (1994) Contributions to chemical dietary reconstruction to the assessment of adaption by ancient highland immigrants (also Ramirez) to coastal conditions at Pisagua, North Chile. J. Archaeol. Sci., 21: 515-524. [26] Aursand M., Mabon F. & Martin G. J. (1997) High-resolution 1H and 2H NMR spectroscopy of pure essential fatty acids for plants and animals. Magn. Reson. Chem., 35: $91-$100. [5] Aursand M., Mabon F. & Martin G. J. (2000) Characterization of farmed and wild salmon (Salmo salar) by a combined use of compositional and isotopic analysis. J.A.O.C.S., 77(6): 659-666. [5] Ayliffe L. K., Lister A. M. & Chivas A. R. (1992) The preservation of glacial-interglacial climatic signatures in the oxygen isotopes of elephant skeletal phosphate. Palaeogeogr., Palaeoclimat., Palaeoecol., 99: 179-191. [22] Baer D., Gupta M., O'Keefe A. & Paul J. (2002) Novel ultra-sensitive instrumentation for trace gas measurements. Presentation at the 16th Int. Forum Process Anal. Chem., San Diego. [34] Baertschi P. (1950) Isotopic composition of the oxygen in silicate rocks. Nature, 166: 112-113. [20] Baertschi P. & Silverman S. R. (1951) The determination of relative abundances of oxygen isotopes in silicate rocks. Geochim. Cosmochim. Acta, 1: 317-328. [20, 25] Baertschi P. (1953) Uber die relativen Unterschiede in H2180-Gehalt nati~rlicher Wasser. Helv. Chim. Acta, 36: 1352-1369. [1] Baertschi P. (1976) Absolute 180 content of standard mean ocean water. Earth Planet. Sci. Lett., 31: 341344. [4, 40, 43, 44] Bailey S. A. & Smith J. w. (1972) Improved method for the preparation of sulphur dioxide from barium sulphate for isotope ratio studies. Anal. Chem., 44: 1542-1543. [26] Bailey K., Chen C. Y., Du X., Li Y. M., Lu Z.-T., O'Connor T. P. & Young L. (2000) ATTA - A new method of ultrasensitive isotope trace analysis. Nucl. Instrum. Methods Phys. Res., B 172: 224-227. [34] Bains-Sahota S. K. & Thiemens M. H. (1987) Mass-independent oxygen isotope fractionation in a microwave-plasma. J. Phys. Chem., 91: 4370-4374. [18, 46] Bains-Sahota S. K. & Thiemens M. H. (1989) A mass-independent sulfur isotope effect in the nonthermal formation of $2F10. J. Chem. Phys., 90: 6099-6109. [18] Bakel A. J., Ostrom P. H. & Ostrom N. E. (1994) Carbon isotopic analysis of individual n-alkanes: evaluation of accuracy and application to marine particulate organic material. Org. Geochem., 21: 595-602. [8] Baker L., Franchi I. A., Maynard J., Wright I. P. & Pillinger C. T. (2002) A technique for the determination of 180/160 and 170/160 isotopic ratios in water from small liquid and solid samples. Anal. Chem., 74: 1665-1673. [1] Balabane M., Galimov E., Hermann M. & L6tolle R. (1987) Hydrogen and carbon isotope fractionation during experimental production of bacterial methane. Geochim. Cosmochim. Acta, 51: 115-119. [46] Ball J. D., Crowley S. F. & Steele D. F. (1996) Carbon and oxygen isotope ratio analysis of small carbonate samples by conventional phosphoric acid digestion: sample preparation and calibration. Rapid. Commun. Mass Spectrom., 10(8): 987-995. [12] Ball J. W. (1996) Thermodynamic and isotopic systematics of chromium chemistry. PhD Thesis, University of Arizona, Tucson, Arizona, USA. [29] Ballentine D. C., Macko S. A., Turekian V. C., Gilhooly W. P. & Matincingh B. (1996) Transport of biomass burning products through compound specific stable isotope analysis. Org. Geochem., 25: 97104. [9] Balzar A., Gleixner G., Grupe G., Scmidt H.-L., Schramm S. & Turban-Just S. (1997) In vitrio decomposition of bone collagen by soil bacteria: the implications for stable isotope analysis in archaeometry. Archaeometry, 39: 415-429. [9] Bandura D. R., Baranov V. I. & Tanner S. D. (2001) Reaction chemistry and collisional processes in multiple devices for resolving isobaric interferences in ICP-MS. Fresenius J. Anal. Chem., 370: 454-470. [31]

1058

References

Banfield J. F. & Nealson K. H. (1997) Geomicrobiology: Interactions between Microbes and Minerals. Mineralogical Society of America, Reviews in Mineralogy, Vol. 35:448 p. [46] Bao H. & Koch P. L. (1999) Oxygen isotope fractionation in ferric oxide-water systems: Low temperature synthesis. Geochim. Cosmochim. Acta, 63(5): 599-613. [46] Bao H. & Thiemens M. H. (2000) Generation of 02 from BaSO4 using a CO2-1aser fluorination system for simultaneous analysis of 8180 and 8170. Anal. Chem., 72(17): 4029-4032. [20] Bao H., Koch P. L. & Thiemens M. H. (2000a) Oxygen isotopic composition of ferric oxides from recent soil, hydrologic, and marine environments. Geochim. Cosmochim. Acta, 64: 2221-2231. [46] Bao H., Thiemens M. H., Farquhar J., Campbell D. A., Lee C. C.-W., Heine K. & Loope D. B. (2000b) Anomalous 170 compositions in massive sulphate deposits on the Earth. Nature, 406: 176-178. [18] Barbeni M., Cisero M. & Fuganti C. (1997) Natural abundance 2H nuclear magnetic resonance study of the origin of (Z)-3-hexenol. ]. Agric. Food Chem., 45: 237-241. [5] Barbieri B., Beverini N. & Sasso A. (1990) Optogalvanic spectroscopy. Rev. Mod. Phys., 62: 603-644. [34] Bariac T., Jusserand C. & L6tolle R. (1982) A routine method for 180 analysis of water vapor. Int. ]. AppI. Radiat. Isot., 33: 1365-1368. [1] Bariac T., Deleens E., Gerbaud A. & Mariotti A. (1991) La composition isotopique (180, 2H) de la vapeur d'eau transpir6e: 6tude en conditions asservies. Geochim. Cosmochim. Acta, 55: 3391-3402. [46] Bariac T., Gonzalez-Dunia J., Tardieu F., Tessier D. & Mariotti A. (1994) Variabilit6 spatiale de la composition isotopique del'eau (180, 2H) au sein des organes des plantes a6riennes: 1. Approche en conditions contr616es. Chem Geol., 115: 307-315. [21] Barker C. & Torkelson B. E. (1975) Gas adsorption on crushed quartz and basalt. Geochim. Cosmochim. Acta, 39: 212-218. [3] Barnes H. L. (1963) Ore solution chemistry. I. Experimental determination of mineral solubilities. Econ. Geol., 58: 1054-1060. [46] Barnes I. L., Garner E. L., Gramlich J. W., Moore L. J., Murphy T. J., Machlan L. A., Shileds W. R., Tatsumoto M. & Knight R. J. (1973) Determination of lead, uranium, thorium and thallium in silicate glass standard materials by isotope dilution mass spectrometry. Anal. Chem., 45: 880-885. [35] Barnes I. L., Moore L. J., Machlan L. A., Murphy T. J. & Shields W. R. (1975) Absolute isotopic abundance ratios and atomic weight of a reference sample of silicon. ]. Research Nat. Bureau Stand., 79A: 727-735. [25] Barnes J. H. IV, Sperline R., Denton M. B., Barinaga C. J., Koppenaal D. W., Young E. T. & Hieftje G. M. (2002) Characterization of a Focal Plane Camera Fitted to a Mattauch-Herzog Geometry Mass Spectrograph: 1. Use with a Glow-Discharge Source. Anal. Chem., 74: 5327-5332. [35] Barnes L., Downes C. T. & Hulston R. J. (1978) Warm Springs, South Island, New Zealand, and their potential to yield laumontite. Am. J. Sci., 278: 1412-1427. [10] Barnes R. M. (1998) Plasma source mass spectrometry in experimental nutrition. In: Mathematical modelling in experimental nutrition. A. J. Clifford & H.-G. M/iller (eds.),Vol. 46, Plenum Press: 379396. [31] Barnhart J. (1997) Chromium chemistry and implications for environmental fate and toxicity. In: Chromium in soil; perspectives in chemistry, health, and environmental regulation. D. M. Proctor, B. L. Finley, D. J. Harris, D. J. Paustenbach & D. Rabbe (eds.). Lewis Publishers: 561-568. [29] Baronnet A. (1982) Ostwald ripening in solution. The case of calcite and mica. Estudios Geol. (Madrid), 38: 185-198. [46] Barrer R. M. (1978) Zeolites and clay minerals as sorbents and molecular sieves. Academic Press: 497 p. [36]. Barrie A. & Coward W. A. (1985) A rapid analytical technique for the determination of enery expenditure by the doubly labelled water method. Biomed. Mass Spectrom., 12: 535-541. [1] Barrie A. (1991) New methodologies in stable isotope analysis. In: Stable isotopes in plant nutrition, soil fertility and environmental studies. IAEA-SM-313 / 106, Proceedings Series, IAEA, Vienna: 3-25. [15, 24] Barrie A. & Prosser S. J. (1996) Automated analysis of light-element stable isotopes by isotope ratio mass spectrometry. In: Mass spectrometry of soils. T. W. Boutton & S. Y. Yamasaki (eds.). Marcel Dekker Inc, NY: 1-46. [1, 9, 15]

References

1059

Barshick C. M., Smith D. H., Wade J. W. & Bayne C. K. (1994) Isotope dilution glow discharge mass spectrometry as applied to the analysis of lead in waste oil samples. J. Anal. At. Spectrom., 9: 83-87. [351 Barth S. (1993) Boron isotope variations in nature: a synthesis. Geol. Rundschau, 82: 640-651. [7] Barth S. (1997) Boron isotopic analysis of natural fresh and saline waters by negative thermal ionization mass spectrometry. Chem Geol., 143: 255-261. [7] Barth S. (1998) Application of boron isotopes for tracing sources of anthropogenic contamination in groundwater. Water Res., 32: 685-690. [7] Bartholomew R. M., Brown F. & Lounsbury M. (1954) Chlorine isotope effect in reactions of tert-butyl chloride. Can. J. Chem., 32: 979-983. [28] Bartoli F., Burtin G. & Herbillon A. J. (1991) Disaggregation and clay dispersion of oxisols: Na resin, a recommended methodology. Geoderma, 49: 301-317. [2] Bates A. L., Spiker E. C., Orem W. H. & Burnett W. C. (1993) Speciation and isotopic composition of sulfur in sediments from Jellyfish Lake, Palau. Chem. Geol., 106: 63-76. [26] Bates D. R. (1986) Heavy ozone in the stratosphere. Geophys. Res. Lett., 13: 664-666. [18] Bates D. R. (1988) Suggested explanation of heavy ozone. Geophys. Res. Lett., 15: 13-16. [18] Bates D. R. (1990) Isotopic fractionation in the formation of ozone and carbon dioxide. J. Chem. Phys., 93: 8739-8744. [18] Battagliarin M., Sentimenti E. & Scattolin R. (1995) An innovative sample preparation procedure for trace and ultra-trace analysis on non-conducting powders by direct current glow discharge mass spectrometry. Spectrochim. Acta, B 50: 13-25. [35] Baylis S. A., Hall K. & Jumeau E. J. (1994) The analysis of the C1-C5 components of natural-gas samples using gas-chromatography combustion isotope ratio mass-spectrometry. Org. Geochem., 21: 777-785.

[8]

Beard B. L. & Johnson C. M. (1999) High-precision iron isotope measurements of terrestrial and lunar materials. Geochim. Cosmochim. Acta, 63: 1653-1660. [29, 31] Beard B. L., Johnson C. M., Cox L., Sun H., Nealson K. & Aguilar C. (1999) Iron isotope biosignatures. Science, 285: 1889-1892. [29, 31, 46] Beard B. L., Johnson C. M., Skulan J. L., Nealson K. H., Cox L. & Sun H. (2003) Aplication of Fe isotopes to tracing the geochemical and biological cycling of Fe. Chem. Geol., 195: 87-117. [46] Bearhop S., Thompson D. R. & Waldron S. (1999) Stable isotopes indicate the extent of freshwater feeding by cormorants Phalacrocorax carbo shot at inland fisheries in England. J. Appl. Ecology, 36: 75-84. [91 Bearhop S., Teece M. A., Waldron S. & Waldron S. F. (2000) The influence of lipid and uric acid upon 613C and h15N values of avain blood: Implications for trophic studies. The Auk, 117: 504-507. [9] Beaudoin G. & Taylor B. E. (1993) MILES laser microprobe. Part 2: Preliminary assessment of precision and accuracy of sulphur isotope analysis. Geol. Surv. Canada, Current Res., Part D, Paper 93-ID: 199204. [20] Beaudoin G. & Taylor B. E. (1994) High precision and spatial resolution sulfur isotope analysis using MILES laser microprobe. Geochim. Cosmochim. Acta, 58(22): 5055-5063. [20] Beaudoin G., Taylor B. E., Rumble D. III& Thiemens M. (1994) Variations in the sulfur isotope composition of troilite from Canon Diablo iron meteorite. Geochim. Cosmochim. Acta, 58: 4253-4255. [26, 45] Beaudoin G. & Taylor B. E. (1995) Comparison of the environmental-SO2 and the laser-SF6 methods: implications for the sulfur isotope scale. In: Reference and intercomparison materials for stable isotopes of light elements. Proceedings of a Consultants Meeting held in Vienna, 1-3 Dec., 1993. IAEA, Vienna. IAEA-TECDOC-825: 49-50. [40] Beaudoin G., Taylor B. E., Rumble D. III & Thiemens M. (1995) Variations in the sulfur isotope composition of CDT. In: Reference and intercomparison materials for stable isotopes of light elements. Proceedings of a Consultants Meeting held in Vienna, 1-3 Dec., 1993. IAEA, Vienna. IAEATECDOC-825: 47-48. [40] Beaufort D., Dudoignon P., Proust D., Parneix J. C. & Meunier A. (1983) Microdrilling in thin section: a useful method for the identification of clay minerals in situ. Clay Minerals, 18: 219-222. [2] Beaumont V. & Robert F. (1999) Nitrogen isotope ratios of kerogens in Precambrian cherts: a record of the evolution of atmospheric chemistry? Precambrian Res., 96: 63-82. [13]

1060

References

Bebout G. E. & Fogel M. L. (1990) Techniques for extraction of nitrogen from silicate minerals and whole-rock samples for isotopic analysis. Ann. Rept. Dir. Geophys. Lab. 1989-1990, No. 2200. Carnegie Institution, Washington: 123-128. [36] Bebout G. E. & Fogel M. L. (1992) Nitrogen-isotope compositions of metasedimentary rocks in the Catalina Schist, California: implications for metamorphic devolatilization history. Geochim. Cosmochim. Acta, 56: 2839-2849. [16] Bebout G. E. & Barton M. D. (1993) Metasomatism during subduction: Products and possible paths in the Catalina Schist, California. Chem Geol., 108: 61-92. [16] Bebout G. E. (1995) Impact of subduction-zone metamorphism on mantle-ocean chemical cycling. Chem. Geol., 126: 191-218. [16] Bebout G. E. (1997) Nitrogen-isotope tracers of high-temperature fluid-rock interactions: case study of the Catalina Schist, California. Earth Planet. Sci. Lett., 151: 77-95. [16] Bebout G. E., Cooper D. C., Bradley A. D. & Sadofsky S. J. (1999a) Nitrogen-isotope study of fluid-rock interactions in the Skiddaw Aureole and Granite, English Lake District. Am. Mineralogist, 84: 14951505. [16] Bebout G. E., Ryan J. G., Leeman W. P. & Bebout A. E. (1999b) Fractionation of trace elements by subduction-zone metamorphism: Effect of convergent margin thermal evolution. Earth Planet. Sci. Lett., 171: 63-81. [16] Becchi M., Aguilera R., Farizon Y., Flament M., Casabianca H. & James P. (1994) Gas-chromatography combustion isotope ratio mass-spectrometry analysis of urinary steroids to detect misuse of testosterone in sport. Rapid Commun. Mass Spectrom., 8: 304-308. [8] Bechtel A. & Hoernes S. (1985) Sauerstoffisotopenuntersuchungen an OH-haltigen Silikaten: Fraktionierung zwischen OH- und Silikat-Sauerstoff. Fortschr. Mineral., 63(Suppl. 1): 18. [20] Bechtel A. & Hoernes S. (1990) Oxygen isotope fractionation between oxygen of different sites in illite minerals: a potential single-mineral thermometer. Contrib. Mineral. Petrol., 104: 463-470. [2, 20, 46] Bechtel A. & Hoernes S. (1993) Stable isotopic variations of clay minerals: A key to the understanding of Kupferschiefer-type mineralization, Germany. Geochim. Cosmochim. Acta, 57: 1799-1816. [2] Becker J. F., Sauke T. B. & Loewenstein M. (1992) Stable isotope analysis using diode laser spectroscopy. Appl. Opt., 31: 1921-1927. [34] Becker J. S., Breuer U., Westheide J., Saprykin A. I., Holzbrecher H., Nickel H. & Dietze H.-J. (1996) Trace and surface analysis of ceramic layers of solid oxide fuel cells by mass spectrometry. Fresenius J. Anal. Chem., 335: 626-632. [35] Becker J. S. & Dietze H.-J. (1998) Ultratrace and precise isotope analysis by double-focusing sector field inductively coupled plasma mass spectrometry. J. Anal. At. Spectrom., 13: 1057-1063. [35] Becker J. S. & Dietze H.-J. (2003) State-of-the-art in inorganic mass spectrometry for analysis of highpurity materials. Int. J. Mass Spectrom., 228: 127-150. [35] Becker R. H. & Clayton R. N. (1975) Nitrogen abundances and isotopic compositions in lunar samples. 6th Lunar Sci. Conf., Proceedings: 2131-2149. [17] Becker R. H. & Clayton R. N. (1977) Nitrogen isotopes in igneous rocks. EOS Trans. AGU, 58:536 (Abstract). [16] Becker R. H. & Epstein S. (1982) Carbon, hydrogen and nitrogen isotopes in solvent extractable organic matter from carbonaceous chondrites. Geochim. Cosmochim. Acta, 34: 257-278. [11] Begley I. S. & Sharp B. L. (1994) Occurence and reduction of noise in inductively coupled plasma mass spectrometry for enhanced precision in isotope ratio measurement. J. Anal. At. Spectrom., 9: 171-176. [32] Begley I. S. & Scrimgeour C. M. (1996) On-line reduction of H20 for 62H and 6180 measurement by continuous-flow isotope ratio mass spectrometry. Rapid Comm. Mass Spectrom., 10: 969-973. [1, 8] Begley I. S. & Scrimgeour C. M. (1997) High-precision 62H and 6180 measurement for water and volatile organic compounds by continuous flow pyrolysis isotope ratio mass spectrometry. Anal. Chem., 69: 1530-1535. [1, 4, 8, 23] Begley I. S. & Sharp B. L. (1997) Characterisation and correction of instrumental bias in inductively coupled plasma quadrupole mass spectrometry for accurate measurement of lead isotope ratios. J. Anal. At. Spectrom., 12: 395-402. [32]

References

1061

Begun G. M. & Landau L. (1961) Mass spectra and metastable transitions in isotopic nitrous oxide. J. Chem. Phys., 35: 547-551. [19] Behne W. (1953) Untersuchungen zur Geochemie des Chlor und Brom. Geochim. Cosmochim. Acta, 3: 186214. [28] Belshaw N. S., O'Nions R. K., Martel D. J. & Burton K. W. (1994) High resolution SIMS analysis of common lead. Chem. Geol., 112: 57-70. [35] Belshaw N. S., Freedman P. A., O'Nions R. K., Frank M. & Guo Y. (1998) A new variable dispersion double-focusing plasma mass spectrometer with performance illustrated for Pb-isotopes. Int. J. Mass Spectrom., 181: 51-58. [31] Belshaw N. S., Zhu X. K., Guo Y. & O'Nions R. K. (2000) High precision measurement of iron isotopes by plasma source mass spectrometry. Int. J. Mass Spectrom., 197: 191-195. [31] Belsky T. & Kaplan I. R. (1971) Light hydrocarbon gases, 13C and origin of organic matter in carbonaceous chondrites. Geochim. Cosmochim. Acta, 34: 257-278. [11] Belyaev S. S., Wolkin R., Kenealy W. R., DeNiro M. J., Epstein S. & Zeikus J. G. (1983) Methanogenic bacteria from the Bondyuzhskoe oil field: general characterization and analysis of stable-carbon isotopic fractionation. Applied Environ. Microbiol., 45: 691-697. [46] Ben-David M., Flynn R. W. & Schell D. M. (1997) Annual and seasonal changes in diets of martens: evidence from stable isotope analysis. Oecologia, 111: 280-291. [9] Bender M., Labeyrie L. D., Raynaud D. & Lorius C. (1985) Isotopic composition of atmospheric 02 in ice linked with deglaciation and global oprimary productivity. Nature, 318: 349-352. [21] Bender M., Sowers T. & Labeyrie L. (1994) The Dole effect and its variation during the last 130,000 years as measured in the Vostok ice core. Global Biogeochem. Cycles, 8(3): 363-376. [21] Benner R., Fogel M. L., Sprague E. K. & Hodson R. E. (1987) Depletion of 13C in lignin and its implications for stable isotope studies. Nature, 329: 708-710. [9] Benninghoven A., Riidenauer F. G. & Werner H. W. (1987) Secondary ion mass spectrometry: Basic concepts, instrumental aspects, applications and trends. Chemical Analysis Monograph, Vol. 86. Wiley & Sons, NY: 1216 p. [30, 39] Berden G., Peeters P. & Meijer G. (2000) Cavity ring-down spectroscopy: Experimental schemes and applications. Int. Rev. Chem. Phys., 19: 565-607. [34] Bergamaschi P., Schupp M. & Harris G. W. (1994) High-precision direct measurements of 13CH4/ 12CH4 and 12CH3D/12CH4 ratios in atmospheric methane sources by means of a long-path tunable diode laser absorption spectrometer. Applied Optics, 33(33): 7704-7716. [34] Bernard C., Bocquier G. & Javoy M. (1976) D6termination des rapports isotopiques 180/160 de la gibbsite et de la boehmite pr6sentes dans diff6rentes bauxites. C. R. Acad. Sci. Paris, 282: 1089-1092. [46] Berndt M. E., Seal R. R. II, Shanks W. C. III & Seyfried W. E. Jr. (1996) Hydrogen isotope systematics of phase separation in submarine hydrothermal systems: experimental calibration and theoretical models. Geochim. Cosmochim. Acta, 60: 1595-1604. [46] Berner R. A., Petsch S. T., Lake J. A., Beerling D. J., Popp B. N., Lane R. S., Laws E. A., Westley M. B., Cassar N., Woodward F. I. & Quick W. P. (2000) Isotopic fractionation and atmospheric oxygen: implications for Phanerozoic 02 evolution. Science, 287: 1630-1633. [46] Berner U. & Faber E. (1996) Empirical carbon isotope/maturity relationships for gases from algal kerogens and terrigenous organic matter, based on dry, open-system pyrolysis. Org. Geochem., 24: 947-955. [8] Bert D., Leavitt S. W. & Dupouey J.-L. (1997) Variations of wood ~13C and water-use efficiency of Abies alba during the last century. Ecology, 78: 1588-1596. [9] Bettany J. R., Stewart J. w. B. & Saggar S. (1979) The nature and forms of sulfur in organic matter fractions of soils selected along an environmental gradient. Soil Sci. Soc. Am. J., 43: 981-985. [26] Betti M., Giannarelli S., Hiernaut T., Rasmussen G. & Koch L. (1996a) Detection of trace radioisotopes in soil, sediment, and vegetation by glow-discharge mass spectrometry. Fresenius J. Anal. Chem., 355: 642-646. [35] Betti M., Rasmussen G. & Koch L. (1996b) Isotopic abundance measurements on solid nuclear-type samples by glow discharge mass spectrometry. Fresenius J. Anal. Chem., 355: 808-812. [35] Betz G. (1980) Surface sputtering. Surface Sci., 92: 283-309. [35]

1062

References

Betz G. & Wien K. (1994) Energy and angular distributions of sputtered particles. Int. J. Mass Spectrom. Ion Proc., 140: 1-110. [35] Bhattacharya S. K. & Thiemens M. H. (1989a) New evidence for symmetry dependent isotope effects: O + CO reaction. Z. Naturforsch., 44a: 435-444. [18] Bhattacharya S. K. & Thiemens M. H. (1989b) Effect of isotopic exchange upon symmetry dependent fractionation in the O + CO ~ CO2 reaction. Z. Naturforsch., 44a: 811-813. [18] Bickle M. J., Chapman H. J. & You C. F. (2000) Measurement of lithium isotopic ratios as lithium tetraborate ions. Int. J. Mass Spectrom., 202: 273-282. [6] Bidigare R. R., Fluegge A., Freeman K. H., Hanson K. L., Hayes J. M., Hollander D., Jasper J. P., King L. L., Laws E. A., Milder J., Millero F. J., Pancrost R., Popp B. N., Steinberg P. A. & Wakeham S. G. (1997) Consistent fractionation of 13C in nature and in the laboratory: Growth-rate effects in some haptophyte algae. Global Biogeochem. Cycles, 11: 279-292. [9] Biederbeck V. O. (1978) Soil organic sulfur and fertility. In: Soil Organic Matter. M. Schnitzer & S. U. Khan (eds.). Elsevier, New York: 273-310. [26] Bigeleisen J., Perlman M. L. & Prosser H. C. (1952) Conversion of hydrogenic materials to hydrogen for isotopic analysis. Anal. Chem., 24: 1356-1357. [1, 2, 3, 21] Bigeleisen J. & Wolfsberg M. (1958) Theoretical and experimental aspects of isotope effects in chemical kinetics. Adv. Chem. Phys., 1: 15-76. [18, 45] Bigeleisen J. (1965) Chemistry of isotopes. Science, 147: 463-471. [46] Bigeleisen J., Lee M. W. & Mandel F. (1973) Equilibrium isotope effects. Ann. Rev. Phys. Chem., 24: 407440. [46] Bijma J., Hemleben C., Huber B. T., Erlenkeuser H. & Kroon D. (1998) Experimental determination of the ontogenetic stable isotope variability in two morphotypes of Globigerinella siphonifera (d'Orbigny). Marine Micropaleo., 35: 141-160. [46] Birck J. L. & All6gre C. J. (1988) Manganese-chromium isotope systematics and the development of the early solar system. Nature, 331(No. 6157): 579-584. [29] Birck J. L. & Lugmair G. W. (1988) Nickel and chromium isotopes in Allende inclusions. Earth Planet. Sci. Lett., 90(2): 131-143. [29] Bird D. K., Cho M., Janik C. J., Liou J. G. & Caruso L. J. (1988) Compositional, order/disorder, and stable isotope characteristics of A1-Fe epidote, state 2-14 drill hole, Salton Sea geothermal system. J. Geophys. Res., 93(Bll): 13135-13144. [46] Bird M. I., Chivas A. R. & Andrew A. S. (1989) A stable-isotope study of lateritic bauxites. Geochim. Cosmchim. Acta, 53: 1411-1420. [2, 46] Bird M. I., Chivas A. R. & Andrews A. S. (1990) Reply to comments by C.-H. Chen, Liu K.-K. and Shieh Y.-N. on "A stable-isotope study of lateritic bauxites". Geochim. Cosmochim. Acta, 54: 1485-1486. [46] Bird M. I., Longstaffe F. J., Fyfe W. S. & Bildgen P. (1992) Oxygen-isotope systematics in a multiphase weathering system in Haiti. Geochim. Cosmochim. Acta, 56: 2831-2838. [2] Bird M. I., Longstaffe F. J. & Fyfe W. S. (1993) Oxygen-isotope fractionation in titanium-oxide minerals at low temperature. Geochim. Cosmochim. Acta, 57: 3083-3091. [2, 46] Bird M. I., Summons R. E., Gagan M. K., Roksandic Z., Dowling L., Head J., Fifield L. K., Cresswell R. G. & Johnson D. P. (1995) Terrestrial vegetation change inferred from n-alkane 613C analysis in the marine environment. Geochim. Cosmochim. Acta, 59: 2853-2857. [9] Bishop P. K. (1990) Precipitation of dissolved carbonate species from natural waters for 613C analysis A critical appraisal. Chem. Geol. (Isot. Geosci. Sect.), 80: 251-259. [10] Bizarro M., Simoneitti A., Stevenson R. K. & Kurszlaukis S. (2003) In situ 87Sr/86Sr investigation of igneous apatites and carbonates using laser-ablation MC-ICP-MS. Geochim. Cosmochim. Acta, 67: 289302. [20] Bjoroy M., Hall K., Gillyon P. & Jumeau J. (1991) Carbon isotope variations in n-alkanes and isoprenoids of whole oils. Chem. Geol., 93: 13-20. [9] Bjoroy M., Hall K. & Moe R. P. (1994) Stable carbon-isotope variation of n-alkanes in central graben oils. Org. Geochem., 22: 355-381. [8] Blair N., Leu A., Munoz E., Olsen J., Kwong E. & des Marais D. (1985) Carbon isotopic fractionation in heterotrophic microbial metabolism. Appl. Environm. Microbiol., 50: 996-1001. [9]

References

1063

Blake R. E., O'Neil J. R. & Garcia G. A. (1997) Oxygen isotope systematics of biologically mediated reactions of phosphate: I. Microbial degradation of organophosphoris compounds. Geochim. Cosmochim. Acta, 61: 4411-4422. [22, 46] Blake R. E., O'Neil J. R. & Garcia G. A. (1998) Effects of microbial activity on the 6180 of dissolved inorganic phosphate and textural features of synthetic apatites. Am. Mineralogist, 83: 1516-1531. [22] Blamart D., Graham C. M. & Sheppard S. M. F. (1986) The system NaA1Si308 (melt)-H20-seawater: solubility of water and hydrogen isotope fractionation factor at 3, 5 and 8 kb. Intern. Symp. Experimental Mineral. Geochem., Nancy, France: 23-24 (Abstract). [46] Blanchar R. W. (1986) Measurement of sulfur in soils and plants. In: Sulfur in Agriculture. M. A. Tabatabai (ed.). American Society of Agronomy - Crop Science Society of America - Soil Science Society of America, Madison, Wisconsin, USA. Monograph No. 27: 455-490. [26] Blatt H. & Brown V. M. (1974) Prophylactic separation of heavy minerals. J. Sedim. Petrol., 44: 260-261. [44] Blatt H. (1987) Oxygen isotopes and the origin of quartz. J. Sediment. Petrol., 57: 373-377. [2] Blattner P. (1973) Oxygen from liquids for isotopic analysis and a new determination of RCO2-H20 at 25~ Geochim. Cosmochim. Acta, 37: 2691-2693. [1] Blattner P. (1975) Oxygen isotope composition of fissure-grown quartz, adularia and calcite from Broadlands geothermal fields, New Zealand with an appendix on quartz-K-feldspar-calcitemuscovite oxygen isotope geothermometers. Am. J. Sci., 275: 785-800. [46] Blattner P. & Hulston J. R. (1978) Proportional variations of geochemical 6180 scales - an interlaboratory comparison. Geochim. Cosmochim. Acta, 42: 59-62. [43] Blattner P., Braithwaite R. B. & Glover R. B. (1983) New evidence on magnetite oxygen isotope geothermometers at 175~ and 112~ in Waikarei steam pipelines (New Zealand). Isot. Geosc., 1: 195-204. [29, 46] Blau K. & Halket J. (1993) Handbook for derivatives of chromatography. John Wiley & Sons, West Sussex, England. [9] Blaum K., Bushaw B. A., Geppert C., Miiller P., N/~rtersh~iuser W., Schmitt A., Trautmann N. & Wendt K. (1998) Diode-laser-based resonance ionization mass spectrometry of gadoliunium. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AlP Press Woodbury, NY. AIP Conference Proceedings 454:275 p. [33] Bligh E. G. & Dyer W. J. (1959) A rapid method of total lipid extraction and purification. Can. J. Biochem. Physiol., 37: 911-917. [9] Blum J. S., Bindi A. B., Buzzelli J., Stolz J. & Oremland R. S. (1998) Bacillus arsenicoselenatis, sp. nov., and Bacillus selenitireducens, sp. nov.: two haloalkaliphiles from Mono Lake, California that respire oxyanions of selenium and arsenic. Arch. Microbiol., 171: 19-30. [29] Blunier T., Chappellaz J. A., Schwander J., Barnola J. M., Desperts T., Stauffer B. & Raynaud D. (1993) Atmospheric methane, record from a Greenland ice core over the last 1000 year. Geophys. Res. Lett., 20(20): 2219-2222. [14] Bogaerts A. & Gijbels R. (1996) Relative sensitivity factors in glow discharge mass spectrometry. The role of charge transfer ionization. J. Anal. At. Spectrom., 11: 841-847. [35] Bogaerts A. & Gijbels R. (1998) Fundamental aspects and applications of glow discharge spectrometric techniques. Spectrochim. Acta, B 53: 1-42. [35] Bogaerts A. & Gijbels R. (2000) Behavior of the sputtered copper atoms, ions and excited species in a radio-frequency and direct current glow discharge. Spectrochim. Acta, B 55: 279-297. [35] Bogaerts A., Chen Z. & Gijbels R. (2003) Glow discharge modelling: from basic understanding towards applications. Surf. Interface Anal., 35: 593-603. [35] Bohn H. L., Barrow N. J., Rajan S. S. S. & Parfitt R. L. (1986) Reactions of inorganic sulfur in soils. In: Sulfur in Agriculture. M. A. Tabatabai (ed.). American Society of Agronomy - Crop Science Society of America - Soil Science Society of America, Madison, Wisconsin, USA. Monograph No. 27: 233-249. [26] Bol R. A. & Harkness D. D. (1995) The use of zeolite molecular sieves for trapping low concentrations of CO2 from environmental atmospheres. Radiocarbon, 37: 643-647. [36] Borella S. (1994) Jahrringstudien von 13C und 180 in Hinsicht auf Klimarekonstruktionen. Diploma Thesis, Universit~it Bern, Switserland. [24]

1064

References

Borella S. (1998) Stable isotope measurements of oxygen and carbon in tree rings: Method improvements and synergy of both parameters for ecological and climatological studies. PhD Thesis, University of Bern, Switserland. [24] Borella S., Leuenberger M., Saurer M. & Siegwolf R. (1998) Reducing uncertainties in ~)13Canalysis of tree rings: Pooling, milling, and cellulose extraction. J. Geophys. Res., 103(D16): 19519-19526. [24] Borella S., Leuenberger M. & Saurer M. (1999) Analysis of 6180 in tree rings: Wood-cellulose comparison and method dependent sensitivity. J. Geophys. Res., 104(D16): 19519-19526. [23, 24] Borggaard O. K. (1982) Selective extraction of amorphous iron oxides by EDTA from selective silicates and mixtures of amorphous and crystalline iron oxides. Clay Minerals, 17: 365-368. [2] Borthwick J. & Harmon R. S. (1982) A note regarding C1F3 as an alternative to BrF5 for oxygen isotope analysis. Geochim. Cosmochim. Acta, 46: 1665-1668. [2, 20] Boschetti A., Bassi D., Iacob E., Ianetti S., Ricci L. & Scotoni M. (2002) Resonant photoacoustic simultaneous detection of methane and ethylene by means of a 1.63 ~m diode laser. Appl. Phys., B 74: 273278. [34] Bottinga Y. (1969) Calculated fractionation factors for carbon and hydrogen isotope exchange in the system calcite-carbon dioxide-graphite-methane-water vapour. Geochim. Cosmochim. Acta, 33: 49-64. [13] Bottinga Y. & Craig H. (1969) Oxygen isotope fractionation between CO2 and water, and the isotopic composition of marine atmospheric CO2. Earth Planet. Sci. Lett., 5: 285-295. [1] Bottinga Y. & Javoy M. (1973) Comments on oxygen isotope geothermometry. Earth Planet. Sci. Lett., 20: 250-265. [46] Bottinga Y. & Javoy M. (1975) Oxygen isotope partitioning among the minerals in igneous and metamorphic rocks. Rev. Geophys. Space Physics, 13: 401-418. [46] Bottomley D. J., Veizer J., Nielsen H. & Moczydlowska M. (1992) Isotopic composition of disseminated sulfur in Precambrian sedimentary rocks. Geochim. Cosmochim. Acta, 56: 3311-3322. [26] Bottomley D. J., Katz A., Chan L. H., Starinsky A., Douglas M., Clark I. D. & Raven K. G. (1999) The origin and evolution of Canadian Shield brines: evaporation or freezing of seawater: New lithium isotope and geochemical evidence from the Slave craton. Chem. Geol., 155: 295-320. [6] Bottomley D. J., Chan L. H., Katz A, Starinsky A. & Clark I. D (2003) Lithium isotope geochemistry and origin of Canadian shield brines. Ground Water, in press. [6] Bottrell S. H., Carr L. P. & Dubessy J. (1988) A nitrogen-rich metamorphic fluid and coexisting minerals in slates from North Wales. Mineral. Magazine, 52: 451-457. [16] Bottrell S. H., Louie P. K. K., Timpe R. C. & Hawthorne S. B. (1994) The use of stable isotope ratio analysis to access selectivity of chemical analyses and extractions of forms of sulphur in coal. Fuel, 73: 1578-1582. [26] Bottrell S. H. & Novak M. (1997) Sulphur isotopic study of two pristine Shagnum bogs in the western British Isles. J. Ecology, 85: 125-132. [26] Bouman C., Vroon P. Z., Elliott T. R., Schwieters J. B. & Hamester M. (2002) Determination of lithium isotope compositions by MC-ICPMS (Thermo Finnigan MAT Neptune). Geochim. Cosmochim. Acta, 66:A97 (Abstract). [6] Bourg C., Stievenard M. & Jouzel J. (2001) Hydrogen and oxygen isotopic composition of aqueous salt solutions by gas-water equilibration method. Chem. Geol., 173: 331-337. [46] Boutton T. W. (1991) Stable carbon isotope ratios of natural materials 1. Sample preparation and mass spectrometer analysis. In: Carbon isotope techniques. D. C. Coleman& B. Fry (eds.). Academic Press, San Diego, USA: 155-171. [9] Bowling D. R., Sargent S. D., Tanner B. D. & Ehleringer, J. R. (2003) Tunable diode laser absorption spectroscopy for stable isotope studies of ecosystem-atmosphere CO2 exchange. Agric. Forest Meteorol., 118: 1-19. [34] Boyd J. R. & England J. L. (1960) Apparatus for phase-equilibrium measurements at pressures up to 50 kilobars and temperatures up to 1750~ J. Geophys. Res., 65: 741-748. [46] Boyd S. R., Mattey D. P., Pillinger C. T., Milledge H. J., Mendelssohn M. & Seal M. (1987) Multiple growth events during diamond genesis: an integrated study of carbon and nitrogen aggregation state in coated stones. Earth Planet. Sci. Lett., 86: 341-353. [16]

References

1065

Boyd S. R., Wright I. P., Franchi I. A. & Pillinger C. T. (1988) Preparation of sub-nanomole quantities of nitrogen gas for stable isotopic analysis. J. Phys. E: Sci. Instrum., 21: 876-885. [13, 17, 36] Boyd S. R. & Pillinger C. T. (1990) Determination of the abundance and isotopic composition of nitrogen within organic compounds: a sealed tube technique for use with static vacuum mass spectrometers. Meas. Sci. Technol., 1: 1176-1183. [13, 16] Boyd S. R. & Pillinger C. T. (1991) Rubidium sulfate - ammonium sulfate solid solution: a standard for use during the determination of the abundance and isotopic composition of nitrogen, at the ppm level, by static vacuum mass spectrometry. Anal. Chem., 63: 1332-1335. [13] Boyd S. R., Hall A. & Pillinger C. T. (1993a) The measurement of 615N in crustal rocks by static vacuum mass spectrometry: application to the origin of the ammonium in the Cornubian batholith, southwest England. Geochim. Cosmochim. Acta, 57: 1339-1347. [13, 16] Boyd S. R., Wright I. P. & Pillinger C. T. (1993b) Accurate determination of nitrogen concentrations by static vacuum mass spectrometry. Meas. Sci. Technol., 4: 1000-1005. [13, 17] Boyd S. R., Rejou-Michel A. & Javoy M. (1994) Non-cryogenic purification of nanomole quantities of nitrogen gas for isotopic analysis. Anal. Chem., 66: 1396-1402. [13] Boyd R. S., Rejou-Michel A. & Javoy M. (1995) Improved techniques for the extraction, purification and quantification of nanomole quantities of nitrogen gas: the nitrogen content of diamond. Meas. Sci. Technol., 6: 297-305. [13] Boyd S. R. (1997) Determination of the ammonium content of potassic rocks and minerals by capacitance manometry: a prelude to the calibration of FTIR microscopes. Chem. Geol., 137: 57-66. [13, 16] Boyd S., Wright I. P. & Pillinger C. T. (1997) Stepped-heating of carbonates and carbon-bearing quartz grains. Chem. Geol., 134: 303-310. [13] Boyd S. R. & Philippot P. (1998) Precambrian ammonium biogeochemistry: a study of the Moine metasediments, Scotland. Chem. Geol., 144: 257-268. [13, 16] Boyd S. R., Wright I. P., Alexander C. M. O. D. & Pillinger C. T. (1998) High resolution steppedcombustion mass spectrometry: Application to the detection and analysis of fine-grained diamond in meteorites and rocks. Geostandards Newsl., 22: 71-83. [13] Boyer P. D., Graves D. J., Suelter C. H. & Dempsey M. E. (1961) Simple procedure for conversion of oxygen of orthophosphate or water to carbon dioxide for oxygen-18 determination. Anal. Chem., 33: 1906-1909. [1] Biihlke J. K., Gwinn C. J. & Coplen T. B. (1993) New reference materials for nitrogen-isotope-ratio measurements. Geostand. Newsl., 17(1): 159-164. [15, 40] Biihlke J. K. & Coplen T. B. (1995) Interlaboratory comparison of reference materials for nitrogenisotope-ratio measurements. In: Reference and intercomparison materials for stable isotopes of light elements. Proceedings of a Consultants Meeting held in Vienna, 1-3 Dec 1993. IAEA, Vienna. IAEATECDOC-825: 51-66. [15, 40, 43] Biihlke J. K. & Denver J. M. (1995) Combined use of ground-water dating, chemical, and isotopic analyses to resolve the history and fate of nitrate contamination in two agricultural watersheds, atlantic coastal plain, Maryland. Water Resour. Res., 31: 2319-2339. [15] Biihlke J. K., Mroczkowski S. J. & Coplen T. B. (2003) Oxygen isotopes in nitrate: new reference materials for 180:170:160 measurements and observations on nitrate-water equilibration. Rapid Commun. Mass Spectrom., 17(16): 1835-1846. [15, 40] Bi3hm F., Joachimski M. M., Dullo W.-C., Eisenhauer A., Lehnert H., Reitner J. & WiJrheide G. (2000) Oxygen isotope fractionation in marine aragonite of coralline sponges. Geochim. Cosmochim. Acta, 64: 1695-1703. [46] Bi~ttcher J., Strebel O., Voerkelius S. & Schmidt H.-L. (1990) Using isotope fractionation of nitratenitrogen and nitrate-oxygen for evaluation of microbial denitrification in a sandy aquifer. J. Hydrol., 114(3-4): 413-424. [15, 29] B/Sttcher G., Brumsack H.-J., Heinrichs H. & Pohlmann M. (1997a) A new high-pressure squeezing technique for pore fluid extraction from terrestrial soils. Water, Air and Soil Pollution, 94: 289-296. [2] B/Sttcher M. E., Rusch A., HiJpner T. & Brumsack H.-J. (1997b) Stable sulfur isotope effects related to local intense sulfate reduction in a tidal sandflat (Southern North Sea): Results from loading experiments. Isot. Environ. Health Stud., 33: 109-129. [27]

1066

References

B/~ttcher M. E., Gehlken P.-L., Fernandez-Gonzalez A. & Prieto M. (1997c) Characterization of syn-

thetic BaCO3-SrCO3 (witherite-strontianite) solid-solutions by Fourier transform infrared spectoscopy. Eur. J. Mineral., 9: 519-528. [46] B/ittcher M. E., Brumsack H.-J. & de Lange G. J. (1998a) Sulfate reduction and related stable isotope (34S, 180) variations in interstitial waters of the eastern Mediterranean. In: Proceedings of the Ocean Drilling Program (ODP). A. H. F. Robertson, K.-C. Emeis & A. Camerlenghi (eds.). Sci. Res., Vol. 160: 365-373. [27] Bi~ttcher M. E., Oelschltiger B., Hi~pner T., Brumsack H.-J. & Rullkiitter J. (1998b) Sulfate reduction related to the early diagenetic degradation of organic matter and "black spot" formation in tidal sandflats of the German Wadden Sea: Stable isotope (13C, 34S, 180) and other geochemical results. Org. Geochem., 29: 1517-1530. [27] Bi~ttcher M. E., Smock A. M. & Cypionka H. (1998c) Sulfur isotope fractionation during experimental precipitation of iron(II) and manganese(II) sulfide at room temperature. Chem. Geol., 146: 127-134. [27] Bi~ttcher M. E. & Lepland A. (2000) Biogeochemistry of sulfur in a sediment core from the west-central Baltic Sea: Evidence from stable isotopes and pyrite textures. J. Mar. Sys., 25: 299-312. [27] Bi~ttcher M. E., Thamdrup B. & Vennemann T. W. (2001) Oxygen and sulphur isotope fractionation during anaerobic bacterial disproportionation of elemental sulfur. Geochim. Cosmochim. Acta, 65: 1573-1581. [27, 46] Brand W. A., Tegtmeyer A. R. & Hilkert A. (1994) Compound-specific isotope analysis: extending toward 15N/14N and 180/160. Org. Geochem., 21: 585-594. [1, 8, 23] Brand W. A. (1995a) Precon: A fully automated interface for the pre-GC concentration of trace gases in air for isotopic analysis. Isot. Environ. Health Stud., 31: 277-284. [8] Brand W. (1995b) High precision isotope ratio monitoring techniques in mass spectrometry. J. Mass Spectrom., 31: 225-235. [16] Brand W. A., Avak H., Seedorf R., Hofmann D. & Conradi T. (1996) New methods for fully automated isotope ratio determination from hydrogen at the natural abundance level. Isot. Environ. Health Stud., 32: 263-273. [1] Brand W. A. (1996) High precision isotope ratio monitoring techniques in mass spectrometry. J. Mass Spectrom., 31: 225-235. [8, 38, 43, 45] Brand W. A. (1998) Isotope ratio mass spectrometry: Precision from transient signals. In: Adv. Mass Spectrom., Vol. 14. E. J. Karjalainen, A. E. Hesso, J. E. Jalonen & U. P. Karjalainen(eds.): 655-680. [38] Brand W. A. & Coplen T. B. (2001) An interlaboratory study to test instrument performance of hydrogen dual-inlet isotope-ratio mass spectrometers. Fresenius J. Anal. Chem., 370: 358-362. [38, 43] Brandes J. A. & Devol A. H. (1997) Isotopic fractionation of oxygen and nitrogen in coastal marine sediments. Geochim. Cosmochim. Acta, 61: 1793-1801. [9] Brandriss M. E., O'Neil J. R., Edlund M. B. & Stoermer E. F. (1998) Oxygen isotope fractionation between diatomaceous silica and water. Geochim. Cosmochim. Acta, 62: 1119-1125. [2, 46] Brantley S. L., Liermann L. & Bullen T. D. (2001) Fractionation of Fe isotopes by soil microbes and organic acids. Geology, 29: 535-538. [29] Bray A. M. (2001) The geochemistry of boron and lithium in mid-ocean ridge hydrothermal vent fluids. PhD Thesis, University of New Hampshire, Durham: 125 p. [6] Br~iuer K. & Strauch G. (2000) An alternative procedure for the 180 measurement of nitrate oxygen. Chem. Geol., 168: 283-290. [15] Breck D. H. (1974) Zeolite molecular sieves: Structure, chemistry, and use. Wiley, NY. [36] Breit G. N., Simmons E. C. & Goldhaber M. B. (1985) Dissolution of barite for the analysis of strontium isotopes and other chemical and isotopic variations using aqueous sodium carbonate. Chem. Geol., 52: 333-336. [26] Bremner J. M. (1965a) Organic forms of nitrogen, n: Methods of soil analysis, Part 2. Chemical and microbiological properties. C. A. Black, D. D. Evans, L. E. Ensminger, J. L. White & F. E. Clark (eds.). American Society of Agronomy Inc., Madison, Wisconsin. Agronomy, Vol. 9: 1238-1255. [15] Bremner J. M. (1965b) Inorganic forms of nitrogen. In: Methods of soil analysis, Part 2. Chemical and microbiological properties. C. A. Black, D. D. Evans, L. E. Ensminger, J. L. White & F. E. Clark (eds.). American Society of Agronomy Inc., Madison, Wisconsin. Agronomy, Vol. 9: 1179-1237. [15, 28]

References

1067

Bremner J. M. (1965c) Isotope-ratio analysis of nitrogen in nitrogen-15 tracer investigations. In: Methods of soil analysis, Part 2. Chemical and microbiological properties. C. A. Black, D. D. Evans, L. E. Ensminger, J. L. White & F. E. Clark (eds.). American Society of Agronomy Inc., Madison, Wisconsin. Agronomy, Vol. 9 1256-1286. [15] Bremner J. M. (1965d) Total nitrogen. In: Methods of soil analysis, Part 2. Chemical and microbiological properties. C. A. Black, D. D. Evans, L. E. Ensminger, J. L. White & F. E. Clark (eds.). American Society of Agronomy Inc., Madison, Wisconsin. Agronomy, Vol. 9: 1149-1178. [15] Bremner J. M. & Edwards A. P. (1965) Determination and isotope-ratio analysis of different forms of nitrogen in soils: I. Apparatus and procedure for distillation and determination of ammonium. Soil Sci. Soc. Am. Proc., 29: 504-507. [15] Bremner J. M. & Keeney D. R. (1965) Steam distillation methods for determination of ammonium, nitrate and nitrite. Anal. Chim. Acta, 32: 485-495. [15] Bremner J. M. & Mulvaney C. S. (1982) Nitrogen - Total. In: Methods of soil analysis, Part 2. Chemical and microbiological properties. A. L. Page, R. H. Miller & D. R. Keeney (eds.). American Society of Agronomy Inc. - Soil Science Society of America Inc., Madison, Wisconsin. Agronomy Series, 2nd edition, No. 9: 595-624. [15] Bremner J. M. (1996) Nitrogen-Total. In: Methods of soil analysis. Part 3. Chemical methods. D. L. Sparks et. al. (eds.). Soil Science Society of America and American society of Agronomy, Madison, Wisconsin: 1085-1121. [15] Brendel O., Iannetta P. P. M. & Steward D. (2000) A rapid and simple method to isolate pure alphacellulose. Phytochem. Anal., 11: 7-10. [24] Brenna J. T. & Morrison G. H. (1984) Low temperature ashing for elemental sensitivity enhancement in secondary ion mass spectrometry. Anal. Chem., 56: 2791-2797. [2] Brenna J. T. (1994) High-precision gas isotope ratio mass spectrometry: Recent advances in instrumentation and biomediacal applications. Acc. Chem. Res., 27: 340-346. [8] Brenna J. T., Corso T. N., Tobias H. J. & Caimi R. J. (1997) High-precision continuous-flow isotope ratio mass spectrometry. Mass Spectrom. Rev., 16: 227-258. [8] Brenninkmeijer C. A. M. & Mook W. G. (1981) A batch process for direct conversion of organic oxygen and water to CO2 for 180/160 analysis. Int. J. Appl. Rad. Isot., 32: 137-141. [1] Brenninkmeijer C. A. M. (1983) Deuterium, oxygen-18 and carbon-13 in tree rings and peat deposits in relation to climate. PhD Thesis, University of Groningen, Groningen, the Netherlands: 146 p. [23, 24] Brenninkmeijer C. A. M. & Morrison P. D. (1987) An automated system for isotopic equilibration of CO2 and H20 for 180 analysis of water. Chem. Geol. (Isot. Geosci. Sect.), 66: 21-26. [1] Brenninkmeijer C. A. M. & Ri~ckmann T. (1999) Mass spectrometry of the intramolecular nitrogen isotope distribution of environmental nitrous oxide using fragment-ion analysis. Rapid Commun. Mass Spectrom., 13: 2028-2033. [19] Brewer A. K. (1936) The isotope effect in the evaporation of lithium, potassium and rubidium ions. J. Chem. Phys., 4: 350-353. [6] Br6as O., Guillou C., Reniero F., Sada E. & Tanet G. (1996) Deuterium/hydrogen measurements in wines and fruit juices by platinum catalysed equilibration method. Rapid Commun. Mass Spectrom., 10:246-249 [1] Br4as O., Guillou C., Reniero F., Sada E. & Angerosa F. (1998) Oxygen-18 measurement by continuous flow pyrolysis / isotope ratio mass spectrometer of vegetable oils. Rapid Commun. Mass Spectrom., 12: 188-192. [1, 23] Brilliant D. R., Franchi I. A. & Pillinger C. T. (1994) Nitrogen components in lunar 12023: Complex grains are not the carrier of isotopically light nitrogen. Meteoritics, 29:718-723 [13] Brimmer S. P., Fawcett W. R. & Kulhavy K. A. (1987) Quantitative reduction of selenate ion to selenite in aqueous samples. Anal. Chem., 59: 1470-1471. [29] Brindle I. D., Alarabi H., Karshman S., Le X. & Zheng S. (1992) Combined generator / separator for continuous hydride generation: Application to on-line pre-reduction of arsenic(V) and determination of arsenic in water by atomic emission spectrometry. Analyst, 117: 407-411. [29] Brindley G. W. & Lemaitre J. (1987) Thermal, oxidation and reduction reacions of clay minerals. In: Chemistry of clays amd clay minerals. A. C. D. Newman (ed.). Mineralogical Society, London: 319370. [2]

1068

References

Bronk D. A. & Glibert P. M. (1991) A 15N tracer method for the measurement of dissolved organic nitrogen release by phytoplankton. Mar. Ecol. Prog. Ser., 77: 171-182. [15] Brooks J. R., Flanagan L. B., Buchmann N. & Ehleringer J. R. (1997) Carbon isotope composition of boreal plants: functional grouping of life forms. Oecologia, 110: 301-311. [9] Brooks P. D., Stark J. M., McInteer B. B. & Preston T. (1989) Diffusion method to prepare soil extracts for automated nitrogen-15 analysis. Soil Sci. Soc. Am. J., 53: 1707-1711. [15] Brooks P. D. & Atkins G. J. (1992) Rapid isotopic analysis of trace gases at atmospheric levels. Isotopenpraxis, 28: 106-109. [8] Brown F., Gillies A. & Stevens W. H. (1953) A note on the preparation of chlorine gas containing 36C1. Can. J. Chem., 31: 768. [28] Brown K. A. (1986) Formation of organic sulphur in anaerobic peat. Soil Biol. Biochem., 18: 131-140. [26] Brown L., Rajan R. S., Roberts R. B., Tera R. & Whitford D. J. (1978) A new method for determining the isotopic composition of lithium. Nucl. Inst. Method, 156: 541-546. [6] Brown M. (1993) Deduction of emissions of source gases using an objective inversion algorithm and a chemical-transport model. J. Geophys. Res., 98(D7): 12639-12660. [14] Brumsack H.-J. (1980) Geochemistry of Cretaceous black shales from the Atlantic ocean (DSDP Legs 11, 14, 36 and 41). Chem. Geol., 31: 1-25. [27] Brumsack H. & Zuleger E. (1992) Boron and boron isotopes in pore waters from ODP Leg 127, Sea of Japan. Earth Planet. Sci. Lett., 113: 427-433. [7] Brunn4e C. (1987) The ideal mass analyzer: fact or fiction? Int. J. Mass Spectrom. Ion Proc., 76: 121-237. [32] Burns D. A. & Kendall C. (2002) Analysis of sources of 15N and 180 to differentiate NO3- sources in runoff in two watersheds in the Catskills Mountain of New York. Water Resour. Res., 38(5): doi 10.1029 / 2001WR000292,2002. [15] Briichert V. & Pratt L. M. (1996) Contemporaneous early diagenetic formation of organic and inorganic sulfur in estuarine sediments from St. Andrew Bay, Florida, USA. Geochim. Cosmochim. Acta, 60(13): 2325-2332. [29] Bryant J. D., Koch P. L., Froelich P. N., Showers W. J. & Genna B. J. (1996) Oxygen isotope partitioning between phosphate and carbonate in mammalian apatite. Geochim. Cosmochim. Acta, 60: 5145-5148. [46] Buchmann N., Brooks J. R., Rapp K. D. & Ehleringer J. R. (1996) Carbon isotope composition of C4 grasses is influenced by light and water supply. Plant, Cell Environm., 19: 392-402. [9] Buchmann N., Guehl J.-M., Barigah T. S. & Ehleringer J. R. (1997) Interseasonal comparison of CO2 concentrations, isotopic composition, and carbon dynamics in an Amazonian rainforest (French Guiana). Oecology, 110: 120-131. [9] Budd P., Lythgoe P., McGill R. A. R., Pollard A. M. & Scaife B. (1999) Zinc isotope fractionation in liquid brass (Cu-Zn) alloy: potential environmental and archeological applications. In: Geoarcheology: exploration, environments, resources. A. M. Pollard (ed.). Geological Society, Geological Society Special Publications, Vol. 165: 147-153. [31] Buhay W. M., Wolfe B. B., Elgood R. J. & Edwards T. W. D. (1995) A report on the 6180 measurements of the IAEA cellulose intercomparison material IAEA-C3. In: Reference and intercomparison materials for stable isotopes of light elements. IAEA-TECDOC-825, IAEA, Vienna: 85-86. [23] Bullen T. D., White A. F., Childs C. W., Vivit D. V. & Schulz M. S. (2001) A demonstration of significant abiotic Fe isotope fractionation in nature. Geology, 29: 699-702. [29] Bunch T. E. & Chang S. (1980) Carbonaceous chondrites II: Carbonaceous chondrites phyllosilicates and light element geochemsitry as indicators of parent body processes and surface conditions. Geochim. Cosmochim. Acta, 44: 1543-1557. [11] Bunn S. E., Davies P. M. & Kellaway D. M. (1997) Contributions of sugar cane and invasive pasture grass to the aquatic food web of tropical lowland stream. Mar. Freshwater Res., 48: 173-179. [9] Burdett J. W., Arthur M. A. & Richardson M. (1989) A Neogene seawater sulfur isotope age curve from calcareous pelagic microfossils. Earth Planet. Sci. Lett., 94: 189-198. [26] Burgoyne T. W., Hieftje G. M. & Hites R. A. (1997) Space charge evaluation in a plasma-source mass spectrograph. Anal. Chem., 69: 485-489. [32] Burgoyne T. W. & Hayes J. M. (1998) Quantitative production of H2 by pyrolysis of gas chromatographic effluents. Anal. Chem., 70: 5136-5141. [3, 14, 38]

References

1069

Burk R. L. & Stuiver M. (1981) Oxygen isotope ratios in tree cellulose reflect mean annual temperature and humidity. Science, 211: 1417-1419. [23] Burkhardt S., Riebesell U. & Zondervan I. (1999) Effects of growth rate, CO2 concentration, and cell size on the stable carbon isotope fractionation in marine phytoplankton. Geochim. Cosmochim. Acta, 63: 3729-3741. [46] Burnham C. W. & Jahns R. H. (1962) A method for determining the solubility of water in silicate melts. Am. J. Sci., 260: 721-745. [46] Burns D. A. & Kendall C. (2002) Analysis of sources of 15N and 180 to differentiate NO3- sources in runoff in two watersheds in the Catskills mountain of NewYork. Water Resour. Res., 38(5): 10.1029/ 2001WR000292. [15] Bushaw B. A., Juston F., Ni~rdh/iuser W., Trautmann N., Voss-de Haan P. & Wendt K. (1996) Multiple resonance RIMS, measurements of calcium isotopes using diode lasers. Resonance Ionization Spectroscopy 1996, Eight International Symposium. N. Winograd & J.E. Parks (eds.). AIP Press Woodbury, NY, AIP Conference Proceedings 388: 115-118. [33] Bushaw B. A. & Cannon B. D. (1998) Diode-laser-based RIMS measurements of 90St. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AlP Press Woodbury, NY, AIP Conference Proceedings 454: 177-182. [33] Buzek F. (1983) A rapid procedure for preparing oxygen-18 determination in water samples. Isotopenpraxis, 19: 70-72. [1] Cai W. J., Wang Y.C., Krest J. & Moore W. S. (2003) The geochemistry of dissolved inorganic carbon in a surficial groundwater aquifer in North Inlet, South Carolina, and the carbon fluxes to the coastal ocean. Geoch. Cosmochim. Acta, 67: 631-639. [10] Caimi R. J., Huoghton L. A. & Brenna J. T. (1994) Condensed-phase carbon isotopic standards for compound-specific isotope analysis. Anal. Chem., 66: 2989-2991. [8] Caimi R. J. & Brenna J. T. (1996) Direct analysis of carbon isotope variability in albumins by liquid flowinjection isotope ratio mass spectrometry. J. Am. Soc. Mass Spectrom., 7: 605-610. [8] Caimi R. J. & Brenna J. T. (1997) Quantitative evaluation of carbon isotopic fractionation during reversed-phase high-performance liquid chromatography. J. Chromatogr., 757A: 307-310. [8] Calhoun J. A. (1990) Chemical and isotopic methods for understanding the natural marine sulfur cycle. PhD Thesis, University of Washington: 86 p. [26] Calhoun J. A., Bates T. S. & Charlson R. J. (1991) Sulfur isotope measurements of submicrometer sulfate aerosol particles over the Pacific Ocean. Geophys. Res. Lett., 18(10): 1877-1880. [26] Callis E. L. & Abernathey R. M. (1991) High-precision isotopic analyses of uranium and plutonium by total sample volatilization and signal integration. Int. J. Mass Spec. Ion Proc., 103: 93-105. [35] Calvert S. E., Thode H. G., Yeung D. & Karlin R. E. (1996) A stable isotope study of pyrite formation in Late Pleistocene and Holocene sediments of the Black Sea. Geochim. Cosmochim. Acta, 60: 1261-1270. [27] Cameron A. E. (1955) Variation in the natural abundance of the lithium isotopes. J. Am. Chem. Soc., 77: 2731-2733. [6] Cameron A. E. & Lippert E. L. Jr. (1955) Isotopic composition of bromine in nature. Science, 121: 136-137. [281 Cameron A. E., Herr W., Herzog W. & Lund6n A. (1956) Isotopen-Anreicherung beim Brom durch electrolytische 0berftihrung in geschmolzenem Bleibromid. Z. Naturforsch., 11a: 203-205. [28] Campbell D. J. & Kaufmann R. S. (1984) Fractionation of chlorine isotopes by flow through semipermeable membranes. EOS Trans. AGU, 65: 882. [28] Campbell D. J., Kendall C., Chang C. C. Y., Silva S. R. & Tonnesen K. A. (2002) Pathways for nitrate release from an alpine watershed: determination using 6180 and 615N. Water Resour. Res., 38(5): doi 10.1029 / 2001WR000294. [15] Cane G. & Clark I. D. (1998) Tracing ground water recharge in an agricultural watershed with isotopes. Ground Water, 27: 133-139. [10] Canfield D. E., Raiswell R., Westrich J. T., Reaves C. M. & Berner R. A. (1986) The use of chromium reduction in the analysis of reduced inorganic sulphur in sediments and shales. Chem. Geol., 54: 149155. [26]

1070

References

Canfield D. E. & Thamdrup B. (1994) The production of 34S depleted sulfide during bacterial disproportionation of elemental sulfur. Science, 266: 1973-1975. [26] Canfield D. E. (2001a) Biogeochemistry of sulfur isotopes. Rev. Mineral. Geochem., 43: 607-636. [46] Canfield D. E. (2001b) Isotope fractionation by natural populations of sulphate-reducing bacteria. Geochim. Cosmochim. Acta, 65: 1117-1124. [46] Cantrell C. A., Shetter R. E., McDaniel A. H., Calvert J. A., Davidson J. A., Lowe D. C., Tyler S. C., Cicerone R. J. & Greenberg J. P. (1990) Carbon kinetic isotope effect in the oxidation by the hydroxyl radical. J. Geophys. Res., 95(D13): 22455-22462. [14] Canuel E. A., Freeman K. H. & Wakeham S. G. (1997) Isotopic compositions of lipid biomarker compounds in estuarine plants and surface sediments. Limnol. Oceanogr., 42: 1570-1583. [9] Caraco N. E & Cole J. J. (1999) Human impact on nitrate export: an analysis using major wolrd rivers. Ambio, 28: 167-170. [15] Carle R., Beyer J., Cheminat A. & Krempp E. (1992) 2H NMR determination of site-specific natural isotope fractionation in (-)-c~-bisabolols. Phytochem., 31(1): 171-174. [5] Carmody R., Plummer L. N., Busenberg E. & Coplen T. B. (1998) Methods for collection of dissolved sulfate and sulfide and analysis of their sulfur isotopic composition. U. S. Geol. Surv., Reston, Virginia. Open-File Report 97-234: 91p. [26] Carmody R. W. & Seal R. R. II (1999) Evaluation of the sulfur isotopic composition and homogeneity of the Soufre de Lacq reference material. Chem. Geol., 153: 289-295. [40] Carpenter S. J. & Lohmann K. C. (1995) 8180 and 813C values of modern brachiopod shells. Geochim. Cosmochim. Acta, 59: 3749-3764. [46] Carr R. H., Wright I. P., Joines A. W. & Pillinger C. T. (1986) Measurement of carbon stable isotopes at the nanomole level: A static mass spectrometer and sample preparation technique. J. Phys. E: Sci. Instrum., 19: 798-808. [3, 13] Casabianca H., Graff J. B., Jame P., Perrucchietti C. & Chastrette M. (1995) Application of hyphenated techniques to the chromatographic authentication of flavours in food-products and perfumes. HRCJ. High Resolut. Chromatogr., 18: 279-285. [8] Casciotti K. L., Sigman D. M., Galanter Hastings M., Biihlke J. K. & Hilkert A. (2002) Measurement of the oxygen isotopic composition of nitrate in seawater and freshwater using the denitrifier method. Anal. Chem., 74: 4905-4912. [15] Casey W. H. & Phillips B. L. (2001) Kinetics of oxygen exchange between sites in the GaO4A112(OH)24(H20)127+(aq) molecule and aqueous solution. Geochim. Cosmochim. Acta, 65: 705714. [46] Cassagnab6re A. (1998) Caract6risation et interpr6tation de la transition kaolinite-dickite dans les r6servoirs a hydrocarbures de Froy et Rind (Mer du Nord, Norv6ge). Th6se d'Universit6, Universit6 de Poitiers, Poitiers, France: 238 p. [2] Catanzaro E. J., Murphy T. J., Garner E. L. & Shields W. R. (1964) Absolute isotopic abundance ratio and the atomic weight of bromine. J. Res. Nat. Bur. Stand., 68A: 593-599. [28] Catanzaro E. J. & Murphy T. J. (1966) Magnesium isotope ratios in natural samples. J. Geophys. Res., 71: 1271-1274. [31] Catanzaro E. J., Murphy T. J., Garner E. L. & Shields W. R. (1966) Absolute isotopic abundance ratios and atomic weights of magnesium. J. Res. Nat. Bureau Stand., 70A: 453-458. [31] Catanzaro E. J., Champion C. E., Garner E. L., Marineneko G., Sappenfield K. M. & Shields W. R. (1970) Standard Reference Material: Boric Acid; Isotopic, and assay of standard reference materials. Natl. Bur. Stand. (US), Spec. Publ. (Febr. 1970), 260-17:70 p. [7, 40] Cerling T. E., Brown F. H. & Bowman J. R. (1985) Low-temperature alteration of volcanic glass: hydration, Na, K, 180 and Ar mobility. Chem. Geol., 52: 281-293. [46] Ceding T. E. (1992) Use of carbon isotopes in paleosols as an indicator of the p(CO2) of the paleoatmosphere. Global Biogeochem. Cycles, 6: 307-314. [9] Cerling T. E. & Sharp Z. D. (1996) Stable carbon and oxygen isotope analysis of fossil tooth enamel using laser ablation. Palaeogeogr. Palaeoclimatol. Palaeoecol., 126: 173-186. [20, 45] Chacko T., Mayeda T. K., Clayton R. N. & Goldsmith J. R. (1991) Oxygen and carbon isotope fractionations between CO2 and calcite. Geochim. Cosmochim. Acta, 55: 2867-2882. [46]

References

1071

Chacko T., Hu X., Mayeda T. K., Clayton R. N. & Goldsmith J. R. (1996) Oxygen isotope fractionations in muscovite, phlogopite, and rutile. Geochim. Cosmochim. Acta, 60(14): 2595-2608. [46] Chacko T., Riciputi L. R., Cole D. R. & Horita J. (1999) A new technique for determining equilibrium hydrogen isotope fractionation factors using the ion microprobe: Application to the epidote-water system. Geochim. Cosmochim. Acta, 63(1): 1-10. [46] Chacko T., Cole D. R. & Horita J. (2001) Equilibrium oxygen, hydrogen and carbon isotope fractionation factors applicable to geological systems. Rev. Mineral. Geochem., 43: 1-81. [46] Chae Y. M. & Krouse H. R. (1986) Alteration of sulfur-34 natural abundance in soil by application of feedlot manure. Soil Sci. Soc. Am. J., 50: 1425-1430. [26] Chai B. H. T. (1974) Mass transfer of calcite during hydrothermal recrystallization. In: Geochemical Transport and Kinetics. A. W. Hoffmann, B. J. Giletti, H. S. Jr. Yoder & R. A. Yund (eds.). Carnegie Inst. Wash. Pubt., Vol. 634: 205-218. [46] Chamberlain C. P., Blum J. D., Holmes R. T., Feng X., Sherry T. W. & Graves G. R. (1997) The use of isotope tracers for identifying populations of migratory birds. Oecologica, 109: 132-141. [9] Chan L. H. (1987) Lithium isotope analysis by thermal ionization mass spectrometry of lithium tetraborate. Anal. Chem., 59: 2662-2665. [6, 7, 31] Chan L.-H. & Edmond J. M. (1988) Variation of lithium isotope composition in the marine environment: a preliminary report. Geochim. Cosmochim. Acta, 52: 1711-1717. [6] Chan L. H., Edmond J. M., Thompson G. & Gillis K. (1992) Lithium isotopic composition of submarine basalts: Implications for the lithium cycle in the oceans. Earth Planet. Sci. Lett., 108: 151-160. [6, 46] Chan L. H., Edmond J. M. & Thompson G. (1993) A lithium isotope study of hot springs and metabasalts from mid-ocean ridge hydrothermal systems. J. Geophys. Res., 98: 9653-9659. [6] Chan L. H., Gieskes J. M., You C. F. & Edmond J. M. (1994) Lithium isotope geochemistry of sediments and hydrothermal fluids of the Guaymas Basin, Gulf of California. Geochim. Cosmochim. Acta, 58: 4443-4454. [6] Chan L. H., Leeman W. P. & You C. F. (1999) Lithium isotopic composition of Central American Volcanic Arc lavas: Implications for modification of subarc mantle by slab-derived fluids. Chem. Geol., 160: 255-280. [6] Chan L. H. & Kastner M. (2000) Lithium isotopic compositions of pore fluids and sediments in the Costa Rica subduction zone: implications for fluid processes and sediment contribution to the arc volcanoes. Earth Planet. Sci. Lett., 183: 275-290. [6] Chan L. H., Leeman W. P. & Tonarini S. (2001) Lithium isotope compositions of South Sandwich Arc and Southwest Washington Cascades: A comparative study of arc processes. EOS Trans. AGU, 82: F1293 (Abstract). [6] Chan L. H., Alt J. C. & Teagle D. A. H. (2002a) Lithium and lithium isotope profiles through the upper oceanic crust: a study of seawater-basalt exchange at ODP Sites 504B and 896A. Earth Planet. Sci. Lett., 201: 187-201. [6] Chan L. H., Leeman W. P., Tonarini S. & Singer B. (2002b) Lithium and boron isotopes in Aleutian Islands: Contribution of marine sediments to island arc magmas. EOS Trans. AGU, 83, F1428 (Abstract). [6] Chan L. H., Leeman W. P. & You C. F. (2002c) Lithium isotopic composition of Central American Volcanic Arc lavas: Implications for modification of subarc mantle by slab-derived fluids, correction. Chem. Geol., 182: 293-300. [6] Chan L. H., Starinsky A. & Katz A. (2002d) The bahavior of lithium and its isotopes in oilfield brines; Evidence from the Heletz-Kokhav field, Israel. Geochim. Cosmochim. Acta, 66: 615-623. [6] Chan L. H. & Frey F. A. (2003) Lithium isotope geochemistry of the Hawaiian Plume: Results from the Hawaii Scientific Drilling Project and Koolau Volcano. Geochem. Geophys. Geosyst., 4(3), 8707, doi 10.1029 / 2002GC000365. [6] Chang C. C. Y., Langston J., Riggs M., Campbell D. H., Silva S. R. & Kendall C. (1999) A method for nitrate collection for 615N and 6180 analysis from waters with low nitrate concentrations. Can. J. Fish Aquat. Sci., 56: 1856-1864. [9, 15] Chang C. C. Y., Kendall C., Silva S. R., Battalgin W. A. & Campbell D. H. (2002) Nitrate stable isotopes: tools for determining nitrate sources among different land uses in the Mississippi River Basin. Can. J. Fish. Aq. Sci., 59: 1874-1885. [15]

1072

References

Chanton J. P. & Martens C. S. (1985) The effects of heat and stannous chloride addition on the active

distillation of acid volatile sulfide from pyrite-rich marine sediment samples. Biochem., 1: 375-383. [26] Chanton J. P., Martens C. S. & Goldhaber M. B. (1987) Biogeochemical cycling in an organic-rich coastal marine basin. 8. A sulfur isotope budget balanced by differential diffusion across the sediment-water interface. Geochim. Cosmochim. Acta, 51: 1201-1208. [27] Chapelle F. H. & Knobel L. L. (1985) Stable carbon isotopes of HCO3- in the Aquiua aquifer, Maryland: evidence for an isotopically heavy source of CO2. Ground Water, 23: 592-599. [10] Chapelle F. H., Morris J. T., McMahon P. B. & Zelibor J. L. (1988) Bacterial metabolism and the 613C composition of groundwater, Floridan aquifer system, South Carolina. Geology, 16: 117-121. [10] Chapelle F. H. & McMahon P. B. (1991) Geochemistry of dissolved inorganic carbon in a Coastal Plain aquifer. 1. Sulfate from confining beds as an oxidant in microbial CO2 production. J. Hydrol., 127: 85108.[10] Chapman B. (1980) Glow discharge processes: Sputtering and plasma etching. Wiley, NY: 406 p. [35] Chapman R. (1972) Versatile wide range electron current regulator. Sci. Instrum., 43: 1536-1538. [39] Charef A. & Sheppard S. M. F. (1984) Carbon and oxygen isotope analysis of calcite or dolomite associated with organic matter. Isotope Geosci., 2: 325-333. [2] Chartier F. & Tabarant M. (1997) Comparison of the performance of a laboratory-built high resolution glow discharge mass spectrometer with that of a quadrupole inductively coupled plasma (glow discharge) mass spectrometer for boron and gadolinium isotopic analysis. J. Anal. At. Spectrom., 12: 1187-1193. [35] Chau Y. K. & Riley J. P. (1965) The determination of selenium in sea water, silicates, and marine organisms. Anal. Chim. Acta, 33: 36-49. [29] Chaussidon M. & Lorand J.-P. (1990) Sulphur isotope composition of orogenic spinel lherzolite massifs from Ari6ge (North-Eastern Pyr6n6es, France): an ion microprobe study. Geochim. Cosmochim. Acta, 54: 2835-2846. [20] Chaussidon M. & Albar6de F. (1992) Secular boron isotope variations in the continental crust: an ion microprobe study. Earth Planet. Sci. Lett., 108: 229-241. [46] Chaussidon M. & Jambon A. (1994) Boron content and isotopic composition of oceanic basalts: Geochemical and cosmochemical implications. Earth Planet. Sci. Lett., 121: 277-291. [30] Chaussidon M. & Uiterdijk Appel P. W. (1997) Boron isotopic composition of tourmalines from the 3.8Ga-old Isua supracrustals, West Greenland: implications on the 611B value of early Archean seawater. Chem. Geol. (Isot. Geosci. Sect.), 136: 171-180. [30] Chaussidon M., Robert F., Mangin D., Hanon P. & Rose E. F. (1997) Analytical procedures for the measurement of boron isotope compositions by ion microprobe in meteorites and mantle rocks. Geostand. Newslett., 21: 7-16. [30] Chaussidon M. & Robert F. (1998)7Li/6Li and 11B/10B variations in chondrules from the Semarkona unequilibrated chondrite. Earth Planet. Sci. Lett., 164: 577-589. [6, 30] Chaussidon M. & Robert F. (1999) Lithium nucleosynthesis in the Sun inferred from the solar wind 7Li/ 6Li ratio. Nature, 402: 270-273. [30] Chaux R. & Lavorel B. (2000) Relative line intensity measurement in absorption spectra using tunable diode laser at 1.6 ~m: application to the determination of 13CO2/12CO2isotope ratio. Appl. Phys., B 72: 237-240. [34] Chazot G. (1997) t~tude de la composition isotopique de l'oxyg6ne des surfaces cuirass6es du bouclier guyanais. BRGM, Report R39465:33 p. [2] Chen C.-H., Liu K.-K. & Shieh Y.-N. (1988) Chemical and isotopic studies of bauxitization in the Tatun Volcanic area, northern Taiwan. Chem. Geol., 68: 41-56. [46] Chen C.-H., Liu K.-K. & Shieh Y.-N. (1990) Comments on "A stable-isotope study of lateritic bauxites" by M. I. Bird, A. R. Chivas and A. S. Andrews. Geochim. Cosmochim. Acta, 54:. 1483-1484. [46] Chen C. Y., Li Y. M., Bailey K., O'Connor T. P., Young L. & Lu Z.-T. (1999) Ultrasensitive isotope trace analysis with a magneto-optical trap. Science, 286: 1139-1141. [34]

References

1073

Cheng H. H. & Bremner J. M. (1965) Gaseous forms of nitrogen. In: Methods of soil analysis, Part 2. Chemical and microbiological properties. C. A. Black, D. D. Evans, L. E. Ensminger, J. L. White & F. E. Clark (eds.). American Society of Agronomy Inc., Madison, Wisconbsin, USA. Vol 9: 1287-1323. [36] Chen X. & Houk R. S. (1996) Spatially-resolved measurements of ion density behind the skimmer of an inductively coupled plasma mass spectrometer. Spectrochim. Acta, B 51: 41-54. [35] Chen Huang T. S. C. & Chin H.-C. (1973) Infrared methods for analysis of very low and very high concentration of deuterium oxide in water. Internal Report INER-62-D-0117. Inst. Nucl. Energy Res., Atomic Energy Council, Taiwan. [7] Chiba H. & Sakai H. (1985) Oxygen isotope exchange rate between dissolved sulfate and water at hydrothermal temperatures. Geochim. Cosmochim. Acta, 49: 993-1000. [26, 46] Chiba H., Sakai H. & Yasutake M. (1985) Automatic measurement of oxygen isotope ratio of water samples. Hot Spring Institute, Okayama University, Okayama, Japan, Report No. 56:27-34 (In Japanese with English abstract). [1] Chiba H., Chadko T., Clayton R. N. & Goldsmith J. R. (1989) Oxygen isotope fractionations involving diopside, forsterite, magnetite, and calcite: Application to geothermometry. Geochim. Cosmochim. Acta, 53: 2985-2995. [46] Chihara K., Suzuki M. & Kawazoe K. (1978) Adsorption rate on molecular sieving carbon by chromatography. AIChE Journal, 24: 237-246. [36] Chou I. M. (1986) Permeability of precious metals to hydrogen at 2 kb total pressure and elevated temperatures. Am. J. Sci., 286: 638-658. [46] Chou I. M. (1987) Oxygen buffer and hydrogen sensor techniques at elevated pressures and temperatures. In: Hydrothermal experimental techniques. G. C. Ulmer & H. L. Barnes (eds.). WileyInterscience, NY: 61-99. [46] Chu X. L., Zhao R., Zang W. X., Lei J. J. & Tang u G. (1994) Extraction of various sulfurs in coal and sedimentary rock and preparation of samples for sulfur isotopic analysis. Chin. Sci. Bull., 39: 140145. [26] Chukhrov F. V., Ermilova L. P., Churikov V. S. & Nosik L. P. (1980) The isotopic composition of plant sulfur. Org. Geochem., 2: 69-75. [26] Ciais P., Tans P., White J. W., Francey R., Trolier M., Berry J., Randall D., Sellers P., Collatz J. & Schimrnel D. (1995) The global carbon budget inferred from a new worldwide data set of 613C measurements in atmospheric CO2. J. Geophys. Res., 100(D3): 5051-5070. [14] Ciais P., Denning A. S., Tans P. P., Berry J. A., Randall D. A., Collatz G. J., Sellers P. J., White J. W. C., Trolier M., Meijer H. A. J., Francey R. J., Monfray P. & Heimann M. (1997) A three-dimensional synthesis study of 6180 in atmospheric CO2. 1. Surface fluxes. J. Geophys. Res., 102: 5857-5872. [14] Cicerone R. J. & McCrumb J. L. (1980) Photodissociation of isotopically heavy 02 as a source of atmospheric 03. Geophys. Res. Lett., 7: 251-254. [18] Cicerone R. J. & Oremland R. S. (1988) Biogeochemical aspects of atmospheric methane. Global Biogeochem. Cycles, 2: 299-327. [14] Cifuentes L. A., Sharp J. H. & Fogel M. L. (1988) Stable carbon and nitrogen isotope biogeochemistry in the Delaware estuary. Limnol. Oceanogr., 33: 1102-1115. [9] Cifuentes L. A., Fogel M. L., Pennock J. R. & Sharp J. H. (1989) Biogeochemical factors that influence the stable nitrogen isotope ratio of dissolved ammonium in the Delaware Estuary. Geochim. Cosmochim. Acta, 53: 2713-2721. [9, 46] Clayton D. D., Obradovich M., Guha M. & Brown L. E. (1991) Silicon and titanium isotopes in SiC from AGB stars. Lunar Planet. Sci., XXII: 221-222 (Abstract). [25] Clayton R. N. & Epstein S. (1958) The relationship between 180/16 0 ratios in coexisting quartz, carbonate, and iron oxides from various geological deposits. J. Geol., 66: 352-373. [20] Clayton R. N. (1959) Oxygen isotope fractionation in the system calcium carbonate - water. J. Chem. Phys., 30: 1246-1250. [46] Clayton R. N. (1961) Oxygen isotope fractionation between calcium carbonate and water. J. Chem. Phys., 34: 724-726. [46] Clayton R. N. & Mayeda T. K. (1963) The use of bromine pentafluoride in the extraction of oxygen from oxides and silicates for isotopic analysis. Geochim. Cosmochim. Acta, 27: 43-52. [2, 20, 25, 36]

1074

References

Clayton R. N., O'Neil J. R. & Mayeda T. K. (1972) Oxygen isotope exchange between quartz and water. J. Geophys. Res., 77: 3057-3067. [46]

Clayton R. N., Mayeda T. K. & Hurd J. M. (1974) Loss of oxygen, silicon, sulfur and potassium from the lunar regolith. Proc. Fifth Lunar Sci. Conf.: 1801-1809. [25]

Clayton R. N., Goldsmith J. R., Karel K. J., Mayeda T. K. & Newton R. C. (1975) Limits on the effect of pressure on isotopic fractionation. Geochim. Cosmochim. Acta, 39: 1197-1201. [46]

Clayton R. N., Mayeda T. K. & Epstein S. (1978) Isotopic fractionation of Si in Allende inclusions. Proceedings of the 9th Lunar Sci. Conf., Houston. Vol. 1: 129-130. [20, 25] Clayton R. N. & Mayeda T. K. (1983) Oxygen isotopes in eucrites, shergottites, nakhlites, and chassignites. Earth Planet. Sci. Lett., 62: 1-6. [36] Clayton R. N. (1986) High temperature isotope effects in the early solar system. In: Stable isotopes in high temperature geological processes. J. W. Valley, H. P. Jr. Taylor & J. R. O'Neil (eds.). Mineral. Soc. America, Reviews in Mineralogy, Vol. 16: 129-140. [25] Clayton R. N., Goldsmith J. R. & Mayeda T. K. (1989) Oxygen isotope fractionation in quartz, albite, anorthite and calcite. Geochim. Cosmochim. Acta, 53: 725-733. [46] Clayton R. N. & Kieffer S. W. (1991) Oxygen isotope thermometer calibrations. In: Stable Isotope Geochemistry: A tribute to Samuel Epstein. H. P. Jr. Taylor, J. R. O'Neil & I. R. Kaplan (eds.). The Geochemical Society, Spec. Publ. 3: 3-10. [46] Clayton R. N. (1992) Silica-carbonate isotopic temperature calibration. Science, 258: 1162-1163. [46] Clement S. W. & Compston W. (1989) SIMS at high sensitivity and high mass resolution. Proc. of the 7th International Conference on Secondary Ion Mass-Spectroscopy. J. Wiley & Sons. [25] Cliff S. S. & Thiemens M. H. (1997) The 180/160 and 170/160 ratios in atmospheric nitrous oxide: A mass-independent anomaly. Science, 278: 1774-1776. [18, 19] Cliff S. S., Thiemens M. H. & Brenninkmeijer C. A. M. (1999) First measurement of the 180/160 and 170/160 ratios in stratospheric nitrous oxide: A mass independent anomaly. J. Geophys. Res., 104: 16171-16175.[19] Cline J. D. & Kaplan I. R. (1975) Isotopic fractionation of dissolved nitrate during denitrification in the eastern tropical North Pacific ocean. Marine Chem., 3: 271-299. [15] Close K. J. & Yarwood J. (1972) A precision electron emission regulator. Vacuum, 22: 45-46. [39] Coburn J. W. & Kay E. (1971) A new technique for the elemental analysis of thin surface layers of solids. Phys. Lett., 19: 350-352. [35] Coburn J. W. & Harrison W. W. (1981) Plasma sources in analytical mass spectrometry. Appl. Spec. Rev., 17: 95-164. [35] Cohn M. & Urey H. C. (1938) Oxygen exchange reactions of organic compounds and water. J. Am. Chem. Soc., 60: 679-682. [1] Cole D. R., Ohmoto H. & Lasaga A. C. (1983) Isotopic exchange in mineral-fluid systems. I. Theoretical evaluation of oxygen isotopic exchange accompanying surface reactions and diffusion. Geochim. Cosmochim. Acta, 47: 1681-1693. [46] Cole D. R. & Ohmoto H. (1986) Kinetics of isotopic exchange at elevated temperatures and pressures. In: Stable isotopes in high temperature geological processes. J. W. Valley, H. P. Jr. Taylor & J. R. O'Neil (eds.). Mineral. Soc. America, Reviews in Mineralogy, Vol. 16: 41-90. [46] Cole D. R. (1992) Influence of solution composition and pressure on the rates of oxygen isotope exchange in the system: calcite-H20-NaC1 at elevated temperatures. Chem. Geol. (Isot. Geosc. Sect.), 102: 199216. [46] Cole D. R. & Ripley E. M. (1998) Oxygen isotope fractionation between chlorite and water from 170 to 350~ a preliminary assessment based on partial exchange and fluid/rock experiments. Geochim. Cosmochim. Acta, 63: 449-457. [46] Cole D. R. (2000) Isotopic exchange in mineral-fluid systems. IV. The crystal chemical controls on oxygen isotope exchange rates in carbonate-H20 and layer silicate-H20 systems. Geochim. Cosmochim. Acta, 64: 921-931. [46] Cole D. R. & Chakraborthy S. (2001) Rates and mechanisms of isotopic exchange. Rev. Mineral. Geochem., 43: 83-223. [46] Coleman D. C. & Fry B. (1991) Carbon isotope techniques. Isotopic techniques in plant, soil, and aquatic biology series. E. A. Paul & J. M. Melillo (series eds.). Academic Press, San Diego: 274 p. [9]

References

1075

Coleman D. D., Risatti J. B. & Schoell M. (1981) Fractionation of carbon and hydrogen isotopes by methane-oxidizing bacteria. Geochim. Cosmochim. Acta, 45: 1033-1037. [46] Coleman J. J., Xu x., Thiemens M. H. & Trogler W. C. (1996) Photopolymerization and mass independent fractionations in carbon disulfide. Science, 273: 774-776. [18] Coleman M. L. & Moore M. P. (1978) Direct reduction of sulfates to sulfur dioxide for isotopic analysis. Anal. Chem., 50: 1594-1595. [26, 44] Coleman M. L. (1980) Corrections for mass spectrometer analysis of SO2.Stable Isotope Report, Inst. Geol. Sci., London: 12 p. [44, 45] Coleman M. L., Shepherd T. J., Durham J. J., Rouse J. E. & Moore G. R. (1982) Reduction of water with zinc for hydrogen isotope analysis. Anal. Chem., 54: 993-995. [1, 3, 21, App C] Coleman R. G. & Delevaux M. H. (1957) Occurrence of selenium in sulfides from some sedimentary rocks of the western United States. Econ. Geol., 52(5): 499-527. [29] Collister J. W., Rieley G., Stern B., Eglinton G. & Fry B. (1994) Compound-specific 812C analyses of leaf lipids from plants with differing carbon dioxide metabolisms. Org. Geochem., 21: 619-627. [9] Colman A. S., Karl D. M., Fogel M. L. & Blake R. E. (2000) A new technique for the measurement of phosphate oxygen isotopes of dissolved inorganic phosphate in natural waters. EOS Trans. AGU, 81(48, Fall Meeting Suppl.): 176 (Abstract). [22] Compston W. (1996) SHRIMP: origins, impact and continuing evolution. J. Royal Soc. West. Aust., 79: 109-117. [35] Condon A. G., Richards R. A. & Farquhar G. D. (1992) The effect of variations in soil water availability, vapour pressure deficit and nitrogen nutrition on carbon isotope discrimination in wheat. Aust. J. Agric. Res., 43: 935-947. [46] Connie J. M. & Currie V. A. (1996) The isotopic characterization of methane, non-methane hydrocarbons and formaldehyde in the troposphere. Atmosph. Environ., 30(4): 621-638. [14] Conway C. M. & Taylor H. P. Jr. (1969) O18/O16 and C 13/C 12 ratios of coexisting minerals in the Oka and Magnet Cove carbonatite bodies. J. Geol., 77: 618-626. [20] Conway E. J. (1947) Microdiffusion analysis and volumetric error. Crosby, Lockwood, London. [15] Conway T. J., Tans P. P., Waterman L. S., Thoning K. W., Kitzis D. R., Massarie K. A. & Zhang N. (1994) Evidence for interannual variability of the carbon cycle for the National Oceanic and Atmospheric / Climate monitoring diagnostics Laboratory Global Network. J. Geophys. Res., 99(Dll): 22831-22855. [14] Cooper D. E., Martinelli U., Carlisle C. B., Riris H., Bour D. B. & Menna R. J. (1993) Measurement of 12CO2:13CO2 ratios for medical diagnostics with 1.6-~m distributed-feedback semiconductor diode lasers. Appl. Opt., 32: 6727-6731. [34] Cooper G. W., Thiemens M. H., Jackson T. L. & Chang S. (1997) Sulfur and hydrogen isotope anomalies in meteorite sulfonic acids. Science, 277: 1072-1074. [11, 18] Cooper L. W., DeNiro M. J. & Keeley J. E. (1991) The relationship between stable oxygen and hydrogen isotope ratios of water in astomatal plants. In: Stable Isotope Geochemistry: A tribute to Samuel Epstein. H. P. Jr. Taylor, J. R. O'Neil & I. R. Kaplan (eds.). The Geochemical Society, Spec. Publ. 3: 247255. [46] Cooper W. C. & Glover J. R. (1974) The toxicology of selenium and its compounds. In: Selenium. R. A. Zingaro & W. C. Cooper (eds). Van Nostrand Reinhold Co: 654-674. [29] Coplen T. B. & Hanshaw B. B. (1973) Ultrafiltration by a compacted clay membrane - 1. Oxygen and hydrogen isotopic fractionation. Geochim. Cosmochim. Acta, 37: 2295-2310. [2] Coplen T. B. & Kendall C. (1982) Preparation and stable isotope determination of NBS-16 and NBS-17 carbon dioxide reference samples. Anal. Chem., 54: 2611-2612. [1, 40, 43] Coplen T. B., Kendall C. & Hopple J. (1983) Comparison of stable isotope reference samples. Nature, 302: 236-238. [12, 14, 40] Coplen T. B. (1988) Normalization of oxygen and hydrogen isotope data. Chem. Geol. (Isot. Geosc. Sect.), 72: 293-297. [1, 38] Coplen T. B., Wildman J. D. & C h e n J. (1991) Improvements in the gaseous hydrogen-water equilibration technique for hydrogen isotope ratio analysis. Anal. Chem., 63: 910-912. [1, 21] Coplen T. B., Krouse H. R. & Biihlke J. K. (1992) Reporting of nitrogen-isotope abundances. Pure Appl. Chem., 64(6): 907-908. [43]

1076

References

Coplen T. B. (1994) Reporting of stable hydrogen, carbon, and oxygen isotopic abundances. Pure & Appl. Chem., 66(2): 273-276. [21] Coplen T. B. & Harper I. T. (1994) An improved technique for the 2H/1H analysis of urines from diabetic volunteers. Biol. Mass Spectrom., 23: 437-439. [1] Coplen T. B. (1996a) New guidelines for reporting stable hydrogen, carbon, and oxygen isotope-ratio data. Geochim. Cosmochim. Acta, 60(17): 3359-3360. [14, 43] Coplen T. B. (1996b) Atomic weights of the elements 1995. Pure Appl. Chem., 68: 2339-2359. [40] Coplen T. B. (1996c) Editorial: More uncertainty than necessary. Paleoceanography, 11(4): 369-370 [40] Coplen T. B., De Bii~vre P., Krouse H. R., Vocke R. D. J., Griining M. & Rozanski K. (1996) Ratios for light-element isotopes standardized for better interlaboratory comparison. EOS Trans. AGU, 77(27): 255-256. [40] Coplen T. B. & Krouse H. R. (1998) Sulphur isotope data consistency improved. Nature, 392: 32. [26, 27] Coplen T. B. (2000) Laboratory information management system (LIMNS) for light stable isotopes. U.S. Geological Survey. Open-File Report 00-345:121 p. [15] Coplen T. B. (2001a) Manual for operation of an isotope hydrology laboratory. IAEA Guidelines, IAEA, Vienna. [1] Coplen T. B. (2001b) Atomic weights of the elements 1999. Pure Appl. Chem., 73: 667-683. [40] Coplen T. B., Hopple J. A., Bi~hlke J. K., Peiser H. S., Rieder S. E., Krouse H. R., Rosman K. J. R., Ding T., Vocke R. D. Jr., R4v6sz K. M., Lamberty A., Taylor P. D. P. & De Bi6vre P. (2002) Compilation of minimum and maximum isotope ratios of selected elements in naturally occurring terrestrial materials and reagents. U.S. Department of the Interior- U.S. Geological Survey, Reston, Virginia. WaterResources Investigations Report 01-4222:98 p. [26, 40] Cormie A. R., Luz B. & Schwarcz H. P. (1994) Relationship between the hydrogen and oxygen isotopes of deer bone and their use in the estimation of relative humidity. Geochim. Cosmochim. Acta, 58: 34393449. [46] Cornwell J. C. & Morse J. W. (1987) The characterization of iron sulfide minerals in anoxic marine sediments. Marine Chem., 22: 193-206. [26] Corso N. T. & Brenna J. T. (1997) High-precision position-specific isotope analysis. Proc. Natl. Acad. Sci., U.S.A., 94: 1049-1053. [8] Cortecci G. & Longinelli A. (1970) Isotopic composition of sulfate in rain water, Pisa, Italy. Earth Planet. Sci. Lett., 8: 36-40. [26] Council N. R. (1981) Prudent practices for handling hazardous chemicals in laboratories. National Academic Press, Washington, DC. [9] Cowie G. L. & Hedges J. I. (1984) Determination of neutral sugars in plankton, sediments, and wood by capillary gas chromatography of equilibrated isomeric mixtures. Anal. Chem., 56: 497-504. [9] Craig H. (1953) The geochemistry of the stable carbon isotopes. Geochim. Cosmochim. Acta, 3: 53-92. [40] Craig H. (1957) Isotopic standards for carbon and oxygen and correction factors for mass-spectrometric analysis of carbon dioxide. Geochim. Cosmochim. Acta, 12: 133-149. [1, 14, 15, 19, 23, 38, 40, 43, 44] Craig H. (1961) Standard for reporting concentrations of deuterium and oxygen-18 in natural waters. Science, 133: 1833-1834. [40, 44] Craig H. & Gordon L. I. (1965) Deuterium and oxygen-18 variations in the ocean and the marine atmosphere. In: Stable isotopes in oceanographic studies and palaeotemperatures. E. Tongiorgio (ed.). Spoleto Conference, Italy. Laboratorio di Geologia Nucleare, Pisa: 9-130. [1, 21] Craig H., Marti K. & Wiens R. (1993) A static mass spectrometer with triple collection for nitrogen and neon isotopes. Scripps Inst. Oceanogr., San Diego. Techn. Rep. No93-11. [17] Creaser R. A., Papanastassiou D. A. & Wasserburg G. J. (1991) Negative thermal ion mass spectrometry of osmium, rhenium, and iridium. Geochim. Cosmochim. Acta, 55(1): 397-401. [29] Cremonini M. A., Tacconi D., Clementi V. & Luchinat C. (1998) Accurate determination of deuterium/ hydrogen ratios in natural organic compounds through a nuclear magnetic resonance time-domain reference convolution method: application to ethanol from three botanical sources and critical analysis of systematic inaccuracies in previous method. J. Agric. Food Chem., 46: 3943-3946. [5] Criss R. E., Gregory R. T. & Taylor H. P. Jr. (1987) Kinetic theory of oxygen isotope exchange between minerals and water. Geochim. Cosmochim. Acta, 51: 1099-1108. [46] Criss R. E. (1999) Principles of Stable Isotope Distribution. Oxford University Press, Oxford, UK. [46]

References

1077

Croce de Cobos A. E. & Troe J. (1984) High-pressure range of the recombination O + 02 --+ 03. Int. J. Chem. Kinet., 16: 1519-1530. [18] Cronin J. R., Pizzarello S. & Cruikshank D. P. (1988) Organic matter in carbonaceous chondrites, planetary satelites, asteroids and comets. In: Meteorites and the early solar system. J. F. Kerridge & M. S. Mathews (eds.). Univ. Arizona Press: 819-857. [11] Cronin J. R. & Pizzarella S. (1990) Aliphatic hydrocarbons of the Murchison meteorite. Geochim. Cosmochim. Acta, 54: 2859-2868. [11] Cronin J. R. & Chang S. (1993) Organic matter in meteorites: Molecular and isotopic analyses of the Murchison meteorite. In: Chemistry of life's origins. J. M. Greenburg & V. Pirronello (eds.). Kluwer: 209-258. [11] Cronin J. R., Pizzarello S., Epstein S. & Krishnamurthy R. V. (1993) Molecular and isotopic analyses of the hydroxy acids, dicarboxylic acids and hydroxylicarboxylic acids of the Murchison meteorite. Geochim. Cosmochim. Acta, 57: 4745-4752. [11] Cross W. G. (1951) Two-directional focussing of charged particles with a sector-shaped, uniform magnetic field. Rev. Sci. Instrum., 22: 717-722. [38] Crosson E. R., Ricci K. N., Richman B. A., Chilese F. C., Owano T. G., Provencal R. A., Todd,M. W., Glasser J., Kachanov A. A. & Paldus B. A. (2002) Stable Isotope Ratios Using Cavity Ring-Down Spectroscopy: Determination of 13C/12C for Carbon Dioxide in Human Breath. Anat. Chem., 74: 2003-2007. [34] Crowe D. E., Valley J. W. & Baker K. L. (1990) Micro-analysis of sulfur-isotope ratios and zonation by laser microprobe. Geochim. Cosmochim. Acta, 54: 2075-2092. [20, 26, 45] Crowe D. E. & Vaughan R. G. (1996) Characterization and use of isotopically homogeneous standards for in situ laser microprobe analysis of 34S/32S ratios. Am. Mineralagist, 81: 187-193. [45] Crowson R. A., Showers W. J., Wright E. K. & Hoering T. C. (1991) Preparation of phosphate samples for oxygen isotope analysis. Anal. Chem., 63: 2397-2400. [20, 22] Cuevas J., Villar M. V., Fern~ndes A. M., G6mez P. & Martin P. L. (1997) Pore waters extracted from compacted bentonite subjected to simultaneous heating and hydration. Appl. Geochem., 12: 473-481.

[2]

Cumrning G. L. (1973) Propagation of experimental errors in lead isotope ratio measurements using the double spike method. Chem. Geol., 11(3): 157-165. [29] Curie I. (1921) Sur le poids atomique du chlore dans quelques mineraux. C. R. Acad. Sci. Paris, 172: 10251028. [28] Czamanske G. K. & Rye R. O. (1974) Experimentally determined sulfur isotope fractionations between sphalerite and galena in the temperature range 600~ to 275~ Econ. Geol., 69:17-25 [46]. D'Allessandro M., Bidoglio G., Mousty F., Sala Benito J. V. & Yllera De Llano A. (1997) Laboratory and in situ investigation of the isotopic ratio 79Br/81Br as a tracer of water migration in crystalline rocks. J. Hydrol., 193: 351-362. [28] D'Angela D. & Longinelli A. (1990) Oxygen isotopes in living mammal's bone phosphate: Further results. Chem. Geol. (Isot. Geosci. Sect.), 86: 75-82. [22, 46] Dahnke H., Kleine D., Urban W., GHering P. & Miirtz M. (2001) Isotopic ratio measurement of methane in ambient air using mid-infrared cavity leak-out spectroscopy. Appl. Phys., B 72: 121-125. [34] Dai S., Ren D., Tang Y. G., Shao L. & Li S. (2002) Distribution, isotopic variation and origin of sulfur in coals in the Wuda coalfield, Inner Mongolia, China. Int. J. Coal Geology, 51: 237-250. [26] Dallwitz M. J. (1970) Fractionation corrections in lead isotopic analysis. Chem. Geol., 6(4): 311-314. [29] Damon P. E. & Kulp J. L. (1958) Excess helium and argon in beryl and other minerals. Am. Mineralogist, 43: 433-459. [16] Danho D., Naulet N. & Martin G. J. (1992) Deuterium, carbon and nitrogen analysis of natural and synthetic caffeines. Authentication of coffee and coffee extracts. Analusis, 20: 179-184. [5] Dansgaard W. (1953) The abundabce of O18 in atmospheric water vapour. Tellus, 5: 461-469. [1] Darimont A., Burke E. & Touret J. (1988) Nitrogen-rich metamorphic fluids in Devonian metasediments from Bastogne, Belgium. Bull. Mineralogie, 111: 321-330. [16]

1078

References

Datta B. P., Khodade P. S., Parab A. R., Goyal A. H. & Chitambar S. A. (1992) Thermal ionization mass spectrometry of Li2BO2+ ions: determination of the isotopic abundance ratio of lithium. Int. J. Mass Spectrom. Ion Proc., 116: 87-114. [6] Dauphas N. & Marty B. (1999) Heavy nitrogen in carbonatites of the Kola peninsula: A possible signature of hte deep mantle. Science, 286: 2488-2490. [17] Dauphas N., Reisberg L. & Marry B. (2001) Solvent extraction, ion chromatography, and mass spectrometry of molybdenum isotopes. Anal. Chem., 73: 2613-2616. [31] David M. B., Mitchell M. J., Aldcorn D. & Harrison R. B. (1989) Analysis of sulfur in soil, plant and sediment material: sample handling and use of an automated analyzer. Soil Biol. Biochem., 21: 119124. [26] Davies S., Spanel P. & Smith D. (1997) Quantitative analysis of ammonia on the breath of patients in end-stage renal failure. Kidney Int., 52: 223-228. [4] Davies S., Spanel P. & Smith D. (2001a) A new 'online' method to measure increased exhaled isoprene in end-stage renal failure. Nephrol. Dial. Transplant., 16: 836-839. [4] Davies S., Spanel P. & Smith D. (2001b) Rapid measurement of deuterium content of breath following oral ingestion to determine body water. Physiol. Meas., 22: 651-659. [4] Davis A. & Olsen R. L. (1995) The geochemistry of chromium migration and remediation in the subsurface. Ground Water, 33: 759-768. [29] Dawson P.H. (1995) Quadrupole mass spectrometry and its application. American Institute of Physiscs, New York, USA: p. 143. [32] De Bi6vre P. J. & Debus G. H. (1969) Absolute isotope ratio determination of a natural boron standard. Int. J. Mass Spectrom. Ion Phys., 2: 15-23. [40] De Bi6vre P. & Barnes I. L. (1985) Table of the isotopic composition of the elements as determined by mass spectrometry. Int. J. Mass Spectrom. Ion Proc., 65: 211-230. [35] De Bi6vre P. (1990) Isotope dilution mass spectrometry: what can it contribute to accuracy in trace analysis? Fresenius J. Anal. Chem., 337: 766-771. [37] De Bi6vre P. (1994) Isotope dilution mass spectrometry. In: Trace element analysis in biological specimens. R. F. M. Herber & M. Stoeppler (eds.). Elsevier, Vol. 15: 169-183. [32, 37, 42] De Bi6vre P. & Valkiers S. (1994) Molar mass determination for NA (Avogadro) projects. Present status and potential for the future. Metrologia, 31: 245-249. [25] De Bi6vre P., Valkiers S. & Peiser H. S. (1994) New values of silicon reference materials, certified for isotope abundance ratios. J. Res. Nat. Ist. Stand. Techn., 99(2): 201-202. [25] De Bi6vre P., Lenaers G., Murphy T. J., Peiser H. S. & Valkiers S. (1995) The chemical preparation and characterization of specimens for "absolute" measurements of the molar mass of an element on silicon, for redeterminations of the Avogadro constant. Metrologia, 32: 103-110. [25, 42] De Bolle W., Alonso A., Verbruggen A. & Wellum R. (1999) The blending of UF6 for the preparation of isotopic reference materials with specified 234U, 235U and 236U abundances. In: Symposium on safeguards and nuclear materials management. C. Foggi & E. Petraglia (EC-JRC-Ispra)(eds.). Proc. of the 21st ESARDA meeting. EC-JRC publication EUR-18963 EN: 585-587. [32] De Endredy A. S. (1963) Estimation of free iron oxides in soils and clays by photolytic method. Clay Minerals Bull., 5: 209-217. [2] De Gendt S., Schelles W., Van Grieken R. & Muller V. (1995a) Quantitative analysis of iron-rich and other oxide-based samples by means of glow discharge mass spectrometry. J. Anat. At. Spectrom., 10: 681-687. [35] De Gendt S., Van Grieken R., Hang W. & Harrison W. W. (1995b) Comparison between direct current and radiofrequency glow discharge mass spectrometry for the analysis of oxide-based samples. J. Anat. At. Spectrom., 10: 689-695. [35] DeLa Rocha C. L., Brzezinski M. A. & DeNiro M. J. (1996) Purification, recovery, and laser-driven fluorination of silicon from dissolved and particulate silica for the measurement of natural stable isotope abundances. Anal. Chem., 68(21): 3746-3750. [20, 25] De La Rocha C. L., Brzezinski M. A. & DeNiro M. J. (1997) Fractionation of silicon isotopes by marine diatoms during biogenic silica formation. Geochim. Cosmochim. Acta, 61: 5051-5056. [46] De La Rocha C. L., Brzezinski M. A. & DeNiro M. J. (2000) A first look at the distribution of the stable isotopes of silicon in natural waters. Geochim. Cosmochim. Acta, 64(14): 2467-2477. [20, 46]

References

1079

De La Rocha C. L. & DePaolo D. J. (2000) Isotopic evidence for variations in the marine calcium cycle over the Cenozoic. Science, 289: 1176-1178. [46] De La Rocha C. L. (2003) Silicon isotope fractionation by marine sponges and the reconstruction of the silicon isotope composition of ancient deep water. Geology, 31: 423-426. [20] De Laeter J. R. & Kennedy A. K. (1998) A double focussing mass spectrometer for geocghronology. Int. J. Mass Spectrom. Ion Proc., 178: 43-50. [35] De Laeter J. R. (2001) Applications of inorganic mass spectrometry. John Wiley & Sons, NY. [39] De Wit J. C., Van der Straaten C. M. & Mook W. G. (1980) Determination of the absolute hydrogen isotopic ratio of V-SMOW and SLAP. Geostand. Newslett., 4(1): 33-36. [40] de Wit M. J., Hart R., Martine A. & Abbott P. (1982) Archean abiogenic and probable biogenic structures associated with mineralised hydrothermal vent systems and regional metasomatism, with implications for greenstone belt studies. Econ. Geol., 77: 1783-1803. [13] Debeer E., Delange C. A. & Westwood N. P. C. (1992) Resonance-enhanced multiphoton-ionization photoelectron-spectroscopy of np and nf Rydberg states of atomic nitrogen. Phys. Rev., A 46: 56535665. [33] Dechan M. (1886) Detection and estimation of iodine, bromine, and chlorine. J. Chem. Soc., 49: 682-685.

[28]

Decitre S., Deloule E., Reisberg L., James R., Agrinier P. & Mevel C. (2002) Behavior of Li and its isotopes during serpentinization of oceanic peridotites. Geochem. Geophys. Geosyst., 3(1): doi 10.1029/ 2001GC000178. [6, 30] Dedina J. & Tsalev D. L. (1995) Hydride generation atomic absorption spectrometry. Wiley. [29] Deiana M., Corongiu F. P., Dessi M. A., Scano P., Casu M. & Lai A. (2001) NMR determination of sitespecific deuterium distribution (SNIF-NMR) in squalene from different sources. Magn. Reson. Chem., 39: 29-32. [5] Deines P. D., Langmuir D. & Harmon R. S. (1974) Stable carbon isotope ratios and the existence of a gas phase in the evolution of carbonate ground waters. Geochim. Cosmochim. Acta, 38: 1147-1164. [10] DeLong W. A. & Lowe L. E. (1961) Note on carbon-bonded sulphur in soil. Can. J. Soil Sci., 42: 223. [26] Deloule E., All6gre C. J. & Doe B. (1986) Lead and sulphur isotope microstratigraphy in galena crystals from Mississippi-valley type deposits. Econ. Geol., 81: 1307-1321. [26] Deloule E., France-Lanord C. & Albar6de F. (1991) D / H analysis of minerals by ion probe. In: Stable Isotope Geochemistry: A tribute to Samuel Epstein. H. P. Jr. Taylor, J. R. O'Neil & I. R. Kaplan (eds.). The Geochemical Society, Spec. Publ. 3: 53-62. [30] Dem6ny A. (1995) H isotope fractionation due to hydrogen-zinc reactions and its implications on D / H analysis of water samples. Chem. Geol., 121: 19-25. [1, App C] Demir I. & Harvey R. D. (1991) Variation of organic sulfur in macerals of selected Illinois Basin coals. Org. Geochem., 17: 525-533. [26] Demmelmair J. & Schmidt H.-L. (1993) Precise 613C-determination in the range of natural abundance of amino acids from protein hydrolysates by gas chromatography-mass spectrometry. Isot. Environ. Health Stud., 29: 237-250. [8] Dempster A. J. (1935) New methods in mass spectroscopy. Proceed. Am. Phil. Soc., 75: 755-767. [35, 38] Demtri3der W. (1982) Laser spectroscopy. Basic concepts and instrumentation. 2nd edition. Springer Verlag, Berlin. [34] DeNiro M. J. & Epstein S. (1978) Influence of diet on the distribution of stable isotopes in animals. Geochim. Cosmochim. Acta, 42: 495-506. [46] DeNiro M. J. & Epstein S. (1979) Relationship between the oxygen isotope ratios of terrestrial plant cellulose, carbon dioxide and water. Science, 204: 51-53. [21, 46] DeNiro M. J. (1981) The effects of different methods of preparing cellulose nitrate on the determination of the D / H ratios of non-exchangeable hydrogen of cellulose. Earth Planet. Sci. Lett., 54: 177-185. [24] DeNiro M. J. & Epstein S. (1981a) Isotopic composition of cellulose from aquatic organisms. Geochim. Cosmochim. Acta, 45: 1885-1894. [46] DeNiro M. J. & Epstein S. (1981b) Influence of diet on the distribution of nitrogen isotopes in animals. Geochim. Cosmochim. Acta, 45: 341-351. [46]

1080

References

DeNiro M. J. & Cooper L. W. (1989) Post-photosynthetic modification of oxygen isotope ratios of carbohydrates in the potato: implications for paleoclimatic reconstruction based upon isotopic analysis of wood cellulose. Geochim. Cosmochim. Acta, 53: 2573-2580. [46] Dennis P. F., Rowe P. J. & Atkinson T. C. (1998) Stable isotope composition of palaeprecipitation and palaeogroundwaters from speleothem fluid inclusions. In: Isotope techniques in the study of environmental change. IAEA, Vienna: 663-671. [3] Dennis P. F., Rowe P. J. & Atkinson T. C. (2001) The recovery and isotopic measurement of water from fluid inclusions in speleothem. Geochim. Cosmochim. Acta, 65(6): 871-884. [3] Deqiu J. & Zhengxin W. (1985) Determination of the deuterium content in water with Mg-Pt reduction method. Acta Scient. Natur. Univers. Pekin., 1:42-46 (in Chinese; English translation by R. Krouse). [1] Dequasie H. L. & Grey D. C. (1970) Stable isotopes applied to pollution studies. Am. Lab., 2: 19-27. [26] Derdau D., Freer R. & Wright K. (1998) Oxygen diffusion in anhydrous sanidine feldspar. Contrib. Mineral. Petrol., 133: 199-204. [46] Derwish G. A. W., Galli A., Giardini-Guidoni A. & Volpi G. G. (1964) High-pressure mass spectrum of N20. J. Chem. Phys., 40: 3450-3451. [19] Desage M., Guilluy R., Brazier J. L., Chaudron H., Girard J., Cherpin H. & Jumeau J. (1991) Gaschromatography with mass-spectrometry or isotope-ratio mass-spectrometry in studying the geographical origin of heroin. Anal. Chim. Acta, 247: 249-254. [8] Desaulniers D. E., Kaufmann R. S., Cherry J. A. & Bentley H. W. (1986) 37C1-35C1 variations in a diffusion-controlled groundwater system. Geochim. Cosmochim. Acta, 50: 1757-1764. [28, 46] DesMarais D. J. (1978a) Carbon, nitrogen, and sulfur in Apollo 15, 16, and 17 rocks. Proceedings of the 9th Lunar Planet. Sci. Conf.: 1451-2467. [13] DesMarais D. J. (1978b) Variable-temperature cryogenic trap for the separation of gas mixtures. Anal. Chem., 50: 1405-1406. [20] DesMarais D. J., Mitchell J. M., Meinschein W. G. & Hayes J. M. (1980) The carbon isotope biogeochemistry of the individual hydrocarbons in bat guano and the ecology of the insectivorous bats in the egion of Carlsbad, New Mexico. Geochim. Cosmochim. Acta, 44: 2075-2086. [9] DesMarais D. J. & Moore J. G. (1984) Carbon and its isotopes in mid-oceanic basaltic glasses. Earth Planet. Sci. Lett., 69: 43-57. [3] Dewar W. A. & McDonald P. (1961) Determination of dry matter in silage by distillation with toluane. J. Sci. Food Agric., 12: 790-795. [21] Dews J. R. (1966) The isotopic composition of lithium in chondrules. J. Geophys. Res., 71: 4011-4020. [6] Deyhle A. (2001) Improvements of boron isotope analysis by positive thermal ionization mass spectrometry using static multicollection of Cs2BO2+ ions. Int. J. Mass Spectrom., 206: 79-89. [40] Dhamala B. R. & Mitchell M. J. (1992) Utilization of 35S in analyzing sulfur transformations and fluxes in forest soils. Proceedings of NHRI Symposium on Sulphur Transformations in Soil Ecosystems. Saskatoon (Canada): 297-314. [26] Dhamenlincourt P., Beny J. M., Dubessy J. & Poty B. (1979) Analysis of fluid inclusions with the MOLE Raman microprobe. Bull. Mindralogie, 102: 600-610. [3] Dickman M. & Thode H. (1990) Sulphur bacteria and sulphur isotope fractionation in a meromictic lake near Toronto, Canada. In: Facets of modern biogeochemistry. V. Ittekot, S. Kempe, W. Michaelis & A. Spitzy (eds.). Springer Verlag: 225-241. [29] Dickson F. W., Blount C. W. & Tunell G. (1963) Use of hydrothermal solution equipment to determine the solubility of anhydrite in water from 100~ to 275~ and from I bar to 1000 bars pressure. Am. J. Sci., 261: 61-78. [46] Dickson J. A. D. (1991) Disequilibrium carbon and oxygen isotope variations in natural calcite. Nature, 353: 842-844. [22] Diemer J., Qu4tel C. R. & Taylor P. D. P. (2002) Comparison of the performance of different ICP-MS instruments on the measurement of Cu in a water sample by ICP-IDMS. J. Anal. At. Spectrom., 17: 1137-1142. [32] Dimeff, J. (1972) Nondispersive gas analyzing method and apparatus wherein radiation is serially passed through a reference and unknown gas. U.S. Patent 3,679,899. [34]

References

1081

Ding T.-P., Li-Hong, Zhang-Guobing, Li-u & Fan-Tianyi (1985) Sulphur isotope determination by using the SF6 method. Chinese Academy of Geological Sciences, Beijing, China. Annual Rept, p.21: 163164. [20] Ding T.-P., Li-Hong, Zhang-Guobing, Li-Yanhe & Li-Jincheng (1987) A study on determining 33S/32S, 34S/ 32S and 36S/ 32S by SF6 method. Mineral Deposits, 6: 81-93. [20] Ding T., Wan D., Li J., Song H., Li Y. & Liu Z. (1988) The analytical method of silicon isotopes and its geological application. Mineral Deposits, 7(4): 90-95 (in Chinese). [20, 25] Ding T.-P. & Zhang Q.-L. (1989) Analysis of the reference material NBS-123 and the atomic weight of sulfur. Chin. Sci. Bull., 34(13): 1086-1089. [45] Ding T., Wan D., Jiang S. & Liu Z. (1990) A study on silicon compositions of some rocks and minerals. ICOG-7, Canberra: 27 (Abstract). [25] Ding T. (1991) A study on analytic method of silicon isotope and its geological application. 2nd ChinaJapan joint symposium on geochronology, cosmochronology and isotope geology, Guangzhou, China: 11 (Abstract). [25] Ding T., Jiang S., Wan D. & Li Y. (1994) The silicon isotope variations in geochemical circulation. ICOG8, Berkeley. U.S. Geol. Surv. Circular 1107:82 (Abstract). [25] Ding T., Jiang S., Was D., Li Y., Li J., Song H., Liu Z. & Yao X. (1996) Silicon isotope geochemistry. Geological Publishing House, Beijing, China: 125 p. [20, 25] Ding T., Wang C., Zhang F. & Wan D. (1998) Silicon isotope fractionation in some surface processes (A study on river water and plants). ICOG-9, Beijing, Chin. Sci. Bull., 43 (Supplem.): 33 (Abstract). [25] Ding T., Bai R., Li Y., Wan D., Zou X. & Zhang Q. (1999) Determination of the absolute 32S/34S ratio of IAEA-S-1 reference material and V-CDT sulfur isotope standard. Sci. China, Ser. D (Earth Sciences), 42: 45-51. [40, 44] Ding T., Valkiers S., Kipphardt H., Qu6tel C., De Bi6vre P., Taylor P. D. P. & Gonfiantini R. (2000) Preparation of two synthetic mixtures for the calibration of isotope amount ratio measurements of sulfur. Int. J. Mass Spectrom., 197: 131-137. [42] Ding T., Valkiers S., Kipphardt H., De Bi6vre P., Taylor P. D. P., Gonfiantini R. & Krouse R. (2001) Calibrated sulfur isotope abundance ratios of three IAEA sulfur isotope reference materials and VCDT with a reassessment of the atomic weight of sulfur. Geochim. Cosmochim. Acta, 65(15): 2433-2437. [26, 27, 40, 42] Ding T., Wan D., Wang C. & Zhang F. (2002) Si, O, H isotopic compositions of water and suspended materials in the Yangtze River. Geochim. Cosmochim. Acta, 66:A185 (Abstract). [20] Ding T., Wan D., Wang C. & Zhang F. (2003) Silicon isotope compositions of dissolved silicon and suspended matter in the Yangtze River, China. Geochim. Cosmochim. Acta, 67: in press. [20] Dittel A. I., Epifanio C. E., Cifuentes L. A. & Kirchman D. L. (1997) Carbon and nitrogen sources for shrimp postlarvae fed natural diets from a tropical mangrove system. Est. Coast. Shelf Sci., 45: 629637. [9] Divis A. F. & Clark J. R. (1978) Exploration for blind ore deposits and geothermal reservoirs by lithium isotope thermometry- atomic absorption "mass spectrometry". Assoc. Explor. Geochem. Proceedings of the Seventh International Geochemical Exploration Symposium, Rexdale, Ontario, Canada: 233241. [6] Dixon P. R., Perrin R. E., Rokop D. J., Maeck R., Janecky D. R. & Banar J. P. (1993) Measurements of iron isotopes (54Fe, 56Fe, 57Fe, 58Fe) in sub microgram quatities of iron. Anal. Chem., 65: 2125-2130. [29] Dlugokencky E. J., Masarie K. A., Lang P. M. & Tans P. P. (1998) Continuing decline in the growth rate of the atmospheric methane burden. Nature, 393(No. 6684): 447-450. [14] Dobson P. F., Epstein S. & Stolper E. M. (1989) Hydrogen isotope fractionation between coexisting vapor and silicate glasses and melts at low pressure. Geochim. Cosmochim. Acta, 53: 2723-2730. [46] Dodson M. H. (1970) Simplified equations for double-spiked isotopic analyses. Geochim. Cosmochim. Acta, 34(11): 1241-1244. [29] Doering W. v. E. & Dorfman E. (1953) Mechanism of the peracid ketone-ester conversion. Analysis of organic compounds for oxygen-18. J. Am. Chem. Soc., 75: 5595-5598. [23] Dongmann G., Niirnberg H. W., Fi~rstel H. & Wagener K. (1974) On the enrichment of H2180 in the leaves of transpiring plants. Radiat. Environm. Biophys., 11: 41-52. [21]

1082

References

D o n o h u e D. L. & Harrison W. W. (1975) Radiofrequency cavity ion source in solids mass spectrometry.

Anal. Chem., 47: 1528-1531. [35] Donohue D. L. & Petek M. (1991) Isotopic measurement of palladium metal containing protium and deuterium by glow discharge mass spectrometry. Anal. Chem., 63: 740-744. [35] Dontsova E. I. (1959) Method of determination of oxygen isotope ratios in igneous rocks and minerals. Geokhimiya, 8:669-678 (in Russian). [20] Dore J. E., Popp B. N., Karl D. M. & Sansone F. J. (1998) A large source of atmospheric nitrous oxide from subtropical North Pacific surface waters. Nature, 396: 63-66. [19] Dostrovsky I. & Klein F. S. (1952) Mass spectrometric determination of oxygen in water samples. Anal. Chem., 24: 414-415. [1] Doucett R. R., Hooper W. & Power G. (1999) Identification of anadromous and nonamadromous adult brook trout and their progeny in the Tabusintac River, New Brunswick, by means of multiple stable isotope analysis. Trans. Am. Fish. Soc., 128: 278-288. [9] Douglas D. J. & Tanner S. D. (1998) Fundamental considerations in ICPMS. In: Inductively coupled plasma mass spectrometry. A. Montaser (ed.). Wiley-VCH: 615-679. [31] Dourson M., Stern B., Griffin S. & Bailey K. (1991) Impact of risk-related concerns on the U.S. Environmental Protection Agency Programs. In: Nitrate contamination:Exposure, consequence, and control. I. Bogardi & R. D. Kuzelka (eds.). Springer Vertag, Berlin. NATO SI Ser. G: Ecological Sciences, Vol. 30: 477-487. [15] Douthitt C. B. (1982) The geochemistry of the stable isotopes of silicon. Geochim. Cosmochim. Acta, 46: 1449-1458. [20, 25] Dowdle P. R. & Oremland R. S. (1998) Microbial oxidation of elemental selenium in soil slurries and bacterial cultures. Environ. Sci. Technol., 32: 3749-3755. [29] Dowling L. M., Boreham C. J., Hope J. M., Murray A. P. & Summons R. E. (1995) Carbon isotopic composition of hydrocarbons in ocean-transported bitumens from the coastline of Australia. Org. Geochem., 23: 729-737. [8] Downey S. W., Emerson A. B., Kopf R. F. & Schubert E. F. (1992) Resonance ionization massspectrometry of device materials. Inst. Phys. Conf. Series, 128: 255-258. [33] Downs M. R., Michener R. H., Fry B. & Nadelhoffer K. J. (1999) Routine measurement of dissolved inorganic 15N in streamwater. Environ. Monitoring Assessment, 55: 211-220. [15] Dreibus G., Spettel B. & W~inke H. (1979) Halogens in meteorites and their primordial abundances. Origin and distribution of the elements. Proceedings of the second Symposium, Paris 1977. Pergamon Press: 33-38. [28] Drever J. I. (1982) The Geochemistry of Natural Waters. Prentice-Hall, Englewood Cliffs, NJ: 388 p. [26] Driesner T. (1997) The effect of pressure on deuterium-hydrogen fractionation in high-temperature water. Science, 277: 791-794. [46] Driesner T. & Seward T. M. (2000) Experimental and simulation study of salt effects and pressure/ density effects on oxygen and hydrogen stable isotope liquid-vapor fractionation for 4-5 molal aqueous NaC1 and KC1 solutions to 400~ Geochim. Cosmochim. Acta, 64(10): 1773-1784. [46] Du Y.-C., Nose A., Wasano K. & Uchida Y. (1998) Responses to water stress of enzyme activities and metabolite levels in relation to sucrose and starch synthesis, the Calvin cycle and the C4 pathway in sugarcane (Saccharumsp.) leaves. Aust. J. Plant Physiol., 25: 253-260. [9] Duan W.-M., Coleman M. L. & Pye K. (1997) Determination of reduced sulphur species in sediments an evaluation and modified technique. Chem. Geol., 141: 185-194. [26] Dubinina E. O. & Lakshtanov L. Z. (1997) A kinetic model of isotopic exchange in dissolution-precipitation processes. Geochim. Cosmochim. Acta, 61: 2265-2273. [46] Dubois A. (1985) An accurate zinc charcoal reduction system for D / H measurements of water and cellulose. Bull. Belg. Veren. Geol., 94: 113-115. [1] DuChateau N. L., Verbruggen A., Hendrickx F. & De Bi6vre P. (1986) Rapid accurate isotopic measurements on boron in boric acid and boron carbide. Talanta, 33: 291-294. [7] Duckworth D. C., Barshick C. M., Bostick D. A. & Smith D. H. (1993) Direct measurement of uranium isotopic ratios in soils by glow discharge mass spectrometry. Appt. Spectrosc., 47: 243-245. [35]

References

1083

Dugan J. P., Borthwick J., Harmon R. S., Gagnier M. A., Glahn J. E., Kinsel E. P., McLeod S., Viglino J. A. & Hess J. W. (1985) Guadinine hydrochloride method for determination of water oxygen isotope ratios and the oxygen-18 fractionation between carbon dioxide and water at 25~ Anal. Chem., 57: 1734-1736. [1, 3, 20] Dugan J. P. & Borthwick J. (1986) Carbon dioxide-water oxygen isotope fractionation factor using chlorine trifluoride and guanidine hydrochloride techniques. Anat. Chem., 58: 3052-3054. [1] Duit W., Jansen J. B. H., van Bremen A. & Bos A. (1986) Ammonium micas in metamorphic rocks as exemplified by Dome de L'Agout (France). Am. J. Sci., 286: 702-732. [16] Dumke I., Faber E. & Poggenburg J. (1989) Determination of stable carbon and hydrogen isotopes of light hydrocarbons. Anal. Chem., 61(19): 2149-2154. [43] Dungan R. S. & Frankenberger W. T. Jr. (1998) Reduction of selenite to elemental selenium by Enterobacter cloacae SLDla-1. J. Env. Qual., 27: 1301-1306. [29] Dunstan L. P., Gramlich J. W., Barnes I. L. & Purdy W. C. (1980) Absolute isotopic abundance and the atomic weight of a reference sample of thallium. J. Res. Natl. Bur. Stand., 85: 1-10. [31] Durakiewicz T. & Halas S. (1995) A novel electronic circuit for a Pirani gauge. Vacuum, 46: 101-103. [39] Durakiewicz T. (1996) Electron emission controller with pulsed heating of filament. Int. J. Mass Spectrom. Ion Proc., 156: 31-40. [39] Durand B., Espitali6 J., Nuaise G. & Combaz A. (1972) ]~tude de la mati6re organique insoluble (kerogene) des argiles du Joacien du bassin de Paris. Premiere partie: ]~tude par les proced6s optiques, analyse elementaire, 6tude en microscopie et diffraction 61ectroniques. Rev. L'Inst. Franfais du Petrol, 27: 865-884. [26] Durand B. & Nicaise G. (1980) Procedures for kerogen isolation. In: Kerogen- Insoluable organic matter from sedimentary rocks. G. Durand (ed.). Edition Technip, Paris: 113-142. [2] Durka W., Schulze E. D., Gebauer G & Voerkelius S. (1994) Effects of forest decline on uptake and leaching of deposited nitrate determined from 15N and 180 measurements. Nature, 372: 765-767. [15] Durrant S. F. (1999) Laser ablation inductively coupled plasma mass spectrometry: achievements, problems and prospects. J. Anal. At. Spectrom., 14: 1385-1403. [20] Eakin P. A., Fallick A. E. & Gerc J. (1992) Some instrumental effects in the determination of stable carbon isotope ratios by gas chromatography - isotope ratio mass spectrometry. Chem. Geol. (Isot. Geosc. Sect.), 101: 71-79. [8, 9] Eastoe C. J., Guilbert J. M. & Kaufmann R. S. (1989) Preliminary evidence for fractionation of stable chlorine isotopes in ore forming hydrothermal systems. Geology, 17: 285-288. [28] Eastoe C. J. & Guilbert J. M. (1992) Stable chlorine isotopes in hydrothermal systems. Geochim. Cosmochim. Acta, 56: 4247-4255. [28] Eberl D. D., Srodon J., Kralik M., Taylor B. E. & Peterman Z. E. (1990) Ostwald ripening of clays and metamorphic minerals. Science, 248: 474-477. [46] Eckelmann W. R., Broecker W. S., Whitlock D. W. & Allsup J. R. (1962) Implications of carbon isotopic composition of total organic carbon of some recent sediments and ancient oils. AAPG Bull., 46: 699704. [40] Ecker K. H. & Pritzkow W. (1994) Measurement of isotope ratios in solids by glow discharge mass spectrometry. Fresenius J. Anal. Chem., 349: 207-208. [35] Edgar A. D. (1973) Experimental Petrology: Basic Principles and Techniques. Oxford University Press, Oxford, UK: 217 p. [46] Edwards T. W. D. & Fritz P. (1986) Assessing meteoric water composition and relative humidity from 180 and 2H in wood cellulose: paleoclimatic implications for southern Ontario, Canada. Appl. Geochem., 1: 715-723. [46] Edwards T. W. D., Buhay W. M., Elgood R. J. & Jiang H. B. (1994) An improved nickel-tube pyrolysis method for oxygen isotope analysis of organic matter and water. Chem. Geol. (Isot. Geosci. Sect.), 114: 179-183. [1, 23] Eggenkamp H. G. M. (1994) 637C1. The geochemistry of chlorine isotopes. PhD Thesis, University of Utrecht, Utrecht, the Netherlands. Geologica Ultraiectina, No. 116:149 p. [28] Eggenkamp H. G. M., Kreulen R. & Koster van Groos A. F. (1995) Chlorine stable isotope fractionation in evaporites. Geochim. Cosmochim. Acta, 59: 5169-5175. [28, 46]

1084

References

Eggenkamp H. G. M. & Schuiling R. D. (1995) 637C1 variations in selected minerals: a possible tool for exploration. J. Geochem. Explor., 55: 249-255. [28] Eggenkamp H. G. M. & Koster van Groos A. F. (1997) Chlorine stable isotopes in carbonatites: evidence for isotopic heterogeneity in the mantle. Chem. Geol., 140: 137-143. [28, 46] Eggenkamp H. G. M. & Coleman M. L. (2000) The application of classical methods for the measurement of stable bromine isotopes in natural samples. Chem. Geol., 167: 393-402. [28] Ehhalt D. H. (1974) Vertical profiles of HTO, HDO, and H20 in the troposphere. NCAR Technical Note, NCAR-TN / STR-100. [1] Ehleringer J. R. & Osmond C. B. (1989) Stable isotopes. In: Plant physiological ecology. R. W. Pearcy, J. R. Ehleringer, H. A. Mooney & P. W. Rundel (eds.). Chapman and Hall, London: 281-299. [21] Ehleringer J. R. (1991) 13C/12C fractionation and its utility in terrestrial plant studies. In: Carbon isotope techniques. D. C. Coleman & B. Fry (eds.). Academic Press: 187-200. [9] Ehleringer J. R. & Dawson T. E. (1992) Water uptake by plants: perspective from stable isotope composition. Plant, Cell Environ., 15: 1073-1082. [21] Ehleringer J. R., Hall A. E. & Farquhar G. D. (1993) Stable isotopes and plant carbon-water relations. Academic Press, San Diego: 555p. [9] Ehleringer J. R. & Cook C. S. (1998) Carbon and oxygen isotope ratios of ecosystem respiuration along an Oregon conifer transect: preliminary observations based on small-flask sampling. Tree Physiol., 18: 513-519. [8] Ehleringer J. R., Casale J. F., Loft M. J. & V.L. Ford V. L. (2000) Tracing the geographical origin of cocaine. Nature, 408: 311-312, 2000. [8] Ehrlich H. L. (1996) Geomicrobiology. Marcel Dekker. [29] Eiler J. M., Baumgartner L. P. & Valley J. W. (1992) Intercrystalline stable isotope diffusion: a fast grain boundary model. Contrib. Mineral. Petrol., 112: 543-557. [46] Eiler J. M., Valley J. W. & Baumgartner L. P. (1993) A new look at stable isotope thermometry. Geochim. Cosmochim. Acta, 57: 2571-2583. [46] Eiler J. M., Graham C. W. & Valley J. W. (1997) SIMS analysis of oxygen isotopes: matrix effects in complex minerals and glasses. Chem. Geol., 138: 221-244. [30] Eiler J. M., Kitchen N. & Rahn T. A. (2000) Experimental constraints on the stable-isotope systematics of CO2 ice / vapor systems and relevance to the study of Mars. Geochim. Cosmochim. Acta, 64: 733-746. [46] Eiler J. M. & Kitchen M. (2001) Hydrogen-isotope analysis of nanomole (picomole) quantities of H20. Geochim. Cosmochim. Acta, 65: 4467-4479. [1] Eldridge C. S., Compston W., Williams I. S., Walshe J. L. & Both R. A. (1987) In situ microanalysis for 34S/ 32S ratios using the ion microprobe SHRIMP. Int. J. Mass Spectrom. Ion Proc., 76: 65-83. [20, 26, 30] Eldridge C. S., Compston W., Williams I. S., Both R. A., Walshe J. L. & Ohmoto H. (1988) Sulfur isotope variability in sediment hosted massive sulfide deposits as determined using the ion microprobe SHRIMP. I. An example from the Rammelsberg ore body. Econ. Geol., 83: 443-449. [26] Eldridge C. S., Compston W., Williams I. S. & Walshe J. L. (1989) Sulfur isotopic analyses on the SHRIMP ion microprobe. In: New frontiers in stable isotopic research: Laser probes, ion probes, and small-sample analysis. W. C. Shanks III & R. E. Criss (eds.). U.S. Geol. Surv., Bull. 1890: 163-174. [25] Ellis A. S., Johnson T. M. & Bullen T. D. (2002) Cr isotopes and the fate of hexavalent chromium in the environment. Science, 195: 2060-2062. [29] Ellis A. S., Johnson T. M., Bullen T. D. & Herbel M. J. (2003) Stable isotope fractionation of selenium by natural microbial consortia. Chem. Geol., 195: 119-129. [29] Ellis L. & Fincannon A. L. (1998) Analytical improvements in irm-GC/MS analyses: Advanced techniques in tube furnace design and sample preparation. Org. Geochem., 29: 1101-1117. [8] Elphick S. C., Dennis P. F. & Graham C. M. (1986) An experimental study of the diffusion of oxygen in quartz and albite using an overgrowth technique. Contrib. Mineral. Petrol., 92: 322-330. [46] Elphick S. C., Graham C. M. & Dennis P. F. (1988) An ion microprobe study of anhydrous oxygen diffusion in anorthite: a comparison with hydrothermal data and some geological implications. Contrib. Mineral. Petrol., 100: 490-495. [46] Elrashidi M. A., Adriano D. C., Workman S. M. & Lindsay W. L. (1987) Chemical equilibria of selenium in soils; a theoretical development. Soil Sci., 144(2): 141-152. [29]

References

1085

Elsenheimer D. & Valley J. W. (1992) In situ oxygen isotope analysis of feldspar and quartz by Nd:YAG laser micro probe. Chem. Geol. (Isot. Geosci. Sect.), 101: 21-42. [20] Elsenheimer D. & Valley J. W. (1993) Sub-millimeter scale zonation of 8180 in quartz and feldspar, Isle of Skye, Scotland. Geochim. Cosmochim. Acta, 57: 3669-3676. [20] Emiliani C. (1955) Pleistocene temperatures. J. Geol., 63: 538-578. [46] Engel M. H., Macko S. A. & Silfer J. A. (1990) Carbon isotope composition of individual amino acids in the Murchison meteorite. Nature, 348: 47-49. [9, 11] Engel M. H. & Macko S. A. (1997) Isotopic evidence for extraterrestrial non-racemic amino acids in the Murchison meteorite. Nature, 389: 265-268. [9, 11] Engeln R., yon Helden G., Berden G. & Meijer G. (1996) Phase shift cavity ring down absorption spectrometry. Chem. Phys. Lett., 262: 105-109. [34] Engeln R., Berden G., Peeters P. & Meijer G. (1998) Cavity enhanced absorption and cavity magnetic rotation apectroscopy. Rev. Sci. Instrum., 69: 3763-3769. [34] England J. G., Zindler A., Reisberg L. C., Salters V., Marcantonio F., Bourdon B., Brueckner H., Turner P. J., Weaver S. & Read P. (1992) The Lamont-Doherty Geological Observatory Isolab 54 isotope ratio mass spectrometer. Int. J. Mass Spectrom. Ion Proc., 121: 201-240. [30] Engrand C., McKeegan K. D. & Leshin L. A. (1999) Oxygen isotopic compositions of individual minerals in Antarctic micrometeorites: Further links to carbonaceous chondrites. Geochim. Cosmochim. Acta, 63(17): 2623-2636. [30] Ensminger L. E. (1954) Some factors affecting the adsorprtion of sulfate by Alabama soils. Soil Sci. Soc. Am. Proc., 18: 259-264. [26] Entwistle D. C. & Reeder S. (1993) New apparatus for pore fluid extraction from mudrocks for geochemical analysis. In: Geochemistry of clay- pore fluid interaction. D. A. C. Manning, P. L. Hall & C. R. Hughes (eds.). The Mineralogical Society Series 4, Chapman & Hall, London: 365-388. [2] Epstein S., Buchsbaum R., Lowenstam H. & Urey H. C. (1951) Carbonate-water isotopic temperature scale. Geol. Soc. Am. Bull., 62: 417-425. [46] Epstein S., Buchsbaum R., Lowenstam H. A. & Urey H. C. (1953) Revised carbonate-water isotopic temperature scale. Geol. Soc. Am. Bull., 64: 1315-1325. [40, 46] Epstein S. & Mayeda T. K. (1953) Variations of 180 of waters from natural sources. Geochim. Cosmochim. Acta, 4: 213-224. [1, 2, 20, 40] Epstein S. & Taylor H. P. Jr. (1970) Stable isotopes, rare gases, solar wind, and spallation products. 180 / 160, 30Si/28Si, D/H, 13C/12C studies of lunar rocks and minerals. Science, 167: 533-535. [25] Epstein S. & Taylor H. P. Jr. (1971) 180/160, 30Si/28Si, D / H and 13C/12C ratios in lunar samples. Proc. Second Lunar Sci. Conf., 2: 1421-1441. [20] Epstein S., Yapp C. J. & Hall J. H. (1976) The determination of the D / H ratio of non-exchangeable hydrogen in cellulose extracted from aquatic and land plants. Earth Planet. Sci. Lett., 30: 241-251. [24] Epstein S., Krishnamurthy R. V., Cronin J. R., Pizzarello S. & Yuen G. U. (1987) Unusual stable isotope ratios in amino acid and carboxylic extract from the Murchison meteorite. Nature, 326: 477-479. [11] Erd61yi M., Richter D. & Tittel F. K. (2002) 13CO2/12CO2 isotopic ratio measurements using a difference-frequency-based sensor operating at 4.35 ~m. Appl. Phys., B 75: 289-295. [34] Erdmann N., Griining C., Trautmann N., Waldek A., Huber G., Kunz P., Nennemann M. & Passler G. (1998) Trace analysis of plutonium by resonance ionization mass spectrometry. J.C. Vickerman, I. Lyon, N.P. Lokyer & J.E. Parks (eds.). Resonance Ionization Spectroscopy, Ninth International Symposium. AIP Press Woodbury, NY, AlP Conference Proceedings 454: 279-282. [33] Erez J. & Luz B. (1983) Experimental paleotemperature equation for plaktonic foraminifera. Geochim. Cosmochim. Acta, 47: 1025-1031. [46] Eriksen T. E. (1972) Sulfur isotope effects I. The isotopic exchange coefficient for the sulfur isotopes 34S32S in the system SO2-g-HSO3-aq at 25, 35, and 45~ Acta Chem. Scandinavia, 26: 573-580. [26] Esat T. M. (1984) A 61 cm radius multi-detector mass spectrometer at the Australian National University. Nucl. Instrum. Meth. Phys. Res., B5: 545-553. [30] Esat T. M. (1995) Charge collection thermal ionization mass spectrometry of thorium. Int. J. Mass Spectrom. Ion Proc., 148: 159-170. [30]

1086

References

Escande M. (1983) Echangeabilit6 et fractionnement isotopique de l'oxig6ne des smectites magnesiennes de synth6ses. Etablissement d'un g6othermom6tre. Th6se, Univ. Paris XI, Paris, France: 150 p. [2] Escande M., Decarreau A. & Labeyrie L. (1984) t~tude exp6rimentale de l'6changeabilit6 des isotopes de l'oxyg6ne des smectites. C. R. Acad. Sc. Paris, 299(series II, No. 11): 707-710. [2] Esler M. B., Griffith D. W. T., Turatti F., Wilson S. R., Rahn T. A. & Zhang H. (2000a) N20 concentration and flux measurements and complete isotopic analysis by FTIR spectroscopy. ChemosphereGlobal Change Sci., 2: 445-454. [19, 34] Esler M. B., Griffith D. W. T., Wilson S. R. & Steele L. P. (2000b) Precision trace gas analysis by FT-IR spectroscopy. 2. The 13C/12Cisotope ratio of CO2. Anal. Chem., 72: 216-221. [34] Eslinger E. V. (1971) Mineralogy and oxygen isotope ratios of hydrothermal and low grade metamorphic argillaceous rocks. PhD Thesis, Case Western Reserve University, Cleveland, Ohio, USA: 205 p. [2] Eslinger E. V. & Savin S. (1973) Mineralogy and oxygen isotope geochemistry of the hydrothermally altered rocks of the Ohaki-Broadlands, New Zealand geothermal area. Am. J. Sci., 273: 240-267. [46] Estep M. L. F. & Hoering T. C. (1980) Biogeochemistry of the stable hydrogen isotopes. Geochim. Cosmochim. Acta, 44: 1197-1206. [46] Estep M. L. F. & Hoering T. C. (1981) Stable hydrogen isotope fractionations during autotrophic and mixotrophic growth of microalgae. Plant Physiol., 67: 474-477. [46] Etheridge D. M., Pearman G. I. & Fraser P. J. (1992) Changes in tropospheric methane between 1841 and 1978 from a high accumulation-rate Antarctic ice core. Tellus, 44(Ser. B; 4): 282-294. [14] Etheridge D. M., Steele L. P., Francey R. J. & Langenfelds R. L. (1998) Atmospheric methane between 1000 AD and present: Evidence of anthropogenic emissions and climatic variability. J. Geophys. Res., 103(D13): 15979-15993. [14] Eugster H. P. (1957) Heterogeneous reactions involving oxidation and reduction at high pressures and temperatures. J. Chem. Phys., 26: 1760-1761. [46] EURACI-IEM (2000) Quantifying Uncertainty in Analytical Measurement. EURACHEM / CITAC Guide, 2nd edition: 120 p. [40, 42] Eugster O., Tera F. & Wasserburg G. J. (1969) Isotopic analyses of barium in meteorites and in terrestrial samples. J. Geophys. Res., 74(15): 3897-3908. [29] Evershed R. P., Arnot K. I., Collister J., Eglinton G. & Charters S. (1994) Application of isotope ratio monitoring gas-chromatography mass-spectrometry to the analysis of organic residues of archaeological origin. Analyst, 119: 909-914. [8] Evershed R. P., Stott A. W., Raven A., Dudd S. N., Charters S. & Leyden A. (1995) Formation of longchain ketones in ancient-pottery vessels by pyrolysis of acyl lipids. Tetrahedron Lett., 36: 8875-8878.

[81 Evershed R. P., Mottram H. R., Dudd S. N., Charters S., Stott A. W., Lawrence G. J., Gibson A. M., Conner A., Blinkhorn P. W. & Reeves V. (1997) New criteria for the identification of animal fats preserved in archaeological pottery. Naturwissensch., 84: 402-406. [8] Ewald A. H. (1985) The effect of pressure on oxygen isotope exchange in silicates. Chem. Geol., 49: 179185. [46] Exley R. A., Mattey D. P., Clague D. A. & Pillinger C. T. (1986) Carbon isotope systematics of a mantle "hotspot": a comparison of Loihi Seamount and MORB glasses. Earth Planet. Sci. Lett., 78: 189-199. [3] Fagan R. & Longstaffe F. J. (1997) Hydrogen-isotope exchange in smectite? 11th International Clay Conference, Ottawa, Canada (Abstract). [2] Fahey A. J. (1998) Measurements of dead time and characterization of ion counting systems for mass spectrometry. Rev. Sci. Instrum., 69: 1282-1288. [4] Fairbank W. M. Jr., Spaar M. T., Parks J. E. & Hutchinson J. M. R. (1989) Measurement of the odd-even effect in the resonance ionization of tin as a function of laser intensity. Phys. Rev., A40: 2195-2198. [33] Fairbank W. M. Jr. (1996) A seat-of-the-pants approach to RIS theory. N. Winograd & J.E. Parks (eds.). Resonance Ionization Symposium 1996, Eighth International Symposium. AIP Press Woodbury, NY, AIP Conference Proceedings 388: 457-464. [33] Fairbanks R. G. (1982) The origin of continental shelf and slope water in the New York Bight and Gulf of Maine: evidence from H2180/H2160 ratio measurements. J. Geophys. Res., 87: 5796-5808. [1]

References

1087

Fairbanks R. G., Sverdlove M., Free R., Wiebe P. H. & B6 A.W.H. (1982) Vertical distribution and isotopic fractionation of living planktonic foraminifera from the Panama Basin. Nature, 298: 841-844. [46] Fang D. & Marcus R. K. (1993) Fundamental plasma processes. In: Glow discharge spectroscopies. R. K. Marcus (ed.). Plenum Press, NY: 17-66. [35] Fang J., Barcelona M. J., Krishnamurthy R. V. & Atekwana E. A. (2000) Stable carbon biogeochemistry of a shallow sand aquifer contaminated with fuel hydrocarbons. Appl. Geochem., 15: 157-169. [10] Fantle M. S., Dittel A. I., Schwalm S. M., Epifano C. E. & Fogel M. L. (1999) A food web analysis of the juvenile blue crab, Callinectus, sapidus, using stable isotopes in whole animals and individual amino acids. Oecologia, 120: 416-426. [9] Farmer V. C. & Mitchell B. D. (1963) Occurence of oxalates in soil clays following hydrogen peroxide treatment. Soil Science, 96(4): 221-229. [2] Farquhar G. D., Ehleringer J. R. & Hubick K. T. (1989a) Carbon isotope discrimination and photosynthesis. Annu. Rev. Plant Physiol. Plant Mol. Biol., 40: 503-537. [9, 46] Farquhar G. D., Hubick K. T., Condon A. G. & Richards R. A. (1989b) Carbon isotope fractionation and plant water-use efficiency. In: Stable isotopes in ecological research. P. W. Rundel, J. R. Ehleringer & K. A. Nagy (eds.). Springer Verlag, NY, Vol. 68: 21-40. [21] Farquhar G. D. & Lloyd J. (1993) Carbon and oxygen isotope effects in the exchange of carbon dioxide between terrestrial plants and the atmosphere. In: Stable isotopes and plant carbon-water relations. J. R. Ehleringer, A. E. Hall & G. D. Farquhar (eds.). Academic Press: 47-70. [21]

Farquhar G. D., Lloyd J., Taylor J. A., Flanagan L. B., Syvertsen J. P., Hubrick K. T., Wong S. C. & Ehleringer J. R. (1993) Vegetation effects on the isotope composition of oxygen in atmospheric CO2. Nature, 363: 439-443. [14, 21]

Farquhar G. D., Henry B. K. & Styles J. M. (1997) A rapid on-line technique for determination of oxygen isotope composition of nitrogen-containing organic matter and water. Rapid Commun. Mass Spectrom., 11: 1554-1560. [1, 8, 15, 21, 23] Farquhar G. D., Barbour M. M. & Henry B. K. (1998) Interpretation of oxygen isotope composition of leaf material. In: Stable isotopes. H. Griffiths (ed.). BIOS Scinetific, Oxford: 27-62. [21] Farquhar J. & Rumble D. III (1998) Comparison of oxygen isotope data obtained by laser fluorination of olivine with KrF excimer laser and CO2 laser. Geochim. Cosmochim. Acta, 62(18): 3141-3149. [20] Farquhar J. & Thiemens M. H. (2000) A 33S enrichment in ureilite meteorites; evidence for a nebular sulfur component. Geochim. Cosmochim. Acta, 64: 1819-1825. [20] Farquhar J., Bao H. & Thiemens M. H. (2000a) Atmospheric influence of Earth's earliest sulfur cycle. Science, 289: 756-758. [20, 26] Farquhar J., Savarino J., Jackson T. & Thiemens M. (2000b) Evidence of atmospheric sulphur in the martian regolith from sulphur isotopes in meteorites. Nature, 404: 50-52. [18] Farquhar J., Savarino J., Airieau S. & Thiemens M. H. (2001) Observation of wavelength-sensitive mass-independent sulfur isotope effects during SO2 photolysis: Implications for the early atmosphere. J. Geophys. Res., 106: 32829-32839. [26] Farver J. R. & Yund R. A. (1995) Interphase boundary diffusion of oxygen and potassium in K-feldspar / quartz aggregates. Geochim. Cosmochim. Acta, 59: 3697-3705. [46] Farver J. R. & Yund R. A. (1999) Oxygen bulk diffusion measurements and TEM characterization of a natural ultramylonite: implications for fluid transport in mica-bearing rocks. J. Metamorphic Geol., 17: 669-683. [46] Fassett J. D. & Paulsen P. J. (1989) Isotope dilution mass spectrometry for accurate elemental analysis. Anal. Chem., 61(10): 643A-649A. [37] Fauhl C. & Wittkowski R. (1996) On-line 1H-NMR to facilitate tube preparation in SNIF-NMR analysis. Z. Lebensm. Unters. Forsch., 203: 541-545. [5] Faure G. (1986) Principles of isotope geology. Wiley & Sons, 2nd edition: 589 p. [29] Fayek M., Harrison T. M., Grove M., McKeegan K. D., Coath C. D. & Boles J. R. (2001) In situ isotopic evidence for protracted and complex carbonate cementation in a petroleum reservoir, North Coles Leveem, San Joaquin Basin, California, U.S.A.J. Sedim. Res., 71(3): 444-458. [30]

1088

References

Feldmann I., Jakubowski N. & Stuewer D. (1999a) Application of a hexapole collision and reaction cell in ICP-MS; Part I: Instrumental aspects and operational optimization. Fresenius J. Anal. Chem., 365: 415-421. [311 Feldmann I., Jakubowski N., Thomas C. & Stuewer D. (1999b) Application of a hexapole collision and reaction cell in ICP-MS; Part II: Analytical figures of merit and first applications. Fresenius J. Anal. Chem., 365: 422-428. [31] Felgett P. B. & Usher M. J. (1980) Fluctuation phenomena in instrument science. J. Phys. E., Sci. Instrum., 13: 1041-1046. [39] Feng F. & Horlick G. (1994) Analysis of aluminum alloys using inductively coupled plasma and glow discharge mass spectrometry. J. Anal. At. Spectrom., 9: 823-831. [35] Feng P., Bu X. & Stucky G. D. (1997) Hydrothermal syntheses and structural characterization of zeolite analogue compounds based on cobalt phosphate. Nature, 388: 735-740. [36] Feng X. & Savin S. M. (1991) Oxygen isotope studies of zeolites: Stilbite, analcime, heulandite, and clinoptilolite- I. Analytical technique. In: Stable Isotope Geochemistry: A tribute to Samuel Epstein. H.P. Jr. Taylor, J. R. O'Neil & I. R. Kaplan (eds.). The Geochemical Society, Spec. Publ. 3: 271- 283. [46] Feng X., Krishnamurthy R. V. & Epstein S. (1993) Determination of D / H ratios of nonexchangeable hydrogen in cellulose: A method based on the cellulose-water exchange reaction. Geochim. Cosmochim. Acta, 57: 4249-4256. [24] Feng X. & Savin S. M. (1993a) Oxygen isotope studies of zeolites- stilbite, analcime, heulandite, and clinoptilolite: II Kinetics and mechanisms of isotopic exchange between zeolites and water vapour. Geochim. Cosmochim. Acta, 57: 4219-4238. [46] Feng X. & Savin S. M. (1993b) Oxygen isotope studies of zeolites: stilbite, analcime, heulandite, and clinoptilolite; III, Oxygen isotope fractionation between stilbite and water or water vapor. Geochim. Cosmochim. Acta, 57(17): 4239-4247. [46] Feng X. & Epstein S. (1995a) Carbon isotopes of trees from arid environments and implications for reconstructing atmospheric CO2 concentrations. Geochim. Cosmochim. Acta, 59: 2599-2608. [46] Feng X. & Epstein S. (1995b) Climatic temperature records in 6D data from tree rings. Geochim. Cosmochim. Acta, 59: 3029-3037. [46] Ferhi A. & L4tolle R. (1979) Relation entre le milieu climatique et les teneurs en oxyg6ne-18 de la cellulose des plantes terrestres. Physiol. Vdg., 17: 107-117. [46] Ferhi A., Bariac T., Jusserand C. & L6tolle R. (1983) An integrated method for isotopic analysis of oxygen from organic compounds, air water vapor and leaf water. Int. J. Appl. Radiat. Isot., 34: 14511457.[1] Ferretti R. J., Lowe D. C., Martin R. J. & Brailsford G. B. (2000) A new gas chromatograph-isotope ratio mass spectrometry technique for high precision, N20-free analysis of 613C and 6180 in atmospheric CO2 from small air samples. J. Geophys. Res., 105(D5): 6709-6718. [14] Fessenden J. E., Cook C. S., Loft M. J. & Ehleringer J. R. (2002) Rapid 180 analysis of small water and CO2 samples using a continuous-flow isotope ratio mass spectrometer. Rapid Commun Mass Spectr., 16: 1257-1260. [1] Fetter C. W. (1994) Applied Hydrogeology. Prentice Hall. [29] Feuerstein T. P., Ostrom P. H. & Ostrom N. E. (1997) Isotopic biogeochemistry of dissolved organic nitrogen: a new technique and application. Org. Geochem., 27: 363-370. [15] Fiedler R. & Proksch G. (1972) The determination of the nitrogen content and nitrogen-15 abundance by means of the nitrogen gas generated from inorganic and organic materials. Anal. Chim. Acta, 60: 277-285. [15] Field E. M., Switsur V. R. & Waterhouse J. S. (1994) An improved technique for the determination of oxygen stable isotope ratios in wood cellulose. Int. J. Appl. Radiat. Isot., 45: 177-181. [23] Filer C. N. (1999) Isotopic fractionation of organic compounds in chromatography. J. Labelled Compd. Radiopharm., 42: 169-197. [8] Finlay J. C. (2003) Controls of streamwater dissolved inorganic carbon dynamics in a forested watershed. Biogeochemistry, 62: 231-252. [10] Firsching F. H. (1961) Precipitation of silver phosphate from homogeneous solution. Anal. Chem., 33: 873-878. [20, 22]

References

1089

Fischer T. P., Giggenbach W. F., Sano Y. & Williams S. N. (1998) Fluxes and sources of volatiles dis-

charged from Kudryavy, a subduction zone volcano, Kurile Islands. Earth Planet. Sci. Lett., 160: 8196. [16] Fitzsimons I. C. W., Harte B. & Clark R. M. (2000) SIMS stable isotope measurement counting statistics and analytical precision. Mineral. Magazine, 64: 59-83. [30] Flesch G. D., Anderson A. R. Jr. & Svec H. J. (1973) A secondary isotopic standard for 6Li/7Li determination. Intl. J. Mass Spectrom. Ion Phys., 12: 265-272. [6, 40] Flintrop C., Hohlmann B., Jasper T., Korte C., Podlaha O. G., Scheele S. & Veizer J. (1996) Anatomy of pollution: rivers of North-Westphalia, Germany. Am. J. Sci., 296: 58-98. [10] Florkowsky T. (1985) Sample preparation for hydrogen isotope analysis by mass spectrometry. Int. J. Appl. Radiat. Isot., 36(12): 991-992. [1] Focken U. & Becker K. (1998) Metabolic fractionation of stable carbon isotopes: implications of different proximate compositions for studies of the aquatic food webs using 613C data. Oecologia, 115: 337-343.

[9]

Fogel M. L., Aguilar C., Cuhel R., Hollander D. J., Willey J. D. & Paerl H. W. (1999) Biological and isotopic changes in coastal waters induced by Hurricane Gordon. Limnol. Oceanogr., 44: 1359-1369.

[9]

Fogel M. L. & Tuross N. (1999) Transformations of plant biochemicals to geological macromolecules during early diagenesis. Oecologia, 120: 336-346. [9] Forrest J. & Newman L. (1973) Sampling and analysis of atmospheric sulfur compounds for isotope ratio studies. Atmos. Environm., 7: 561-573. [26] Fortier S. M. & Giletti B. J. (1991) Volume self-diffusion of oxygen in biotite, muscovite, and phlogopite micas. Geochim. Cosmochim. Acta, 55: 1319-1330. [46] Fortier S. M. (1994) An on-line experimental/analytical method for measuring the kinetics of oxygen isotope exchange between CO2 and saline/hypersaline salt solutions at low (25-50~ temperatures. Chem. Geol. (Isot. Geosci. Sect.), 116: 155-162. [1] Fortier S. M., Liittge A., Satir M. & Metz P. (1994) Oxygen isotope fractionation between fluorphlogopite and calcite: an experimental investigation of temperature dependence and F-/OH- effects. Eur. J. Mineral., 6: 53-65. [2, 46] Fortier S. M., Cole D. R., Wesolowski D. J., Riciputi L. R., Paterson B. A., Valley J. W. & Horita J. (1995) Determination of the magmatic-water equilibrium oxygen isotope fractionation factor at 350~ A comparison of ion microprobe and laser fluorination techniques. Geochim. Cosmochim. Acta, 59(18): 3871-3875. [46] Fortier S. M. & Liittge A. (1995) An experimental calibration of the temperature dependence of oxygen isotope fractionation between apatite and calcite at high temperatures (350-800~ Chem. Geol. (Isot Geosci. Sect.), 125: 281-290. [20, 22] Fossing H. & Jorgensen B. B. (1989) Measurement of bacterial sulfate reduction in sediments: Evaluation of a single-step chromium reduction method. Biogeochem., 8: 205-222. [26] Fossing H. (1995) 35S radiolabeling to probe biogeochemical cycling of sulfur. In: Geochemical transformations of sedimentary sulphur. M. A. Vairavamurthy, M. A. A. Schoonen, T. I. Eglinton, G. W. Luther, and B. Manowitz (eds.). American Chemical Society, Washington. ACS Symposium Series 612: 349-364. [26] Fouillac C., Fouillac A. M. & Criaud A. (1990) Sulphur and oxygen isotopes of dissolved sulphur species in formation water from the Dogger geothermal aquifer, Paris Basin, France. Appl. Geochem., 5: 415-429. [26] Fouillac A.-M. & Girard J.-P. (1996) Laser oxygen isotope analysis of silicate/oxide grain separates: Evidence for a grain effect? Chem. Geol., 130: 31-54. [20] France-Lanord C. & Sheppard S. M. F. (1992) Hydrogen isotope composition of pore waters and interlayer water in sediments from the Central Western Pacific, LEG 129. Proceed. ODP, Sci. Results, 129: 295-302. [2, 46] France-Lanord C. & Derry L. A. (1994) 613C of organic carbon in the Bengal Fan: Source evolution and transport of C3 and C4 plant carbon to marine sediments. Geochim. Cosmochim. Acta, 58: 4809-4814.

[9]

1090

References

Francey R. J. (1986) Carbon isotope measurements in baseline air, forest canopy air, and plants. In: The changing carbon cycle. J. R. Trabalka & D. E. Reichle (eds.). Springer Verlag: 160-174. [24] Francey R. J. & Tans P. P. (1987) The latitudinal variation in oxygen-18 of atmospheric CO2. Nature, 327: 495-497. [141 Francey R. J., Allison C. E., Steele L. P., Langenfelds R. L., Welch E. D., White J. W. C., Trollier M., Tans P. P. & Massarie K. A. (1994) Intercomparison of stable isotope measurements of CO2. In: Climate monitoring and diagnostics laboratory (CMDL). J. T. Peterson & R. M. Rosson (eds.). US Department of Commerce, Boulder, Colorado, USA. Summary report 1993, No. 22: 106-110. [14] Francey R. J., Tans P., Allison C. E., Enting I. G., White J. W. C & Trolier M. (1995) Changes in oceanic and terrestrial carbon uptake since 1982. Nature, 373: 326-330. [14] Franchi I. A., Wright I. P., Gibson E. K. Jr. & Pillinger C. T. (1986) The laser microprobe: A technique for extracting carbon, nitrogen and oxygen from solid samples for isotopic measurements. J. Geophys. Res., Supplem. 91 (B4; Proceedings of the sixteenth lunar and planetary science conference, Part. 2): D514-D524. [17, 20] Franchi I. A., Akagi T. & Pillinger C. T. (1993a) Laser fluorination of silicates for 6170 and 6180 using 02 gas. TERRA Nova, 5 (Abstr. Suppl. No.l): 370. [20] Franchi I. A., Wright I. P. & Pillinger C. T. (1993b) Constraints on the formation conditions of iron meteo-rites based on concentration and isotopic compositions of nitrogen. Geochim. Cosmochim. Acta, 57: 3105-3121. [13] Franconi J. M., Naulet N. & Martin G. J. (1989) Natural factors of site-specific isotope fractionation of hydrogen in the fermentation products (ethanol and water) of barley. J. Cereal Sci., 10: 69-80. [5] Frazer J. w. & Crawford R. (1963) Modifications in the simultaneous determination of carbon, hydrogen, and nitrogen. Mikrochim. Acta, 1963(3): 561-566. [15] Freeman K. H., Hayes J. M., Trendel J. M. & Albrecht P. (1990) Evidence from carbon isotope measurements for diverse origins of sedimentary hydrocarbons. Nature, 343: 254-256. [9] Freeman K. H., Wakeham S. G. & Hayes J. M. (1994) Predictive isotopic biogeochemistry: Hydrocarbons from anoxic marine basins. Org. Geochem., 21: 629-644. [9] Freer R. & Dennis P. F. (1982) Oxygen diffusion studies. I. A preliminary ion microprobe investigation of oxygen diffusion in some rock-forming minerals. Mineral. Magazine, 45: 179-192. [46] Freer R., Wright K., Knoll H. & Gi~ttlicher J. (1997) Oxygen diffusion in sanidine feldspars and a critical appraisal of oxygen isotope-mass-effect measurements in non-cubic materials. Phil. Magazine, A 75: 485-503. [46] Freney J. R. (1961) Some observations on the nature of organic sulphur compounds in soil. Aust. J. Agric. Res., 12: 424-432. [26] Freney J. R., Melville G. E. & Williams C. H. (1970) The determination of carbon bonded sulfur in soil. Soil. Sci., 109: 310-318. [26] Freney J. R. (1986) Forms and reactions of organic sulfur compounds in soils. In: Sulfur in Agriculture. M. A. Tabatabai (ed.). American Society of Agronomy - Crop Science Society of America - Soil Science Society of America, Madison, Wisconsin, USA. Monograph No. 27: 207-232. [26] Frezzotti M. L., Dallai L. & Sharp Z. D. (2000) Fluid-inclusion and stable-isotope evidence for fluid infiltration and veining during metamorphism in marbles and metapelites. Eur. J. Mineral., 12(1): 231-246. [3] Frick U. & Pepin R. O. (1981) Microanalysis of nitrogen isotope abundances: Association of nitrogen with noble gas carriers in Allende. Earth Planet. Sci. Lett., 56: 64-81. [13, 17] Fricke H. C., O'Neil J. R. & Lynnerup N. (1995) Oxygen isotope composition of human tooth enamel from medieval Greenland: Linking climate and society. Geology, 23: 869-872. [22] Fricke H. C. & O'Neil J. R. (1996) Inter- and intra-tooth variation in the oxygen isotope composition of mammalian tooth enamel phosphate: implications for palaeoclimatological and palaeobiological research. Palaeogeogr., Palaeoclim., Palaeoecol., 126: 91-99. [22] Fricke H. C., Clyde W. C., O'Neil J. R. & Gingerich P. D. (1998a) Evidence for rapid climate change in North America during the latest Paleocene thermal maximum: oxygen isotope compositions of biogenic phosphate from the Bighorn Basin (Wyoming). Earth Planet. Sci. Lett., 160: 193-208. [22]

References

1091

Fricke H. C., Clyde W. C., O'Neil J. R. & Gingerich P. D. (1998b) Intra-tooth variations in ~)180 (PO4) of mammalian tooth enamel as a record of seasonal variations in continental climate variables. Geochim. Cosmochim. Acta, 62: 1839-1850. [22] Friedheim C. & Meyer R. J. (1892) Ober die quantitative Trennung und Bestimmung von Chlor, Brom und Jod. Z. anorg. Chemie, 1: 407-422. [28] Friedman I. (1953) Deuterium content of natural waters and other substances. Geochim. Cosmochim. Acta, 4:89-103.[1] Friedman I. & Smith R. L. (1958) The deuterium content of water in some volcanic glasses. Geochim. Cosmochim. Acta, 15: 218-228. [3, 46] Friedman I. (1970) Some investigations of the deposition of travertine from hot springs, I. Geochim. Cosmochim. Acta, 34: 1303-1315. [10] Friedman I. & Hardcastle K. (1970) A new technique for pumping hydrogen gas. Geochim. Cosmochim. Acta, 34: 125-126. [1] Friedman I. & O'Neil J. R. (1977) Compilation of stable isotope fractionation factors of geochemical interest. U. S. Geol. Survey Prof. Paper, 440-KK. [40, App B] Friedman I., O'Neil J. & Cebula G. (1982) Two new carbonate stable isotope standards. Geostand. Newslett. 6, 11-12. [40] Friedman L. & Bigeleisen J. (1950) Oxygen and nitrogen isotope effects in the decomposition of ammonium nitrate. J. Chem Phys., 18: 1325-1331. [19] Fritz P. & Fontes J. C. (1966) Fractionnement isotopique pendant l'attaque acide des carbonates naturels: r61e de la granulom6trie. C. R. Acad. Sc. Paris, t. 263: 1345-1348. [12] Fritz P., Drimmie R. J. & Nowicki V. K. (1974) Preparation of sulfur dioxide for mass spectrometer analysis by combustion of sulfides with copper oxide. Anal. Chem., 46: 164-166. [26] Fritz P., Frape S. K., Drimmie R. J. & Heemskerk A. R. (1986) Reply to comments by Grobczak et al. on "Water-rock interaction and chemistry of groundwaters from the Canadian Shield". Geochim. Cosmochim. Acta, 50: 1561-1563. [1] Fronza G., Fuganti C., Grasselli P., Barbeni M. & Cisero M. (1993) Natural abundance 2H nuclear magnetic resonance study of the origin of (R)-h-decanolide. J. Agric. Food Chem., 41: 235-239. [5] Fronza G., Fuganti C., Grasselli P., Servi P., Zucchi G., Barbeni M. & Cisero M. (1995) Natural abundance 2H nuclear magnetic resonance study of the origin of 2-phenylethanol and 2-phenylethyl acetate. J. Agric. Food Chem., 43: 439-443. [5] Fronza G., Fuganti C., Guillou C., Reniero F. & Joulain D. (1998) Natural abundance 2H nuclear magnetic resonance study of the origin of raspberry ketone. J. Agric. Food Chem., 46(1): 248-254. [5] Fry B. & Sherr E. B. (1984) 813C measurements as indicators of carbon flow in marine and freshwater ecosystems. Contrib. Mar. Sci., 27: 13-47. [9] Fry B. (1986) Stable sulfur isotopic distributions and sulfate reduction in lake sediments of the Adirondack Mountains, New York. Biogeochemistry, 2: 329-343. [26] Fry B., Gest H. & Hayes J. M. (1986) Sulfur isotope effects associated with protonation of HS- and volatilization of H2S. Chem. Geol. (Isotope Geosci. Sec.), 58: 253-258. [46] Fry B., Ruf W., Gest H. & Hayes J. M. (1988) Sulfur isotope effects associated with oxidation of sulphide by 02 in aqueous solution. Chem. Geol., 73: 205-210. [46] Fry B. (1991) Stable isotope diagrams of freshwater food webs. Ecology, 72: 2293-2297. [9] Fry B., Brand W., Mersch F. J., Tholke K. & Garritt R. (1992) Automated analysis system for coupled 813C and 815N measurement. Anal. Chem., 64: 288-291. [9, 45] Fry B., Garritt R., Tholke K., Neill C., Michener R. H., Mersch F. J. & Brand W. (1996) Cryoflow: cryofocussing nanomole amounts of CO2, N2 and SO2 from an elemental analyzer for stable isotopic analysis. Rapid Commun. Mass Spectrom., 10: 953-958. [15, 45] Fry B., Silva S. R., Kendall C. & Anderson R. K. (2002) Oxygen isotope corrections for online 834S analysis. Rapid Commun. Mass Spectrom., 16: 854-859. [26, 45] Fung I., John J., Lerner E., Matthews E., Prather M., Steele L. P. & Fraser P. J. (1991) Dimensional model synthesis of the global methane cycle. J. Geophys. Res., 96(D7): 13033-13065. [14] Fusch C. (1985) Quantitative Analyse von D20-Konzentrationen im Serum mit der Infrarot-Spektroskopie. Inaugural-Dissertation (Inauguration Thesis), UniversiNit T/,ibingen. [34]

1092

References

Fusch C., Spririg N. & Moeller H. (1993) Fourier transform infrared spectroscopy measures 1H/2H

ratios of native water with a precision comparable to that of isotope ratio mass spectrometry. Eur. J. Clin. Chem. Clin. Biochem., 31: 639-644. [34] Gaillardet J. & All6gre C. J. (1995) Boron isotopic compositions of corals: seawater or diagenesis record? Earth Planet. Sci. Lett., 136: 665-676. [7] Gale N. H. (1970) A solution in closed form for lead isotopic analysis using a double spike. Chem. Geol., 6: 305-310. [29] Gale N. H., Woodhead A. P., Stos-Gale Z. A., Walder A. & Bowen I. (1999) Natural variations detected in the isotopic composition of copper: possible applications to archaeology and geochemistry. Int. J. Mass Spectrom., 184: 1-9. [31] Galy A., Young E. D., Ash R. D. & O'Nions R. K. (2000) The formation of chondrules at high gas pressure in the solar nebula. Nature, 290: 1751-1753. [31] Galy A., Belshaw N. S., I-Ialicz L. & O'Nions R. K. (2001) High-precision measurement of magnesium isotopes by multiple-collector inductively coupled plasma mass spectrometry. Int. J. Mass Spectrom., 208: 89-98. [31] Games L. M. & Hayes J. M. (1976) Isotopic and quantitative analysis of the major carbon fractions in natural water samples. Anal. Chem., 48: 130-135. [10] Games L. M. & Hayes J. M. (1977) Carbon isotopic study of the landfill leachate in groundwater. J. Water Poll. Control Fed., 49: 668-677. [10] Games L. M., Hayes J. M. & Gunsalus R. P. (1978) Methane-producing bacteria: natural fractionations of the stable carbon isotopes. Geochim. Cosmochim. Acta, 42: 1295-1297. [46] Ganther H. E. (1974) The biochemistry of selenium. In: Selenium. R. A. Zingaro& W. C. Cooper (eds.). Van Nostrand Reinhold Co: 546-614. [29] Gao X. & Thiemens M. H. (1991) Systematic study of sulfur isotopic composition in iron meteorites and the occurence of excess 33S a n d 36S. Geochim. Cosmichim. Acta, 55: 2671-2679. [20, 45] Gao X. & Thiemens M. H. (1993a) Isotopic composition and concentration of sulfur in carbonaceous chondrites. Geochim. Cosmochim. Acta, 57: 3159-3169. [45] Gao X. & Thiemens M. H. (1993b) Variations of the isotopic composition of sulfur in enstatite and ordinary chondrites. Geochim. Cosmochim. Acta, 57: 3171-3176. [45] Gao Y. Q. & Marcus R. A. (2001) Strange and unconventional isotope effects in ozone formation. Science, 293: 259-263. [18, 46] Gao Y. Q. & Marcus R. A. (2002) On the theory of the strange and unconventional isotope effects in ozone formation. J. Chem. Phys., 116(1): 137-154. [18] Garcia J.-P., Sheppard S. M. F., Picard S., L4cuyer C. & Ernig C. (1997) Carbon and oxygen isotopic signals in living brachiopods: prerequisites for the record of oceanic secular variations. Terra Nova, 9 (Abstracts suppl. No. 1): 271 (Abstract). [46] Garlick G. D. & Epstein S. (1967) Oxygen isotope ratios in coexisting minerals of regionally metamorphosed rocks. Geochim. Cosmochim. Acta, 31: 181-214. [20] Garner E. L. & Raspberry S. D. (1993) What's new in traceability. J. Test. Eval., 21: 505-509. [43] Garrels R. M. & Christ C. L. (1965) Solutions, minerals and equilibria. Harper and Row, NY. [10] Garten C. T. Jr. (1992) Nitrogen isotope composition of ammonium and nitrate in bulk precipitation and forest throughfall. Intern. J. Environ. Anal. Chem., 47: 33-45. [15] Gaudette H. E. (1990) Chlorine and boron isotopic analysis of Antarctic ice and snow: Indicators of marine and volcanic atmospheric inputs. GSA Ann. Mtg. Abstr. Prog.: 173. [28] Gaunt J. (1956) The analysis of heavy water by infra-red spectrometry. Spectrochim. Acta, 8: 57-65. [34] Gazi D., Spiro B. & Greenwood P. B. (1993) Experiments on laser fluorination of silicate and oxide minerals. Terra Nova, 5 (Abstr. Suppl. No.l): 371. [20] Gearing J. N. (1991) The study of diet and trophic relationships through natural abundance 13C. In: Carbon isotope techniques. D. C. Coleman & B. Fry (eds.). Academic Press Inc, San Diego: 201-218. [9] Gehre M., I-Ioefling R. & Kowski P. (1996a) Methodical studies for D/H-isotope analysis - A new technique for the direct coupling of sample preparation to an IRMS. Isot. Environ. Health Stud., 32: 335-340. [1] Gehre M., Hoefling R., Kowski P. & Strauch G. (1996b) Sample preparation device for quantitative hydrogen isotope analysis using chromium metal. Anal. Chem., 68(24): 4414-4417. [1]

References

1093

Gellene G. I. (1992) Symmetry restrictions in diatom/diatom reactions. I. Group theoretical analysis. J. Chem. Phys., 96: 4387-4402. [18] Gellene G. I. (1993) Symmetry-dependent isotope fractionation in the formation of He2 +. J. Chem. Phys., 97: 34-39. [18] Gellene G. I. (1996) An explanation for symmetry-induced isotopic fractionation in ozone. Science, 274: 1344-1346. [18] Gemery P. A., Trolier M. & White J. W. C. (1996) Oxygen isotope exchange between carbon dioxide and water following atmospheric sampling using glass flasks. J. Geophys. Res., 101: 14415-14420. [14] Genoni L., Iacumin P., Nikolaev V., Gribchenko Y. & Longinelli A. (1998) Oxygen isotope measurements of mammoth and reindeer skeletal remains: an archive of late Pleistocene environmental conditions in Eurasian Arctic. Earth Planet. Sci. Lett., 160: 587-592. [22] Gensho R. & Honda M. (1971) Measurements of the isotopic ratio of boron. Shitsuryo Bunseki (Mass Spectrom.), 19:131-143 (in Japanese, with English abstract). [7] Genty D., Plagnes V., Causse C., Cattani O. & Stievenard M. (2002) Fossil water in large stalagmite voids as a tool for paleoprecipitation stable isotope composition reconstitution and paleotemperature calculation. Chem. Geol., 184: 83-95. [3] Gerstenberger H. & Herrmann M. (1983) Report on the intercomparison for the isotope standards Limestone KH2 and Polyethylene Foil PEF 1. Zfl-Mitteilungen, Zentralinstitut f/Jr Isotopen- und Strahlenforschung. Leipzig, Germany. Vo160: 67-83. [40] Gessler R. & v. Gehlen K. (1986) Investigation of sulfur isotope fractionation between H2S gas and aqueous solutions. Fresenius Z. Anal. Chem., 324: 130-136. [46] Gianfrani L., Gagliardi G., van Burgel M. & Kerstel E. (2003) Isotope analysis of water by means of near-infrared dual-wavelength diode laser spectroscopy. Opt. Express, 11:1566-1576 (http:// www.opticsexpress.org / abstract.cfm?URI=OPEX-11-13-1566). [34] Giesemann A., J/iger H.-J., Norman A. L., Krouse H. R. & Brand W. A. (1994) On-line sulfur-isotope determination using an elemental analyzer coupled to a mass spectrometer. Anal. Chem., 66: 28162819. [26, 27, 45] Gifford S., Bentley H. W. & Graham D. L. (1985) Chlorine isotopes as environmental tracers in Columbia River basalt groundwaters. Proc. 17th Int. Congr. I.A.H.. Hydrogeology of rocks of low permeability. Vol. 17: 417-429. [28] Giggenbach W. F. (1982) Carbon-13 exchange between CO2 and CH4 under geothermal conditions. Geochim. Cosmochim. Acta, 46:159-165 [46]. Giggenbach W. F. (1997) Relative importance of thermodynamic and kinetic processes in governing the chemical and isotopic composition of carbon gases in high-heat flow sedimentary basins. Geochim. Cosmochim. Acta, 61: 3763-3785. [46] Giletti B. J., Semet M. P. & Yund R. A. (1978) Studies in diffusion. III. Oxygen in feldspar: an ion microprobe determination. Geochim. Cosmochim. Acta, 42: 45-57. [46] Giletti B. J. & Yund R. A. (1984) Oxygen diffusion in quartz. J. Geophys. Res., 89: 4039-4046. [46] Giletti B. J. (1985) The nature of oxygen transport within minerals in the presence of hydrothermal water and the role of diffusion. Chem. Geol., 53: 197-206. [46] Giletti B. J. (1986) Diffusion effects on oxygen isotope temperatures of slowly cooled igneous and metamorphic rocks. Earth Planet. Sci. Lett., 77: 218-228. [46] Gilg H. A. & Sheppard S. M. F. (1995) Hydrogen isotope fractionation between smectite and water. Terra Nova, Abstracts suppl. No. 1:329 (Abstract.) [46] Gilg H. A. & Sheppard S. M. F. (1996) Hydrogen isotope fractionation between kaolinite and water revisited. Geochim. Cosmochim. Acta, 60: 529-533. [46] Gilg H. A. & Sheppard S. M. F. (1999) Stability of H and O isotope ratios in natural hydrated silicic glasses: estimation of diffusion coefficient for water at ambient temperatures. EUG 10, Conference Abstracts, 4:531 (Abstract). [46] Gillet P., McMillan P., Schott J., Badro J. & Grzechnik A. (1996) Thermodynamic properties and isotopic fractionation of calcite from vibrational spectroscopy of 180-substituted calcite. Geochim. Cosmochim. Acta, 60: 3471-3485. [46]

1094

References

Gillson G. R., Douglas D. J., Fulford J. E., Halligan K. W. & Tanner S. D. (1988) Nonspectroscopic interelement interferences in inductively coupled mass spectrometry. Anal. Chem., 60: 1472-1474. [6,

35]

Gilmour I. & Pillinger C. T. (1994) Isotopic compositions of individual polycyclic aromatic hydrocarbons from the Murchison meteorite. Mon. Not. R. Astron. Soc., 269: 235-240. [11] Gilmour I., Johnston M. A., Pillinger C. T., Pond C. M., Mattacks C. A. & Prestrud P. (1995a) The carbon isotopic composition of individual fatty acids as indicators of dietary history in arctic foxes on Svalbard. Phil. Trans. R. Soc. Lond., 349(B): 135-142. [9] Gilmour J. D., Hewett S. M., Lyon I. C., Stringer M. & Turner G. (1991) A resonance ionization mass spectrometer for xenon. Meas. Sci. Technol., 2: 1-7. [33] Gilmour J. D., Ash R. D., Lyon I. C. & Turner G. (1994a) I-Xe dating using the RELAX mass spectrometer. Resonance Ionization Spectroscopy 1994 (RIS 94). AIP Press Woodbury, NY, AIP Conference Proceedings 329: 233-236. [33] Gilmour J. D., Johnston W. A., Lyon I. C. & Turner G. (1994b) RELAX-l: An ultrasensitive resonance ionization mass spectrometer for xenon. Rev. Sci. Instrum., 65(3): 617-625. [33] Gilmour J. D., Ash R. D., Lyon I. C. & Turner G. (1995b) Iodine-xenon dating of Bjub61e and Parnallee using RELAX. Meteoritics, 30: 405-411. [33] Gim6nez-Miralles J. E., Salazar D. M. &Solana I. (1999) Regional origin assignment of red wines from Valencia (Spain) by 2H-NMR and 13C IRMS stable isotope analysis of fermentative ethanol. J. Agric. Food Chem., 47(7): 2645-2652. [5] Giral S. (1994) Variations des rapports isotopiques 180/160 des kaolinites de deux profils lat6ritiques amazoniens: signification pour la p6dologie et la pal6oclimatologie. Thbse Universit6, Universit6 Aix-Marseille III, Aix en Province-Marseille, France: 198 p. [2] Giral-Kacmarcik S., Savin S. M., Nahon D., Girard J.-P., Lucas Y. & Abel L. (1998) Oxygen isotope geochemistry of kaolinite in laterite-forming processes, Manaus, Amazonas, Brazil. Geochim. Cosmochim. Acta, 62: 1865-1879. [2] Girard G. (1994) The third periodic verification of national prototypes of the kilogram (1988-1992). Metrologia, 31: 317-336. [42] Girard J.-P., Deynaux M. & Nahon D. (1989) Diagenesis of the upper Proterozoic siliciclastic sediments of the Taoudeni Basin (West Africa) and relation to diabase emplacement. J. Sediment. Petrol., 59: 233248. [2] Girard J.-P. & Fouillac A.-M. (1995) G6ochimie isotopique de l'oxyg6ne et de l'hydrog6ne des argiles: examples d'application aux domaines diag6n6tique et g6othermique. Bull. Centres Rech. Explor.-Prod. ElfAquit., 19(1): 167-195. [2] Girard J.-P. & Savin S. M. (1996) Intracrystalline fractionation of oxygen isotopes between hydroxyl and non-hydroxyl sites in kaolinite measured by thermal dehydroxylation and partial fluorination. Geochim. Cosmochim. Acta, 60: 469-487. [2, 20] Girard J.-P., Freyssinet P. & Chazot G. (2000) Unraveling climatic changes from intra-profile variation in oxygen and hydrogen isotopic composition of goethite and kaolinite: An integrated study from Yaou, French Guyana. Geochim. Cosmochim. Acta, 64: 409-426. [2] Giraudon S., Danzart M. & Merle H. (2000) Deuterium nuclear magnetic resonance spectroscopy and stable carbon isotope ratio analysis/mass spectrometry of certain monofloral honeys. J. AOAC Int., 83(6): 1401-1409. [5] Giuliani G., France-Lanord C., Zimmermann J. L., Arboleda C., Charoy B., Coget P., Fontan F. & Giard D. (1997) Fluid composition, 6D of channel H20 and 6180 of lattice oxygen in beryls: genetic implications for Brazilian, Columbian, and Afghanistani emerald deposits. Intern. Geol. Rev., 39: 400424. [16] Giussani A., Roth P., Werner E., Schramel P., Wendler I. & Nusslin F. (1997) Thermal ionization and inductively coupled plasma mass spectrometry: Potential for the application to studies on the biokinetics of molybdenum in humans. Isot. Environ. Health Stud., 33: 207-215. [31] Gleason J. D., Friedman I. & Hanshaw B. B. (1969) Ectraction of dissolved carbonate species from natural water for carbon-isotope analysis. U.S. Geol. Surv. Prof. Paper, 650-D: D248-D250. [10] Gleditsch E. & Sandahl B. (1922) Radioactivit6 sur le poids atomique de chlore dans un mineral ancien, l'apatede Balme. C. R. Acad Sci. Paris, 174: 746-748. [28]

References

1095

Gleit C. E. & Holland W. D. (1962) Use of electrically excited oxygen for the low temperature decomposition of organic substances. Anal. Chem., 34: 1454-1457. [2] Glibert P. M., Lipschultz E, McCarthy J. J. & Altabet M. A. (1982) Isotope dilution models of uptake and remineralization of ammonium by marine plankton. Limnol. Oceanogr., 27: 639-650. [15] Godfrey J. D. (1962) The deuterium content of hydrous minerals from the East-Central Sierra Nevada and Yosemite National Park. Geochim. Cosmochim. Acta, 26: 1215-1245. [1] Goerges T. & Dittert K. (1998) Improved diffusion technique for 15N:14N analysis of ammonium and nitrate from aqueous samples by stable isotope spectrometry. Comm. Soil Sci. Plant Anal., 29(3&4): 361-368. [15] Goldhaber M. B. & Kaplan I. R. (1974) The sulphur cycle. In: The Sea. E. D. Goldberg (ed.). Interscience Wiley, NY, Vol. 5: 569-655. [46] Goldhaber M. B. & Kaplan I. R. (1975) Controls and consequences of sulfate reduction rates in recent marine sediments. Soil Sci., 119: 42-55. [27] Goldman A., Murcray F. J., Murcray D. G., Kosters J. J., Rinsland C. P., Flaud J. M., Camy-Peyret C. & Barbe A. (1989) Isotopic abundances of stratospheric ozone from balloon-borne high resolution infrared solar spectra. J. Geophys. Res., 94: 8467-8473. [18] Gonfiantini R. (1977) Consultants' meeting on stable isotope standards and intercalibration in hydrology and geochemistry. 8-10 Sept., 1976. IAEA, Vienna. Internal Report, March 1977. [40] Gonfiantini R. (1978) Standards for stable isotope measurements in natural compounds. Nature, 271(No. 5645): 534-536. [1, 40] Gonfiantini R. (1981) The h-notation and the mass-spectrometric measurement techniques. In: Stable isotope hydrology, deuterium and oxygen-18 in the water cycle. J. R. Gat & R. Gonfiantini (eds.). IAEA, Vienna. Technical Report Series, No. 210: 35-84. [1, 14, 15, 40, 43] Gonfiantini R. (1984) Advisory group meeting on stable isotope reference samples for geochemical and hydrological investigations. 19-21 Sept, 1983. Report to Director General, IAEA, Vienna: 77 p. [21, 40] Gonfiantini R., Stichler W. & Rozanski K. (1995) Standards and intercomparison materials distributed by the International Atomic Energy Agency for stable isotope measurements. In: Reference and intercomparison materials for stable isotopes of light elements. Proceedings of a Consultants Meeting held in Vienna, 1-3 Dec 1993. IAEA-TECDOC-825, IAEA, Vienna: 13-29. [40, 42, 43, 44] Gonfiantini R., Valkiers S., Taylor P. D. P. & De Bi6vre P. (1997) Using isotopic disequilibrium of CO2 to model gas adsorption in mass spectrometric measurements. Int. J. Mass Spectrom. Ion Proc., 161: 15-26. [43] Gonfiantini R., Tonarini S., Griining M., Adorni-Braccesi A., A1-Ammar A. S., Astner M., Btichler S., Barnes R. M., Bassett R. L., Cocherie A., Deyhle A., Dini A., Ferrara G., Gaillardet J., Grimm J., Guerrot C., Kr/ihenbiihl U., Layne G., Lemarchand D., Meixner A., Northington D. J., Pennisi M., Reitznerov~i E., Rodushkin I., Sugiura N., Surberg R., Tonn S., Wiedenbeck M., Wunderli S., Xiao Y. & Zack T. (2003) Intercomparison of boron isotope and concentration measurements. Part Ih Evaluation of the results obtained. Geostand. Newslett., 27(1): 41-57. [40] Gonzalez J., Jamin E., Remaud G., Martin Y. L., Martin G. G. & Martin M. L. (1998) Authentication of lemon juices and concentrates by a combined multi-isotope approach using SNIF-NMR and IRMS. J. Agric. Food Chem., 46: 2200-2205. [5] Gofii M. A. & Eglinton T. I. (1996) Stable carbon isotopic analyses of lignin-derived CuO oxidation products by isotope ratio monitoring-gas chromatography-mass-spectrometer (irm-GC-MS). Org. Geochem., 24: 601-615. [9] Goodman K. J. & Brenna J. T. (1994) Curve fitting for restoration od accuracy for overlapping peaks in gas chromatography/combustion isotope ratio mass spectrometry. Anal. Chem., 66: 1294-1301. [8] Goodman K. J. & Brenna J. T. (1995) High-precision gas-chromatography combustion isotope ratio mass-spectrometry at low signal levels. J. Chromatogr., A 689: 63-68. [8] Goodman K. J. (1998) Hardware modifications to an isotope ratio mass spectrometer continuous-flow interface yielding improved signal, resolution, and maintenance. Anal. Chem., 70: 833-837. [8] Grabczak J., Rozanski K. & Zuber A. (1986) Comments on "Water-rock interaction and chemistry of groundwaters from the Canadian Shield" by Frape, Fritz and McNutt. Geochim. Cosmochim. Acta, 50: 1559-1560. [1]

1096

References

Graber E. R. & Aharon P. (1991) An improved microextraction technique for measuring dissolved inorganic carbon (DIC), 613CDIC and 618OH20 from milliliter-size water samples. Chem. Geol., 94(2): 137144. [1, 10] Grachev A. M. & Severinghaus J. P. (2003) Laboratory determination of thermal diffusion constants for 29N2/28N2 in air at temperatures from -60 to 0~ for reconstruction of magnitudes of abrupt climate changes using the ice core fossil-air paleothermometer. Geochim. Cosmochim. Acta, 67: 345-360. [46] Gradsztajn E., Salome M., Yaniv A. & Bernas R. (1967) Isotopic analysis of lithium in the Holbrook meteorite and in terrestrial samples with a sputtering ion mass spectrometer. Earth Planet. Sci. Lett., 3: 387-393. [6] Gradsztajn E. & Guez M. (1969) Mass discrimination introduced by the optical properties of a sputtering ion source during the isotopic analysis of lithium in stony meteorites. Int. J. Mass Spectrom. Ion Phys., 4: 305-313. [6] Grady M. M., Verchovsky A., Wright I. P. & Pillinger C. T. (1996) Fully automated stepped combustion analysis of C, N, and Ar in QUE 94201: Comparison with other shergottites. Meteoritics, 31: A52-A53. [13] Graff J. & Rittenberg D. (1952) Microdetermination of deuterium in organic compounds. Anal. Chem., 24: 878-881. [1] Graham C. M. & Sheppard S. M. F. (1980) Experimental hydrogen isotope studies, II. Fractionations in the system epidote - NaC1 - H20, epidote - CaC12- H20 and epidote - seawater, and the hydrogen isotope composition of natural epidotes. Earth Planet. Sci. Lett., 49: 237-251. [46] Graham C. M., Sheppard S. M. F. & Heaton T. H. E. (1980) Experimental hydrogen isotope studies - I. Systematics of hydrogen isotope fractionation in the system epidote - H20, zoisite - H20 and A10(OH) - H20. Geochim. Cosmochim. Acta, 44: 353-364. [46] Graham C. M. (1981) Experimental hydrogen isotope studies: III. Diffusion of hydrogen in hydrous minerals, and stable isotope exchange in metamorphic rocks. Contrib. Mineral. Petrol., 76: 216-228. [46] Graham C. M., Harmon R. S. & Sheppard S. M. F. (1984) Experimental hydrogen isotope studies: hydrogen isotope exchange between amphibole and water. Am. Mineralogist, 69: 128-138. [46] Graham C. M., Viglino J. A. & Harmon R. S. (1987) Experimental study of hydrogen-isotope exchange between aluminous chlorite and water and of hydrogen diffusion in chlorite. Am. Mineralogist, 72: 566-579. [46] Graham C. M., Valley J. W. & Winter B. L. (1996) Ion microprobe analysis of 180 / 160 in authigenic and detrital quartz in the St. Peter Sandstone, Michigan Basin and Wisconsin Arch, USA: Contrasting diagenetic histories. Geochim. Cosmochim. Acta, 60(24): 5101-5116. [30] Graham R. P., MacNamara J., Crocker I. H. & MacFarlene R. B. (1951) The isotopic constitution of germanium. Can. J. Chem., 29: 89-102. [28] Grassineau N. V., Mattey D. P. & Lowry D. (2001) Sulfur isotope analysis of sulfide and sulfate minerals by continuous flow-isotope ratio mass spectrometry. Anal. Chem., 73(2)" 220-225. [26, 45] Gray J. & Tompson P. (1976) Climatic information from 180 / 160 ratios of cellulose in tree rings. Nature, 262: 481-482. [23] Gray J. & Thompson P. (1977) Climatic information from 180/160 analysis of cellulose, lignin and whole wood from tree rings. Nature, 270: 708-709. [24] Gray J., Yonge C. & Thompson P. (1984) Simultaneous determination of 180/160 and D / H in compounds containing C, H and O. Int. J. Appl. Radiat. Isot., 35(6)" 525-530. [1] Green J. W. (1963) Wood cellulose. In: Methods of carbohydrate chemistry. R. L. Whistler (ed.). Academic Press, NY: 9-21. [24] Green L. W., Leppinen J. J. & Elliot N. L. (1988) Isotopic analysis of lithium as thermal dilithium fluoride ions. Anal. Chem., 60: 34-37. [6] Green M. D., Rosman K. J. R. & De Laeter J. R. (1986) The isotopic composition of germanium in terrestrial samples. Int. J. Mass Spectrom Ion Proc., 68: 15-24. [31] Greenwood J. P., Mojzsis S. J. & Coath C. D. (2000) Sulfur isotopic compositions of individual sulfides in martian meteorites ALH84001 and Nakhla: Implications for crust-regolith exchange on Mars. Earth Planet. Sci. Lett., 184: 23-35. [30] Greenwood N. N. & Earnshaw A. (1997) Chemistry of the Elements. Butterworth & Heinemann. [26]

References

1097

Gregory R. T. & Taylor H. P. Jr. (1986) Possible non-equilibrium oxygen isotope effects in mantle nodules an alternative to the Kyser-O'Neil-Carmichael 180/16 0 geothermometer. Contrib. Mineral. Petrol., 93: 114-119. [46] Grey D. C. & Jensen M. L. (1972) Bacteriogenic sulfur in air pollution. Science, 177: 1099-1100. [26] Gr6goire D. C., Acheson B. M. & Taylor R. P. (1996) Measurement of lithium isotope ratios by inductively coupled plasma mass spectrometry: Application to geological materials. J. Anat. Atom. Spectrom., 11: 765-772. [6, 31] Griffith K. S. & Gellene G. I. (1992) Symmetry restrictions in diatom/diatom reactions. II. Nonmassdependent isotope effects in the formation of 04 +. J. Chem. Phys., 96: 4403-4411. [18] Griffiths H. (1998) Stable isotopes: Integration of biological ecological and geochemical processes. Bios Scientific Publishers, Oxford, UK: 438 p. [9, 24] Grimm W. (1968) Eine neue Glimmentladungslampe ftir die Optische Emissionspsektralanalyse. Spectrochim. Acta, B 23: 443-450. [35] Grinenko V. A., Grinenko L. N. & Zagryazhskaya G. D. (1969) Kinetic isotope effect in high temperature reduction of sulfate. Geokhimiya, 4: 484-491. [46] Grinenko V. A., Mineyev S. D., Devirts A. L. & Laguntina Y. P. (1987) Hydrogen isotope fractionation in the lizardite-water system at 100~ and I atm. Geokhimiya, 4: 574-579. [46] Grinsted M. J., Wilson A. T. & Ferguson C. W. (1979) 13C / 12C ratio variations in Pinus Iongaeva (Bristlecone pine) cellulose during the last millenium. Earth Planet. Sci. Lett., 42: 251-253. [24] Grob K. (1986) Making and manipulating capillary columns for gas chromatography. Dr. Alfred Huethig Verlag, Heidelberg, Germany. [8] Grossman E. L. & Ku T.-L. (1986) Oxygen and carbon isotope fractionation in biogenic aragonite: Temperature effects. Chem. Geol., 59: 59-74. [46] Grove M. (1993) Thermal histories of southern California basement terranes. PhD Thesis, University of California, Los Angeles, California, USA. [16] Grove M. & Bebout G. E. (1995) Cretaceous tectonic evolution of coastal southern California: insights from the Catalina Schist. Tectonics, 14: 1290-1308. [16] GriJning M., Fr/Jhlich K., De Regge P. P. & Danesi P. R. (1999) Intended use of the IAEA reference materials Part II: Examples on reference materials for stable isotope composition. In: The use of matrix reference materials in environmental analytical processes. A. Fajgelj & M. Parkany (eds.). Royal Society of Chemistry, Cambridge. Spec. Pub. No. 238: 81-92. [40] Gr/Jning M., van Duren M., Andreescu L., Tanweer A., Jaklitsch M. & Gattin M. (2004) Quality assurance for stable isotope analysis of water samples by mass spectrometry. Accr. Qual. Assur., (in press). [40] Guenther J., Krankowski D. & Mauersberger K. (2000) Third-body dependence of rate coefficients for ozone formation in 160-180 mixtures. Chem. Phys. Lett., 324: 31-36. [18] Guillou C. (1986) Application de la mesure du fractionnement isotopique de l'hydrogbne a la d6termination de l'enrichissement des vins. PhD Thesis, University of Nantes, Nantes, France. [5] Guillou C., Trierweiler M. & Martin G. J. (1988) Repeatability and reproducibility of site-specific isotope ratios in quantitative 2H NMR. Magn. Reson. Chem., 26: 491-496. [5] Guillou C. (1991) L'analyse isotopique des vins. Analusis Magazine, 19(10): M29-M34. [5] Guillou C. & Martin G. J. (1993) Characterization of reference tetramethylurea (TMU) for the determination of H / D ratio in alcohols by SNIF-NMR. BCR information - chemical analysis. Report CDNA-14396-EN-C. Commission of the European Communities, Directorate XIII, Luxembourg: 43 p. [5] Guillou C., Remaud G. & Lees M. (2001) EUR 20064- The certification of the carbon-13 and deuterium isotopic ratio content and the alcoholic grade of ethanol from wine, BCR-656 (96% vol). The carbon13 isotopic content of sugar, BCR-657. The oxygen-18 isotopic ratio content of water from wine, BCR-658 (7% vol) BCR-659 (12% vol). The carbon-13 and deuterium isotopic ratio content and the alcoholic grade of ethanol from wine BCR-660 (12% vol). European Commission, BCR information series, Office for Official Publications of the European Communities. Luxembourg: 94 p. (also available under URL: http://www.irmm.jrc.be/mrm.html). [40] Gulens J., Olmstead W. J., Denovan A. S., Gale K. L. & Rolston J. H. (1985) Determination of oxygen isotopes in water by direct oxidation and oxygen mass spectrometry. Int. J. Appl. Radiat. Isot., 36: 663667. [1]

1098

References

GUM (1995) Guide to the expression of uncertainty in measurement. IFCC, ISO, IUPAC, IUPAP, OIML. International Organization for Standardization, Geneva: 101 p. ISBN 92-67-10188-9. [32, 42] Gunter B. D. & Gleason J. D. (1971) Isotope fractionation during gas chromatographic separations. J. Chromat. Sci., 9: 191-192. [38] Guo J. & Qian Y. (1997) Hydrogen isotope fractionation and hydrogen diffusion in the tourmalinewater system. Geochim. Cosmochim. Acta, 61: 4679-4688. [46] Guzowski J. P. Jr., Broekaert J. A. C. & Hieftje G. M. (2000) Electrothermal vaporization for sample introduction into a gas sampling glow discharge time-of-flight mass spectrometer. Spectrochim. Acta, B 55: 1295-1314. [35] Giinther D., Frischknecht R. & Heinrich C. (1977) Capabilities of a 193 nm Arf excimer laser for LAICP-MS micro analysis of geological materials. J. Anal. At. Spectrom., 12: 939-944. [20] Giinther D. & Heinrich C. A. (1999) Comparison of the ablation behaviour of 266 nm Nd:YAG and 193 nm ArF excimer lasers for LA-ICP-MS analysis. J. Anal. At. Spectrom., 14: 1369-1374. [20] Gwiazda R., Woolard D. & Smith D. (1998) Improved lead isotope ratio measurements in environmental and biological samples with a double focusing magnetic sector inductively coupled plasma mass spectrometer. J. Anal. At. Spectrom., 13: 1233-1238. [35] Gygli A. (1993) Microdetermination of oxygen in organic compounds using a glassy carbon pyrolysis tube and nondispersive infrared detection. Mikrochim. Acta, 111: 37-43. [23] Haberhauer G. & Blochberger K. (1999) A simple cleanup method for the isolation of nitrate from natural water samples for O isotope analysis. Anat. Chem., 71: 3587-3590. [15] I-Iabfast K. (1960) Praezisionsbestimmung von Isotopenhaufigkeiten mit dem Massenspektrometer. Z. Instrumentenk., 68: 82-86. [39] I-Iabfast K. (1997) Advanced isotope ratio mass spectrometry I: Magnetic isotope ratio mass spectrometers. In: Modern isotope ratio mass spectrometry. I. T. Platzner (ed.). John Wiley & Sons, Chichester, UK: 11-82 [31, 38] I-Iabfast K. (1998) Fractionation correction and multiple collectors in thermal ionization isotope ratio mass spectrometry. Int. J. Mass Spectrom., 176: 133-148. [31] Habicht K. S. & Canfield D. E. (1997) Sulfur isotope fractionation during bacterial sulfate reduction in organic-rich sediments. Geochim. Cosmochim. Acta, 61: 5351-5361. [46] Habicht K. S., Canfield D. E. & Rethmeier J. (1998) Sulfur isotope fractionation during bacterial sulfate reduction and disproportionation of thiosulfate and sulfite. Geochim. Cosmochim. Acta, 62: 2585-2595. [46] Haendel D., Muhle K., Nitzsche H.-M., Stiehl G. & Wand U. (1986) Isotopic variations of the fixed nitrogen in metamorphic rocks. Geochim. Cosmochim. Acta, 50: 749-758. [16] Hagedorn M. L. (1992) Differentiation of natural and synthetic benzaldehydes by 2H nuclear magnetic resonance. J. Agric. Food Chem., 40: 634-637. [5] Hagemann R., Nief G. & Roth E. (1970) Absolute isotopic scale for deuterium analysis of natural waters. Absolute D / H ratio for SMOW. Tellus, 22: 712-715. [40, 43] Hagemann R. & Lohez P. (1978) Twin mass spectrometers for simultaneous isotopic analysis of hydrogen and oxygen in water. Adv. Mass Spectrom., 7A: 504-508. [1] Hagiwara Y. (2000) Selenium isotope ratios in marine sediments and algae - A reconaissance study. MSc Thesis, Univ. of Illinois at Urbana-Champaign, Urbana, IL, USA: 62 p. [29] Haglund R. F. Jr. & Itoh N. (1994) Electronic processes in laser ablation of semiconductors and insulators. In: Laser ablation. J. C. Miller (ed.). Springer-Verlag. Springer series in material sciences, Vol. 28: 11-52. [20] Haimson M. & Knauth L. P. (1983) Stepwise fluorination- a useful approach for the isotopic analysis of hydrous minerals. Geochim. Cosmochim. Acta, 47: 1589-1595. [20] Haisch M., Hering P., Fabinski W. & Z/~chbauer M. (1996) Isotopenselektive Kozentrationsmessungen an Atemgasen mit einem NDIR-Spektrometer. Technisches Messen, 63: 322-328. [34] Halas S. (1979) An automatic inlet system with pneumatic changeover valves for isotope ratio mass spectrometer. J. Phys. E: Sci. Instrum., 12: 418-420. [26] Halas S. & Skorzynski Z. (1980) An inexpensive device for digital measurements of isotopic ratios. J. Phys., E: Sci. Instrum., 13: 346-349. [39]

References

1099

Halas S. & Wolacewicz W. P. (1981) Direct extraction of sulfur dioxide from sulfates for isotopic analysis. Anal. Chem., 53: 686-689. [26, 45] Halas S. (1987) On bias in 34S/34S data obtained using SO2 gas in mass spectrometry. Advisory Group Meeting on Studies on Sulphur Isotope Variations in Nature. IAEA, Vienna: 105-111 [26]. Halas S. & Sikora J. (1987) Precise m / e monitoring in a magnetically scanned mass spectrometer by using two Hall probes. J. Phys. E., Sci. Instrum., 20: 662-664. [39] I-Ialas S. & Sikora J. (1990) Electron emission stabiliser with double negative feedback loop. Meas. Sci. Technol., 1: 980-982. [39] Halas S., Kaminski A. & Durakiewicz T. (1993) A bridge-based temperature controller with a resister as temperature sensor and heater. Meas. Sci. Technol., 4: 1208-1212. [39] Halas S. & Durakiewicz T. (1995) Improved adsorbant trap for gases. Chem. Geol. (Isot. Geosci. Sect.), 122: 271-273. [36] Halas S. & Kaminski A. (1995) An instrumental bridge. Patent description PL 1667134 B1, (in Polish). [39] Halas S. & Jasinska B. (1996) Platinized magnesium as reducing agent for hydrogen isotope analysis. Isot. Environm. Health Stud., 32(1): 105-109. [1] Halas S., Szaran J. & Niezgoda H. (1997) Experimental determination of carbon isotope equilibrium fractionation between dissolved carbonate and carbon dioxide. Geochim. Cosmochim. Acta, 61: 26912695. [46] Halas S. & Durakiewicz T. (1998a) Temperature distribution along a metal filament heated in vacuum by DC current. Vacuum, 49: 331-336. [39] Halas S. & Durakiewicz T. (1998b) Filament temperature stabilizer for a thermal ionization mass spectrometer. Int. J. Mass Spectrom. Ion. Proc., 181: 167-171. [39] Halas S. & Szaran J. (1999) Low-temperature thermal decomposition of sulfates to SO2 for on-line 34S/ 32S analysis. Anal. Chem., 71: 3254-3257. [26, 44] I-lalas S., Szaran J. & Niezgoda H. (2000) Reply to the comment by M. P. Verma on "Experimental determination of carbon isotope equilibrium fractionation between dissolved carbonate and carbon dioxide". Geochim. Cosmochim. Acta, 64: 575-576. [46] Halas S. & Szaran J. (2001) Improved thermal decomposition of sulfates to SO2 and mass spectrometric determination of 634S of IAEA SO-5, IAEA SO-6 and NBS-127 sulfate standards. Rapid Commun. Mass Spectrom., 15: 1618-1620. [26, 40] Halas S., W6jtowicz A., Nowak J. & Durakiewicz T. (2001) Temperature controller for thermal ionization mass spectrometry. Rapid Commun. Mass Spectrom., 15: 1-4. [39] I-Ialicz L., Galy A., Belshaw N. S. & O'Nions, R. K. (1999) High-precision measurement of calcium isotopes in carbonates and related materials by multiple collector inductively coupled plasma mass spectrometry (MC-ICP-MS). J. Anal. Atom. Spectrom., 14: 1835-1838. [31] Hall A. (1999) Ammonium in granites and its petrogenitic significance. Earth Sci. Revs., 45: 145-165. [16] Hall G. E. M., Pelchat J.-C. & Loop J. (1988) Separation and recovery of various sulphur species in sedimentary rocks for stable isotopic determination. Chem. Geol., 67: 35-45. [26] Hall J. A., Barth J. A. C. & Kalin R. M. (1999) Routine analysis by high precision gas chromatography mass selective detector isotope ratio mass spectrometry to 0.1 parts per mil. Rapid Commun. Mass Spectrom., 13: 1231-1236. [8] Halliday A. N., Lee D.-C., Christensen J. N., Walder A. J., Freedman P. A., Jones C. E., Hall C. M., Yi W. & Teagle D. (1995) Recent developments in inductively coupled plasma magnetic sector multiple collector mass spectrometry. Intl. J. Mass Spec. Ion Proc., 146 / 147: 21-33. [31, 35] Halliday A. N., Lee D.-C., Christensen J. N., Rehk~imper M., Yi W., Luo X., Hall C. M., Ballentine C. J., Pettke T. & Stifling C. (1998) Applications of multiple collector-ICPMS to cosmochemistry, geochemistry, and paleoceanography. Geochim. Cosmochim. Acta, 62(6): 919-940. [31] Halliday A. N., Christensen J. N., Lee D.-C., Rehktimper M., Hall C. M. & Luo X. (2000a) Multiplecollector inductively coupled plasma mass spectrometry. In: Inorganic mass spectrometry. C. M. Barshick, D. C. Duckworth & D. H. Smith (eds.). Marcel Dekker: 291-328. [31] Halliday A. N., Rehk/imper M., Schi~nb/ichler M., Oberli F., Freedman P. A., Frank M., Miiller W., Teutsch N., Baur H. & Wiechert U. (2000b) MC-ICPMS - the good, the small and the massive. J. Conf. Abstr., 5:474 (Abstract). [31]

11 O0

References

Halsted R. E. & Nier A. O. (1950) Gas flow through the mass spectrometer viscous leak. Rev. Sci.

Instrum., 21: 1019-1021. [38] Hamelin B., Manhes G., Albar6de F. & All6gre C. J. (1985) Precise lead isotope measurements by the double spike technique; a reconsideration. Geochim. Cosmochim. Acta, 49(1): 173-182. [29] Hamester M., Wiederin D., Wills J., Kerl W. & Douthitt C. B. (1999) Strategies for isotope ratio measurements with a double focusing sector field ICP-MS. Fresenius J. Anal. Chem., 364: 495-497. [31] Hamza M. S. & Broecker W. S. (1974) Surface effect on the isotopic fractionation between CO2 and some carbonate minerals. Geochim. Cosmochim. Acta, 38: 669-681. [46] Hamza M. & Epstein S. (1980) Oxygen isotopic fractionation between different oxygen sites in hydrous silicate minerals. Geochim. Cosmochim. Acta, 44: 173-182. [2, 20] Hang W., Walden W. O. & Harrison W. W. (1996) Microsecond pulsed glow discharge as an analytical spectroscopic source. Anal. Chem., 68: 1148-1152. [35] Hang W., Yan X., Wayne D. M., Olivares J. A., Harrison W. W. & Majidi V. (1999) Glow discharge source interfacing to mass analyzers: Theoretical and practical considerations. Anal. Chem., 71: 32313237. [35] Hanneguelle S., Thibault J. N., Naulet N. & Martin G. J. (1992) Authentication of essential oils containing linalool and linalyl acetate by isotopic methods. J. Agric. Food Chem., 40: 81-87. [5] Hanschmann G. (1981) Berechnung von isotopeneffekten auf quantenchemischer Grundlage am Beispiel stickstoffhaltiger Molekule. Zitl.-Mitt., 41: 19-39. [16] Hansen C. S., Calaway W. F., Pellin M. J., Wiens R. C. & Burnett D. S. (1996) Three-color resonance ionization spectrometry of Zn in Si. Resonance Ionization Spectroscopy 1996, Eighth International Symposium N. Winograd & J.E. Parks (eds.). AIP Press Woodbury, NY, AIP Conference Proceedings 388: 215-218. [33] Hanson D. & Mauersberger K. (1986) The vapor pressure of solid and liquid ozone. J. Chem. Phys., 85: 4669-4672. [18] Hansson S., Hobbie J. E., Elmgren R., Larsson U., Fry B. & Johansson S. (1997) The stable nitrogen isotope ratio as a marker of food-web interactions and fish migration. Ecology, 78: 2249-2257. [9] Hardcastle K. G. & Friedmann I. (1974) A method for oxygen isotope analysis of organic material. Geophys. Res. Lett., 1: 165-167. [23] Hare P. E., Fogel M. L., Stafford T. W., Mitchell A. D. & Hoering T. C. (1991) The isotopic composition of carbon and nitrogen in individual amino aciuds isolated from modern and fossil proteins. J. Archaeol. Sci., 18: 277-292. [9] Harkins W. D. & Stone S. B. (1926) The isotopic composition and atomic weight of chlorine from meteorites and from minerals of non-marine origin. J. Am. Chem. Soc., 48: 938-949. [28] Harmon R. S., Schwarcz H. P & O'Neil J. R. (1979) D / H ratios in speleothem fluid inclusions; a guide to variations in the isotopic composition of meteoric precipitation. Earth Planet. Sci. Lett., 42: 254-266. [3] Harmon R. S. & I-Iinton R. W. (1992) Frontiers in isotope geosciences, part II. Chem. Geol., 102: 169-199. [30] Harren F. & Reuss J. (1997) Photoacoustic spectroscopy. In: Encyclopedia of applied physics. G. L. Trigg (ed.).Weinheim, Vol. 19: 413-435. [34] Harren F. J. M., Cotti G., Oomens J. & te Lintel Hekkert S. (2000) Photoacoustic spectroscopy in trace gas monitoring. In: Encyclopedia of analytical chemistry. R. A. Meyers (ed.). John Wiley Ltd, Chichester: 2203-2226. [34] Harrington R. R., Kennedy B. P., Chamberlain C. P., Blum J. D. & FoR C. L. (1998) 15N enrichment in agricultural catchments: field patterns and applications to tracking Atlantic salmon (Salmo salar). Chem. Geol., 147: 281-294. [15] Harris C., Compton J. S. & Bevington S. A. (1999) Oxygen and hydrogen isotope composition of kaolinite deposits, Cape Peninsula, South Africa: Low-temperature, meteoric origin. Econ. Geol., 94: 13531366. [2] Harris N. B. W., Jackson D. H., Mattey D. P., Santosh M. & Bartlett J. (1993) Carbon-isotope constraints on fluid advection during contrasting examples of incipient charnockite formation. J. Metam. Geol., 11: 833-843. [3]

References

1101

Harrison A. G. & Thode H. G. (1957) The kinetic isotope effect in chemical reduction of sulphate. Trans. Faraday Soc., 53: 1648-1651. [46] Harrison A. G. & Thode H. G. (1958) Mechanism of the bacterial reduction of sulphate from isotope fractionation studies. Trans. Faraday Soc., 54: 84-92. [46] Harrison W. G. (1978) Experimental measurements of nitrogen remineralization in coastal waters. Limnol. Oceanogr., 23: 684-694. [15] Harrison W. W. & Magee C. W. (1974) Hollow cathode ion source for solids mass spectrometry. Anal. Chem., 46: 461-464. [35] Harrison W. W., Hess K. R., Marcus R. K. & King F. L. (1986) Glow discharge mass spectrometry. Anat. Chem., 58: 341A-356A. [35] Harrison W. W. (1988) Glow discharge mass spectrometry. In: Chemical analysis. F. Adams, R. Gijbels & R. Van Grieken (eds.). Wiley-Interscience, NY, Vol. 95: 85-123. [35] Harrison W. W. & Bentz B. L. (1988) Glow discharge mass spectrometry. Prog. Analyt. Spectrosc., 11: 53110.[35] Harrison W. W. (1998) Pulsed glow discharge as a solid analysis source. J. Anat. At. Spectrom., 13: 10511056. [35] Hart S. R. & Zindler A. (1989) Isotope fractionation laws: a test using calcium. Int. J. Mass Spectrom. Ion Proc., 89: 287-301. [31] Harte B., Fitzsimons I. C. W., Harris J. W. & Otter M. L. (1999) Carbon isotope ratios and nitrogen abundances in relation to cathodoluminescence characteristics for some diamonds from the Kaapvaal Province, S. Africa. Mineral. Magazine, 63(6): 829-856. [30] Hartley P. E. (1980) Mass spectrometric determination of deuterium in water at natural levels. Anat. Chem., 52: 2232-2234. [1] Hartmann M. & Nielsen H. (1969) 634S-Werte in rezenten Meeressedimenten und ihre Deutung am Beispiel einiger Sedimentprofile aus der westlichen Ostsee. Geol. Rundschau, 58: 621-655. [27] Hashimoto I. & Jackson M. L. (1960) Rapid dissolution of allophane and kaolinite-halloysite after dehydration. Clays Clay Minerals, 7: 102-113. [2] Hashizume K. & Sugiura N. (1990) Precise measurement of nitrogen isotopic composition using a quadrupole mass spectrometer. Mass Spectroscopy, 38: 269-286. [13, 17] Hashizume K. & Sugiura N. (1992) Measurement of cosmogenic nitrogen using a static massspectrometry system and its implication. Geochim. Cosmochim. Acta, 56: 1625-1631. [17] Hashizume K. & Sugiura N. (1995) Nitrogen isotopes in bulk ordinary chondrites. Geochim. Cosmochim. Acta, 59: 4057-4069. [16] Hashizume K., Chaussidon M., Marty B. & Robert F. (2000) Solar wind record on the moon: Deciphering presolar from planetary nitrogen. Science, 290: 1142-1145. [17] Hashizume K., Marty B. & Weiler R. (2002) Analyses of nitrogen and argon in single lunar grains: towards a quantification of the asteroidal contribution to planetary surfaces. Earth Planet. Sci. Lett., 202: 201-216. [17] Hassan A. A. (1982) Methodologies for extraction of dissolved inorganic carbon for stable carbon isotope studies evaluation and alternatives. U. S. Geol. Survey, Water Resour. Invest. Report No. 82-6: 51p. [10] Hathorn B. C. & Marcus R. A. (1999) An intramolecular theory of the mass-independent isotope effect for ozone. I. J. Chem. Phys., 111: 4087-4100. [15, 18] Hathorn B. C. & Marcus R. A. (2000) An intramolecular theory of the mass-independent isotope effect for ozone. II. Numerical implementation at low pressure using a loose transition state. J. Chem. Phys., 113(21): 9497-9509. [18] Hattori K. & Sakai H. (1979) D / H ratios, Origins, and Evolution of the Ore-Forming Fluids for the Neogene Veins and Kuroko Deposits of Japan. Econ. Geol., 74: 535-555. [3] I-Iauck R. D. & Bremner J. M. (1976) Use of tracers for soil and fertilizer nitrogen research. Adv. Agron., 28: 219-266. [15] Hauck R. D. (1982) Nitrogen-isotope-ratio analysis. In: Methods of soil analysis, Part 2. Chemical and microbiological properties. A. L. Page, R. H. Miller & D. R. Keeney (eds.). American Society of Agronomy Inc. - Soil Science Society of America Inc., Madison, Wisconsin. Agronomy Series, 2nd edition, No. 9: 735-779. [15]

1102

References

Haur A., Hladikov~i J. & Smejkal V. (1973) Procedure of direct conversion of sulfates into SO2 for mass spectrometric analysis of sulfur. Isotopenpraxis, 9: 329-331. [26] Hayatsu R. & Anders E. (1981) Organic compounds in meteorites and their origins. Top. Curr. Chem., 99: 1-37.[11] Hayes J. M. & Schoeller D. A. (1977) High precision pulse counting: Limitations and optimal conditions. Anal. Chem., 49(2): 305-311. [30] Hayes J. M. & Baker K. L. (1986) Reagent and procedure for preparation of H2 for hydrogen-isotopic analysis of water. Indiana University, Bloomington. Intern. Rpt. (unpublished). [1] Hayes J. M. & Johnson M. W. (1988) Reagent and procedure for preparation of H2 for hydrogen-isotopic analysis of water. Indiana University, Bloomington. Intern. Rpt. (unpublished). [1] Hayes J. M., Freeman K. H., Popp B. N. & Hoham C. H. (1990) Compound-specific isotopic analyses, a novel tool for reconstruction of ancient biogeochemical processes. Org. Geochem., 16: 1115-1128. [9, 26, 45] Hays P. A., Remaud G. S., Jamin E. & Martin Y. L. (2000) Geographic origin determination of heroin and cocaine using site-specific isotopic ratio deuterium NMR. J. Forensic Sci., 45(3): 552-562. [5] Heathcote J. A. (1981) Hydrochemical aspects of the Gipping Chalk salinity investigation. PhD Thesis, University of Birmingham. [10] Heathcote J. A. (1985) Carbonate chemistry of recent Chalk groundwater in part of East Anglia. J. Hydrol., 78: 215-227. [10] I-Ieaton T. H. E. & Collett G. M. (1985) The analysis of 15N / 14N ratios in natural samples, with emphasis on nitrate and ammonium in precipitation. Conc. Sci. Ind. Res., Pretoria. Res. Report: 624. [15] Heaton T. H. E. (1999) Spatial, species, and temporal variations in the 13C / 12C ratios of C3 plants: implications for palaeodiet studies. J. Archaeological Soc., 26: 637-649. [46] Heidenreich J. E. III & Thiemens M. H. (1986) A non-mass-dependent oxygen isotope effect in the production of ozone from molecular oxygen: The role of molecular symmetry in isotope chemistry. J. Chem. Phys., 84: 2129-2136. [18, 46] Heier K. S. & Billings G. K. (1972) Lithium. In: Handbook of Geochemistry. K. H. Wedepohl (ed.). Springer-Verlag, Vol. II-I, Chap. 3. [6] Hein R., Crutzen P. J. & Heimann M. (1997) An inverse modeling approach to investigate the global atmospheric methane cycle. Global Biogeochem. Cycles, 11(1): 43-76. [14] I-Ieinrichs H. & Herrmann A. G. (1990) Praktikum der analytischen Geochemie. Springer-Verlag. [27] Helle J. F., Hillaire-Marcel C. & Rondeau B. (2002) Seasonal changes in the sources and fluxes of dissolved inorganic carbon through the St. Lawrence River- isotopic and chemical constraint. Chem. Geol., 186: 117-138. [10] Helliker B. R. & Ehleringer J. R. (2000) Establishing a grassland signature in veins: 180 in the leaf water of C3 and C4 grasses. Proceed. Nat. Acad. Sci., 97(14): 7894-7898. [21] Helliker B. R., Roden J. S., Cook, C. & Ehleringer J. R. (2002) A rapid and precise method for sampling and determining the oxygen isotope ratio of atmospheric water vapor. Rapid Commun Mass Spectr., 16: 929-932. [1] Hemming N. G. & Hanson G. N. (1992) Boron isotopic composition and concentration in modern marine carbonates. Geochim. Cosmochim. Acta, 56: 537-543. [7, 46] Hemming N. G. & Hanson G. N. (1994) A procedure for the isotopic analysis of boron by negative thermal ionization mass spectrometry. Chem. Geol. (Isot. Geosci. Sect.), 114: 147-156. [7, 29] Henchman M. J., Smith D. & Adams N. G. (1982) Estimation of the enthalpy changes in several ionmolecule reactions involving H / D exchange from zero-point-energy considerations. Int. J. Mass Spectrom. Ion Phys., 42: 25-32. [4] Henderson J. H., Clayton R. N., Jackson M. L., Syers J. K., Rex R. W., Brown J. L. & Sachs I. B. (1972) Cristobalite and quartz isolation from soils and sediments by hydrofluosilicic acid treatment and heavy liquid separation. Soil Sci. Soc. Am. Proc., 36: 830-835. [2] Hener U., Brand W. A., I-Iilkert A. W., Juchelka D., Mosandl A. & Podebrad F. (1998) Simultaneous online analysis of 180/16 0 and 13C/12C ratios of organic compounds using GC-pyrolysis-IRMS. Food Res. Technol. Z. Lebensm. Unters. Forsch., A206(3): 230-232. [8, 23] Herbel M. J. (1997) Stable selenium isotopic ratios of saline environmental samples from the western San Joaquin Valley of California. PhD Thesis, University of California, Davis. [29]

References

1103

Herbel M. J., Johnson T. M., Oremland R. S. & Bullen T. D. (1998) Selenium stable isotope fractionation during bacterial dissimilatory reduction of selenium oxyanions. Geochim. Cosmochim. Acta, 64: 37013709. [291 Herbel M. J., Johnson T. M., Oremland R. S. & Bullen T. D. (2000) Fractionation of selenium isotopes during bacterial respiratory reduction of selenium oxyanions. Geochim. Cosmochim. Acta, 64: 37013709. [29, 46] Herbelin J. M., McKay J. A., Kwok M. A., EUenten R. H., Urevig D. S., Spencer D. J. & Bernard D. J. (1980) Sensitive measurement of photon lifetime and true reflectances in an optical cavity by a phase-shift method. Appl. Opt., 19: 144-147. [34] Herbert B. K. (1976) A circuit for stabilizing the electron current to the anode of a hot-filament device. Vacuum, 26: 363-369. [39] Herczeg A. L. (1987) A stable isotopic study of dissolved inorganic carbon cycling in a softwater lake. Biogeochemistry, 4: 231-263. [10] Hermann A. (1999) Determination of site-specific D / H isotope ratios of glycerol from different sources by 2H-NMR spectroscopy. Z. Lebensm. Unters. Forsch., A 208: 194-197. [5] Hermann A. (2001) Determination of D / H isotope ratio in acetic acid from vinegars and pickled products by 2H-NMR-spectroscopy. Eur. Food Res. Technol., 212: 683-686. [5] Herriott D. R., Kogelnik H. & Kompfner R. (1964) Off-axis paths in spherical interferometers. Appl. Opt., 3: 523-526. [34] Hersh C. K. (1961) Molecular sieves. Reinhold, NY: 129 p. [36] Herut B., Spiro B., Starinsky A. & Katz A. (1995) Sources of sulfur in rainwater as indicated by isotopic h34S data and chemical composition, Israel. Atmos. Environ., 29: 851-857. [26] Hervig R. L. (1992) Oxygen isotope analysis using extreme energy filtering. Chem. Geol. (Isot. Geosci. Sect.), 101: 185-186. [30] Hervig R. L., Williams P., Thomas R. N., Schauer S. N. & Steele I. M. (1992) Microanalysis of oxygen isotopes in insulators by secondary ion mass spectrometry. Int. J. Mass Spec. Ion Proc., 120: 45-63. [30] Herzog G. F., Xue S., Hall G. S., Nyquist L. E., Shih C.-Y., Wiesmann H. & Brownlee D. E. (1999) Isotopic and elemental composition of iron, nickel, and chromium in type I deep-sea spherules: Implications for origin and composition of the parent micrometeoroids. Geochim. Cosmochim. Acta, 63(9): 1443-1457. [29] Herzog R. (1934) Ionic- and electro-optical cylindrical lenses. Z. Phys., 89: 447-. [38] Herzog W. & Di~rnenburg E. (1958) Die Verwendung von Methylchlorid zur massenspektrometrischen Isotopenanlyse von Chlor. Z. Naturforsch., 13a: 51-52. [28] Herzog W. & Klemm A. (1958) Die Temperaturabh~ingigkeit des Isotopie-Effektes bei der elektrolytischen Wanderungen der Chlorionen in geschmolzenem Thallium(I)-chlorid. Z. Naturforsch., 13a: 7-16. [28] Hesslein R. H., Capel M. J., Fox D. E. & Hallard K. A. (1991) Stable isotopes of sulfur, carbon, and nitrogen as indicators of trophic level and fish migration in the lower Mackenzie River Basin, Canada. Can. J. Fish Aquat. Sci., 48: 2258-2265. [26] Hesslein R. H., Hallard K. A. & Ramlal P. (1993) Replacement of sulfur, carbon, and nitrogen in tissue of growing broad whitefish (Coregonus nasus) in response to a change in diet traced by h34S, 613C, and 615N. Can. J. Fish. Aquat. Sci., 50: 2071-2076. [9] Heumann K. G. & Zeininger H. (1985) Boron trace determination in metals and alloys by isotope dilution mass spectrometry with negative thermal ionization. Int. J. Mass Spectrom. Ion Proc., 67: 237252. [7] Heumann K. G. & R~idlein N. (1989) Negative thermal ionization mass spectrometry of selenium, Part 3. Selenium trace determination in food samples. Fresenius Z. Anal. Chem., 335: 751-754. [29] Heumann K. G. & Wachsmann M. (1989) Negative thermal ionization mass spectrometry of selenium, Part 4. Selenium trace determination in sediments and related samples. Fresenius Z. Anal. Chem., 335: 755-758. [29] Heumann K. G., Eisenhut S., Gallus S., Hebeda E. H., Nusko R., Vengosh A. & Walczyk T. (1995) Recent developments in thermal ionization mass spectrometric techniques for isotope analysis: a review. Analyst, 120: 1291-1299. [7]

1104

References

Heumann K. G., Gallus S. M., Radlinger G. & Vogl J. (1998) Precision and accuracy in isotope ratio measurements by plasma source mass spectrometry. J. Anal. At. Spectrom., 13: 1001-1008. [26, 32, 35] Heuser E. (1944) Cellulose chemistry. J. Wiley & Sons. [24] Heymann D., Yancey T. E., Wolbach W. S., Thiemens M. H., Johnson E. A., Roach D. & Moecker S. (1998) Geochemical markers of the Cretaceous-Tertiary boundary event at Brazos River, Texas, USA. Geochim. Cosmochim. Acta, 62(1): 173-181. [45] Hildebrand P. & Beglinger C. (1997) Nondispersive infrared spectrometry. A new method for the detection of Heliobacter pylori infection with the 13C-urea breath test. Clin. Infect. Dis., 25: 1003-1005. [34] Hilkert A. W., Douthitt C. B., Schliiter H. J. & Brand W. A. (1999) Isotope ratio monitoring gas chromatography/mass spectrometry of D / H by high temperature conversion isotope ratio mass spectrometry. Rapid Commun. Mass Spectrom., 13: 1226-1230. [1, 14, 38] Hill J. W. & Fry A. (1962) Chlorine isotope effects in the reactions of benzyl and substituted benzyl chlorides with various nucleophiles. J. Am. Chem. Soc., 84: 2763-2769. [28] Hillier S., Fallick A. E. & Matter A. (1996) Origin of pore-lining chlorite in the aeolian Rotliegend of northern Germany. Clay Minerals, 31: 153-171. [2] Hinners T. A., Hughes R., Outridge P. M., Davis W. J., Simon K. & Woolard D. R. (1998) Interlaboratory comparison of mass spectrometric methods for lead isotopes and trace elements in NIST SRM 1400 Bone Ash. J. Anal. At. Spectrom., 13: 963-970. [35] Hinton R. W., Long J. V. P., Fallick A. E. & Pillinger C. T. (1983) Ion probe measurements of D / H ratios in meteorites. Lunar Planet. Sci., 14: 313-314. [30] Hinton R. W. (1995) Ion microprobe analysis in geology. In: Microbe techniques in Earth sciences. P. J. Potts, J. F. W. Bowles, S. J. B. Reed & M. R. Cave (eds.). Chapman & Hall, London: 235-289. [30] Hintze C. (1915) Handbuch der mineralogie. Verlag von Veit & Comp., Leipzig. [28] I-Iippler H., Rahn R. & Troe J. (1990) Temperature and pressure dependence of ozone formation rates in the range 1-1000 bar and 90-370 K. J. Chem. Phys., 93: 6560-6569. [18] Hirata T. (1996) Lead isotopic analyses of NIST standard reference materials using multiple collector inductively coupled plasma mass spectrometry coupled with a modified external correction method for mass discrimination effect. Analyst, 121: 1407-1411. [31] I-Iirata T. (1997) Isotopic variations of germanium for iron and stony iron meteorites. Geochim. Cosmochim. Acta, 61: 4439-4448. [31] Hirata T. & Ohno T. (2001) In-situ isotopic ratio analysis of iron using laser ablation-multiple collectorinductively coupled plasma mass spectrometry (LA-MC-ICP-MS). J. Anal. At. Spectrom., 16: 487-491. [31] Hirata T. (2003) Chemically assisted laser ablation ICPO mass spectrometry. Anal. Chem., 75: 228-233. [20] Hitchcock D. R. & Black M. S. (1984) 34S/32Sevidence of biogenic sulfur oxides in a salt marsh atmosphere. Atmos. Environ., 18: 1-17. [26] Hitchon B. (1974) Aplication of geochemistry to the search for crude oil and natural gas. In: Introduction to Exploration Geochemistry. A. A. Levinson (ed.). Applied Publishing Ltd.: 509-545 [26] Hobson K. A. & Clark R. G. (1992) Assessing avian diets using stable isotopes II: factors influencing diet-tissue fractionation. The Condor, 94: 189-197. [9] Hobson K. A., Hughes K. D. & Ewins P. J. (1997) Using stable isotope analysis to identify endogenous and exogenous sources of nutrients in eggs of migratory birds: applications to Great Lakes contaminants research. The Auk, 114: 467-478. [26] I-Iobson K. A. & Schell D. M. (1998) Stable carbon and nitrogen isotope patterns in baleen from eastern Artic bowhead whales (Balaena mysticetus). Can. J. Fish. Aquat. Sci., 55: 2601-2607. [9] Hobson K. A. (1999) Tracing origins and migration of wildlife using stable isotopes: a review. Oecologia, 120: 314-326. [26] Hobson K. A., Drever M. C. & Kaiser G. W. (1999) Norway rats as predators of burrow-nesting seabirds: insights from stable isotope analyses. J. Wildl. Manage., 63: 14-25. [26] Hockey J. A. & Pethica B. A. (1961) Surface hydration of silicas. Trans. Faraday Soc., 57: 2247-2262. [3] Hoefs J. & Stalder H. A. (1977) Die C-Isotopenzusammensetzung von CO2-haltigen Flfissigkeitseinschlfisse in Kluftquartzen der Zentralalpen. Schweiz. Miner. Petrol. Mitt., 57: 329-347. [3]

References

1105

Hoefs J. & Morteani G. (1979) The carbon isotopic composition of fluid inclusions in Alpine fissure quartzes from the western Tauern Window (Tyrol, Austria). Neues Jhb. Mineral., 3: 123-134. [3] Hoefs J. (1980) Stable Isotope Geochemistry. Springer-Verlag. 2nd edition: 203 p. [45] Hoefs J. (1997) Stable Isotope Geochemistry. Springer-Verlag. 4th completely revised, updated, and enlarged edition: 201 p. [App B] Hoefs J. & Sywall M. (1997) Lithium isotope composition of Quaternary and Tertiary biogene carbonates and a global lithium isotope balance. Geochim. Cosmochim. Acta, 61: 2679-2690. [6] Hoekstra H. R. & Katz J. J. (1953) Direct determination of oxygen in less familiar metal oxides. Anal. Chem., 25: 1608-1612. [20] Hoering T. (1955) Variations of nitrogen fifteen abundance in naturally occuring substances. Science, 122: 1233-1234. [16] I-Ioering T. (1957) The isotopic composition of ammonia and nitrate ion in rain. Geochim. Cosmochim. Acta, 12: 97-102. [15] Hoering T. C. & Parker P. L. (1961) The geochemistry of the stable isotopes of chlorine. Geochim. Cosmochim. Acta, 23: 186-199. [28] Hoering T. C. (1990) Development of the sulfur hexafluoride method for sulfur isotope analysis. Carnegie Institution, Washington. Ann. Rept. Dir. Geophys. Lab. 1989-1990, No. 2200: 128-131. [20] Hofmann A. (1971) Fractionation corrections for mixed-isotope spikes of Sr, K, and Pb. Earth Planet. Sci. Lett., 10(4): 397-402. [29] Hofmann D., Jung K., Segschneider H.-J., Gehre M. & Schiiiirmann G. (1995) 15N/14N analysis of amino acids with GC-C-IRMS - methodical investigations and ecotoxicological applications. Isot. Environ. Health Stud., 31: 367-375. [8] Hogg A. J. C., Pearson M. J. & Fallick A. E. (1993) Pretreatment of Fithian illite for oxygen isotope analysis. Clay Minerals, 18: 149-152. [2] Holloway J. M. & Dahlgren R. A. (2002) Nitrogen in rock: occurrences and biogeochemical implications. Global Biogeochem. Cycles, 14(4): doi 10.1029/2002GB001862,2002. [15] Holloway J. R. & Wood B. J. (1988) Simulating the Earth: Experimental Geochemistry. Unwin Hyman, Boston, USA: 196 p. [46] Holmden C., Papanastassiou D. A. & Wasserburg G. J. (1997) Negative thermal ion mass spectrometry of oxygen in phosphates. Geochim. Cosmochim. Acta, 61: 2253-2264. [22] Holmes R. M., McLelland J. W., Sigman D. M., Fry B. & Peterson B. J. (1998) Measuring 15N-NH4 + in marine, estuarine and fresh waters: An adaption of the ammonia diffusion method for samples with low ammonium concentrations. Marine Chem., 60: 235-243. [9, 15] Holmes R. M., Aminot A., Kerouel R., Hooker B. A. & Peterson B. J. (1999) A simple and precise method for measuring ammonium in marine and freshwater ecosystems. Can. J. Fish. Aquat. Sci., 56: 1801-1808. [9] Holt B. D. & Engelkemeir A. G. (1970) Thermal decomposition of barium sulfate to sulfur dioxide for mass spectrometric analysis. Anal. Chem., 42: 1451-1453. [26, 44, 45] Holt B. D., Engelkemeir A. G. & Venters A. (1972) Variations of sulfur isotope ratios in samples of water and air near Chicago. Environ. Sci. Tech., 6: 338-341. [26] Holt B. D. (1975) Determination of stable isotope ratios in the environment. Prog. Nucl. Energy Series, Anal. Chem., 12: 11-26. [26] Holt B. D., Sturchio N. C., Arehart G. B. & Bakel A. J. (1995) Ultrasonic vacuum extraction of gases from water for chemical and isotopic analysis. Chem. Geol. (Isot. Geosci. Sect.), 122: 275-284. [10] Holt B. D., Sturchio N. C., Abrajano T. A. & Heraty L. J. (1997) Conversion of chlorinated volatile organic compounds to carbon dioxide and methyl chloride for isotopic analysis of carbon and chlorine. Anal. Chem., 69: 2727-2733. [28] Honig R. E. (1945) Gas flow in the mass spectrometer. J. Appl. Phys., 16: 646-654. [38] Honma H. & Itihara Y. (1981) Distribution of ammonium in minerals of metamorphic and granitic rocks. Geochim. Cosmochim. Acta, 45: 983-988. [16] Hopple J. A., Hannon J. E. & Coplen T. B. (1998) Comparison of two stable hydrogen isotope-ratio measurement techniques on Antarctic surface-water and ice samples. Chem. Geol., 152: 321-323. [1] Horibe Y. & Kobayakawa M. (1960) Deuterium abundance of natural waters. Geochim. Cosmochim. Acta, 20: 273-283. [1, 40]

1106

References

Horibe Y. & Craig H. (1995) D / H fractionation in the system methane-hydrogen-water. Geochim. Cosmochim. Acta, 59(24): 5209-5217. [46] Horita J. & Matsuo S. (1986) Extraction and isotopic analysis of fluid inclusions in halites. Geochem. J., 20: 261-272. [1, 3] Horita J. (1988) Hydrogen isotope analysis of natural waters using an H2-water equilibration method: A special implication to brines. Chem. Geol. (Isot. Geosci. Sect.), 72: 89-94. [1, 3, 38] I-Iorita J. & Gat J. R. (1988) Procedure for the hydrogen isotope analysis of water from concentrated brines. Chem. Geol. (Isot. Geosci. Sect.), 72: 85-88. [1] Horita J. & Gat J. R. (1989) Deuterium in the Dead Sea: Remeasurement and implications for the isotopic activity correction in brines. Geochim. Cosmochim. Acta, 53: 131-133. [1] Horita J., Ueda A., Mizukami K. & Takatori I. (1989) Automatic 6D and 6180 analyses of multi-water samples using H2- and CO2-water equilibration methods with a common equilibration set-up. Appl. Radiot. Isot., 40(9): 801-805. [1, 3] Horita J. (1989a) Analytical aspects of stable isotopes in brines. Chem. Geol. (Isot. Geosci. Sect.), 79: 107112.[1] Horita J. (1989b) Stable isotope fractionation factors of water in hydrated saline mineral-brine systems. Earth Planet. Sci. Lett., 95: 173-179. [46] Horita J. (1990) Stable isotope paleoclimatology of brine inclusions in halite; modeling and application to Searles Lake, California. Geochim. Cosmochim. Acta, 54: 2059-2073. [3] Horita J., Cole D. R. & Wesolowski D. J. (1993a) The activity-composition relationship of oxygen and hydrogen isotopes in aqueous salt solutions: II. Vapor-liquid water equilibration of mixed salt solutions from 50 to 100~ and geochemical implications. Geochim. Cosmochim. Acta, 57: 4703-4711. [1, 46] Horita J., Wesolowski D. J. & Cole D. R. (1993b) The activity-composition relationship of oxygen and hydrogen isotopes in aqueous salt solutions. I. Vapor-liquid equilibration of single salt solutions from 50 to 100~ Geochim. Cosmochim. Acta, 57: 2797-2817. [1, 46] I-Iorita J. & Wesolowski D. J. (1994) Liquid-vapor fractionation of oxygen and hydrogen isotopes of water from the freezing to the critical temperature. Geochim. Cosmochim. Acta, 58: 3425-3437. [1] Horita J., Cole D. R. & Wesolowski D. J. (1995) The activity-composition relationship of oxygen and hydrogen isotopes in aqueous salt solutions: III. Vapor-liquid water equilibration of NaC1 solutions to 350~ Geochim. Cosmochim. Acta, 59(6): 1139-1151. [46] I-Iorita J., Dreisner T. & Cole D. R. (1999) Pressure effect on hydrogen isotope fractionation between brucite and water at elevated temperatures. Science, 286: 1545-1547. [46] Horita J. (2001) Carbon isotope exchange in the system CO2-CH4 at elevated temperatures. Geochim. Cosmochim. Acta, 65: 1907-1919. [46] Horita J., Cole D. R., Polyakov V. B. & Driesner T. (2002) Experimental and theoretical study of pressure effects on hydrogen isotope fractionation in the system brucite-water at elevated temperatures. Geochim. Cosmochim. Acta, 66: 3769-3788. [46] I-Iorlick G. & Montaser A. (1998) Analytical characteristics of ICPMS. In: Inductively coupled plasma mass spectrometry. A. Montaser (ed.).Wiley-VCH: 503-613. [31] Hornibrook E. R. C. & Longstaffe F. J. (1996) Berthierine from the Lower Cretaceous Clearwater formation, Alberta, Canada. Clays Clay Minerals, 44: 1-21. [2] I-Iouerou G., Kelly S. D. & Dennis M. J. (1999) Determination of the oxygen-18/oxygen-16 isotope ratios of sugar, citric acid and water from single strength orange juice. Rapid Commun. Mass Spectrom., 13: 1257-1262. [1] Hough R. M., Gilmour I., Pillinger C. T., Arden J. W., Gilkes K. W. R., Yuan J. & Milledge H. J. (1995) Diamonds and silicon carbide in impact melt rock from the Ries impact crater. Nature, 378: 41-44. [13] Houston P. L., Suits A. G. & Toumi R. (1996) Isotopic enrichment of heavy ozone in the stratosphere. J. Geophys. Res., 101: 18829-18834. [18] Houweling S., Kaminski T., Dentener F., Lelieveld J. & Heimann M. (1999) Inverse modeling of methane sources and sinks using the adjoint of a global transport model. J. Geophys. Res., 104(D21): 26173-26160.[14] Howald R. A. (1960) Ion pairs. I. Isotope effects shown by chloride solutions in glacial acetic acid. J. Am. Chem. Soc., 82: 20-24. [28]

References

1107

H s i e h J. C. C. (1997) An oxygen isotopic study of soil water and pedogenic clays in Hawaii. PhD Thesis,

California Institute of Technology, Pasadena, California, USA: 181 p. [2] Hsieh J. C. C., Savin S. M., Kelley E. F. & Chadwick O. A. (1998) Measurement of soil-water 6180 values by direct equilibration with CO2. Geoderma, 82: 255-268. [1, 2] Hsieh J. C. C. & Yapp C. J. (1999) Hydrogen isotope exchange in halloysite. Clays Clay Minerals, 47: 811816. [2] Hughes W. B., Holba A. G. & Dzou L. I. P. (1995) The ratios of dibenzothiophene to phenanthrene and pristane to phytane as indicators of depositional environment and lithology of petroleum source rocks. Geochim. Cosmochim. Acta, 59: 3581-3598. [8] Huh Y., Chan L. H., Zhang L. & Edmond J. M. (1998) Lithium and its isotopes in major world rivers: implications for weathering and the oceanic budget. Geochim. Cosmochim. Acta, 62: 2039-2052. [6] Huh Y., Chan L. H. & Edmond J. M. (2001) Lithium isotopes as a probe of weathering process: Orinoco River. Earth Planet. Sci. Lett., 194: 189-199. [6] Hulston J. R. & Thode H. G. (1965a) Cosmic-ray-produced 36S and 33S in the metallic phase of iron meteorites. J. Geophys. Res., 70(18): 4435-4442. [45] Hulston J. R. & Thode H. G. (1965b) Variations in the S33, S 34, and S36 contents of meteorites and their relation to chemical and nuclear effects. J. Geophys. Res., 70: 3475-3484. [20, 26, 45] Humbert F., Libourel G., France-Lanord C., Zimmermann L. &Marty B. (2000) CO2-1aser extractionstatic mass spectrometry analysis of ultra-low concentrations of nitrogen in silicates. Geostand. Newslett., 24: 255-260. [16, 17] Huneke J. C., Armstrong J. T. & Wasserburg G. J. (1983) Fun with Panurge: High mass resolution ion microprobe measurements of Mg in Allende inclusions. Geochim. Cosmochim. Acta, 47: 1635-1650. [25] Hunter K. L. & Gray J. W. (1993) ETH Technical Publication, TP007. [37] Hunkeler D., Hohener P., Bernasconi S. & Zeyer J. (1999) Engineered in situ bioremediation of a petroleum hydrocarbon-contaminated aquifer: assessment of mineralization based on alkalinity, inorganic carbon and stable carbon isotope balances. J. Cont. Hydrolgeol. 37: 201-223. [10] Hurai V., Simon D. & Bezik V. (1997) Contrasting chemistry and H, O, C isotope composition of greenschist-facies, Hercynian and Alpine metamorphic fluids (Western Carpathians). Chem. Geol., 136: 281-293. [3] Hurst G. S. & Payne M. G. (1988) Principles and applications of resonance ionization spectroscopy. Adam Hilger, Boston & Philadelphia. [33] Hurtgen T. M., Arthur M. A., Suits N. S. & Kaufman A. J. (2002) The sulfur isotopic composition of Neoproterozoic seawater sulfate: implications for a snowball Earth? Earth Planet. Sci. Lett., 203: 413429. [26] Husek P. (1995) Simultaneous profile analysis of plasma amino and organic-acids by capillary gaschromatography. J. Chromatogr. Biomed. Appl., 669: 352-357. [8] Husek P. & Matucha P. (1997) Simple approach for analysis of plasma oxo-, hydroxy- and dicarboxylic acids. J. Chromaogr., B 693: 499-502. [8] Husek P. (1998) Chloroformates in gas chromatography as general purpose derivatizing agents. J. Chromatogr., B 717: 57-91. [8] Hut G. (1987) Consultants' group meeting on stable isotope reference samples for geochemical and hydrological investigations. Scientific Meeting 16-18 Sept., 1985. IAEA, Vienna. Internal Report: 42 p. [1, 40, 43] Hutchinson D. A. (1954) Natural abundance of the lithium isotopes. Phys. Rev., 96: 1018-1021. [6] Huygen C. (1963) The sampling of sulfur dioxide in air with impregnated filter paper. Anal. Chim. Acta, 28: 349-360. [26] Hyeong, K. & Capuano, R. M. (2000) The effect of organic matter and the H202 organic-matter-removal method on the 6D of smectite-rich samples. Geochim. Cosmochim. Acta, 64: 3829-3837. [2] IAEA (2001) New approaches for stable isotope ratio measurements. Proceedings of an Advisory Group Meeting, held in Vienna, 20-23 Sept., 1999. IAEA, Vienna. TECDOC-1247:151 p. [34, 40] Iacumin P., Bocherens H., Mariotti A. & Longinelli A. (1996) Oxygen isotope analyses of co-existing carbonate and phosphate in biogenic apatite: a way to monitor diagenetic alteration of bone phosphate? Earth Planet. Sci. Lett., 142: 1-6. [46]

1108

References

Igumnov S. A. (1976) Sulfur isotope exchange between sulfide and sulfate in hydrothermal solutions. Geokhimiya, 4: 497-503. [46] Igumnov S. A., Grinenko V. A. & Porter N. B. (1977) Temperature dependence of 634S for H2S and dissolved sulfate over the range 260-400~ Geokhimiya, 7: 1085-1087. [46] Ihinger P. D., Hervig R. L. & McMillan P. F. (1994) Analytical methods for volatiles in glasses. In: Volatiles in magma. M. R. Carrol & J. R. Holloway (eds.). MSA Rev. Mineral., 30: 67-121. [1] Ikezoe Y., Matsuoka S., Takebe M. & Viggiano A. (1987) Gas phase ion-molecule reaction rateconstants through 1986. Maruzen, Japan. [4] Ingraham N. L. & Shadel C. (1992) A comparison of the toluene distillation and vacuum/heat methods for extracting soil water for stable isotopic analysis. J. Hydrol., 140: 371-387. [1, 2, 21] Inoue H. & Sugimura Y. (1985) Carbon isotopic fractionation during the CO2 exchange process between air and sea water under equilibrium and kinetic conditions. Geochim. Cosmochim. Acta, 49: 2453-2460. [46] IPCC (Intergovernmental Panel on Climate Change) (1990) Scientific assessment of climate change. Rept. Working Group h J.T. Houghton, G.J. Jenkins & J.J. Ephraums. Cambridge University Press, Cambridge, UK: 365 p. [14] IPCC (Intergovernmental Panel on Climate Change) (1995) Climate change 1995: The scientific basis contribution of Working Group I to the Third Assessment Report of the IPCC. J. T. Houghton, L. G. Meiro Filho, B. A. Callender, N. B. W. Harris, A. Kattenberg & K. Maskell (eds.). Cambridge University Press, Cambridge, UK: 572 p. [14] IPCC (Intergovernmental Panel on Climate Change) (2001) Climate change 2001: The scientific basis contribution of Working Group I to the Third Assessment Report of the IPCC. J. T. Houghton, Y. Ding, D. J. Griggs, M. Noguer, P. J. van der Linden & D. Xiaosu (eds.). Cambridge University Press, Cambridge, UK: 944 p. [14], 241-287. [19] Ireland T. R. (1986) Isotope systematics of refractory inclusions from carbonaceous chondrites. PhD Thesis, The Australian National University (unpublished). [30] Ireland T. R. (1995) Ion microprobe mass spectrometry: Techniques and applications in cosmochemistry, geochemistry, and geochronology. In: Advances in analytical geochemistry. M. Hyman & M. Rowe (eds.). JAI Press, Vol. 2: 1-118. [30] Irion F. W., Gunson M. R., Rinsland C. P., Yung Y. L., Abrams M. C., Chang A. Y. & Goldman A. (1996) Heavy ozone enrichments from ATMOS infrared solar spectra. Geophys. Res. Lett., 23: 2377-2380. [18] IRMM (1993) IRMM Certificate: Isotopic reference material IRMM-015. Institute for Reference Materials and Measurements, Geel, Belgium: 2 p. [40] IRMM (1997) IRMM Certificate: Isotopic reference material IRMM-016. Institute for Reference Materials and Measurements, Geel Belgium: 2 p. [40] IRMM (2001) Certificate: Isotopic reference material IRMM-011. Institute for Reference Materials and Measurements, Geel, Belgium: 2 p. [40] Irving C. S., Klein P. D., Navratil P. R. & Boutton T. W. (1986) Measurement of 13CO2/12CO2abundance by nondispersive heterodyne ratiometry as an alternative to gas isotope mass spectrometry. Anat. Chem., 58: 2172-2178. [33, 34] Ishikawa T. & Nakamura E. (1990) Suppression of boron volatilization from a hydrofluoric acid solution using a boron-mannitol complex. Anal. Chem., 62: 2612-2616. [7] Ishikawa T. & Nakamura E. (1993) Boron isotope systematics of marine sediments. Earth Planet. Sci. Lett., 117: 567-580. [7] Ishikawa T. & Nakamura E. (1994) Origin of the slab component in arc lavas from across-arc variation of B and Pb isotopes. Nature, 370: 205-208. [7] ISO Guide 35 (1989) Certification of reference materials - General and statistical principles. International Organization for Standardization, Geneva: 32 p. [43] ISO (1992) Quantities and units, Part 0: General principles, subclause 2.3.3, ISO 31-0. International Organization for Standardization, Geneva: 21 p. [43] ISO Guide 33 (1998) Uses of Certified Reference Materials. Revision of ISO Guide 33 (1989). International Organisation for Standardisation, Geneva. [40] ISO (1999) General requirements for the competence of testing and calibration laboratories. International Organization for Standardization, Geneva. IEC 17025. [42]

References

1109

ISO Guide 31 (2000a) Reference materials - Contents of certificates and labels. International Organization for Standardization, Geneva: 7 p. [43] ISO Guide 33 (2000b) Uses of certified reference materials. International Organization for Standardization, Geneva: 23 p. [43] ISO Guide 34 (2000c) General requirements for the competence of reference material producers. International Organization of Standardization, Geneva: 22 p. [40, 43] ISO/BIPM (1995) Guide to the Expression of Uncertainty in Measurement. International Organisation for Standardisation, Geneva, Switzerland: 110 p. [40, 43] ISO/REMCO (1998) Reference Materials - Contents of Certificates and Labels. ISO/REMCO No. 488, Apr. 1998, Revision of ISO Guide 31. International Organisation for Standardisation, Geneva. [40] Itoh S. & Hasegawa R. (1998) Measurements of isotope ratios of boron and silicon by glow-discharge mass-spectrometry using a small quantity of sample. J. Japan Inst. Metals, 62: 289-292. [35] IUPAC-CAWIA (1997) Reporting of relative sulfur isotope-ratio data. Pure Appl. Chem., 69: 293-295. [40] Iyengar V. & Wolf W. (1998) Global activities in hte world of reference materials including the needs of developing countries. Fresenius J. Anal. Chem., 360: 282-286. [43] Jackson C. H. N. (1933) On the true density of tsetse flies. J. Animal Ecol., 2: 204-209. [37] Jackson D. H., Mattey D. P. & Harris N. B. W. (1988) Carbon isotope compositions of fluid inclusions in charnockites from southern India. Nature, 333: 167-170. [3] Jackson M. C. & Young W. A. P. (1973) A capacitive integration system for the precise measurement of isotopic ratios in a mass spectrometer. Rev. Sci. Instrum., 44: 32-34. [39] Jackson M. L., Savin M. & Clayton R. N. (1976) Hexafluorosilicic acid reagent modification for quartz isolation. Soil Sci. Soc. Am. J., 40: 958-960. [2] Jackson M. L. (1979) Soil chemical analysis- advanced course. Madison, Wisconsin. Published by the author. 2nd edition. [2] Jackson S. E., Griffin W. L. & Pearson N. J. (1999b) Determination of high-precision isotopic ratios by laser ablation MC-ICP-MS. Ninth Annual V.M. Goldschmidt Conference, Abstr. #7361. [31] Jackson S. E. & Ryde N. (2001) The application of Nd:YAG lasers in LA-ICP-MS. In: Laser-AblationICPMS in the Earth Sciences. P. Sylvester (ed.). Mineral. Assoc. of Canada. Short Course Vol. Series, Vol. 29, Chapter 3: 29-45. [20] Jackson S. E. & Giinther, D. (2003) The nature and sources of laser induced isotopic fractionation in laser ablation-multicollector-inductively coupled plasma-masspectrometry. J. Anal. At. Spectrom., 18: 205-212. [20] Jackson S. M., Morgan G. H., Morse A. D., Butterworth A. L. & Pillinger C. T. (1999a) The use of static mass spectrometry to determine the combined stable isotopic composition of small samples of atmospheric methane. Rapid Commun. Mass Spectrom., 13: 1329-1333. [36] Jackson T. A. (2001) Variations in isotope composition of mercury in a freshwater sediment sequence and food web. Can. J. Fish. Aquat. Sci., 58: 185-196. [31] Jacob E. & Christie K. O. (1977) A simple method for the purification of fluorine. J. Fluorine Chem., 10: 169-172. [20] James B. R. (1994) Hexavalent chromium solubility and reduction in alakaline soils enriched with chromite ore processing residue. J. Environ. Qual., 23: 227-233. [29] James R. H., Rudnicki M. D. & Palmer M. R. (1999) The alkali element and boron geochemistry of the Escanaba trough sediment-hosted hydrothermal system. Earth Planet. Sci. Lett., 171: 157-169. [6] James R. H. & Palmer M. R. (2000) The lithium isotope composition of international rock standards. Chem. Geol., 166: 319-326. [6] Jameson R. F. (1959) The composition of the "strong" phosphoric acids. J. Chem. Soc., 1: 752-759. [12] Jamin E., Naulet N. & Martin G. J. (1997) Multi-element and multi-site isotopic analysis of nicotine from tobacco leaves. Plant Celt Environ., 20: 589-599. [5] Jancso G. & Van Hook W. A. (1974) Condensed phase isotope effects. Chem. Rev., 74: 689-710. [4] Janghorbani M., Davis T. A. & Ting B. T. G. (1988) Measurement of stable isotopes of bromine in biological fluids with inductively coupled plasma mass spectrometry. Analyst, 113: 405-411. [28] Janssen C., Guenther J., Krankowski D. & Mauersberger K. (1999) Relative formation rates of 5003 and 5203 in 160-180 mixtures. J. Chem. Phys., 111: 7179-7182. [18]

1110

References

Jaques A. L., Hall A. E., Sheraton J. W., Smith C. B., Sun S.-S., Drew R. M., Foudoulis C. & Ellingsen K. (1989) Comparison of crystalline inclusions and C-isotopic composition of Argyle and Ellendale diamonds. In: Kimberlites and related rocks. J. Ross (ed.). Blackwell (Geol. Soc. Austr. Spec. Publ., 14): 966-989. [13] Jarman W. M., Hilkert A., Bacon C. E., Collister J. W., Ballschmiter K. & Risebrough R. W. (1998) Compound-specific carbon isotopic analysis of arochlors, clophens, kaneclors, and phenoclors. Environm. Sci. Technol., 32: 833-836. [9] Jarvis K. E., Gray A. L. & Houk R. S. (1992) Handbook of inductively coupled plasma mass spectrometry. Blackie and Son - Chapman and Hall (USA): 380 p. [31] Jasper J. P. (2001) Quantitative estimates of precision for molecular isotopic measurements. Rapid Commun. Mass Spectrom., 15: 1554-1557. [34] Javoy M. & Fourcade S. (1973) 160/180 ratios in lunar fines and rocks. Earth Planet. Sci. Lett., 211: 377382. [20] Javoy M., Marette J. & Pineau F. (1973) 160/180 ratios in Luna 16 fines. Geochim. Cosmochim. Acta, 37: 2017-2019. [20] Javoy M., Pineau F. & Iiyama I. (1978) Experimental determination of the isotopic fractionation between gaseous CO2 and carbon dissolved in tholeiitic magma: apreliminary study. Contrib. Mineral. Petrol., 67: 35-39. [46] Javoy M., Pineau F. & Delorme H. (1986) Carbon and nitrogen isotopes in the mantle. Chem. Geol., 57: 41-62. [16] Jenden P. D., Kaplan I. R., Poreda R. J. & Craig H. (1988) Origin of nitrogen-rich natural gases in the California Great Valley. Evidence from helium, carbon and nitrogen isotope ratios. Geochim. Cosmochim. Acta, 52: 851-861. [16] Jendrzejewski N., Eggenkamp H. M. G. & Coleman M. L. (1997) Sequential determination of chlorine and carbon isotopic composition in single microlitre samples of chlorinates solvent. Anal. Chem., 69: 4259-4266. [28] Jenkin G. R. T., Fallick A. E. & Leake B. E. (1992) A stable isotope study of retrograde alteration in SW Connemara, Ireland. Contrib. Mineral. Petrol., 110: 269-288. [3] Jenkin G. R. T., Craw D. & Fallick A. E. (1994a) Stable isotopic and fluid inclusion evidence for meteoric fluid penetration into an active mountain belt; Alpine Schist, New Zealand. J. Metam. Geol., 12: 429444. [3] Jenkin G. R. T., Farrow C. M., Fallick A. E. & I-Iiggins D. (1994b) Oxygen isotope exchange and closure temperatures in cooling rocks. J. Metamorphic Geol., 12: 221-235. [46] Jensen E. S. (1991) Evaluation of automated analysis of 15N and total N in plant material and soil. Plant Soil, 133: 83-92. [15] Jensen M. L. & Nakai N. (1962) Sulfur isotope meteorite standards results and recommendations. In: Biogeochemistry of sulfur isotopes. M. L. Jensen (ed.). Yale University, New Haven, Connecticut. Nat. Sci. Found. Symp. Vol.: 30-35. [26, 40] Jeong D.-Y., Rhee Y.-J., Lee J. & Rhee B. K. (1998) CW laser-initiated resonance ionization mass spectrometry of Gd atoms. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AIP Press Woodbury, NY. AIP Conference Proceedings 454: 289-292. [33] Jia Y. J. & Kerrich R. (1999) Nitrogen isotope systematics of mesothermal lode gold deposits: metamorphic, granitic, meteoric water, or mantle origin? Geology, 27: 1051-1054. [16] Jiang S., Ding T. & Wan D. (1992) A study on silicon and oxygen isotope geochemistry of some leadzinc ore depositis in China. Chinese Sci. Bull., 37(12): 1022-1027. [25] Jiang S., Ding T., Wan D. & Li Y. (1993) Silicon isotopic composiitons of Archean banded Si-Fe formation (BIF) in the Gongchangling ore deposit, Liaoning, China. Chinese Sci. Bull., 36: 482-489.

[25]

Jiang S.-Y., Palmer M. R., Ding T.-P. & Wan, D.-F. (1994) Silicon isotope geochemistry of the Sullivan Pb-Zn deposit, Canada: A preliminary study. Econ. Geol., 89: 1623-1629. [20] Johnson B. J., Fogel M. L. & Miller G. H. (1998) Stable isotopes in modern ostrich eggshell: a calibration for paleoenvironmental applications in semi-arid regions of southern Africa. Geochim. Cosmochim. Acta, 62: 2451-2461. [9, 46]

References

1111

Johnson C. M. & Nishita H. (1952) Microestimation of sulfur in plant materials, soils, and irrigation waters. Anal. Chem., 24: 736-742. [26] Johnson C. M. & Ulrich A. (1959) Analytical methods for use in plant analysis. Calif. Agric. Exp. Stn. Bull., 766: 25-78. [26] Johnson C. M. & Beard B. L. (1999) Correction of instrumentally produced mass fractionation during isotopic analysis of Fe by thermal ionization mass spectrometry. Int. J. Mass Spectrom., 193(1): 87-99. [29, 31] Johnson D. G., Jucks K. W., Traub W. A. & Chance K. V. (2000a) Isotopic composition of stratospheric ozone. J. Geophys. Res., 105(D7): 9025-9031. [18] Johnson R. C. & Tieszen L. L. (1994) Variation for water-use efficiency in alfalfa germplasm. Crop Sci., 34: 452-458. [9] Johnson T. C., Bonrud P. A., Dosch T. L., Kilness A. W., Senger K. A., Busch D. C. & Meyer M. R. (1987) Fatal outcome of methemoglobinemia in an infant. J. Am. Med. Assoc., 257: 2796-2797. [15] Johnson T. M., Herbel M. J., Bullen T. D. & Zawislanski P. T. (1999) Selenium isotope ratios as indicators of selenium sources and oxyanion reduction. Geochim. Cosmochim. Acta, 63(18): 2775-2783. [20, 29, 30, 46] Johnson T. M., Bullen T. D. & Zawislanski P. T. (2000b) Selenium stable isotope ratios as indicators of sources and cycling of selenium: Results from the Northern Reach of San Francisco. Environ. Sci. Technol., 34: 2075-2079. [30] Johnson T. M. & Bullen T. D. (2003) Selenium isotope fractionation during reduction by Fe(II)-Fe(III) hydroxide-sulfate (green rust). Geochim. Cosmochim. Acta, 67: 413-419. [46] Johnson T. M. (in press) A review of mass-dependent fractionation of selenium isotopes and implications for other "heavy stable isotopes". Chem. Geol. [29] Johnston A. M., Scrimgeour C. M., Henry M. O. & Handley L. L. (1999) Isolation of NO3--N as 1-phenylazo-2-naphthol (Sudan-I) for measurement of 615N. Rapid Commun. Mass Spectrom., 13: 15311534. [15] Johnston J. C. & Thiemens M. H. (1997) The isotopic composition of troposheric ozone in three environments. J. Geophys. Res., 102(D21): 25395-25404. [18] Johnston J. C., Ri~ckmann T. & Brenninkmeijer C. A. M. (2000) CO2 + O(1D) isotopic exchange: Laboratory and modeling studies. J. Geophys. Res., 105: 15213-15229. [18] Johnston W. H. & Arnold J. R. (1953) The existence of multiple charged molecular ions of HBr, HC1, and DC1. J. Chem. Phys., 21: 1499-1502. [28] Jones A. M., Iacumin P. & Young E. D. (1999) High-resolution 6180 analysis tooth enamel phosphate of by isotope ratio monitoring gas chromatography mass spectrometry and ultraviolet laser fluorination. Chem. Geol., 153: 241-248. [20, 22] Jones B. J. J. & Steyn w. J. A. (1973) Sampling, handling, and analyzing plant and tissue samples. In: Soil testing and plant analysis. L. W. Walsh & J. D. Beaton (eds.). Soil Science Society of America, Madison, Wisconsin, USA: 249-270. [26] Jones D. M., Carter J. F., Eglinton G., Jumeau E. J. & Fenwick C. S. (1991) Determination of 613C values of sedimentary straight chain and cyclic alcohols by gas chromatography/isotope mass spectrometry. Biol. Mass Spectrom., 20: 641-646. [9] Jones R. I., Grey J., Sleep D. & Quarmby C. (1998) An essessment, using stable isotopes, of the importance of allochthonous organic carbon sources to the pelagic food web in Loch Ness. Proc. R. Soc. London, B 265: 105-111. [9] Jones R. L. & Pyle J. A. (1984) Observations of CH4 and N20 by the Nimbus-7 SAMS: a comparison with in situ data and two-dimensional numerical model calculations. J. Geophys. Res., 89(D4): 5263-5279. [14] Jovanovic S. & Reed G.W. Jr. (1976) 196Hg and 202Hg isotopic ratios in chondrites: revisited. Earth Planet. Sci. Lett., 31: 95-100. [31] Juchelka D., Beck T., Hener U., Dettmar F. & Mosandl A. (1998) Multidimensional gas chromatography coupled on-line with isotope ratio mass spectrometry (MDGC-IRMS): Progress in the analytical authentication of genuine flavor components. HRC-J. High Resolut. Chromatogr., 21: 145151. [8]

1112

References

Juillet-Leclerc A. & Labeyrie L. D. (1987) Temperature dependence of the oxygen isotopic fractionation between diatom silica and water. Earth Planet. Sci. Lett., 84: 69-74. [2] Junge F., Seltmann R. & Stiehl G. (1989) Nitrogen isotope characteristics of breccias, granitoids, and greisens from eastern Erzgebirge tin ore deposits, Sadisdorf: Altenberg, GDR. Proc. 5th Working Meeting, Isotopes in Nature: 321-332. [16] Junk G. & Svec H. J. (1958) The absolute abundance of the nitrogen isotopes in the atmosphere and compressed gas from various sources. Geochim. Cosmochim. Acta, 14: 234-243. 15,[40, 43] Jusserand C. (1980) Extraction de l'eau interstitielle des s6diments et des sols - Comparison des valeurs de l'oxyg6ne 18 par diff6rentes m6thodes, premiers r6sultats. Catena, 7: 87-96. [2] Kakihana I-I., Kotaka M., Satoh S., Nomura M. & Okamoto M. (1977) Fundamental studies on the ionexchange separation of boron isotopes. Bull. Chem. Soc. Jap., 50: 158-163. [7] Kakiuchi M. (1994) Temperature dependence of fractionation of hydrogen isotopes in aqueous sodium chloride solutions. J. Sol. Chem., 23: 1073-1087. [46] Kampschulte A., Bruckschen P. & Strauss H. (2001) The sulphur isotopic composition of trace sulphates in Carboniferous brachiopods: implications for coeval seawater, correlation with other geochemical cycles and isotope stratigraphy. Chem. Geol., 175: 165-189. [26] Kaplan I. R., Emery K. O. & Rittenberg S. C. (1963) The distribution and isotopic abundance of sulphur in recent marine sediments off Southern California. Geochim. Cosmochim. Acta, 27: 297-331. [29] Kaplan I. R. & Rittenberg S. C. (1964) Microbiological fractionation of sulfur isotopes. J. Gen. Microbiol., 34: 195-212. [46] Kaplan I. R. (1983) Stable isotopes of sulfur, nitrogen and deuterium in recent marine environments. In Stable isotopes in sedimentary geology. M. A. Arthur, T. F. Anderson, I. R. Kaplan, J. Veizer & L. S. Land (eds.). Society of Sedimentary Geology. Vol. 10: 2.1-2.108. [29] Karasek F. W. & Clement R. E. (1988) Basic gas chromatography - mass spectrometry. Elsevier, Amsterdam. [4] Karhu J. & Epstein S. (1986) The implication of the oxygen isotope record in coexisting cherts and phosphates. Geochim. Cosmochim. Acta, 50: 1745-1756. [20, 22] Karhu J. A. (1997) Catalytic reduction of water to hydrogen for isotopic analysis using zinc containing traces of sodium. Anat. Chem., 69: 4728-4730. [1] Karlsson H. R. & Clayton R. N. (1990) Oxygen isotope fractionation between analcime and water: An experimental study. Geochim. Cosmochim. Acta, 54: 1359-1368. [36, 46] Karlsson H. R. (1995) Application of oxygen and hydrogen isotopes top zeolites. In: Natural zeolites '93: Occurence, properties, use. D. W. Ming & F. A. Mumpton (eds.). Int. Comm. Nat. Zeol., NY: 125-140. [36] Kasemann S., Meixner A., Rocholl A., Vennemann T. W., Rosner M., Schmitt A. K. & Wiedenbeck M. (2001) Boron and oxygen isotopic composition of certified reference materials NIST SRM-610/612 and reference materials JB-2 and JR-2. Geostand. Newslett., 25: 405-416. [40] Kator H., Morris L. J., Wetzel R. L. & Koepfler E. T. (1992) A rapid chromatographic method for recovery of 15NO2- and NO3- produced by nitrification in aqueous samples. Limnol. Oceanogr., 39(4): 900907. [15] Katz S. A. & Salem H. (1994) The biological and environmental chemistry of chromium. VCH Publishers. [29] Katzenberg M. A. & Krouse H. R. (1989) Application of stable isotopes in tissues to problems in identification. Cdn. Soc. Forensic Sci. J., 22: 7-19. [26] Kaufmann R. S., Long A., Bentley H. & Davis S. (1983) Application of chloride stable isotope analyses to hydrogeology. Proc. 1983 meet. Amer. Water Res., Ariz. & Ariz.-Nev. Acad. Sci., Hydr. sect. Hydr. Water Res. Ariz. Southwest.: 85-90. [28] Kaufmann R. S. (1984) Chlorine in groundwater: Stable isotope distribution. PhD Thesis, University of Arizona, Tucson, Arizona, USA: 137 p. [28] Kaufmann R. S. & Long A. (1984) Stable chlorine ratio as a ground-water tracer. EOS Trans. AGU, 65: 886. [28] Kaufmann R. S., Long A., Davis S. & Bentley H. (1984a) Natural variations of chlorine stable isotopes. Geol. Soc. Am. Abstr. w. Prog., 16: 556. [28]

References

1113

Kaufmann R. S., Long A., Bentley H. & Davis S. (1984b) Natural chlorine isotope variations. Nature, 309: 338-340. [28] Kaufmann R. S. & Arn6rsson S. (1986) 37C1/35C1 ratios in Icelandic geothermal waters. Proc. 5th International Symposium on Water-Rock Interaction: 325-327. [28] Kaufmann R. S., Frape S., Fritz P. & Bentley H. (1987) Chlorine stable isotope composition of Canadian Shield brines. In: Saline water and gases in crystalline rocks. P. Fritz & S. K. Frape (eds.). Geol. Ass. Can. Spec. Paper, Vol. 33: 89-93. [28] Kaufmann R. S., Long A. & Campbell D. J. (1988) Chlorine isotope distribution in formation waters, Texas and Louisiana. AAPG Bull., 72: 839-844. [28, 40] Kaufmann R. S. (1989) Equilibrium exchange models for chlorine stable isotope fractionation in high temperature environments. Proc. 6th International Symposium on Water-Rock Interaction: 365-368. [281 Kaufmann R. S., Frape S. K., McNutt R. & Eastoe C. J. (1992) Chlorine stable isotope distribution of Michigan Basin Shield formation waters. Proc. 7th International Symposium on Water-Rock Interaction: 943-946. [28] Kaufmann R. S., Frape S. K., McNutt R. & Eastoe C. J. (1993) Chlorine stable isotope distribution of Michigan basin formation waters. Appl. Geochem., 8: 403-407. [28] Kaye J. A. & Strobel D. F. (1983) Enhancement of heavy ozone in the earth's atmosphere? J. Geophys. Res., 88: 8447-8452. [18] Kaye J. A. (1986) Theoretical analysis of isotope effects on ozone formation in oxygen photochemistry. J. Geophys. Res., 91: 7865-7874. [18] Kazahaya K. & Matsuo S. (1985) A new ball-milling method for extraction and fluid inclusions from minerals. Geochem. J., 19: 45-54. [1, 3] K/irger J. & Ruthven D. M. (1992) Diffusion in zeolites and other microporous solids. Wiley, NY: 605 p. [36] Keeling C. D., Bacastow R. B., Carter A. F., Piper S. C., Whorf T. P., Heimann M., Mook W. G. & Roeloffzen H. (1989) A three-dimensional model of atmospheric CO2 transport based on observed winds: 1. Analysis of observational data. In: Aspects of climate variability in the Pacific and the Western Americas. D. H. Peterson (ed.). Am. Geophys. Union, Geophys. Monogr. 55: 165-236. [14] Keeling C. D., Whorf T. P., Wahlen M. & van der Plicht J. (1995) Interannual extremes in the rate of rise of atmospheric carbon dioxide since 1980. Nature, 375: 666-670. [14] Keene W. C., Pszenny A. A. P., Galloway J. N. & Hawley M. E. (1986) Sea-salt corrections and interpretation of constituent ratios in marine precipitation. J. Geophys. Res., D91: 6647-6658. [26] Keeney D. R. & Nelson D. W. (1982) Nitrogen- Inorganic forms. In: Methods of soil analysis, Part 2. Chemical and microbiological proper. A. L. Page, R. H. Miller & D. R. Keeney (eds.). American Society of Agronomy Inc. - Soil Science Society of America Inc., Madison, Wisconsin. Agronomy Series, 2nd edition, No. 9: 643-698. [15] Keller J. & Krafft M. (1990) Effusive natrocarbonatite activity of Oldoinyo Lengai, June 1988. Bull. Volcanol., 52: 629-645. [28] Kelley D. B. & Friih-Green G. L. (2001) Volatile lines of descent in submarine plutonic environments: Insights from stable isotope and fluid inclusion analyses. Geochim. Cosmochim. Acta, 65(19): 33253346. [3] Kelley K. R., Ditsch D. C. & Alley M. M. (1991) Diffusion and automated nitrogen-15 analysis of lowmass ammonium samples. Soil Sci. Soc. Am. J., 55: 1016-1020. [15] Kelley S. P. & Fallick A. E. (1990) High precision spatially resolved analysis of 634S in sulphides using a laser extraction technique. Geochim. Cosmochim. Acta, 54: 883-888. [20, 26, 45] Kelly J. F. (1999) Stable isotopes of carbon and nitrogen in the study of avian and mammalian ecology. Can. J. Zoology, 78: 1-27. [9] Kelly S. D., Parker I. G., Sharman M. & Dennis M. J. (1998) On-line quantitative determination of 2H/ 1H isotope ratios in organic and water samples using an elemental analyser coupled to an isotope ratio mass spectrometer. J. Mass Spectrom., 33: 735-738. [1] Kemp A. L. W. & Thode H. G. (1968) The mechanism of the bacterial reduction of sulphate and of sulphite from isotope fractionation studies. Geochim. Cosmochim. Acta, 32: 71-91. [46]

1114

References

Kendall C. & Coplen T. B. (1985) Multisample conversion of water to hydrogen by zinc for stable isotope determination. Anal. Chem., 57: 1437-1440. [1, 21] Kendall C. & Grim E. (1990) Combustion tube method for measurement of nitrogen isotope ratios using calcium oxide for total removal of carbon dioxide and water. Anal. Chem., 62: 526-529. [15, 16, 40] Kendall C., Campbell D. H., Burns D. A., Shanley J. B., Silva S. R. & Chang C. C. (1995a) Tracing sources of nitrate in snowmelt runoff using the oxygen and nitrogen isotopic compositions of nitrate. Proceedings of a Boulder Symposium, IAHS Publ. No. 228. [15] Kendall C., Sklash M. G. & Bullen T. D. (1995b) Isotope tracers of water and solute sources in catchments. In: Solute modelling in catchment systems. S. T. Trudgill (ed.). John Wiley & Sons Ltd, NY: 261-303. [1 ] Kendall C. (1998) Tracing nitrogen sources and cycling in catchments. In: Isotope tracers in catchment hydrology. C. Kendall & J. J. McDonnell (eds.). Elsevier Science B.V, Amsterdam: 534-569. [15] Kerrich R., Beckinsale R. D. & Shackleton N. J. (1978) The physical and hydrothermal regime of tectonic vein systems; evidence from stable isotope and fluid inclusion studies. N. Jhb. Miner. Abhand., 131: 225-239. [3] Kerridge J. F. (1985) Carbon, hydrogen, and nitrogen in carbonaceous chondrites: Abundances and isotopic compositions in bulk samples. Geochim. Cosmochim. Acta, 49: 1707-1714. [16] Kerridge J. F. (1993) Long-term compositional variation in solar corpuscular radiation: Evidence from nitrogen isotopes in the lunar regolith. Rev. Geophys., 31: 423-437. [17] Kerstel E. R. T., van Trigt R., Dam N., Reuss J. & Meijer H. A. J. (1999) Simultaneous determination of the 2H/1H, 170/160, and 180/160 isotope abundance ratios in water by means of laser spectrometry. Anal. Chem., 71: 5297-5303. [34] Kerstel E. R. T., van Trigt R., Dam N., Reuss J. & Meijer H. A. J. (2001) Laser spectrometry applied to the simultaneous determination of the 62H, 6170, and 6180 isotope abundances in water. In: New approaches for stable isotope ratio measurements. Proceedings of an Advisory Group meeting held in Vienna, 20-23 September 1999. IAEA, Vienna. IAEA-TECDOC-1247: 7-13. [34, 40] Kerstel E. R. T. (2001) O-17 analysis of Ontario Hydro heavy water. Center for Isotope Research, University of Groningen, Groningen, the Netherlands. Customer Report No. 2001-32. [34] Kerstel E. R. T., Gagliardi G., Gianfrani L., Meijer H. A. J., van Trigt R. & Ramaker R. (2002) Determination of the 2H/1H, 170/160, and 180/160 isotope ratios in water by means of tunable diode laser spectroscopy at 1.39 ~m. Spectrochim. Acta, A 58: 2389-2396. [34] Kester C. L., Rye R. O., Johnson C. A., Schwartz C. & Holmes C. (2001) On-line sulfur isotope analysis of organic material by direct combustion: preliminary results and potential applications. Isot. Environ. Health Stud., 37: 53-65. Ketterer M. E., Peters M. J. & Tisdale P. J. (1991) Verification of a correction procedure for measurement of lead isotope ratios by inductively coupled plasma mass spectrometry. J. Anal. At. Spectrom., 6: 439443. [35] Khalfallah Y., Normand S., Tissot S., Pachiaudi C., Beylot M. & Riou J. P. (1993) Determination of 613C values of plasma lactate by gas chromatography/isotope ratio mass-spectrometry. Biol. Mass Spectrom., 22: 707-711. [8] Khan S. A., Mulvaney R. L. & Mulvaney C. S. (1997) Accelerated diffusion methods of inorganicnitrogen analysis of soil extracts and water. Soil Sci. Soc. Am. J., 61: 936-942. [15] Khan S. A., Mulvaney R. L. & Brooks P. D. (1998) Diffusion methods for automated nitrogen-15 analysis using acidified disks. Soil Sci. Soc. Am. J., 62: 406-412. [15] Khan S. A., Mulvaney R. L. & Hoeft R. G. (2000) Direct-diffusion methods for inorganic-nitrogen analysis of soil. Soil Sci. Soc. Am. J., 64: 1083-1089. [15] Kiba T., Tagaki T., Yoshimura Y. & Kishi I. (1955) Tin (II)-strong phosphoric acid. A new reagent for the determination of sulfate by reduction to hydrogen sulfide. Bull. Chem. Soc. Japan, 28: 641-644. [26, 27] Kiba T., Akaza I. & Sugishita N. (1957) Rapid determination of inorganic sulfur in various forms, particularly in sulfide ores, by the tin(II)-strong phosphoric acid reduction method. Bull. Chem. Soc. Japan, 30: 972-975. [26]

References

1115

Kieffer S. W. (1982) Thermodynamics and lattice vibrations of minerals: 5. Applications to phase equilibria, isotopic fractionation, and high-pressure thermodynamic properties. Rev. Geophys. Space Phys., 20: 827-849. [46] Kilmer V. J. & Nearpass D. C. (1960) The determination of available sulfur in soils. Soil Sci. Soc. Am. Proc., 24: 337-340. [26] Kim C.-K., Seki R., Morita S., Yamasaki S., Tsumura A., Takaku Y., Igarashi Y. & Yamamoto M. (1991) Application of a high resolution inductively coupled plasma mass spectrometer to the measurement of long-lived radionuclides. J. Anal. At. Spectrom., 6: 205-209. [35] Kim K.-R. & Craig H. (1990) Two-isotope characterization of N20 in the Pacific Ocean and constraints on its origin in deep water. Nature, 347: 58-61. [19] Kim K.-R. & Craig H. (1993) Nitrogen-15 and oxygen-18 characteristics of nitrous oxide: A global perspective. Science, 262: 1855-1857. [19] Kim S. A. & O'Neil J. R. (1997) Equilibrium and nonequilibrium oxygen isotope effects in synthetic carbonates. Geochim. Cosmochim. Acta, 61: 3461-3475. [46] Kindness A. & Marr I. L. (1997) Improved infrared spectroscopic method for the measurement of 13C:12C ratios. Appl. Spectrosc., 51(1): 17-21. [34] King E. M., Valley J. W., Davis D. W. & Kowallis B. J. (2001) Empirical determination of oxygen isotope fractionation factors for titanite with respect to zircon and quartz. Geochim. Cosmochim. Acta, 65: 3165-3175. [46] King F. L. & Harrison W. W. (1990) Glow discharge mass spectrometry. An introduction to the technique and its utility. Mass Spec. Rev., 9: 285-317. [35] King F. L. & Harrison W. W. (1993) Glow discharge mass spectrometry. In" Glow discharge spectroscopies. R. K. Marcus (ed.). Plenum Press, NY: 175-214. [35] King S. L., Quay P. D. & Lansdown J. M. (1989) The C-13/C-12 kinetic isotope effect for soil oxidation of methane at ambient atmospheric concentrations. J. Geophys. Res., 94(D15): 18273-18277. [14] Kipphardt H., Valkiers S., Taylor P. D. P. & De Bi6vre P. (2000) "Calibration" in isotopic measurements. Int. J. Mass Spectrom., 198: 71-76. [42] Kipphardt H., Dudzus T., Meier K. A., Recknagel S., Hedrich M. & Matschat R. (2002) Measurement of oxygen and nitrogen in high purity metals used as national standards for element analysis in Germany by classical carrier gas hot extraction (HE) and HE after activation with photons. Mater. Trans. (JIM), 43: 98-100. [42] Kirschner D. L. & Sharp Z. D. (1997) Oxygen isotope analyses of fine-grained minerals and rocks using the laser-extraction technique. Chem. Geol., 137: 109-115. [20] Kirshenbaum I., Smith J. S., Crowell T., Graft J. & McKee R. (1947) Separation of the nitrogen isotopes by the exchange reaction between ammonia and solutions of ammonium nitrate. J. Chem. Phys., 15: 440-446. [15] Kirshenbaum I. (1951) Isotopic analysis of heavy water by the mass spectrometer. In: Physical properties and analysis of heavy water. McGraw-Hill Book Comp. Inc. [1] Kiser R. W. (1965) Introduction to mass spectrometry and its applications. Prentice Hall, NY. [38] Kishima N. & Sakai H. (1980) Oxygen-18 and Deuterium determination on a single water sample of a few milligrams. Anal. Chem., 52: 356-358. [1, 3, 20] Kita I. (1981) A new type of ball mill made of Pyrex glass. Geochem. J., 15: 289-291. [3] Klaue B. & Blum J. D. (2000) Mercury isotopic analyses by single- and multi-collector magnetic sector inductively coupled plasma mass spectrometry. J. Conf. Abstr., 5:591 (Abstract). [31] Klein R. T., Lohmann K. C. & Thayer C. W. (1996) Sr / Ca and 13C / 12C ratios in skeletal calcite of Mytilus trossulus: covariation with metabolic rate, salinity, and carbon isotopic composition of seawater. Geochim. Cosmochim. Acta, 60: 4207-4221. [46] Klemm A. & Lund6n A. (1955) Isotopenanreicherung beim Chlor durch electrolytische Uberffihrung in geschmolzenem Bleichlorid. Z. Naturforsch., 10a: 282-284. [28] Klossa B., Pierre A. & Minster J. F. (1981) Mesures de la composition isotopique du lithium dans les inclusions. Earth Planet. Sci. Lett., 52: 25-30. [6] Kli~tzli U. S. (1992) Negative thermal ionization mass spectrometry: a new approach to boron isotope geochemistry. Chem. geol. (Isot. Geosci. Sect.), 101: 111-122. [7]

1116

References

Knauth L. P. & Epstein S. (1975) Hydrogen and oxygen isotope ratios in silica from the Joides deep drilling project. Earth Planet. Sci. Lett., 25: 1-10. [46] Knauth L. P. & Epstein S. (1976) Hydrogen and oxygen isotope ratios in nodular and bedded cherts. Geochim. Cosmochim. Acta, 40: 1095-1108. [46] Knauth L. P. & Beeunas M. A. (1986) Isotope geochemistry of fluid inclusions in Permian halite with implications for the isotopic history of ocean water and the origin of saline formation waters. Geochim. Cosmochim. Acta, 50: 419-433. [3] Knoll G. F. (2000) Chapter 4: General properties of Radiation detectors. In: Radiation detection and measurement. Wiley & Sons, New York. 3rd edition: 802 p. [32] Knowles R. & Blackburn H. T. (1993) Nitrogen isotope techniques. Academic Press Inc., San Diego: 311 p. [9, 15] Koch P. L., Fogel M. L. & Tuross N. (1994) Tracing the diets of fossil animals using stable isotopes. In: Stable isotopes in ecology and environmental science, methods in ecology. K. Lajtha & R.H. Michener (eds.). Blackwell Scientific Publications, Oxford, UK: 63-92. [9] Koch P. L., Heisinger J., Moss C., Carlson R. W., Fogel M. L. & Behrensmeyer A. K. (1995) Isotopic tracking of change in diet and habitat use in African elephants. Science, 267: 1340-1343. [9] Koehler G. D., Chipley D. & Kyser T. K. (1991) Measurement of the hydrogen and oxygen isotopic compositions of concentrated chloride brines from fluid inclusions in halite. Chem. Geol. (Isot. Geosci. Sect.), 94: 45-54. [1] Koehler G., Wassenaar L. I. & Hendry M. J. (2000) An automated technique for measuring 6D and 6180 values of porewater by direct CO2 and H2 equilibration. Anal. Chem., 72: 5659-5664. [1, 2] Kohlmuller D. & Kochen W. (1993) Is n-pentane really an index of lipid-peroxidation in humans and animals - a methodological reevaluation. Anal. Biochem., 210: 268-276. [8] Kohn M. J., Valley J. W., Elsenheimer D. & Spicuzza M. J. (1993) O isotope zoning in garnet and staurolite. Evidence for closed-system mineral growth during metamorphism. Am. Mineral., 78: 9881001. [20] Kohn M. J. (1996) Predicting animal 6180: accounting for diet and physiological adaptation. Geochim. Cosmochim. Acta, 60: 4811-4829. [20] Kohn M. J., Schoeninger M. J. & Valley J. W. (1998) Variability in oxygen isotope compositions of herbivore teeth: reflections of seasonality or development physiology? Chem. Geol., 152: 97-112. [20, 22] Kohn M. J. & Valley J. W. (1998) Obtaining equilibrium oxygen isotope fractionations from rocks: theory and examples. Contrib. Mineral. Petrol., 132: 209-224. [46] Kohnen M. E. L. (1988) Stabiele chloorisotopen onderzoek. University of Utrecht, Internal Report: 17 p.

[28]

Koirtyohann S. R. (1994) Precise determination of isotopic ratios for some biologically significant elements by inductively coupled plasma mass spectroscopy. Spectrochim. Acta, B 49: 1305-1311. [32] Koletzko S., Haisch M., Seeboth I., Braden B., Hengels K., Koletzko B. & Hering P. (1995) Isotopeselective non-dispersive infrared spectrometry for detection of Helicobacter pylori infection with 13Curea breath test. Lancet, 345: 961-962. [34] Kolodny Y. & Epstein S. (1976) Stable isotope geochemistry of deep sea cherts. Geochim. Cosmochim. Acta, 40: 1195-1209. [46] Kolodny Y., Luz B. & Navon O. (1983) Oxygen isotope variation in phosphate of biogenic apatites. 1. Fish bone apatite - rechecking the rules of the game. Earth Planet. Sci. Lett., 64: 298-404. [20, 22, 46] Kolodny Y. & Raab M. (1988) Oxygen isotopes in phosphatic fish remains from Israel: paleothermometry of tropical Cretaceous and Tertiary shelf waters. Palaeogeogr., Palaeoclimatol., Palaeoecol., 64: 5967. [22] Kolodny Y. & Luz B. (1992) Isotope signatures in phosphate deposits: formation and diagenetic history. In: Isotopic signatures and sedimentary records. S. Chauduri & N. Clauer (eds.). Springer: 69-122. [22] Kolodny Y., Luz B., Sander M. & Clemens W. A. (1996) Dinosaur bones: Fossils or pseudomorphs? The pitfalls of physiology reconstruction from apatitic fossils. Palaeogeogr., Palaeoclimatol., Palaeoecol., 126: 161-171. [22]

References

1117

Komor S. C. & Magner J. A. (1996) Nitrate in groundwater and water sources used by riparian trees in an agricultural watershed: a chemical and isotopic investigation in southern Minnesota. Water Resour. Res., 32: 1039-1050. [15] Konstantinov B. P. & Bakulin E. A. (1965) Separation of chloride isotopes in aqueous solutions of lithium chloride, sodium chloride, and hydrochloric acid. Russ. J. Phys. Chem., 39: 315-318. [28] Kornexl B. E., Gehre M., Hi, fling R. & Werner R. A. (1999a) On-line 6180 measurements of organic and inorganic substances. Rapid Commun. Mass Spectrom., 13: 1685-1693. [1, 9, 15, 20, 22, 23] Kornexl B. E., Werner R. A. & Gehre M. (1999b) Standardization for oxygen isotope ratio measurement - Still an unsolved problem. Rapid Commun. Mass Spectrom., 13: 1248-1251. [1, 23, 43] Kosir I. J., Kocjancic M., Ogrinc N. & Kidric J. (2001) Use of SNIF-NMR and IRMS in combination with chemometric methods for the determination of chaptalisation and geographical origin of wines (the example of Slovenian wines). Anal. Chim. Acta, 429: 195-206. [5] Kosler J., Kucera M. & Sylvester P. (2001) Precise measurement of Li isotopes in planktonic foraminiferal tests by quadrupole ICPMS. Chem. Geol., 181: 169-179. [6] Kosterev A. A., Curl R. F., Tittel F. K., Gmachl C., Capasso F., Sivco D. L., Baillargeon J. N., Hutchinson A. L. & Cho A. Y. (1999) Methane concentration and isotopic composition measure-ments with a mid-infrared quantum-cascade laser. Opt. Lett., 24: 1762-1764. [34] Kotaka M. (1973) Chromatographic separation of boron and nitrogen isotopes using pure water as an eluant. PhD Thesis, Tokyo Institute of Technology, Tokyo, Japan. [7] Kotzer T. G., Kyser T. K., King R. W. & Kerrich R. (1993) An empirical oxygen- and hydrogen-isotope geothermometer for quartz-tourmaline and tourmaline-water. Geochim. Cosmochim. Acta, 57: 34213426. [46] Koumenis I. L., Vestal M. L., Yergey A. L., Abrams S., Deming S. N. & Hutchens T. W. (1995) Quantitation of metal isotope ratios by laser-desorption time-of-flight mass-spectrometry. Anal. Chem., 67: 4557-4564. [33] Koziet J., Rossmann A., Martin G. J. & Johnson P. (1995) Determination of the oxygen-18 and deuterium content of fruit and vegetable juice water. A European inter-laboratory study. Anal. Chim. Acta, 302: 29-37. [1] Koziet J. (1997) Isotope ratio mass spectrometric method for the on-line determinartion of oxygen-18 in organic matter. J. Mass Spectrom., 32: 103-108. [1, 15, 23] Kragten J. (1994) Calculating standard deviations and confidence intervals with a universally applicable spreadsheet technique. Analyst, 119: 2161-2165. [32, 37] Krankowsky D., Bartecki F., Klees G. G., Mauersberger K., Schellenbach K. & Stehr J. (1995) Measurement of heavy isotope enrichment in tropospheric ozone. Geophys. Res. Lett., 22(13): 1713-1716. [18] Krankowsky D., L~immerzahl P. & Mauersberger K. (2000) Isotope measurement of stratospheric ozone. Geophys. Res. Lett., 27(17): 2593-2595. [18] Kretz R. (1983) Symbols for rock forming minerals. Am. Mineralogist, 68(1-2): 277-279. [3] Kreulen R. (1980) CO2-Rich fluids during regional metamorphism on Naxos (Greece): carbon isotopes and fluid inclusions. Am. J. Sci., 280: 745-771. [3] Kreulen R. & Schuiling R. D. (1982) N2-CH4-CO2 fluids during formation of the Dome de L'Agout, France. Geochim. Cosmochim. Acta, 46: 193-203. [3, 16] Kreulen R., Van Breemen A. & Duit W. (1982) Nitrogen and carbon isotopes in metamorphic fluids from the Dome de L'Agout, France. Proc. Fifth Int. Conf. Geochronol. Cosmochronol. Isot. Geol.: 191. [16] Krishnamurthy R. V., Epstein S., Cronin J. R., Pizzarello S. & Yuen G. (1992) Isotopic and molecular analyses of hydrocarbons and monocarboxylic acids of the Murchison meteorite. Geochim. Cosmochim. Acta, 56: 4045-4058. [11] Krishnamurthy R. V., Atekwana E. A. & Guha H. (1997) A simple, inexpensive carbonate - phosphoric acid reaction method for the analysis of carbon and oxygen isotopes of carbonates. Anal. Chem., 69: 4256-4258. [10] Krohn M. D., Kendall C., Evans J. R. & Fries T. L. (1993) Relations of ammonium minerals at several hydrothermal systems in the western U.S.J. Volcanol. Geotherm. Res., 56: 401-413. [16]

1118

References

Krouse H. R. & Thode H. G. (1962) Thermodynamic properties and geochemistry of isotopic compounds of selenium. Can. J. Chem., 40: 367-375. [20, 29, 46] Krouse H. R. & Modzeleski V. E. (1970) 13C/12C abundances in components of carbonaceous chondrites and terrestrial samples. Geochim. Cosmochim. Acta, 34: 459-474. [11] Krouse H. R. (1977) Sulphur isotope studies and their role in petroleum exploration. J. Geochem. Explor., 7: 189-211. [26] Krouse H. R. (1980) Sulphur isotopes in our environment. In: The terrestrial environment, A. P. Fritz & J. C. Fontes (eds.). Elsevier. Handbook of environemental isotope geochemistry, Vol. 1: 435-471. [26] Krouse H. R. & Tabatabai M. A. (1986) Stable sulfur isotopes. In: Sulfur in agriculture. M. A. Tabatabai (ed.). Amer. Soc. Agronomy, Crop Sci. Soc. Amer., Soil Sci. Soc. Amer.: 169-205. [26] Krouse H. R. & Ueda A. (1987) Sulphur isotope analyses of trace sulphide and sulphate in various materials using Kiba reagent. In: Studies on sulphur isotope variations in nature. IAEA, Vienna. Vol. STI/ PUB/747: 113-121. [26] Krouse H. R., Levinson A. A., Piggott D. & Ueda A. (1987a) Further stable isotope investigations of human urinary stones: comparison with other body components. Appl. Geochem., 2: 205-211. [26] Krouse H. R., Ritchie R. G. G. & Roche R. S. (1987b) Sulphur isotope composition of H2S evolved during nonisothermal pyrolysis of sulphur containing materials. J. Anal. Appl. Pyrol., 12: 19-29. [26] Krouse H. R. (1991) Tracing of the fate of emissions from sour gas processing in Alberta, Canada. In: Stable Isotopes in the Assessment of Natural and Anthropogenic Sulphur. H. R. Krouse & V. A. Grinenko (eds.). John Wiley & Sons: 307-326. [26] Krouse H. R., Mayer B. & Schoenau J. J. (1996) Applications of stable isotope techniques to soil sulfur cycling. In: Mass spectrometry of soils. T. W. Boutton & S. Yamasaki (eds.). Marcel Dekker, Inc., New York: 247-284. [26] Krylov D. P., Zagnitko V. N., Hoernes S., Lugovaja I. P. & Hoffbauer R. (2002) Oxygen isotope fractionation between zircon and water; experimental determination and comparison with quartzzircon calibrations. Eur. J. Mineral., 14: 849-853. [20] Kung C. C. & Clayton R. N. (1978) Nitrogen abundances and isotopic compositions in stony meteorites. Earth Planet. Sci. Lett., 38: 421-435. [17] Kuroda P. K. & Sandell E. B. (1953) Chlorine in igneous rocks. Geol. Soc. Am. Bull., 64: 879-896. [28] Kusakabe M., Rafter T. A., Stout J. D. & Collie T. W. (1976) Sulfur isotopic variations in nature. 12. Isotopic ratios of sulfur extracted from some plants, soils and related materials. N. Z. J. Sci., 19: 433440. [26] Kusakabe M. (2001) A simple method for sampling total dissolved carbonate in carbonate-rich natural waters and CO2 preparation for 613C determination. Geochem. J., 35: 459-464. [10] Kush P., Hustralid A. & Tate J. T. (1937) Dissociation processes produced in SbC13, AsC13 and PC13 by electron impact. Phys. Rev., 52: 840-842. [28] Kvenvolden K., Lawless J., Pering K., Peterson E., Flores J., Ponnamperuma C., Kaplan I. & Moore C. (1970) Evidence for extra-terrestrial amino acids and hydrocarbons in the Murchison meteorite. Nature, 228: 926-928. [11] Kwak T. J. & Zedler J. B. (1997) Food web analysis of southern California coastal wetlands using multiple stable isotopes. Oecologia, 110: 262-277. [26] Kyser K., Klassen K. & Mitchell R. (2000) Analysis of N isotopes in micas and amphiboles using EACF-IRMS: problems, solutions and implications for the origin of N in hydrous phases from the mantle. Proc. GeoCanada 2000, Calgary, Canada. [16] Kyser T. K., Carmichael I. S. E. & O'Neil J. R. (1981) Oxygen isotope thermometry of basic lavas and mantle nodules. Contrib. Mineral. Petrol., 81: 88-102. [20, 46] Kyser T. K. (1995) Micro-analytical techniques in stable isotope geochemistry. Can. Mineral., 33: 261-278. [20] Labeyrie L. (1974) New approach to surface seawater paleotemperatures using 180/16 0 ratios in silica diatom frustules. Nature, 248: 40-42. [2, 46] Labeyrie L. D. Jr. & Juillet A. (1982) Oxygen isotopic exchange-ability of diatom valve silica; interpretation and consequences for paleoclimatic studies. Geochim. Cosmochim. Acta, 46: 976-975. [2] Lai A., Casu M., Saba G., Corongiu F. P. & Dessi M. A. (1995) NMR investigation of the intramolecular distribution of deuterium in natural triacylglycerols. Magn. Reson. Chem., 33: 163-166. [5]

References

1119

Lajtha K. & Michener R. H. (1994) Stable isotopes in ecology and environmental science, methods in ecology. Blackwell Scientific Publications, Oxford, UK: 316 p. [9] Lal D., Arnold J. R. & Somayajulu B. L. K. (1964) A method for the extraction of trace elements from seawater. Geochim. Cosmochim. Acta, 7: 1111. [22] Lambert S. J. & Epstein S. (1980) Stable isotope investigations of an active geothermal system in valles Caldera, Jemez Mountains, New Mexico. J. Volcan. Geotherm. Res., 8: 111-129. [46] Lamberty A., Michiels E. & De Bi6vre P. (1987) On the atomic weight of lithium. Int. J. Mass Spectrom. Ion Phys., 79: 311-333. [6] Lancet M. S. & Anders E. (1970) Carbon isotope fractionation in the Fischer-Tropsch synthesis and in meteorites. Science, 170: 980-982. [46] Lancoursi6re J., Meyer M., Nahon L., Morin P. & Larzilli6re M. (1994) Resonance-enhanced twophoton ionization of helium using an Ar+ mode-locked laser synchronized with VUV synchrotron radiation pulses. H.-J. Kluge, J.E. Parks & K. Wendt (eds.). Resonance Ionization Spectroscopy 1994, Seventh International Symposium. AIP Press Woodbury, NY. AIP Conference Proceddings 329: 423426. [33] Landmeyer J. E. & Stone P. A. (1995) Radiocarbon and 513C values related to groundwater recharge and mixing. Ground Water, 33: 227-234. [10] Landon M. K., Delin G. N., Komor S. C. & Regan C. P. (1999) Comparison of the stable-isotopic composition of soil water collected from suction lysimeters wick samples, and cores in a sandy unsaturated zone. J. Hydrot., 224: 45-54. [1] Lang P., Steele L. P. & Martin R. C. (1990) Atmospheric methane data for the period 1986-1988 from the NOAA/CMDL global cooperative flask sampling network. NOAA Climate Monitoring and Diagnostics Laboratory, Boulder, Colorado, USA: 108 p. [14] Lange J. & Brumsack H.-J. (1977) Total sulfur analysis in geological and biological materials by coulometric titration following combustion. Fresenius Z. Anal. Chem., 286: 361-366. [27] Langvad T. (1954) Separation of chlorine isotopes by ion exchange chromatography. Acta Chem. Scandinavia, 8: 526-527. [28] Larras-Regard E. & Mony M.-C. (1997) Improvements in biomedical imaging by SIMS (IMS 3F, 4F, and nanoSIMS). In: SIMS XI. G. Gillen, R. Lareau, J. Bennett & F. Stevie (eds.): 119-122. [30] Larsen N. W., Pedersen T. & Sehested J. (1991) Isotopic study of the mechanism of ozone formation. Int. J. Chem. Kinet., 23: 331-343. [18] Lashof D. A. & Ahuja D. R. (1990) Relative contributions of greenhouse gas emissions to global warming. Nature, 344: 529-531. [14] Lasorsa B. & Casas A. (1996) A comparison of sample handling and analytical methods for determination of acid volatile sulfides in sediment. Marine Chem., 52: 211-220. [26] Lassen J., Benck E. C. & Schuessler H. A. (1994) Ultra-sensitive detection of krypton atoms utilizing step-wise-laser excitation and field ionization. H.-J. Kluge, J.E. Parks & K. Wendt (eds.). Resonance Ionization Spectroscopy 1994, Seventh International Symposium. AIP Press Woodbury, NY. AlP Conference Proceddings 329: 327-330. [33] Lauretta D. S., Bevouard B. & Buseck P. R. (1999) The cosmochemical behavior of mercury. Earth Planet. Sci. Lett., 171: 35-47. [31] Lauretta D. S., Klaue B., Blum J. D. & Buseck P. R. (2001) Mercury abundances and isotopic compositions in the Murchison (CM) and Allende (CV) carbonaceous chondrites. Geochim. Cosmochim. Acta, 65: 2807-2818. [31] Lawrence J. R. (1970) O18/O16 and D / H ratios of soils, weathering zones and clay deposits. PhD Thesis, California Institute of Technology, Pasadena, California, USA: 263 p. [2] Lawrence J. R. & Taylor H. P. Jr. (1971) Deuterium and oxygen-18 correlation: clay minerals and hydroxides in Quaternary soils compared to meteoric waters. Geochim. Cosmochim. Acta, 35: 993-1003. [46] Laws E. A., Popp B. N., Bidigare R. R., Kennicutt M. C. & Macko S. A. (1995) Dependence of phytoplankton carbon isotopic composition on the growth rate and [CO2]aq: theoretical considerations and experimental results. Geochim. Cosmochim. Acta, 59: 1131-1138. [46] Lazar B. & Holland H. D. (1988) The analysis of fluid inclusions in halite. Geochim. Cosmochim. Acta, 52: 485-490. [3]

1120

References

Leaney F. W., Smettem K. R. J. & Chittleborough D. J. (1993) Estimating the contribution of preferential flow to subsurface runoff from a hillslope using deuterium and chloride. J. Hydrol., 147: 83-103. [2] Leavitt S. W. & Long A. (1984) Sampling strategy for stable carbon isotope analysis of tree rings in pine. Nature, 311: 145-147. [24] Leavitt S. W. & Long A. (1991) Seasonal stable isotope variability in tree rings: Possible paleoenvironmental signals. Chem. Geol., 87: 59-70. [24] Leavitt S. W. & Danzer S. R. (1993) Method for batch processing small wood samples to holocellulose for stable-carbon isotope analysis. Anal. Chem., 65: 87-89. [24] Leavitt S. W. (1993a) Environmental information from 13C/12C ratios of Wood. In: Climate change in continental isotopic records. Geophysical Monograph, Vol. 78: 325-331. [24] Leavitt S. W. (1993b) Seasonal 13C/12C changes in tree rings: Species and site coherence, and a possible drought influence. Can. J. Forest Res., 23: 210-218. [24] Leavitt S. W. & Lara A. (1994) South American tree rings show declining 613C trend. Tellus, 46(B): 152157. [24] Leckrone K. J. & Hayes J. M. (1997) Efficiency and temperature dependence of water removal by membrane dryers. Anal. Chem., 69: 911-918. [14] Leckrone K. J. & Hayes J. M. (1998) Water-induced errors in continuous-flow carbon isotope ratio mass spectrometry. Anal. Chem., 70: 2737-2744. [8, 14, 38] Leckrone K. J. & Crowe D. E. (1999) In-situ isotopic analysis of sulfide minerals by UV laser ablationgas chromatography - isotope ratio mass spectrometry. EOS Trans. AGU, 89(17): 51-V. [45] Leder J. J., Swart P. K., Szmant A. M. & Dodge R. E. (1996) The origin of variations in the isotopic record of scleractinian corals: I. Oxygen. Geochim. Cosmochim. Acta, 60: 2857-2870. [46] Lee D.-C. & Halliday A. N. (1995) Precise determinations of the isotopic compositions and atomic weights of molybdenum, tellurium, tin, and tungsten using ICP magnetic sector multiple collector mass spectrometry. Int. J. Mass Spectrom. Ion Proc., 146/147: 35-46. [31, 35] Lee P. S. & Majkowski R. F. (1986) High resolution infrared diode laser spectroscopy for isotope analysis - measurement of isotopic carbon monoxide. Appl. Phys. Lett., 48: 619-621. [34] Lee T., Mayeda T. K. & Clayton R. N. (1980) Oxygen isotopic anomalies in Allende inclusions HAL. Geophys. Res. Lett., 7: 493-496. [20] Lee T. & Tera F. (1986) The meteoritic chromium isotopic composition and limits for radioactive (super 53) Mn in the early solar system. Geochim. Cosmochim. Acta, 50(2): 199-206. [29] Lee-Thorp J. A. & van der Merwe N. J. (1991) Aspects of the chemistry of modern and fossil biological apatites. J. Archaeol. Sci., 18: 343-354. [9] Leeman W. P., Vocke R. D. Jr., Beary E. S. & Paulsen P. J. (1991) Precise boron isotopic analysis of aqueous samples: Ion exchange extraction and mass spectrometry. Geochim. Cosmochim. Acta, 55: 3901-3907. [40] Leeman W. P. & Sisson V. P. (1996) Geochemistry of boron and its implications for crustal and mantle processes. In: Boron: Mineralogy, petrology and geochemistry. E. S. Grew & L. M. Anovitz (eds.). Mineralogical Society of America, Reviews in Mineralogy, Vol. 33: 645-707. [7] Legge A. H., Corbin J., Bogner J. C., Strosher M., Krouse H. R., Laishley E. J., Bryant R. D., Cavey M. J., Prescott C. E., Nosal M., Schellhase H. U., Weidensaul T. C. & Mayo J. (1988) Acidification in a temperate forest ecosystem: the role of sulfur gas emissions and sulfur dust. Phoenix Press Inc., Calgary, Canada: 399 p. [26] Lehmann B., Wahlen M., Zumbrunn R. & Oescher H. (1977) Isotope analysis by infrared laser absorption spectroscopy. Appl. Phys., B 13: 153-158. [34] Lehmann K. K. (1996) Ring-down cavity spectroscopy cell using continuous wave excitation for trace species detection. U.S. Patent 5,528,040. [34] Lehmann M. F., Bernasconi S. M. & McKenzie J. A. (2001) A method for the extraction of ammonium from freshwaters for nitrogen isotope analysis. Anal. Chem., 73: 4717-4721. [15] Lei J. J., Ren D. u Tang Y. G., Chu X. L. & Zhao R. (1994) Sulfur-accumulating model of super-high organosulfur coal from Guiding, China. Chinese Sci. Bull., 39: 1817-1821. [26] Leshin L. A., Rubin A. E. & McKeegan K. D. (1997) The oxygen isotopic composition of olive and pyroxene from CI chondrites. Geochim. Cosmochim. Acta, 61: 835-845. [30]

References

1121

Leshin L. A., McKeegan K. D., Carpenter P. K. & Harvey R. P. (1998) Oxygen isotopic constraints on the genesis of carbonates from martian meteorite ALH84001. Geochim. Cosmochim. Acta, 62: 3-13. [30]

Leshin L. A. (2000) Implications for the origin and history of water on Mars from analyses of martian meteorite QUE94201. Geophys. Res. Lett., 27: 2017-2020. [30]

Leuenberger M., Borella S., Stocker T., Saurer M., Siegwolf R., Schweingruber F. & Matyssek R. (1998) Stable isotopes in tree rings as climate and stress indicators, vdf Hochschulverlag AG an der ETH Zfirich, Zfirch, Switserland. [24] Leuenberger M. & Huber C. (2002) On-line determination of oxygen isotope ratios of water or ice by mass spectrometry. Anal. Chem., 74: 4611-4617. [1] Lewis R. S., Tang M., Wacker J. F., Anders E. & Steel E. (1987) Interstellar diamonds in meteorites. Nature, 326: 160-162. [13] L6cluse C. & Robert F. (1994) Hydrogen isotope exchange reaction rates: origin of water in the inner solar system. Geochim. Cosmochim. Acta, 58: 2927-2939. [46] L6cuyer C., Grandjean P., O'Neil J. R., Cappette H. & Martineau F. (1993) Thermal excursions in the ocean at the Cretaceous-Tertiary boundary (northern Morocco): 6180 record of phosphatic fish debris. Palaeogeog. Palaeoclimatol. Palaeoecol., 105: 235-243. [20, 22] L6cuyer C. & O.Neil J. R. (1994) Stable isotope compositions of fluid inclusions in biogenic carbonates. Geochim. Cosmochim. Acta, 58: 353-363. [3] L6cuyer C., Grandjean P. & Emig C. C. (1996a) Determination of oxygen isotope fractionation between water and phosphate from living lingulids: potential application to palaeoenvironmental studies. Paleogeogr., Palaeoclim., Palaeoecol., 126: 101-108. [22, 46] L6cuyer C., Grandjean P., Paris F., Robardet M. & Robineau D. (1996b) Deciphering "temperature" and "salinity" from biogenic phosphates: the 6180 of coexisting fishes and mammals of the Middle Miocene sea of western France. Palaeogeogr., Palaeoclim., Palaeoecol., 126: 61-74. [22] L6cuyer C., Grandjean P., Barrat J.-A., Nolvak J., Emig C., Paris F. & Robardet M. (1998) 6180 and REE contents of phosphatic brachiopods: A comparison between modern and lower Paleozoic populations. Geochim. Cosmochim. Acta, 62(14): 1429-2436. [22] L6cuyer C., Grandjean P. & Sheppard S. M. F. (1999) Oxygen isotope exchange between dissolved phosphate and water at temperatures M35~ Inorganic versus biological fractionations. Geochim. Cosmochim. Acta, 63(6): 855-862. [22, 46] L6cuyer C., Grandjean P., Reynard B., Albar6de F. & Telouk P. (2000) 11B/ 10B analysis of geochemical materials by ICP-MS Plasma 54: methodology and paleo-pH of seawater. EOS Trans. AGU, 81: F700. [31] Li B. & Shi J. (1983) Continuous determination of stable isotopes of hydrogen-gases and salts in fluid inclusions. Scientia Sinica, Series B: Chemical, Biological, Agricultural, Medical and Earth Sciences (Chung Kuo K'o Hsueh Yuan, Chu Pan), 26(1): 93-102. [3] Li G. Q., Duan Y. X. & Hieftje G. M. (1995) Space-charge effects and ion distribution in plasma source mass spectrometry. J. Mass Spectrom., 30: 841-848. [35] Li W., Ni B., Jin D. & Zhang Q. (1988) Measurement of the absolute abundance of oxygen-17 in VSMOW. Kexue Tongbao (Chinese Sci. Bull.), 33: 1610-1613. [4, 40, 44] Li Y. M. (1996) Sulfur isotope composition of SO2 prepared by laser assited oxidation of sulfur. PhD Thesis, University of Calgary, Calgary, Canada: 149 p. [26] Libby L. M., Pandolfi L. J., Payton P. H., Marshall J., Becker B. & Giertz-Siebenlist V. (1976) Isotope tree thermometers. Nature, 261: 284-288. [23] Lichtfouse 1~., Elbisser B., Balesdent J., Mariotti A. & Bardoux G. (1995) Isotope and molecular evidence for direct input of maize leaf wax n-alkanes into crop soils. Org. Geochem., 22: 349-351. [9] Lichtfouse l~., Bardoux G., Mariotti A., Balesdent J., Ballentine D. C. & Macko S. A. (1997) Molecular, 13C, and 14C evidence for the allochthonous and ancient origin of C16 - C18 n-alkanes in modern soils. Geochim. Cosmochim. Acta, 61: 1891-1898. [9] Lide R. A. (1994) Handbook of chemistry and physics. CRC-Press, Ann Arbor, USA. 75th edition. [36] Liebl H. (1983) Ion gun systems for submicron SIMS. In: SIMS. J. Okano, R. Shimizu & H. W. Werner (eds.). Springer. Vol. IV: 114-118. [30] Liewig N., Caluer N. & Sommer F. (1987) Rb-Sr and K-Ar dating of clay diagenesis in Jurassic sandstone oil reservoir, north Sea. Am. Assoc. Petrol. Geol. Bull., 71: 1467-1474. [2]

1122

References

Lindars E. S., Grimes S. T., Mattey D. P., Collinson M. E., Hooker J. J. & Jones T. P. (2001) Phosphate ~)180 determination of modern rodent teeth by direct laser fluorination: An appraisal of methodology and potential application to palaeoclimate reconstruction. Geochim. Cosmochim. Acta, 65(15): 2535-2548. [20] Linder P., Bermann E. & Gamarnik B. (1996) Characterization of citrus honey by deuterium NMR. J. Agric. Food Chem., 44: 139-140. [5] Lipschultz F., Wofsy S. C. & Fox L. E. (1986) Nitrogen metabolism of the eutrophic Delaware River ecosystem. Limnol. Oceanogr., 31: 701-716. [15] Litaor M. I. (1988) Review of soil solution samplers. Water Resour. Res., 24: 727-733. [2] Liu K. K. & Epstein S. (1984) The hydrogen isotope fractionation between kaolinite and water. Isotope Geosci., 2: 335-350. [46] Liu K. K., Su M. J., I-Isueh C. R. & Gong G. C. (1996) The nitrogen isotopic composition of nitrate in the Kuroshio Water northeast of Taiwan: evidence for nitrogen fixation as a source of isotopically light nitrate. Marine Chem., 54: 273-292. [9] Liu S. C. & Trainer M. (1988) Responses of the tropospheric ozone and odd hydrogen radicals to column ozone change. J. Atmos. Chem., 6: 221-234. [18] Liu W. G., Xiao Y. K., Peng Z. C., An Z. S. & He X. X. (2000) Boron concentration and isotopic composition of halite from experiments and salt lakes in the Qaidam Basin. Geochim. Cosmochim. Acta, 64: 2177-2183. [46] Liu u P. & Mulvaney R. L. (1992) Use of diffusion for automated nitrogen-15 analysis of soil extracts. Commun. Soil. Sci. Plant Anal., 23: 613-629. [15] Livingston S. W. & Spittlehouse D. L. (1996) Carbon isotope fractionation in tree ring early and late wood in relation to intragrowing season water balance. Plant Cell Environm., 19: 768-774. [24] Lloyd R. M. (1968) Oxygen isotope behavior in the sulfate-water system. J. Geophys. Res., 73: 6099-6110. [46] Loader N. J., Robertson I., Barker A. C., Switsur V. R. & Waterhouse J. S. (1997) An improved technique for the batch processing of small wholewood samples to c~-cellulose. Chem. Geol., 136: 313-317. [24] Loader N. J.& Buhay W. M. (1999) Rapid catalytic oxidation of CO to CO2 - On the development of a new approach to on-line oxygen isotope analysis of organic matter. Rapid Commun. Mass Spectrom., 13: 1828-1832. [1] Lober R. W., Reeder J. D. & Porter L. K. (1987) New and conventional evaporitive systems in concentrating nitrogen samples prior to isotope-ratio analysis. Soil. Sci. Soc. Am. J., 51: 951-958. [15] Lockheart M. J., Van Bergen P. F. & Evershed R. P. (1997) Variations in the stable carbon isotope compositions of individual lipids from the leaves of modern angiosperms: implications for the study of higher land plant-derived sedimentary organic matter. Org. Geochem., 26: 137-153. [9] Long A., Eastoe C. J., Kaufmann R. S., Martin J. G., Wirt L. & Finley J. B. (1993) High-precision measurement of chlorine stable isotope ratios. Geochim. Cosmochim. Acta, 57: 2907-2912. [28] Longerich H. P., Fryer B. J. & Strong D. F. (1987) Determination of lead isotope ratios by inductively coupled plasma-mass spectrometry (ICP-MS). Spectrochim. Acta, B 42: 39-48. [31, 32, 35] Longinelli A. (1965) Oxygen isotope composition of orthophosphate from shells of living marine organisms. Nature, 207: 716. [20, 22, 46] Longinelli A. (1966) Ratios of oxygen-18:oxygen-16 in phosphate and carbonate from living and fossil marine organisms. Nature, 211: 923-927. [22, 46] Longinelli A. & Nuti S. (1968a) Oxygen isotope composition of phosphates from marine formations. Earth Planet. Sci. Lett., 5: 13-17. [22] Longinelli A. & Nuti S. (1968b) Oxygen-isotope ratios in phosphate from marine organisms. Science, 160: 879-882. [22] Longinelli A. & Nuti S. (1973a) Revised phosphate-water isotopic temperature scale. Earth Planet. Sci. Lett., 19: 373-376. [20, 22, 46] Longinelli A. & Nuti S. (1973b) Oxygen isotope measurements of phosphate from fish teeth and bones. Earth Planet. Sci. Lett., 20: 337-340. [22] Longinelli A., Bartelloni M. & Cortecci G. (1976) The isotopic cycle of oceanic phosphate, I. Earth Planet. Sci. Lett., 32: 389-392. [22]

References

1123

Longinelli A. (1984) Oxygen isotopes in mammal bone phosphate: A new tool for paleohydrological and paleoclimatological research? Geochim. Cosmochim. Acta, 48: 385-390. [20, 22] Longstaffe F. J. (1986) Oxygen isotope studies of diagenesis in the basal Belly River sandstone, Pembina I-pool, Alberta. J. Sediment. Petrol., 56: 78-88. [2] Longstaffe F. J. & Mizota C. (1999) Rapid changes in the oxygen- and hydrogen-isotope composition of halloysite during laboratory wetting at 23~ Geol. Soc. Am. Abstr. w. Prog., (Abstract). [2] Lory J. & Russelle M. P. (1994) Evaluation of diffusion method for preparing low-nitrogen samples for 15N analysis. Soil Sci. Soc. Am. J., 58: 1400-1404. [15] Lovley D. R., Stolz J. F., Nord G. L. & Phillips E. J. P. (1987) Anaerobic production of magnetite by dissimilatory iron-reducing microorganism. Nature, 330: 252-254. [29] Lovley D. R. & Phillips E. J. P. (1994) Reduction of chromate by Desulfovibrio viulgaris and its C-3 cytochrome. Appl. Env. Microbiol., 60: 726-728. [29] Lowe D. C., Brenninkmeijer C. A. M., Tyler S. C. & Dlugokencky E. J. (1991) Determination of the isotopic composition of atmospheric methane and its application in the Antarctic. J. Geophys. Res., 96(D8): 15455-15467. [14] Lowe D. C., Brenninkmeijer C. A. M., Brailsford G. W., Lassey K. R., Gomez A. J. & Nisbet E. G. (1994) Concentration and C-13 records of atmospheric methane in New-Zealand and Antarctica- Evidence for changes in methane sources. J. Geophys. Res., 99(D8): 16913-16925. [14] Lowe D. C., Manning M. R., Bromley T. B., Allen W., Brailsford G., Ferretti D., Gomez A., Knobben R., Martin R., Mei Z., Moss R., Koshy K. & Maata M. (1999) Shipboard determinations of the distribution of 13C in atmospheric methane in the Pacific. J. Geophys. Res., 104(D21): 26125-26135. [14] Lowe L. E. (1965) Sulphur fractions of selected Alberta soil profiles of the chernozemic and podzolic orders. Can. J. Soil Sci., 45: 297-303. [26] Lowenstam H. A. (1961) Mineralogy, 180/16 0 ratios, and strontium and magnesium contents of recent and fossil brachiopods and their bearing on the history of the oceans. J. Geol., 69: 241-260. [46] Lu Z.-T. & Wendt K. D. A., (2003) Laser-based methods for ultrasensitive trace-isotope analyses. Rev. Sci. Instr., 74: 1169-1407. [34] Lund6n A. & Herzog W. (1956) Isotopenanreicherung bei Chlor durch electrolytische Oberfiihrung in geschmolzenem Zinkchlorid. Z. Naturforsch., 11a: 520. [28] Lund6n A. & Lodding A. (1960) Isotopenanreicherung bei Brom durch electrolytische Oberfiihrung in geschmolzenem Zinkbromid. Z. Naturforsch., 15a: 320-322. [28] Luo Y. H. & Sternberg L. (1992) Spatial D / H heterogeneity of leaf water. Plant Physiol., 99: 348-350. [21] Lusk J. & Krouse H. R. (1997) Comparative stable isotope and temperature investigation of minerals and associates fluids in two regionally metamorphosed (Kuroko-type) volcanogenic massive sulfide deposits. Chem. Geol., 143: 231-253. [3] Luz B., Kolodny Y. & Kovach J. (1984) Oxygen isotope variations in phosphate of biogenic apatites. III. Conodonts. Earth Planet. Sci. Lett., 69: 255-262. [20, 22] Luz B. & Kolodny Y. (1985) Oxygen isotope variation in phosphate of biogenic apatites, IV. Mammal teeth and bones. Earth Planet. Sci. Lett., 75: 29-36. [22] Luz B., Barkan E., Bender M. L., Thiemens M. H. & Boering K. A. (1999) Triple-isotope composition of atmospheric oxygen as a tracer for biospheric productivity. Nature, 400: 547-550. [18] Lyon G. L. & Cox M. A. (1978) Stable isotope analysis of hydrogen and methane mixtures. Inst. Nucl. Sci., DSIR, Lower Hutt, New Zealand. Int. Rept. INS-R--282. [33] Lyon G. L. & Cox M. A. (1980) The reduction of water to hydrogen for D / H ratio analysis, using zinc in a matrix of sand. Inst. Nucl. Sci., DSIR, Lower Hutt, New Zealand. Int. Rept. INS-R-282, November, 1980: 7p. [1] Lyon I. C., Saxton J. M., McKeever P. J., Chatzitheodoridis E. & Van Lierde P. (1995) Precision and reproducibility on quartz obtained using an Isolab 54 ion microprobe. Int. J. Mass Spectrom. Ion Phys., 9(8): 1-16. [30] Lyon T. D. B. & Fell G. S. (1990) Isotopic composition of copper in serum by inductively coupled plasma mass spectrometry. J. Anal. Atom. Spectrom., 5: 135-137. [31] Lyon T. D. B., Fletcher S., Fell G. S. & Patriarca M. (1996) Measurement and application of stable copper isotopes to investigations of human metabolism. Microchem. J., 54: 236-245. [31]

1124

References

Lyons J. R. (2001) Transfer of mass-independent fractionation in ozone to other oxygen-containing radicals in the atmosphere. J. Geophys. Res., 28(17): 3231-3234. [18] Lyons T. W., Zhang C. L. & Romanek C. S. (2003) Isotopic records of microbially mediated processes. Chem. Geol., 195(edited issue): 1-274. [46] MacAvoy S. E., Macko S. A. & Garman G. C. (1998) Tracing marine biomass into tidal freshwater ecosystems using stable sulfur isotopes. Naturwissensch., 85: 544-546. [26] MacAvoy S. E., Macko S. A., McInnich S. P. & Garman G. C. (2000) Marine nutrient contributions to freshwater apex predators. Oecologia, 122: 568-573. [26] MacAvoy S., Macko S. A. & Garman G. (2001) Isotopic turnover in aquatic predators: quantifying the explotation of migratory prey. Can. J. Fish. Aq. Sci., 58: 923-932. [45] MacDaniel C. V. & Maher P. K. (1976) Zeolite stability and ultrastable zeolites. In: Zeolite chemistry and catalysis. J. A. Rabo (ed.). American Chemical Society, Washington. ACS Monograph 171, Chapter 4: 285-331. [36] MacElroy J. M. D., Freidman S. P. & Seaton N. A. (1999) On the origin of transport resistences within carbon molecular sieves. Chem. Eng. Sci., 54: 1015-1027. [36] MacFarlane A. W. & Shimizu N. (1991) SIMS measurement of 634S in sulfide minerals from adjacent vein and stratabound ores. Geochim. Cosmochim. Acta, 55: 525-541. [26, 30] Macko S. (1981) Stable nitrogen isotope ratios as tracers of organic geochemical processes. PhD Thesis, University of Texas, Austin, Texas, USA: 181 p. [15, 16, 36] Macko S. A., Fogel Estep M. L., Engel M. H. & Hare P. E. (1986) Kinetic fractionation of stable nitrogen isotopes during amino acid transamination. Geochim. Cosmochim. Acta, 50: 2143-2146. [46] Macko S. A., Estep M. L. F., Hare P. E. & Hoering T. C. (1987) Isotopic fractionation of nitrogen and carbon in the synthesis of amino acids by microorganisms. Chem Geol. (Isot. Geosc.), 65: 79-92. [9] Macko S. A., Helleur R., Hartley G. & Jackman P. (1990) Diagenesis of organic matter - A study using stable isotopes of individual carbohydrates. Org. Geochem., 16: 1129-1137. [9] Macko S. A., Engel M. H., Hartley G., Hatcher P., Helleur R., Jackman P. & Silfer J. A. (1991) Isotopic compositions of individual carbohydrates as indicators of early diagenesis of organic matter in peat. Chem. Geol., 93: 147-161. [9] MacNamara J. & Thode H. G. (1950) Comparison of the isotopic constitution of terrestrial and meteoritic sulfur. Phys. Rev., 78: 307-308. [26, 40, 43, 44] Macpherson C., Hilton D. R., Newman S. & Mattey D. P. (1999) CO2, 13C/12C and H20 variability in natural basaltic glasses: A study comparing stepped heating and FTIR spectroscopic techniques. Geochim. Cosmochim. Acta, 63: 1805-1813. [36] Madigan M. T., Martinko J. M. & Parker J. (2000) Brock Biology of Microorganisms. Prentice Hall Int. 9th edition. [46] Madorsky S. L. & Strauss S. (1948) Concentration of isotopes of chlorine by the counter-current electromigration method. J. Res. Nat. Bur. Stand., 38: 185-189. [28] Madu B. E., Nesbitt B. E. & Muehlenbachs K. (1990) A mesothermal gold-stibnite-quartz vein occurrence in the Canadian Cordillera. Econ. Geol., 85(6): 1260-1268. [3] Magaritz M. & Taylor H. P. Jr. (1976) 180/160 and D / H studies along a 500 km traverse across the Coast Range batholith and its country rocks, central British Columbia. Can. J. Earth Sci., 13: 1514-1536. [20] Magenheim A. J., Spivack A. J., Volpe C. & Ranson B. (1994) Precise determination of stable chlorine isotope ratios in low-concentration natural samples. Geochim. Cosmochim. Acta, 58: 3117-3121. [28] Magenheim A. J., Spivack A. J., Michael P. J. & Gieskes J. M. (1995) Chlorine stable isotope composition of the oceanic crust: implications for Earth's distribution of chlorine. Earth Planet. Sci. Lett., 131: 427-432. [46] Mahon K. I., Harrison T. M. & McKeegan K. D. (1998) The thermal and cementation histories of a sandstone petroleum reservoir, Elk Hills, California. Part 2: In situ oxygen and carbon isotopic results. Chem. Geol., 152: 257-271. [30] Majoube M. (1971) Fractionnement en oxyg6ne 18 et en deut6rium entre l'eau et sa vapeur. J. Chim. Phys. Physicochim. Biol., 68: 1423-1436. [46] Majzoub M. A. (1966) Une methode d'analyse isotopique de l'oxyg6ne sur des microquantit6s d'eau; determination des coefficients de partage a l'6quilibre de l'oxig6ne 18 entre H20 et CO2; D20 et CO2. J. chim. Phys., 63(4): 565-568. [1, 20]

References

1125

Majzoub M. & Nief G. (1968) A mass spectrometer for the isotopic oxygen analysis of water. Adv. Mass Spectrom., 4: 511-520. [1] Maliva R. G., Dickson J. A. D. & Fallick A. E. (1999) Kaolin cements in limestones: Potential indicators of organic-rich pore waters during diagenesis. J. Sedim. Petrol., 69: 158-163. [2] Mandernack K. W., Bazylinski D. A., Shanks W. C. III& Bullen T. D. (1999) Oxygen and iron isotope studies of magnetite produced by magnetotactic bacteria. Science, 285: 1892-1896. [29, 31, 46] Mandernack K. W., Lynch L., Krouse H. R. & Morgan M. D. (2000) Sulfur cycling in wetland peat of the New Jersey Pinelands and its effect on stream water chemistry. Geochim. Cosmochim. Acta, 64: 3949-3964. [26] Manheim F. T. & Sayles F. L. (1974) Composition and origin of interstitial waters of marine sediments based on deep sea drill cores. In: The Sea. E. D. Goldberg (ed.). Wiley, NY. Vol. 5 : 527-568. [2] Manian S. H., Urey H. C & Bleakney W. (1934) An investigation of the relative abundance of the oxygen isotopes O16:O 18 in stone meteorites. J. Am. Chem. Soc., 56: 2601-2609. [20] Manowitz B., Krouse H. R., Barker C. & Premuzic E. T. (1990) Sulfur isotope data analysis of crude oils from Bolivar Coastal fields (Venezuela). In: Geochemistry of sulfur in fossil fuels. W. L. Orr & C. M. White (eds.). American Chemical Society. Vol. 429: 592-612. [26] Mansuy L., Philp R. P. & Allen J. (1997) Source identification of oil spills based on the isotopic composition of individual components in weathered oil samples. Environm. Sc. & Techn., 31: 3417-3425. [8] Manzano B. K., Fowler M. G. & Machel H. G. (1997) The influence of thermochemical sulphate reduction on hydrocarbon composition in Nisku reservoirs, Brazeau river area, Alberta, Canada. Org. Geochem., 27: 507-521. [26] Marcus R. K. (1993) Operation principles and design considerations for radiofrequency powered glow discharge devices. J. Anal. At. Spectrom., 8: 935-943. [35] Marcus R. K., Harville T. R., Mei Y. & Schick C. R. Jr. (1994) Rf-powdered glow discharges: Elemental analysis across the solids spectrum. Anal. Chem., 66: 902A-911A. [35] Mar6chal C. N., T61ouk P. & Albar6de F. (1999) Precise analysis of copper and zinc isotopic compositions by plasma-source mass spectrometry. Chem. Geol., 156: 251-273. [31, 32, 46] Mar6chal C. N., Nicolas E., Douchet C. & Albar6de F. (2000) Abundance of zinc isotopes as a marine biogeochemical tracer. Geochem. Geophys. Geosys., 1: 1999GC000029. [31] Mar6chal C. & Albari~de F. (2002) Ion-exchange fractionation of copper and zinc isotopes. Geochim. Cosmochim. Acta, 66:1499-1509 (and see GCA, 66: 3675). [46] Mariotti A., German J. C., Hubert P., Kaiser P., Letolle R., Tradieux A. & Tradieux P. (1981) Experimental determination of nitrogen kinetic isotope fractionation: some principles; Illustration for the denitrification and nitrification processes. Plant and Soil, 62: 413-430. [46] Mariotti A. (1983) Atmospheric nitrogen is a reliable standard for natural 15N abundance measurements. Nature, 303: 680-683. [40] Marra P. P., Hobson K. A. & Holmes R. T. (1998) Linking winter and summer events in a migratory bird by using stable-carbon isotopes. Science, 282: 1884-1886. [9] Marsh K. L., Mulvaney R. L. & Khan S. A. (2003) Use of diffusion for enzymatic determination of urea-nitrogen in soil extracts. Commun. Soil Sci. Plant Anal., in press. [15] Marshall D., Pfeiffer H. R. & Sharp Z. D. (1994) A re-evaluation of Raman as a tool for the determination of 12C and 13C in geological fluid inclusions. Analusis, 22(10): 38-41. [3] Martin G. J. & Martin M. L. (1981) Deuterium labelling at the natural abundance level as studied by high field quantitative 2H-NMR. Tetrahedron Lett., 22(36): 3525-3528. [5] Martin G. J., Martin M. L., Mabon F. & Bricout J. (1982a) A new method for the identification of the origin of natural products. Quantitative 2H NMR at the natural abundance level applied to the characterization of anetholes. J. Am. Chem. Soc., 104: 2658-2659. [5] Martin G. J., Martin M. L., Mabon F. & Michon M. J. (1982b) Identification of the origin of natural alcohols by natural abundance hydrogen-2 nuclear magnetic resonance. Anal. Chem., 54(13): 2380-2382.

[51

Martin G. J., Martin M. L., Mabon F. & Michon M. J. (1983a) A new method for the identification of the origin of ethanols in grain and fruit spirits: high field quantitative deuterium nuclear magnetic resonance at the natural abundance level. J. Agric. Food Chem., 31: 311-315. [5]

1126

References

Martin G. J., Martin M. L., Mabon F. & Bricout J. (1983b) Natural Abundance Deuterium NMR, a new method of identification of the origin of food products applied to the recognition of anetholes and estragoles. Sci. Aliments, 3: 147-155. [5] Martin G. J., Benbernou M. & Lantier F. (1985a) Application of site-specific (SNIF NMR) natural isotope fractionation of hydrogen to the characterization of european beers. J. Inst. Brew., 91: 242-249.

[51

Martin G. J., Janvier P. & Mabon F. (1985b) Caract6risation des an6tholes par RMN du deuterium. Analyse factorielle des rapports isotopiques sp6cifiques. Analusis, 13(6): 267-274. [5] Martin G. J., Zhang L. L., Naulet N. & Martin M. L. (1986a) Deuterium transfer in the bioconversion of glucose to ethanol studied by specific isotope labeling at the natural abundance level. J. Am. Chem. Soc., 108(17): 5116-5122. [5] Martin G. J., Guillou C., Naulet N., Brun S., Tep Y., Cabanis J. C., Cabanis M. T. & Sudraud P. (1986b) Contr61e de l'origine et de l'enrichissement des vins par analyse isotopique sp6cifique. Sci. Aliments, 6: 385-405. [5] Martin G. J., Janvier P., Akoka S., Mabon F. & Jurczak J. (1986c) A relation between the site-specific natural deuterium contents in R-pinenes and their optical activity. Tetrahedron Lett., 27(25): 28552858. [51 Martin G. J., Guillou C., Martin M. L., Cabanis M. T., Tep Y. & Aerny J. (1988) Natural factors of isotope fractionation and the characterization of wines. J. Agric. Food Chem., 36: 316-322. [5] Martin G. J. & Naulet N. (1988) Precision, accuracy and referencing of isotope ratios determined by nuclear magnetic resonance. Fresenius Z. Anal. Chem., 332: 648-651. [5] Martin G. J., Danho D. & Vallet C. (1991a) Natural isotope fractionation in the discrimination of sugar origins. J. Sci. Food. Agric., 56: 419-434. [5] Martin G. J., Martin M. L. & Zhang B. L. (1992) Site-specific natural isotope fractionation of hydrogen in plant products studied by nuclear magnetic resonance. Plant Cell Environ., 15: 1037-1050. [5] Martin G., Remaud G. S. & Martin G. J. (1993a) Isotopic methods for control of natural flavours authenticity. Flavour Frag. J., 8: 97-107. [5] Martin G. J., Hanneguelle S. & Remaud G. (1993b) Application of site-specific natural isotope fractionation studied by Nuclear Magnetic Resonance (SNIF-NMR) to the detection of flavour and fragrance adulteration. Ital. J. Food Sci., 3: 191-213. [5] Martin G. J., Trierweiler M., Ristow R., Hermann A. & Belliardo J. J. (1994) The certification of the three reference ethanols for SNIF-NMR. BCR certified materials, EUR 15347 EN. [5] Martin G. J., Koziet J., Rossmann A. & Dennis J. (1996b) Site-specific natural isotope fractionation in fruit juices determined by deuterium NMR. An european inter-laboratory comparison study. Anal. Chim. Acta, 321: 137-146. [5] Martin G. G., Guillou C. & Caisso M. (1991c) Detection of beet sugar addition in fruit juices and concentrates. An overview of isotopic techniques including new development made possible by 2H-NMR. Flfdssiges Obst, 1(2): 18-21. [5] Martin G. G., Hanote V., Lees M. & Martin Y. L. (1996a) Interpretation of combined 2H SNIF/NMR and 13C SIRA/MS analyses of fruit juices to detect added sugar. J. AOAC Int., 79(1): 62-72. [5] Martin G. G., Martin Y. L., Naulet N. & McManus H. J. D. (1996c) Application of 2H SNIF-NMR and 13C SIRA-MS analyses to maple syrup: detection of added sugars. J. Agric. Food Chem., 44: 3206-3213.

[51

Martin M. L. & Martin G. J. (1990) Deuterium NMR in the study of site-specific natural isotope fractionation (SNIF-NMR). In: NMR basic principles and progress. P. Dietl, E. Fluck, H. Gfinter, R. Kosfeld& J. Seelig (eds.). Springer Verlag, Berlin. Vol. 23: 1-61. [5] Martin M. L., Martin G. J. & Guillou C. (1991b) A site-specific and multi-element isotopic approach to origin interference of sugars in foods and beverages. Mikrochim. Acta [Wien], Ih 81-91. [5] Martin Y. L. (1994) A global approach to accurate and automatic quantitative analysis of NMR spectra by complex least-squares curve fitting. J. Magn. Reson., Ser. A, 111: 1-10. [5] Martinelli R. U. (1996) Mid-infrared wavelengths enhance trace-gas sensing. Laser Focus World, 3 (March): 7781-7783. [34] Marry B., Lenoble M. & Vassard N. (1995) Nitrogen, helium and argon in basalt: a static mass spectrometry study. Chem. Geol., 120: 183-185. [13, 17]

References

1127

Marty B. & Humbert F. (1997) Nitrogen and argon isotopes in oceanic basalts. Earth Planet. Sci. Lett., 152: 101-112. [17] Marumo K., Longstaffe F. J. & Matsubaya O. (1995) Stable isotope geochemistry of clay minerals from fossil and active hydrothermal systems, southwestern Hokkaido, Japan. Geochim. Cosmochim. Acta, 59: 2545-2559. [2] Mas F., Beemsterboer B., Veltkamp A. C. & Verweij A. M. A. (1995) Determination of common-batch members in a set of confiscated 3,4-(Methylendioxy)Methylamphetamine samples by measuring the natural isotope abundances - a preliminary-study. Forensic Sci. Int., 71: 225-231. [8] Masarie K. A., Langenfelds R. L., Allison C. E., Conway T. J., Dlugokencky E. J., Francey R. J., Novelli P. C., Steele L. P., Tans P. P., Vaughn B. & White J. W. C. (2001) The NOAA / CSIRO Flask-Air Intercomparison Program: A strategy for directly assessing consistency among atmospheric measurements derived from independent laboratories. J. Geophys. Res., 105(D17): 20445-20464. [14] Mason P. R. D., de Hoog C.-J. & Meffan-Main S. (1999a) A new method for in-situ high precision sulfur isotope ratio measurmeents by laser ablation multiple collector ICP-MS. Ninth Annual V.M. Goldschmidt Conference, Abst. #7393. [31] Mason P., Kaspers K. & Van Bergen M. J. (1999b) Determination of sulphur isotope ratios and concentrations using ICP-MS incorporating hexapole ion optics. J. Anal. At. Spectrom., 14: 1067-1084. [26, 31] Matheney R. K. & Knauth L. P. (1989) Oxygen-isotope fractionation between marine biogenic silica and seawater. Geochim. Cosmochim. Acta, 53: 3207-3214. [2] Mathieu H. J. & Leonard D. (1998) Use of post-ionization techniques to complement SIMS analysis. A review with practical aspects. High Temp. Materials Proc., 17: 29-44. [33] Matschat R., Czerwensky M., Pattberg S. & Heinrich H.-J. (2001) The certification of high purity metals as primary referencve materials - A challenge for multi-element trace analysis. Phys. Stat. Solidi, A 189: 107-122. [42] Matschat R., Czerwensky M., Pattberg S., Heinrich H.-J. & Tutschku S. (2002) High purity metals as primary calibration materials for element analysis - Their importance and their certification. Mater. Trans. (JIM), 43: 90-97. [42] Matsuhisa Y., Goldsmith J. R. & Clayton R. N. (1978) Mechanisms of hydrothermal crystallization of quartz at 250~ and 15 kbar. Geochim. Cosmochim. Acta, 42: 173-182. [45, 46] Matsuhisa Y., Goldsmith J. R. & Clayton R. N. (1979) Oxygen isotopic fractionation in the system quartz-albite-anorthite-water. Geochim. Cosmochim. Acta, 43: 1131-1140. [46] Matsumi Y., Kishigami M., Tanaka N., Kawasaki M. & Inoue G. (1998) Isotope 180 / 160 ratio measurements of water vapor by use of photoacoustic spectroscopy. Appl. Opt., 37: 6558-6562. [34] Matsuo S. (1991) Technical problems related to the precise measurements of isotope ratios of H20 and CO2 extracted from fluid inclusions. International Institute for Geothermal Research, Pisa, Italy: 383-408.

[3]

Mattauch J. & Herzog R. (1934) Uber einen neuen Massenspektographen. Z. Physik, 89: 786-795. [38] Mattey D. P., Carr R. H., Wright I. P. & Pillinger C. T. (1984) Carbon isotopes in submarine basalts. Earth Planet. Sci. Lett., 70: 196-206. [3] Mattey D. P., Exley R. A. & Pillinger C. T. (1989) Isotopic composition of CO2 and dissolved carbon species in basalt glass. Geochim. Cosmochim. Acta, 53: 2377-2386. [3] Mattey D. P. (1991) Carbon dioxide solubility and carbon isotope fractionation in basaltic melt. Geochim. Cosmochim. Acta, 55: 3467-3473. [46] Mattey D. & Macpherson C. (1993) High-precision oxygen isotope microanalysis of ferromagnesian minerals by laser-fluorination. Chem. Geol. (Isot. Geosci. Sect.), 105: 305-318. [20] Mattey D., Jackson D. H., Harris N. B. W. & Kelley S. (1994) Isotopic constraints on fluid infiltration from an eclogite facies shear zone, Holsenoy, Norway. Metamorph. Geol., 12: 311-325. [16] Matthews D. E. & Hayes J. M. (1978) Isotope-ratio-monitoring gas chromatography-mass spectrometry. Anal. Chem., 50: 1465-1473. [11, 20, 22, 38] Matthews A. & Beckinsale R. D. (1979) Oxygen isotope equilibration systematics between quartz and water. Am. Mineralogist, 64: 232-240. [46] Matthews A., Goldsmith J. R. & Clayton R. N. (1983a) On the mechanisms and kinetics of oxygen isotope exchange in quartz and feldspar at elevated temperatures and pressures. Geol. Soc. Am. Bull., 94: 396-412. [46]

1128

References

Matthews A., Goldsmith J. R. & Clayton R. N. (1983b) Oxygen isotope fractionatios involving pyroxenes: the calibration of mineral-pair geothermometers. Geochim. Cosmochim. Acta, 47: 631-644. [46] Matthews A., Palin J. M., Epstein S. & Stolper E. M. (1994) Experimental study of 180 / 160 partitioning between crystalline albite, albitic glass, and CO2 gas. Geochim. Cosmochim. Acta, 58(23): 5255-5266. [46] Matthews A., Ayalon A. & Bar-Matthews M. (2000) D / H ratios of fluid inclusions of Soreq cave (Israel) speleothems as a guide to the Eastern Mediterranean Meteoric Line relationships in the last 120 ky. Chem. Geol., 166: 183-191. [3] Matthews A., Zhu X.-K. & O'Nions K. (2001) Kinetic iron stable isotope fractionation between iron (-II) and (-III) complexes in solution. Earth Planet. Sci. Lett., 192: 81-92. [46] Matucha M., Jockisch W., Verner P. & Anders G. (1991) Isotope effect in gas-liquid-chromatography of labelled compounds. J. Chromatogr., 588: 251-258. [8] Matucha M. (1995). Isotope effects (IEs) in gas chromatography (GC) of labelled compounds (LCs). In: Synthesis and applications of isotopically labelled compounds. J. Allen (ed.). John Wiley & Sons Ltd.: 489-494. [8] Maubert C., Gu6rin C., Mabon F. & Martin G. J. (1988) D6termination de l'origine de la vanilline par analyse multidimensionnelle du fractionnement isotopique naturel sp6cifique de l'hydrog6ne. Analusis, 16(8): 434-439. [5] Mauersberger K. (1977) Mass spectrometer beam system for applications in the stratosphere. Rev. Sci. Instrum., 48: 1169-1173. [18] Mauersberger K. & Finstad R. (1979) Further laboratory studies and stratospheric flight of a mass spectrometer beam system. Rev. Sci. Instrum., 50: 1612-1617. [18] Mauersberger K. (1981) Measurement of heavy ozone in the stratosphere. Geophys. Res. Lett., 8(8): 935937.[18] Mauersberger K. (1987) Ozone isotope measurements in the stratosphere. Geophys. Res. Lett., 14(1): 8083.[18] Mauersberger K., Morton J., Schueler B., Stehr J. & Anderson S. M. (1993) Multi-isotope study of ozone: Implications for the heavy ozone anomaly. Geophys. Res. Lett., 20: 1031-1034. [18] Mauersberger K., Erbacher B., Krankowski D., Giinther J. & Nickel R. (1999) Ozone isotope enrichment: Isotopomer-specific rate coefficients. Science, 283: 370-372. [18] Mauersberger K., L~immerzahl P. & Krankowski D. (2001) Stratospheric ozone isotope enrichments revisited. Geophys. Res. Lett., 28(16): 3155-3158. [18] May W., Parris R., Beck C., Fassett J., Greenberg R., Guenther F., Kramer G., Wise S., Gills T., Colbert J., Gettings R. & MacDonald B. (2000) Definitions of terms and modes used at NIST for value-assignment of reference materials for chemical measurements. NIST, Special Publ. 260-136:16 p. [43] Mayer B. (1993) Untersuchungen zur Isotopengeochemie des Schwefels in Waldb6den und neu gebildetem Grundwasser unter Wald. GSF-Forschungszentrum ffir Umwelt und Gesundheit, Neuherberg, Germany. GSF-Bericht Rep. 2/93. [26] Mayer B., Feger K. H., Giesemann A. & J~iger H.-J. (1995) Interpretation of sulfur cycling in two catchments in the Black Forest (Germany) using stable sulfur and oxygen isotope data. Biogeochemistry, 30: 31-58. [26] Mayer B. & Krouse H. R. (1996) Prospects and limitations of an isotope tracer technique for understanding sulfur cycling in forested and agro-ecosystems. Isotopes Environ. Health Stud., 32: 191-201. [26] Mayer B., Bollwerk S. M., Mansfeldt T., Hiitter B. & Veizer J. (2001) The oxygen isotope composition of nitrate generated by nitrification in acid forest floors. Geochim. Cosmochim. Acta, 65: 2743-2756. [46] Mayer B., Boyer E., Goodale C., Jawoarski N., van Bremen N., Howarth R. W., Seitzinger S., Billen G., Lajtha K., Nadelhoffer K., Van Dam D., Hetling L., Nosil M., Paustian K. & Alexander R. (2002) Sources of nitrate in rivers draining 16 major watersheds in the northeastern US: Isotopic Constraints. Biogeochem., 57/58: 86-90. [15] Maynard D. G. & Addison P. A. (1985) Extraction and colorimetric determination of elemental sulfur in organic horizons of forest soils. Can. J. Soil Sci., 65: 811-813. [26] Mayne Y. (1957) Natural variation in the nitrogen isotope ratio in igneous rocks. Geochim. Cosmochim. Acta, 12: 185-189. [16]

References

1129

Mazany T., Lerman J. C. & Long A. (1980) Carbon-13 in tree-ring cellulose as an indicator of past climates. Nature, 287: 432-435. [24] McArdle N. C. & Liss P. S. (1995) Isotopes and atmospheric sulfur. Atmos. Environ., 29(18): 2553-2556. [26] McArthur J. M., Coleman M. L. & Bremner J. M. (1980) Carbon and oxygen isotopic composition of structural carbonate in sedimentary francolite. J. Geol. Soc. London, 137: 669-673. [22] McArthur J. M. & Herczeg A. (1990) Diagenetic stability of the isotopic composition of phosphateoxygen: palaeoenvironmental implications. In: Phosphorite research and development. A. J. G. Notholt & I. Jarvis (eds.). Geol. Soc. Am., Vol. 52: 119-124. [22]. McBain J. W. (1932) The sorption of gases and vapors by solids. Rutledge and Sons, London. [36] McCarroll D. & Pawellek F. (1998) Stable carbon isotope ratios of latewood cellulose in Pinus sylvestris from northern Finland: variability and signal strength. The Holocene, 8: 693-702. [24] McCarthy M. D. B., Newton R. J. & Bottrell S. H. (1998) Oxygen isotopic compositions of sulphate from coals: implications for primary sulphate sources and secondary weathering processes. Fuel, 77: 677682. [26] McConnaughey T. A., Burdett J., Whelan J. F. & Paull C. K. (1997) Carbon isotopes in biological carbonates: Respiration and photosynthesis. Geochim. Cosmochim. Acta, 61: 611-622. [46] McConville C., Kalin R. M. & Flood D. (1999) Direct equilibration of soil water for 6180 analysis and its application to tracer studies. Rapid Commun. Mass Spectrom., 13: 1339-1345. [1, 2] McCord B. R. & Taylor J. W. (1986) Computer-controlled detection system for high precision isotope ratio measurements. Anal. Chem., 58: 2589-2591. [39] McCrea J. M. (1950) On the isotopic chemistry of carbonates and a paleotemperature scale. J. Chem. Phys., 18: 849-857. [10, 12, 40, 46] McCready R. G. L., Grinenko V. A. & Krouse H. R. (1980) Sulfur isotope fractionation by Proteus vulgaris and Salmonella heidelberg during the reduction of thiosulfate. Can. J. Microbiol., 26: 1173-1177. [26] McIlvane T. C. (1921) A buffer solution for calorimetric comparison. J. Biol. Chem., 49: 183-186. [28] McInteer B. B. & Montoya J. G. (1981) Automation of a mass spectrometer for nitrogen isotope analysis. In: Recent developments in mass spectrometry in biochemistry, medicine and environmental research. A. Frigerio (ed.). Elsevier Scientific PubI., Amsterdam, the Netherlands. Vol. 7: 343-349. [15] McKeegan K. D., Walker R. M. & Zinner E. (1985) Ion microprobe isotopic measurements of individual interplanetary dust particles. Geochim. Cosmochim. Acta, 49: 1971-1987. [25] McKeegan K. D. (1987) Ion microprobe measurements of H, C, O, Mg, and Si isotopic abundances in individual interplanetary dust particles. PhD Thesis (unpublished), Washington University, St. Louis, Missouri, USA: 187 p. [30] McKeegan K. D., Leshin L. A., Russell S. S. & MacPherson G. J. (1998) Oxygen isotopic abundances in calcium-aluminum-rich inclusions from ordinary chondrites: Implications for nebular heterogeneity. Science, 280: 414-418. [30] McKeegan K. D., Chaussidon M. & Robert F. (2000) Incorporation of short-lived 10Be in a calciumaluminum-rich inclusion from the Allende meteorite. Nature, 289: 1334-1337. [30] McKibben M. A. & Eldridge C. S. (1989) Sulfur isotopic variations among minerals and aqueous species in the Salton Sea geothermal system: A SHRIMP ion microprobe and conventional study of active ore genesis in a sediment-hosted environment. Am. J. Sci., 289: 661-707. [26] McKibben M. A. & Eldridge C. S. (1995) Microscopic sulfur isotope variations in ore minerals from the Viburnum Trend, Southeast Missouri: A SHRIMP study. Econ. Geol., 90(2): 228-245. [26] McKibben M. A. & Riciputi L. R. (1998) Sulfur isotopic ratios determined by ion microprobe. In: Applications of microanalytical techniques to understanding mineralizing processes. M. A. McKibben, W. C. Shanks III & W. I. Ridley (eds.). Society of Economic Geologists, Littleton, CO, USA. Reviews in Econ. Geol., Vol. 7: 121-139. [20, 26] McKibben M. A.& Shanks W. C. (1998) Applications of microanalytical techniques to understanding mineralizing processes. Reviews in economic geology. C. M. Lesher (series ed). Pubco, Vol. 7:263 p.

[30]

1130

References

McKinney C. R., McCrea J. M., Epstein S., Allen H. A. & Urey H. C. (1950) Improvements in mass spectrometers for the measurement of small differences in isotopic abundance ratios. Rev. Sci. Instrum., 21: 724-730. [28, 38, 40] McLelland J. W., Valiela I. & Michener R. H. (1997) Nitrogen-stable isotope signatures in estuarine food webs: A record of increasing urbanization in coastal watersheds. Limnol. Oceanogr., 42: 930-937. [9] McMahon P. B. & Chapelle F. H. (1991) Geochemistry of dissolved inorganic carbon in a Coastal Plain aquifer. 2. Modelling carbon sources, sinks and 613C evolution. J. Hydrol., 127: 109-135. [10] McMahon P. B. & Bohlke J. K. (1996) Denitrification and mixing in a stream-aquifer system; effects on nitrate loading to surface water. J. Hydrol., 186(1-4): 105-128. [29] McMahon A. W., Gilmour J. D., Hernandez M. B. & Rateitzak M. (1998) A RIS-TOF instrument for the measurement of ultratrace quantities of uranium and plutonium. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AIP Press Woodbury, AIP Conference Proceedings 454: 269-274. [33] McManus J. B., Zahniser M. S., Nelson D. D., Williams L. R. & Kolb C. E. (2002) Infrared laser spectrometer with balanced absorption for measurement of isotopic ratios of carbon gases. Spectrochim. Acta, A58: 2465-2480. [34] McMullen C. C., Cragg C. B. & Thode H. G. (1961) Absolute ratios of 11B/10B in Searles Lake borax. Geochim. Cosmochim. Acta, 23: 147-149. [7] McNeal J. M. & Balistrieri L. S. (1989) Geochemistry and occurence of selenium; an overview. In: Selenium in agriculture and the environment. L. W. Jacobs (ed.). Soil Science Society of America: 1-13. [29] Mehra O. P. & Jackson M. L. (1960) Iron oxide removal from soils and clays by a dithionite-citrate system buffered by sodium bicarbonate. 7th Natl. Conf. on Clays and Clay Minerals. A. Swineford (ed.). Pergamon Press, Washington DC: 317-327. [2] Mei Y. & Harrison W. W. (1991) Getters as plasma reagents in glow discharge mass spectrometry. Spectrochim. Acta, B 46: 175-182. [35] Meier A. L. (1982) Determination of lithium isotopes at natural abundance levels by atomic absorption spectrometry. Anal. Chem., 54: 2158-2161. [6] Meier A. & Notholt J. (1996) Determination of the isotopic abundances of heavy 03 as observed in arctic ground-based FTIR-spectra. Geophys. Res. Lett., 23: 551-554. [18] Meier L. P. & Menegatti A. P. (1997) A new, effcient, one-step method for the removal of organic matter from clay-containing sediments. Clay Minerals, 32: 557-563. [2] Meier-Augenstein W., Brand W., Hoffmann G. F. & Rating D. (1994a) Bridging the information gap between isotope ratio mass spectrometry and conventional mass spectrometry. Biol. Mass Spectrom., 23: 376-378. [8] Meier-Augenstein W., Hess A., Hoffmann G. F. & Rating D. (1994b) Evaluation of water removal and memory effect in 13CO2 breath tests by isotope ratio mass-spectrometry. Isot. Environ. Health Stud., 30: 349-358. [36] Meier-Augenstein W. (1995) Online recording of 13C / 12C ratios and mass-spectra in one gas-chromatographic analysis. HRC-J. High Resolut. Chromatogr., 18: 28-32. [8] Meier-Augenstein W., Watt P. W. & Langhans C. D. (1996) Influence of gas-chromatographic parameters on measurment of 13C / 12C isotope ratios by gas-liquid-chromatography combustion isotope ratio mass-spectrometry. I. J. Chromatogr., A 752: 233-241. [8] Meier-Augenstein W. (1997a) The chromatographic side of isotope ratio mass spectrometry: Pitfalls and answers. LC/GC, 15: 244-252. [8] Meier-Augenstein W. (1997b) A reference gas inlet module for internal isotopic calibration in high precision gas chromatography/combustion-isotope ratio mass spectrometry. Rapid Comm. Mass Spectrom., 11: 1775-1780. [8] Meier-Augenstein W. (1999) Applied gas chromatography coupled to isotope ratio mass spectrometry. J. Chromatography, A. 842: 351-371. [8] Meier-Augenstein W. (2002) Stable isotope analysis of fatty acids by gas chromatography-isotope ratio mass spectrometry. Anal. Chem. Acta, 465: 63-79. [8]

References

1131

Meijer H. A. J. (1995) The isotopic composition of the Groningen GS-19 and GS-20 pure CO2 standards. In: Reference and intercomparison materials for stable isotopes of light elements. IAEA-TECDOC825, IAEA, Vienna: 81-83. [43] Meijer H. A. J. & Li W. J. (1998) The use of electrolysis for accurate 6170 and 6180 isotope measurements in water. Isot. Environ. Health Stud., 34(1): 349-369. [1, 43, 44, 45] Meijer H. A. J., Neubert R. E. M. & Visser G. H. (2000) Cross contamination in dual inlet isotope ratio mass spectrometers. Int. J. Mass Spectrom., 198: 45-61. [26, 34, 43] Melander L. & Saunders W. (1980) Reaction rates of isotopic molecules. J. Wiley-Interscience, New York.

[8]

Menegatti A. P., Friih-Green G. L. & Stille P. (1999) Removal of organic matter by disodium peroxodisulphate: effects on mineral structure, chemical composition and physicochemical properties of some clay minerals. Clay Minerals, 34: 247-257. [2] Merritt D. A., Brand W. A. & Hayes J. M. (1994) Isotope-ratio-monitoring gas chromatography-mass spectrometry: methods for isotopic calibration. Org. Geochem., 21: 573-583. [8, 38] Merritt D. A. & Hayes J. M. (1994) Factors controlling precision and accuracy in isotope-ratio-monitoring mass spectrometry. Anal. Chem., 66(14): 2336-2347. [8, 14, 20, 38] Merritt D. A., Freeman K. H., Ricci M. P., Studley S. A. & Hayes J. M. (1995a) Performance and optimization of a combustion interface for isotope ratio monitoring gas chromatography/mass spectrometry. Anal. Chem., 67: 2461-2473. [8] Merritt D. A., Hayes J. M. & DesMarais D. J. (1995b) Carbon isotopic analysis of atmospheric methane by isotope-ratio-monitoring gas-chromatography mass-spectrometry. J. Geophys. Res., 100(D1): 1317-1326. [8, 14] Mestres M., Busto O. & Guasch J. (2000) Analysis of organic sulfur compounds in wine aroma. J. Chromatogr., A 881: 569-581. [26] Metges C. C., Petzke K.-J. & Hennig J. (1996) Gas chromatography/combustion/isotope ratio mass spectrometric comparison of N-acetyl- and N-pivaloyl amino acid esters to measure 15N isotopic abundances in physiological samples: a pilot study on amino acid synthesis in the upper gastrointestinal tract of minipigs. J. Mass Spetrom., 31: 367-376. [8, 9] Metges C. C. & Petzke K. J. (1997) Measurement of 15N / 14N isotopic composition in individual plasma free amiono acids of human adults at natural abundance by gas chromatography-combustion isotope ratio mass spectrometry. Anal. Biochem., 247: 158-164. [8] Metges C. C. & Petzke K. J. (1999) The use of GC-C-IRMS for the analysis of stable isotope enrichment in nitrogenous compounds. In: Methods for investigation of amino acid and protein metabolism. A. E. E1-Khoury (ed.). CRC Press, Boca Raton, Florida: 121-134. [8] Meyer M., Lancoursi6re J., Nahon L., Gisselbrecht M., Morin P. & Larzilli6re M. (1996) Resonant 2photon-ionization of Xe. Resonance Ionization Spectroscopy 1996, Eighth International Symposium. N Winograd & J.E. Parks (eds.). AIP Press Woodbury, NY, AIP Conference Proceedings 388: 357-360. [33] Michalski G. M. & Thiemens M. (2000) Mass independent fractionation in nitrate aerosols. EOS, Trans. AGU, 81:F120 (Abstract). [15] Michalski G., Savarino J., BShlke J. K. & Thiemens M. (2002) Determination of the total oxygen isotopic composition of nitrate and the calibration of a A170 nitrate reference material. Anal. Chem., 74: 4989-4993. [15] Michener R. H. & Schell D. M. (1994) Stable isotope ratios as tracers in marine aquatic food webs. In: Stable isotopes in ecology and environmental science. K. Lajtha & R. H. Michener (ed.). Blackwetl Scientic Publications, Oxford: 138-157. [9] Michiels E. & De Bi6vre P. (1983) Absolute isotopic composition and the atomic weight of a natural sample of lithium. Int. J. Mass Spectrom. Ion Phys., 49: 265-274. [6] Milatz J. M., Kluyver J. C. & Hardebol J. (1951) Determination of isotope ratios, e.g. in tracer work, by infrared absorption method. J. Chem. Phys., 19: 887-888. [34] Miller R. L., Suits A. G., Houston P. L., Toumi R., Mack J. A. & Wodtke A. M. (1994) The "ozone decifit" problem: 02 (X, ve 26) + O(3P) from 226-nm ozone photodissociation. Science, 265: 1831-1838. [18]

1132

References

Miller M. & Pillinger C. T. (1997) An appraisal of stepped heating release of fluid inclusion CO2 for isotopic analysis: A preliminary to 613C characterisation of carbonaceous vesicles at the nanomole level. Geochim. Cosmochim. Acta, 61: 193-205. [3, 13] Miller M. F., Franchi I. A., Sexton A. S. & Pillinger C. T. (1999) High precision 6170 isotope measurements of oxygen from silicates and other oxides: Method and applications. Rapid Commun. Mass Spectrom., 13: 1211-1217. [20, 36] Miller J. B., Mack K. A., Dissly R., White J. W. C., Dlugokencky E. J. & Tans P. P. (2002) Development of analytical methods and measurements of 13C/12C in atmospheric CH4 from the NOAA-CMDL global air sampling network. J. Geophys. Res., 107: doi 10.1029/2001JD000630. [14] Mills G. A. & Urey H. C. (1939) Oxygen exchange between carbon dioxide, bicarbonate ion, carbonate ion and water. J. Am. Chem. Soc., 61: 534. [1] Mills A. & Urey H. C. (1940) The kinetics of isotopic exchange between carbon dioxide, bicarbonate ion, carbonate ion, and water. J. Am. Chem. Soc., 62: 1019-1026. [1] Milton D. M. P. & Hutton R. C. (1993) Investigations into the suitability of using a secondary cathode to analyse glass using glow discharge mass spectrometry. Spectrochim. Acta, B 48: 39-52. [35] Minagawa M., Winter D. A. & Kaplan I. R. (1984) Comparison of Kjeldahl and conmbustion method for measurement of nitrogen isotope ratios in organic matter. Anal. Chem., 56: 1859-1861. [16] Mineev S. D. & Grinenko V. A. (1996) The pressure influence on hydrogen isotopes fractionation in the serpentine-water system. J. Conf. Abstr., 1:404 (Abstract). [46] Mitchell B. D. & Smith B. F. L. (1974) The removal of organic matter from soil extracts by bromine oxidation. J. Soil Sci., 25: 239-241. [2] Mitchell S. (1996) Biological Interactions of Sulfur Compounds. Taylor & Francis. [26] Miyajima T., Yamada Y. & Hanba Y. T. (1995) Determinating the stable isotope ratio of total dissolved inorganic carbon in lake water by GC/C/IRMS. Limnol. Oceanogr., 40: 994-1000. [10] Miyake Y., Shimomura K., Mills A. P. Jr., Marangos J. P. & Nagamine K. (1998) Ultra-sensitive detection of hydrogen isotopes by the Lyman-a RIS. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AIP Press Woodbury, NY. AIP Conference Proceedings 454: 47-52. [33] Miyazaki T., Wada E. & Hattori A. (1973) Capacities of shallow waters of Sagami Bay for oxidation and reduction of inorganic nitrogen. Deep-Sea Res., 20: 571-577. [15] Mizota C., Domon Y. & Yoshida N. (1992) Oxygen isotope composition of natural phosphates from volcanic ash soils of the Great Rift Valley of Africa and east Java, Indonesia. Geoderma, 5: 111-123. [2] Mizota C. (1996) Pedogeochemistry of hydrogen and oxygen isotope ratios. In: Mass spectrometry of soils. T. W. Boutton & S.-I. Yamasaki (eds.). Marcel Dekker Inc.: 155-176. [2] Mizutani Y. & Rafter T. A. (1969) Oxygen isotopic composition of sulphates - Part 5. Isotopic composition of sulphate in rain water, Gracefield, New Zealand. N. Z. Jl. Sci., 12: 69-80. [26] Moers M. E. C., Jones D. M., Eakin P. A., Fallick A. E., Griffths H. & Larter S. R. (1993) Carbohydrate diagenesis in hypersaline environments: application of GC-IRMS to the stable isotope analysis of derivatives saccharides from surficial and buried sediments. Org. Geochem., 20: 927-933. [9] Moerth C.-M. & Torssander P. (1995) Sulfur and oxygen isotope ratios in sulfate during an acidification reversal study at Lake Garsjon, Western Sweden. Water Air Soil Pollut., 79: 261-278. [26] Mohler F. L. (1960) Isotopic abundance ratios reported for reference samples stocked at the National Bureau of Standards. National Bureau of Standards, Washington. Tech. Note No. 51:8 p. [40] Mojzsis S. J., Arrhenius G. A., McKeegan K. D., Harrison T. M., Nutman A. P. & Friend R. L. (1996) Evidence of life on Earth before 3,800 million years ago. Nature, 384: 55-59. [30] Mojzsis S. J., Harrison T. M. & McKeegan K. D. (2001) Oxygen isotope evidence from ancient zircons for liquid water at the Earth's surface 4,300 Myr ago. Nature, 409: 178-181. [30] Mojzsis S. J., Coath C. D., Greenwood J. P., McKeegan K. D. & Harrison T. M. (2003) Mass-independent isotope effects in Archean (2.5 - 3.8 Ga) sedimentary sulfides determined by ion microprobe analysis. Geochim. Cosmochim. Acta, 67: 1635-1658. [30] Molini-Velsko C. A. (1983) Isotopic composition of silicon in meteorites. Ph.D. Thesis, University of Chicago, Chicago, USA: 174p. [36]

References

113 3

Molini-Velsko C., Mayeda T. K. & Clayton R. N. (1986) Isotopic composition of silicon in meteorites. Geochim. Cosmochim. Acta 50: 2719-2726. [25] Molini-Velsko C., Mayeda T. K. & Clayton R. N. (1987) Silicon isotope systematics during distillation. Lunar Planet. Sci., 18: 657-658. [46] Monaghan J. M., Scrimgeour C. M., Stein W. M., Zhao F. J. & Evans E. J. (1999) Sulphur accumulation and re-distribution in wheat (Triticum aestivum): a study using stable sulphur isotope ratios as a tracer system. Plant Cell Environ., 22, 831-839. [26] Moncaster S. J. & Bottrell S. H. (1991) Extraction of low-level sulphide from groundwaters for sulphur isotope analysis. Chem. Geol. (Isot. Geosci. Sect.), 94: 79-82. [26] Monetti A., Reniero F. & Versini G. (1994) Interregional and interannual variability of ethanol sitespecific deuterium content in Italian wines. Z. Lebensm. Unters. Forsch., 199: 311-316. [5] Monetti A., Versini G., Dalpiaz G. & Reniero F. (1996) Sugar adulterations control in concentrated rectified grape musts by finite mixture distribution analysis of the myo- and scyllo-inositol content and the D / H methyl ratio of fermentative ethanol. J. Agric. Food Chem., 44: 2194-2201. [5] Monin J. C., Pelet R. & Fevrier A. (1978) Geochemical analysis of organic matter extracted from sedimentary rocks. Part 4. Extraction of small quantities of rock. Revue Inst. Franr Pdtrole, 33: 223-240.

[2]

Monster J. (1972) Homogeneity of sulfur and carbon isotope ratios S34 /S 32 and C 13 / C 12 in petroleum. AAPG Bull., 54: 941-949. [26] Montaser A., Minnich M. G., McLean J. A., Liu H., Caruso J. A. & McLeod C. W. (1998) Sample introduction in ICPMS. In: Inductively coupled plasma mass spectrometry. A. Montaser (ed.). WileyVCH: 83-264 [31] Mook W. G. (1968) Geochemistry of the stable carbon and oxygen isotopes of natural waters in the Netherlands. PhD Thesis, University of Groningen. [10] Mook W. G. & Grootes P. M. (1973) The measuring procedure and corrections for the high precision mass spectrometric analysis of isotopic abundance ratios, especially referring to carbon, oxygen and nitrogen. Int. J. Mass Spectrom. Ion Phys., 12: 273-298. [23, 43, 45] Mook W. G., Bommerson J. C. & Staverman W. H. (1974) Carbon isotope fractionation between dissolved bicarbonate and gaseous carbon dioxide. Earth Planet. Sci. Lett., 22: 169-176. [46] Mook W. G. & Jongsma J. (1987) Measurement of the N20 correction for 13C / 12C ratios of atmospheric CO2 by removal of N20. Tellus, 39B: 96-99. [14] Mook W. G. (2000) Environmental isotopes in the hydrological cycle: Principles and applications, Volume I: Introduction: Theory, Methods, Review. International Hydrological Programme IHP-V. UNESCO, Paris. Technical Documents in Hydrology, No.39, Vol.l: 49-51. [40] Mooney-Slater R. C. L. (1962) Polymorphic forms of bismuth phosphate. Z. Kristallogr., 117: 371-385. [22] Moore G. S. M. & Rose H. E. (1973) The structure of powdered quartz. Nature, 242: 187-190. [3] Moore H. (1977) The isotopic composition of ammonia, nitrogen dioxide and nitrate in the atmosphere. Atmos. Environ., 21: 1239-1243. [15] Moore P. B. & Shen J. (1983) An X-ray structural study of cacoxenite, a mineral phosphate. Nature, 306: 356-358. [36] Moran K. K., Mulvaney R. L. & Khan S. A. (2002) A Technique to facilitate diffusions for nitrogen-isotope analysis by direct combustion. Soil Sci. Soc. Am. J., 66: 1008-1011. [15] Moreau-Le Golvan Y., Michelot J.-L. & Boisson J.-Y. (1997) Stable isotope contents of porewater in a claystone formation (Tournemire, France): assessment of the extraction technique and preliminary results. Appl. Geochem., 12: 739-745. [2] Morgan M. D. & Mandernack K. W. (1996) Biogeochemistry of sulfur in wetland peat following 3.5 y of artificial acidification (HUMEX). Environ. Int., 22: 605-610. [26] Moriguti T. & Nakamura E. (1993) Precise lithium isotope analysis by thermal ionization mass spectrometry using lithium phosphate as an ion source. Proc. Japan Acad. Sci.: 69B: 123-128. [6] Moriguti T. & Nakamura E. (1998a) High-yield lithium separation and precise isotopic analysis for natural rock and aqueous samples. Chem. Geol., 145: 91-104. [6, 31] Moriguti T. & Nakamura E. (1998b) Across-arc variation of Li isotopes in lavas and implications for crust/mantle recycling at subduction zones. Earth Planet. Sci. Lett., 163: 167-174. [6]

1134

References

Morrison J., Brockwell T., Merren T., Fourel F. & Phillips A. M. (2001) On-line high-precision stable hydrogen isotopic analyses on nanolitre water samples. Anal. Chem., 73: 3570-3575. [1] Morse A. D., Wright I. P. & Pillinger C. T. (1993) An investigation into the cause of memory effects associated with the conversion of H20 to H2 for D / H measurement. Chem. Geol. (Isot. Geosc. Sect.), 107: 147-158. [1] Mortimer R. J. G. & Coleman M. L. (1997) Microbial influence on the oxygen isotopic composition of diagenetic siderite. Geochim. Cosmochim. Acta, 61: 1705-1711. [46] Morton J., Schueler B. & Mauersberger K. (1989) Oxygen fractionation of ozone isotopes 4803 through 5403. Chem. Phys. Lett., 154: 143-145. [18] Morton J., Barnes J., Schueler B. & Mauersberger K. (1990) Laboratory studies of heavy ozone. J. Geophys. Res., 95: 901-907. [18] Morton R. D. & Catanzaro E. J. (1964) Stable chlorine isotope abundances in apatites from Odegardens Verk, Norway. Norsk Geol. Tiddsk., 44: 307-313. [28] Morville J., Chenevier M., Kachanov A. A. & Romanini D. (2001) Trace gas detection with DFB lasers and cavity ring-down spectroscopy. Proceedings SPIE 4485 (Optical spectroscopic techniques, remote sensing, and instrumentation for atmospheric and space research IV): 236-243. [34] Morville J., Romanini D., Kachanov A. A. & Chenevier, M. (2003) Two schemes for trace gas detection using cavity ring down spectroscopy. Appl. Phys., B (submitted). [34] Mosandl A. (1995) Enantioselective capillary gas-chromatography and stable-isotope ratio massspectrometry in the authenticity control of flavors and essential oils. Food Rev. Int., 11: 597-664. [8] Mosandl A., Braunsdorf R., Bruche G., Dietrich A., Hener U., Karl V., Kopke T., Kreis P., Lohmann D. & Maas B. (1995) New methods to assess authenticity of natural flavors and essential oils. ACS Symp. Ser., 596: 94-112. [8] Mosandl A. (1997) Progress in the authenticity assessment of wines and spirits. Analusis, 25: M31-M38.

[81

Motz J. E., Edwards T. W. D. & Buhay W. M. (1997) Use of nickel-tube pyrolysis for hydrogen-isotope analysis of water and other compounds. Chem. Geol., 140: 145-149. [1] Moum J. & Rosenqvist I. T. (1958) Hydrogen (protium)-deuterium exchange in clays. Geochim. Cosmochim. Acta, 14: 250-252. [2] Moussa I., Naulet N., Martin M. L. & Martin G. J. (1990) A site-specific and multielement approach to the determination of liquid-vapor isotope fractionation parameters. The case of alcohols. J. Phys. Chem., 94(21): 8303-8309. [5] Mroz E. J. (1993) Deuteromethanes: potential fingerprints of the sources of atmospheric methane. Chemosphere, 26: 45-53. [34] Muehlenbachs K. & Kushiro I. (1974) Oxygen isotope exchange and equilibrium of silicates with CO2 and 02. Carnegie Instn. Washington, Year Book, 74: 232-240. [46] Mulitza S., Wolff T., P~itzold J., Hale W. & Wefer G. (1998) Temperature sensitivity of planktic foraminifera and its influence on the oxygen isotope record. Marine Micropaleo., 33: 223-240. [46] Mullane M. V., Waterhouse J. S. & Switsur v. R. (1988) On the development of a novel method for the determination of stable oxygen isotope ratios in cellulose. Int. J. Appl. Radiat. Isot., 39: 1029-1035. [23] Mullis J., Dubessy J., Poty B. & O'Neil J. R. (1994) Fluid regimes during late stages of a continental collision: physical, chemical and stable isotope measurements of fluid inclusions in fissure quartz from a geotraverse through the Central Alps, Switzerland. Geochim. Cosmochim. Acta, 58: 2239-2267.

[3]

Mulvaney R. L. (1986) Comparison of procedures for reducing cross-contamination during steam distillations in nitrogen-15 tracer research. Soil Sci. Am. J., 50: 92-96. [15] Mulvaney R. L., Fohringer C. L., Bojan V. J., Michlik M. M. & Herzog L. F. (1990) A commercial system for automated nitrogen isotope-ratio analysis by the Rittenberg technique. Rev. Sci. Instrum., 61: 897-903. [15] Mulvaney R. L. & Liu Y. P. (1991) Refinement and evaluation of an automated mass spectrometer for nitrogen isotope analysis by the Rittenberg technique. J. Autom. Chem., 13: 273-280. [15] Mulvaney R. L. (1993) Mass spectrometry. In: Nitrogen isotope techniques. R. Knowles & T. H. Blackburn (eds.). Academic Press, San Diego, CA: 11-57. [15]

References

113 5

Mulvaney R. L., Ashraf M., Azam F., Brooks P. D. & Herman D.J. (1994) Evaluation of methods for nitrogen-15 analysis of inorganic nitrogen in soil extracts: I. Steam-distillation methods. Commun. Soil Sci. Plant Anal., 25: 2187-2200. [15] Mulvaney R. L. (1996) Nitrogen- Inorganic forms. In: Methods of soil analysis. Part3. Chemical Methods. D. L. Sparks et al. (eds.). Soil Science Society of America and American society of Agronomy, Madison, Wisconsin: 1123-1184. [15] Mulvaney R. L., Khan S. A., Stevens W. B. & Mulvaney C. S. (1997) Improved diffusion methods for determination of inorganic nitrogen in soil extracts and water. Biol. Fertil. Soils, 24: 413-420. [15] Mulvaney R. L. & Khan S. A. (1999) Use of diffusion to determine inorganic nitrogen in a complex organic matrix. Soil Sci. Soc. Am. J., 63: 240-246. [15] Mulvaney R. L. & Khan S. A. (2001) Diffusion methods to determine different forms of nitrogen in soil hydrolysates. Soil Sci. Soc. Am. J., 65: 1284-1292. [15] Mulvaney R. L., Khan S. A., Hoeft R. G. & Brown H. M. (2001) A soil organic nitrogen fraction that reduces the need for nitrogen fertilization. Soil Sci. Soc. Am. J., 65: 1164-1172. [15] Murnick D. E. & Peer B. J. (1994) Laser-based analysis of carbon isotope ratios. Science, 263: 945-947. [34] Murnick D. E., Colgan M. J. & Stoneback D. N. (1998) Optogalvanic effect isotope ratio analysis in carbon dioxide. In: Synthesis and applications of isotopes and isotopically labeled compounds. J. R. Heys & D. G. Melillo (eds.). John Wiley and Sons, Chichester: 111-115. [34] Murnick D. E. (2001) Laser assisted ratio analysis - An alternative to GC/IRMS for CO2. In: New approaches for stable isotope ratio measurements. Proceedings of an Advisory Group meeting held in Vienna, 20-23 September 1999. IAEA-TECDOC-1247, IAEA, Vienna: 15-22. [34, 40] Murphy B. F. (1947) The temperature variation of thermal diffusion factors for binary mixtures of H, D and He. Phys. Rev., 72: 834-837. [38] Murty S. V. S. & Goswami J. N. (1992) Nitrogen, noble gases and nuclear tracks in lunar meteorites MAC 88104/105. Proc. Lunar Planet. Sci., 22: 225-237. [17] Musashi M., Markl G. & Kreulen R. (1998) Stable chlorine-isotope analysis of rock samples: new aspects of chlorine extraction. Anal. Chem. Acta, 362: 261-269. [28] Miiller P., Bushaw B. A., Blaum K., Ntirtersh~iuser W., Trautmann N. & Wendt K. (1998) Ultratrace determination of hte long-lived isotope 41Ca by narrowband CW-RIMS. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AIP Press Woodbury, NY. AIP Conference Proceedings 454: 73-78. [33] Myers W. R. & DeLong W. B. (1984) Fluorine corrosion: High-temperature attack on metals by fluorine and hydrogen fluoride. Chem. Eng. Prog., 44: 359-362. [20] Myhre G., Highwood E. J., Shine K. P. & Stordal F. (1998) New estimates of radioactive forcing due to well-mixed greenhouse gases. Geophys. Res. Lett., 25: 2715-2718. [14] Nabelek P. I., O'Neil J. R. & Papike J. J. (1983) Vapor phase exsolution as a controlling factor in hydrogen isotope variation in granitic magmas: the Notch Peak granitic stock, Utah. Earth Planet. Sci. Lett., 66: 137-150. [46] Nadeau S., Pineau F., Javoy M. & Francis-Donald M. (1990) Carbon concentrations and isotopic ratios in fluid-inclusion-bearing upper-mantle xenoliths along the northwestern margin of North America. Chem. Geol., 81: 271-297. [3] Naden J., Kilias S. P., Leng M. J., Cheliotis I. & Sheperd T. J. (2003) Do fluid inclusions preserve 6180 values of hydrothermal fluids in epithermal systems over geological time? Evidence from paleo- and modern geothermal systems, Milos island, Aegean Sea. Chem. Geol., 197: 143-159. [3] Nakai N. & Jensen M. L. (1967) Sources of atmospheric sulfur compounds. Geochem. J., 1: 199-210. [26] Nakamura E., Ishikawa T., Birck J.-L. & All6gre C. J. (1992) Precise boron isotopic analysis of natural rock samples using a boron-mannitol complex. Chem. Geol. (Isot. Geosci. Sect.), 94: 193-204. [40] Nakamura T., Kojima S., Ohta H., Oda H., Ikeda A., Okuno M., Yokota K., Mizutani Y. & Kretschmer W. (1998) Isotopic analyses and cycling of dissloved inorganic carbon at Lake Biwa, central Japan. Radiocarbon, 40: 933-944. [10] Nakano T. & Nakamura E. (2001) Boron isotope geochemistry of metasedimentary rocks and tourmalines in a subduction zone metamorphic suite. Phys. Earth Planet. Int., 127: 233-252. [30]

1136

References

Nakazawa T., Morimoto S., Aoki S. & Tanaka M. (1997) Temporal and spatial variations of the carbon isotopic ratio of atmospheric carbon dioxide in the western Pacific region. J. Geophys. Res., 102: 12711285. [141 Nascimento C., Atekwana E. A. & Krishnamurthy R. V. (1997) Concentrations and isotope ratios of dissolved inorganic carbon in groundwaters of denitrifying environments. Geophys. Res. Lett., 24: 1511-1514. [10] Navon O. & Wasserburg G. J. (1985) Self shielding in 02 - A possible explanation for oxygen isotopic anomalies in meteorites? Earth Planet. Sci. Lett., 73: 1-16. [18] Neal R. H. & Spositom G. (1989) Selenate adsorption on alluvial soils. Soil Sci. Soc. Am. J., 53(1): 70-74. [29] Nealson K. H. (1997) Sediment bacteria: who's there, what are they doing, and what's new? Ann. Rev. Phys. Chem., 25: 403-434. [46] Nehring N. L., Bowen P. A. & Truesdell A. H. (1977) Techniques for the conversion to carbon dioxide of oxygen from dissolved sulfate in thermal waters. Geothermics, 5: 63-66. [20] Nelms S. M., Qu6tel C., Prohaska T., Vogl J. & Taylor P. D. P. (2001) Evaluation of detector dead time calculation models for ICP-MS. J. Anal. At. Spectrom., 16: 333-338. [32, 37] Nelson D. E., Angerbjorn A., Liden K. & Turk I. (1998) Stable isotopes and the metabolism of the European cave bear. Oecologia, 116: 177-181. [9] Nelson S.T. & Dettman D. (2001) Improving hydrogen isotope ratio measurements for on-line chromium reduction systems. Rapid Commun Mass Spectr., 15: 2301-2306. [1] Nesbitt B. E., Muehlenbachs K. & Murowchick J. B. (1989) Genetic implications of stable isotope characteristics of mesothermal Au deposits and related Sb and Hg deposits in the Canadian Cordillera. Econ. Geol., 84: 1489-1506. [3] Nesbitt B. E. & Muehlenbachs K. (1995) Geochemical studies of the origins and effects of synorogenic crustal fluids in the southern Omineca Belt of British Columbia, Canada. Geol. Soc. Am. Bull., 107: 1033-1050. [3] Nesbitt B. E. & Muehlenbachs K. (1997) Paleo-hydrogeology of late proterozoic units of southeastern Canadian Cordillera. Am. J. Sci., 297: 359-392. [3] N e w m a n L., Forrest J. & Manowitz B. (1975) The application of an isotopic ratio technique to a study of the atmospheric oxidation of sulphur dioxide in the plume from coal-fired power plant. Atmos. Environ., 9: 959-968. [26] Newton R. J., Bottrell S. H., Dean S. P., Hatfield D. & Raiswell R. (1995) An evaluation of the use of the chromous chloride reduction method for isotopic analysis of pyrite in rocks and sediment. Chem. Geol. (Isot. Geosci. Sect.), 125: 317-320. [26] Nicolussi G. K., Pellin M. J., Lykke K. R., Trevor J. L., Mencer D. E. & Davis A. M. (1996) Surface analysis by SNMS: Femtosecond laser postionization of sputtered and laser desorbed atoms. Surf. Interface Anal., 24: 363-370. [33] Nicolussi G. K., Davis A. M., Pellin M. J., Lewis R. S., Clayton R. N. & Amari S. (1997a) The isotopic composition of molybdenum in single presolar carbide grains. Lunar and Planetary Science Conference, Ext. Abs. XXVIIh 1021-1022 (Abstract). [33] Nicolussi G. K., Davis A. M., Pellin M. J., Lewis R. S., Clayton R. N. & Amari S. (1997b) S-process zirconium in individual presolar silicon carbide grains. Lunar and Planetary Science Conference, Ext. Abs. XXVIII: 1023-1024 (Abstract). [33] Nicolussi G. K., Pellin M. J., Lewis R. S., Davis A. M., Amari S. & Clayton R. B. (1998) Molybdenum isotopic composiiton of individual presolar silicon carbide grains from the Murchison meteorite. Geochim. Cosmichim. Acta, 62: 1093-1104. [31] Nief G. & Botter R. (1959) Mass spectrometric analysis of simple hydrogen compounds. Adv. Mass Spectrom., 1: 515-525. [1] Nielsen H. (1984) The standard situation for sulfur isotopes. In: Advisory Group Meeting on stable isotope reference samples for geochemical and hydrological investigations, 19-21 Sept., 1983. R. Gonfiantini (ed.). IAEA, Vienna: 73-77. [40] Nier A. O. & Hanson E. E. (1936) A mass-spectrographic analysis of the ions produced in HC1 under electron impact. Phys. Rev., 50: 722-726. [28]

References

1137

Nier A. O. (1940) A mass spectrometer for routine isotope abundance measurements. Rev. Sci. Instrum., 11: 212-216. [38] Nier A. O. (1947) A mass spectrometer for isotope and gas analysis. Rev. Sci. Instrum., 18: 398-411. [28,

38, 40] Nier A. O., Ney E. P. & Inghram M. G. (1947) A null method for the comparison of two ion currents in a mass spectrometer. Rev. Sci. Instrum., 18: 294. [28] Nier A. O. (1950) A redetermination of the relative abundances of the isotopes of carbon, nitrogen, oxygen, argon and potassium. Phys. Rev., 77: 789-793. [40] Nier A. O. (1955) Determination of isotopic masses and abundances by mass spectrometry. Science, 121: 737-744. [28] Niewenhuize J., Maas Y. E. M. & Middelburg J. J. (1994) Rapid analysis of organic carbon and nitrogen in particulate materials. Marine Chem., 45: 217-224. [9] Nirel P. M. V. & Lutz T. M. (1991) Plasma-source mass spectrometry (ICP-MS) application to multielement analysis in environmental samples, a practical evaluation. J. Trace Microprobe Techn., 9: 95106. [28] Nishio Y. & Nakai S. (2002) Accurate and precise lithium isotopic determinations of igneous rock samples using multi-collector inductively coupled plasma mass spectrometry. Anal. Chim. Acta, 456: 271-281. [6] Nishizawa K., Maeda Y., Kawashiro F., Fujii T. & Yamamoto T. (1998) Contribution of nuclear size and shape, nuclear mass, and nuclear spin to enrichment factors of zinc isotopes in a chemical exchange reaction. Sep. Sci. Technol., 33: 2102-2112. [31] NIST Certificate of analysis (1999) Standard reference material 951, boric acid standard. National Institute of Standards and Technology, Gaithersburg, USA: 2 p. [40] NIST Certificate of analysis (2001) Standard reference material 975a, isotopic standard for chlorine. National Institute of Standards and Technology, Gaithersburg, USA: 2 p. [40] Nitz S., Weinreich B. & Drawert F. (1992) Multidimensional gas-chromatography- isotope ratio massspectrometry (Mdgc-irms). A system description and technical requirements. HRC-J. High Resolut. Chromatogr., 15: 387-391. [8] Niu H. & Houk R. S. (1996) Fundamental aspects of ion extraction in inductively coupled plasma mass spectrometry. Spectrochim. Acta, B 51: 779-815. [35] Nomura M., Okamoto M. & Kakihana H. (1973) Determination of boron isotope ratio by the surface ionization method. Shitsuryo Bunzeki, 21:277-281 (in Japanese). [7] Norman A. L. & Krouse H. R. (1992) Stable isotope studies of atmospheric sulphur: comparison of Calgary, Canada and Bermuda. Papers from the 9th World Clean Air Congress: Towards the Year 2000. IU-8.01. [26] Norman D. I. & Sawkins F. J. (1987) Analysis of volatiles in fluid inclusions by mass spectrometry. Chem. Geol., 61: 1-10. [3] Northrop D. A. & Clayton R. N. (1966) Oxygen-isotope fractionations in systems containing dolomite. J. Geology, 74: 174-196. [46] Noto M. & Kusakabe M. (1995) The reactivity of zinc used for preparation of water samples for D / H isotope analysis. Okayama University. Tech Report, Inst. Study Earth's Interior, Ser. B, No. 14 (in Japanese). [1] Novak M., Adamova M. & Milicic J. (2003a) Sulfur metabolism in polluted shagnum peat bogs: a combined 34S-35S-210Pb study. Water Air Soil Pollut., 3: 181-200. [26] Novak M., Buzek F., Harrison A. F., Prechova E., Jackova I. & Fottova D. (2003b) Similarity between C, N and S stable isotope profiles in European spruce forest soils: implications for the use of ~34S as a tracer. Appl. Geochem., 18: 765-779. [26] Ni~rtersh~iuser W., Blaum K., Bushaw B. A., Miiller P., Trautmann N. & Wendt K. (1998) Lineshapes and optical selectivity in double-resonance RIMS measurements. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AIP Press Woodbury, NY. AIP Conference Proceedings 454: 171-176. [33] Norgaard J. V., Valkiers S., Van Nevel L., Vendelbo E., Papadakis I., Br6as O. & Taylor P. D. P. (2002) The international measurement evaluation programme, IMEP-8: carbon and oxygen isotope ratios in C02. Anal. Bioanal. Chem., 374: 1147-1154. [43]

1138

References

Nriagu J. O. & Coker R. D. (1983) Sulphur in sediments chronicles past changes of lake acidification. Nature, 303: 692-694. [26] Nriagu J. O. & Soon u K. (1985) Distribution and isotopic composition of sulfur in lake sediments of Northern Ontario. Geochim. Cosmochim. Acta, 49: 823-834. [26] Nriagu J. O., Holdway D. A. & Coker R. D. (1987) Biogenic sulfur and the acidity of rainfall in remote areas of Canada. Science, 237: 1189-1192. [26] Nriagu J. O., Coker R. D. & Barrie L. A. (1991) Origin of sulphur in Canadian Artic haze from isotope measurements. Nature, 349: 142-145. [26] O'Keefe A. & Deacon D. A. G. (1988) Cavity ring-down optical spectrometer for absorption measurements using pulsed laser sources. Rev. Sci. Instrum., 59: 2544-2551. [34] O'Leary M. H. (1984) Measurement of the isotope fractionation associated with diffusion of carbon dioxide in aqueous solution. J. Phys. Chem., 88: 823-825. [46] O'Malley V. P., Abrajano T. A. Jr. & I-Iellou J. (1994) Determination of the 13C/12C ratios of individual PAH from environmental samples: can PAH sources be apportioned? Org. Geochem., 21: 809-822. [9] O'Neil J. R. & Clayton R. N. (1964) Oxygen isotope thermometry. In: Isotopic and cosmic chemistry. H. Craig, G. J. Miller, and G. J. Wasserburg (eds.). North Holland: 157-168. [29, 46] O'Neil J. R. & Epstein S. (1966a) Oxygen isotope fractionation in the system dolomite-calcite-carbon dioxide. Science, 152: 198-201. [1, 46] O'Neil J. R. & Epstein S. (1966b) A method for oxygen isotope analysis of milligram quantities of water and some of its applications. J. Geophys. Res., 71: 4956-4961. [3, 20] O'Neil J. R. & Taylor H. P. Jr. (1967) The oxygen isotope and cation exchange chemistry of feldspars. Am. Mineralogist, 52: 1414-1437. [46] O'Neil J. R., Clayton R. N & Mayeda T. K. (1969) Oxygen isotope fractionation in divalent metal carbonates. J. Chem. Phys., 51: 5547-5558. [46] O'Neil J. R. & Taylor H. P. Jr. (1969) Oxygen isotope equilibrium between muscovite and water. J. Geophys. Res., 74: 6012-6022. [46] O'Neil J. R. (1986) Theoretical and experimental aspects of isotopic fractionation. In: Stable isotopes in high temperature geological processes. J. W. Valley, H. P. Jr. Taylor & J. R. O'Neil (eds.). Mineral. Soc. America, Reviews in Mineralogy, Vol. 16: 1-40. [46, App B] O'Neil J. R. & Truesdell A. H. (1991) Oxygen isotope fractionation studies of solute-water interactions. In: Stable isotope geochemistry: A tribute to Samuel Epstein. H. P. Jr. Taylor, J. R. O'Neil & I. R. Kaplan (eds.). The Geochemical Society, Spec. Publ. 3: 17-25. [46] O'Neil J. R., Roe L. J., Reinhard E. & Blake R. E. (1994) A rapid and precise method of oxygen isotope analysis of biogenic phosphate. Isr. J. Earth Sci., 43: 203-212. [20, 22] O'Neil J. R., Vennemann T. W. & McKenzie W. F. (2003) Effects of speciation on equilibrium fractionations and rates of oxygen isotope exchange between (PO4)aq and H20. Geochim. Cosmochim. Acta, 67 (Clayton issue): 3135-3144. [46] O'Reilly C., Jenkin G. R. T., Feely M., Alderton D. H. M. & Fallick A. E. (1997) A fluid inclusion and stable isotope study of 200 Ma of fluid evolution in the Galway Granite, Connemara, Ireland. Contrib. Mineral. Petrol., 129: 120-142. [3] Oana S. & Ishikawa H. (1966) Sulfur isotopic fractionation between sulfur and sulfuric acid in the hydrothermal solution of sulfur dioxide. Geochem. J., 1: 45-50. [26] Ohashi S. (1955) Rapid determination of sulfur in organic compounds by the tin (II)-strong phosphoric acid reduction method. Bull. Chem. Soc. Japan, 28(9): 645-649. [26] Ohba T. (1987) 180/160 and D / H determinations for small amounts of water. Geochem. J., 21: 183-186. [1,20] Ohba T., Kazahaya K. & Matsuo S. (1995) Diffusiona1180 loss from inclusion water in a natural hydrothermal quartz from the Kaneuchi tungsten deposit, Japan. Geochim. Cosmochim. Acta, 59: 3039-3047.

[3]

Ohba T. & Hirabayashi J. (1996) Handling of Pt catalyst in H2-H20 equilibration method for D / H measurement of water. Geochem. J., 30: 373-377. [1] Ohba T. & Matsuo S. (1998) Precise determination of hydrogen and oxygen isotope ratios of water in fluid inclusions of quartz and halite. Geochem. J., 22(2): 55-68. [3]

References

1139

Ohkouchi N., Kawamura K., Kajiwara Y., Wada E., Okada M., Kanamatsu T. & Taira A. (1999) Sulfur isotope records around Livello Bonarelli (northern Apennines Italy) black shale at the CenomanianTuronian boundary. Geology, 27: 535-538. [26] Ohmoto H. & Lasaga A. C. (1982) Kinetics of reactions between aqueous sulfates and sulfides in hydrothermal systems. Geochim. Cosmochim. Acta, 46: 1725-1745. [46] Ohmoto H. (1992) Biogeochemistry of sulphur and the mechanisms of sulfide-sulfate minralization in Archean oceans. In: Early organic evolution. M. Schidlowski, S. Golubic, M. M. Kimberley, D. M. McKirdy & P. A. Trudinger (eds.). Springer Verlag, Berlin-Heidelberg: 378-397. [27] Ohmoto H. & Goldhaber M. B. (1997) Sulfur and carbon isotopes. In: Geochemistry of hydrothermal ore deposits. H. L. Barnes (ed.). Wiley & Sons: 570-611. [27] Ohorodnik S. K. & Harrison W. W. (1993) Cryogenic coil for glow discharge sources. Anal. Chem., 65: 2542-2544. [35] Ohsumi T. & Fujino H. (1986) Isotopic exchange technique for preparation of hydrogen gas in mass spectrometric D /H analysis of natural waters. Anal. Sci., 2: 489-490. [1] Oi T., Kawada K., Hosoe M. & Kakihana H. (1991) Fractionation of lithium isotopes in cation-exchange chromatography. Sep. Sci. Tech., 26: 1353-1375. [6] Oi T., Odagiri T. & Nomura M. (1997) Extraction of lithium from GSJ rock reference samples and determination of their lithium isotopic compositions. Anal. Chim. Acta, 340: 221-225. [6] Olive K. A. & Schramm D. N. (1992) Astrophysical 7Li as a product of Big Bang nucleosynthesis and galactic cosmic-ray spallation. Nature, 360: 439-442. [6] Oremland R. S., HoUibaugh J. T., Maest A. S., Presser T. S., Miller L. G. & Culbertson C. W. (1989) Selenate reduction to elemental selenium by anaerobic bacteria in sediments and culture: Biogeochemical significance of a novel, sulfate-independent respiration. Appt. Env. Microbiol., 55(9): 2333-2343. [29] Orlando T. M., Burns A. R., Stechel E. B. & Jennison D. R. (1991) Resonance ionization detection of neutral O(3PJ) produced from stimulated surface processes. J. Vac. Sci. Technol. A - Vac Surf. Films, 9: 1769-1773. [33] Orr W. L. (1974) Changes in sulfur content and isotopic ratios of sulfur during petroleum maturation study of Big Horn Basin Paleozoic oils. Am. Assoc. Pet. Geol. Bull., 58: 2295-2318. [26] Osberghaus O. (1950) Die Isotopenh/iufigkeit des Bors. Massenspektrometrische untersuchung der Elektronenstossprodukte von BF3 und BC13. Z. Physik, 128: 366-377. [28] Ostrom P. H., Colunga-Garcia M. & Gage S. H. (1997) Establishing pathways of energyfow for insect predators using stable isotope ratios: field and laboratory evidence. Oecotogia, 109: 108-113. [9] Ostwald W. (1900) Ober die vermeintliche Isomerie des roten und gelben Quecksilberoxyds und die Oberfl/ichenspannung fester K6rper. Z. Phys. Chem. Stoechiom. Verwandschaftslehre, 34: 495-503. [46] Otsuki A., Ino Y. & Fujii T. (1983) Simultaneous measurements and determinations of stable carbon and nitrogen isotope ratios, and organic carbon and nitrogen contents in biological samples by coupling of a small quadrupole mass spectrometer and modified carbon-nitrogen elemental analyzer. Int. J. Mass Spectrom. Ion Phys., 48: 343-346. [15] Oversby V. M. (1969) Determination of lead isotope compositions in iron meteorites by the double spike technique. Meteoritics, 4(3): 201. [29] Owen H. R. & Schaeffer O. A. (1954) The isotope abundance of chlorine from various sources. J. Am. Chem. Soc., 77: 898-899. [28] Owens N. J. P. & Rees A. P. (1989) Determination of nitrogen-15 at sub-microgram levels of nitrogen using automated continuous-flow isotope ratio mass spectrometry. Analyst, 114: 1655-1657. [45] Pajo L., Tamborini G., Rasmussen G., Mayer K. & Koch L. (2001) A novel isotope analysis of oxygen in uranium oxides: comparison of secondary ion mass spectrometry, glow discharge mass spectrometry, and thermal ionization mass spectrometry. Spectrochim. Acta, B 56: 541-549. [35] Paldus B. A., Harb C. C., Spence T. G., Wilke B., Xie J., Harris J. S. & Zare R. N. (1998) Cavity-locked ring-down spectroscopy. J. Appl. Phys., 83: 3991-3997. [34] Paldus B. A., Spence T. G., Zare R. N., Oomens J., Harren F. J. M., Parker D. H., Gmachl C., Cappasso F., Sivco D. L., Baillargeon J. N., Hutchinson A. L. & Cho A. Y. (1999) Photoacoustic spectroscopy using quantum-cascade lasers. Opt. Lett., 24: 178-180. [34]

1140

References

Palin J. M., Epstein S. & Stolper E. M. (1996) Oxygen isotope partitioning between rhyolitic glass/melt

and CO2: an experimental study at 550-950~ and I bar. Geochim. Cosmochim. Acta, 60: 1963-1973. [46] Palmer D. A. S. & Williams-Jones A. E. (1996) Genesis of the carbonatite-hosted fluorite deposit at Amba Dongar, India; evidence from fluid inclusions, stability isotopes, and whole rock-mineral geochemistry. Econ. Geol., 91: 934-950. [3] Palmer G. H. (1958) The thermal-emission ion source in solid source mass spectrometry. J. Nuclear Energy, 7: 1-12. [7] Palmer M. R., Spivack A. J. & Edmond J. M. (1987) Temperature and pH controls over isotopic fractionation during adsorption of boron on marine clay. Geochim. Cosmochim. Acta, 51: 2319-2323. [46] Palmer M. R. (1991) Boron isotope systematics of Halmahera arc (Indonesia) lavas: Evidence for involvement of the subducted slab. Geology, 19: 215-217. [7] Palmer M. R. & Swihart G. H. (1996) Boron isotope geochemistry: An overview. In: Boron. Mineralogy, petrology and geochemistry. E. S. Grew & L. M. Anovitz (eds.). Mineral. Soc. of America, Reviews in Mineralogy, Vol. 33: 709-744. [7] Palmer M. R., Pearson P. N. & Cobb S. J. (1998) Reconstructing past ocean pH-depth profiles. Science, 282: 1468-1471. [7] Papadakis I., Poulsen E., Taylor P. D. P., Van Nevel L. & De Bi6vre P. (1999) IMEP-9 trace elements in water. IRRM, Report to participants, EUR 18724 or Internat. Rept. (IRRM)GE / R / IM / 15 / 1999. [37] Papadakis I., Taylor P. D. P. & De Bi6vre P. (2001) CCQM-K2 key comparison: cadmium and lead content in natural water. Metrologia, 38: 543-547. [37] Papic P. B., Ciric M. M. & Zmbov K. F. (1979) Isotopic analysis of boron by negative thermal ionization. Glas. Hem. Drus. Beograd, 44: 195-201. [7] Pardue J. W., Scalan R. S., Van Baalen C. & Parker P. L. (1976) Maximum carbon isotope fractionation in photosynthesis by blue-green algae and a green algae. Geochim. Cosmochim. Acta, 40: 309-312. [46] Park H., Yi J.-H., Han J.-M., Rhee Y.-J. & Lee J. (1998) Studies on autoionization states of Sm by 3-step resonance photoionization. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AIP Press Woodbury, NY. AIP Conference Proceedings 454: 257-260. [33] Park R. & Epstein S. (1960) Carbon isotope fractionation during photosynthesis. Geochim. Cosmochim. Acta, 21: 110-126. [46] Parr R. M. & Clements S. A. (1991) Intercomparison of enriched stable isotope reference materials for medical and biological studies. IAEA, Vienna. IAEA Report NAHRES-5:31 p. [40] Passier H. F., Biittcher M. E. & De Lange G. J. (1999) Sulphur enrichment in organic matter of eastern Mediterranean sapropels: A study of sulfur isotope partitioning. Aq. Geochem., 5(1): 99-119. [45] Patris N., Delmas R. J. & Jouzel J. (2000a) Isotopic signatures of sulfur in shallow Antarctic ice cores. J. Geophys. Res., 105: 7071-7078. [26] Patris N., Mihalopoulos N., Baboukas E. D. & Jouzel J. (2000b) Isotopic composition of sulfur in sizeresolved marine aerosols above the Atlantic Ocean. J. Geophys. Res., 105: 14449-14457. [26] Pattberg S. (1999) Multielement-Spurenbestimmung in Reinstkupfer mit HR-ICP-MS. PhD Thesis, Humboldt-University, Berlin, Germany. [42] Pattberg S. & Matschat R. (1999) Determination of trace impurities in high purity copper using sectorfield ICP-MS: continuous nebulization, flow injection analysis and laser ablation. Fresenius J. Anal. Chem., 364: 410-416. [42] Patterson K. Y., Veillon C., Hill A. D., Moser-Veillon P. B. & O'Haver T. C. (1999) Measurement of calcium stable isotope tracers using cool plasma ICP-MS. J. Anal. Atmos. Spectrom., 14: 1673-1677. [31] Patterson W. P., Smith G. R. & Lohmann K. C. (1993) Continental paleothermometry and seasonality using the isotopic composition of aragonitic otoliths of freshwater fishes. In: Climate Change in Continental Isotopic Records. P. K. Swart, K. C. Lohmann, J. McKenzie & S. Savin (eds.). Am. Geophys. Union. Monograph 78: 191-202. [46] Pattison D. R. M. (1994) Are reversed Fe-Mg exchange and solid solution experiments really reversed? Am. Mineralogist, 79: 938-950. [46] Patton C. J. (1995) U.S. Geological Survey nutrient preservation experiment - Nutrient concentration data for surface-, ground-, and municipal-supply water samples and quality-assurance samples. U.S. Geological Survey. Open-File Report 95-141:140 p. [15]

References

1141

Patton C. J., Fischer A. E., Campbell W. H. & Campbell E. R. (2002) Corn leaf nitrate reductace - a nontoxic alternative to cadmium for photometric nitrate determinations in water samples by air-segmented continous-flow analysis. Environ. Sci. Technol., 36(4): 729-735. [15] Pauling L. (1964) College chemistry. W.H. Freeman and Co., San Francisco, 3rd edition: 832 p. [20] Paulsen P. J. & Kelly W. (1984) Determination of sulfur as arsenic monosulfide by isotope dilution thermal mass spectrometry. Anal. Chem., 56: 708-713. [26, 29] Pauwels J., Lamberty A. & Schimmel H. (1998a) Homogeneity testing of reference materials. Accred. Qual. Assur., 3: 51-55. [43] Pauwels J., Lamberty A. & Schimmel H. (1998b) Quantification of the expected shelf-life of certified reference materials. Fresenius J. Anal. Chem., 361: 395-399. [43] Pawellek F. & Veizer J. (1994) Carbon cycle in the upper Danube and its tributaries: 13CDICconstraints. Isr. J. Earth Sci., 43: 187-194. [10] Paytan A., Luz B. & Kolodny Y. (1990) 6180 in dissolved phosphate. Geol. Soc. Am., ICOG 7, Vol. 27:77 (Abstract.) [22] Pearson F. J. Jr., Fisher D. W. & Plummer L. N. (1978) Correction of ground-water chemistry and carbon isotopic composition for effects of CO2 outgassing. Geochim. Cosmochim. Acta, 42: 1799-1807. [10] Peele G. L. & Brent D. (1977) Cleaning of mass spectrometer ion sources by electropolishing. Anal. Chem., 49: 674. [39] Peiser H. S. & De Bi6vre P. (1999) Traceability in measurement - time for an update? Accred. Qual. Assur., 4: 154-155. [43] Penman A. D. & Wright I. A. (1987) Determination of deuterium level in biological fluids by isotope ratio mass spectrometry. Biomed. Environ. Mass Spectrom., 14: 339-342. [1] Penzias A. A. (1980) Nuclear processing and isotopes in the galaxy. Science, 208: 663-669. [11] Perera I. K., Lyon I. C., Turner G. & Van Lierde P. (1994) Continuous wave laser resonance ionization mass spectrometric analysis of strontium. Int. J. Mass Spetrom. Ion Process., 132: 225-238. [33] Perera I. K., Lyon I. C & Turner G. (1995) Isotope ratio measurement in strontium using two-photon two-colour resonance ionization mass spectrometry. J. Anal. Atom. Spectrosc., 10: 273-280. [33] Perlmutter N. M., Lieber M. & Frauenthal H. L. (1963) Movement of waterborne cadmium and hexavalent chromium wastes in South Farmingdale, Nassau County, Long Island. U.S. Geological Survey Professional Paper, Vol. 475C: C170-C184. [29] Petersen B. J., Howarth R. W. & Garritt R. H. (1986) Sulfur and carbon isotopes as tracers of salt-marsh organic matter flow. Ecology, 67: 865-874. [26] Peterson B., Wollheim W., Mulholland P., Webster J., Meyer J., Tank J., Marti E., Bowden W., Valett M., Hershey A., McDowell W., Dodds W., Hamilton S., Gregory S. & Morral D. (2001) Control of nitrogen export from watersheds by headwater streams. Science, 292: 86-90. [15] Petterson D. W. & Hayes J. M. (1978) Contemporary topics in analytical and clinical chemistry. Plenum Press. New York: 217-252. [14] Phillips B. L., Lee A. & Casey W. H. (2003) Rates of oxygen exchange between the A1208A128(OH)56 (H20)2618+(aq)(A130) molecule and aqueous solution. Geochim. Cosmochim. Acta, 67: 2725-2733. [46] Picard S., Garcia J.-P., L6cuyer C., Sheppard S. M. F., Cappetta H. & Emig C. C. (1998) 5180 values of coexisting brachiopods and fish: Temperature differences and estimates of paleo-water depths. Geology, 26(11): 975-978. [22] Pichlmayer F. & Blochberger K. (1988) Isotopenh~iufigkeitsanalyse von Kohlenstoff, Stickstoff and Schwefel mittels Ger~itekopplung Elementaranalysator-Massenspectrometer. Fresenius Z. Anal. Chem., 331: 196-201. [26, 27] Pichlmayer F., Lamprecht G. & Rossmann A. (1998) Sulfur isotope measurements in organic substances. Isotopes Environ. Health Stud., 34: 227-335. [26] Pickthorn W. J. & O'Neil J. R. (1985) 180 relations in alunite minerals: a potential single-mineral thermometer. Geol. Soc. Am. Abstr. w. Progr., 17:689 (Abstract). [20] Pillinger C. T. (1992) New technologies for small sample stable isotope measurement: static vacuum gas source mass spectrometry, laser probes, ion probes and gas chromatography-isotope ratio mass spectrometry. Int. J. Mass Spectrom. Ion Proc., 118/119: 477-501. [20] Pimminger M., Grasserbauer M., Schroll E. & Cerny I. (1984) Microanalysis in galena by secondary ion mass spectrometry for determination of sulfur isotopes. Anal. Chem., 56: 407-411. [30]

1142

References

Pineau F. & Javoy M. (1983) Carbon isotopes and concentrations in mid-oceanic ridge basalts. Earth Planet. Sci. Lett., 62: 239-257. [3] Pineau F., Shilobreeva S., Kadik A. & Javoy M. (1998) Water solubility and D / H fractionation in the system basaltic andesite-H20 at 1250~ and between 0.5 and 3 kbars. Chem. Geol., 147: 173-184. [46] Piperov N. B. & Penchev N. P. (1973) A study on gas inclusions in minerals. Analysis of the gases from micro-inclusions in allanite. Geochim. Cosmochim. Acta, 37: 2075-2097. [3] Piperov N. B. & Penchev N. P. (1982) Deuterium content of the inclusion water from hydrothermal galenas, Madan, Bulgaria; preliminary investigation. Econ. Geol., 77(1): 195-197. [3] Pisonero J., Costa J. M., Pereiro R., Bordela N. & Sanz-Medel A. (2001) A simple glow discharge ion source for direct solid analysis by on-axis time-of-flight mass spectrometry. J. Anal. At. Spectrom., 16: 1253-1258. [35] Pizzarello S., Krishnamurthy R. V., Epstein S. & Cronin J. R. (1991) Isotopic analyses of amino acids from the Murchison meteorite. Geochim. Cosmochim. Acta, 55: 905-910. [11] Pizzarello S., Feng X., Epstein S. & Cronin J. (1994) Isotopic analyses of nitrogenous compounds from the Murchison meteorite: Ammonia, amines, amino acids and polar hydrocarbons. Geochim. Cosmochim. Acta, 58: 5579-5587. [11] Pizzarello S. & Cronin J. (1998) Alanine enantiomers in the Murchison meteorite. Nature, 394: 236. [11] Plas A. & Friih-Green G. L. (1997) Intracrystalline fractionation of oxygen isotopes in serpentine minerals determined by thermal dehydroxylation. Terra Nova, 9(Abs. Suppl. No. 1): 445 (Abstract). [2] Platzner I. T., Habfast K., Walder A. J. & Goetz A. (1997) Modern isotope ratio mass spectrometry. John Wiley & Sons, Chichester, West Sussex, England: 514 p. [1, 32, 38] Podosek F. A., Prombo C. A., Grossman L. & Zinner E. K. (1991) Chromium isotopic compositions of individual spinel crystals from the Murchison meteorite. 54th meeting of the Meteoritical Society. Arizona State University, Center for Meteorite Studies, Vol. 26:385-386 (Abstract). [29] Polyakov V. B. & Kharlashina N. N. (1994) Effect of pressure on equilibrium isotopic fractionation. Geochim. Cosmochim. Acta, 58: 4739-4750. [46] Polyakov V. B. (1998) On anharmonic and pressure corrections to the equilibrium isotopic constants for minerals. Geochim. Cosmochim. Acta, 62: 3077-3085. [46] Pombo S. A., Pelz O., Schroth M. H. & Zeyer J. (2002) Microbial influence in natural and polluted groundwater: Field-scale C-13-1abeling of phospholipid fatty acids (PLFA) and dissolved inorganic carbon: tracing acetate assimilation and mineralization in a petroleum hydrocarbon-contaminated aquifer. Ferns Microbiot. Ecol., 41: 259-267. [10] Pond D. W., Bell M. V., Dixon D. R., Fallick A. E., Segonzac M. & Sargent J. R. (1998) Stable-carbon isotope composition of fatty acids in hydrothermal vent mussels containing methanotrophic thiotrophic bacterial endosymbionts. Appt. Environm. Microbiol., 64(1): 370-375. [9] Porter L. K. & O'Deen W. A. (1977) Apparatus for preparing nitrogen ammonium chloride for 15Ndeterminations. Anal. Chem., 49: 514-516. [15] Poschenrieder W. P., Herzog R. F. & Barrington A. E. (1965) The relative abundance of the lithium isotopes in the Holbrook meteorite. Geochim. Cosmochim. Acta, 29: 1193-1195. [6] Poulson S. R. & Schoonen M. A. A. (1994) Variations of the oxygen isotope fractionation between NaCO3- and water due to the presence of NaC1 at 100-300~ Chem. Geol. (Isot. Geosci. Sect.), 116: 305315. [46] Powell M. D. & Kyser T. K. (1991) Analysis of 613C and 6180 in calcite, dolomite, rhodochrosite and siderite using a laser extraction system. Chem. Geol. (Isot. Geosci. Sect.), 94: 55-66. [22] Presser T. S. (1994a) Geological origin and pathways of selenium from the California Coastal Ranges to the west-central San Joaquin Valley. In: Selenium in the environment.W.T. Jr. Frankenberger & S. Benson (eds.). Marcel Dekker: 139-155. [29] Presser T. S. (1994b) The Kesteron effect. Environ. Manage., 18: 437-454. [29] Preston T. & Owens N. J. P. (1983) Interfacing an automatic elemental analyser with an isotope ratio mass spectrometer: the potential for fully automated total nitrogen and nitrogen-15 analysis. Analyst, 108: 971-977. [15] Preston T. (1992) The measurement of stable isotope natural abundance variations. Plant Cell Environ., 15: 1091-1097. [8]

References

1143

Preston T. & Slater C. (1994) Mass spectrometric analysis of stable-isotope-labelled acid tracers. Proc. Nutr. Soc., 53: 363-372. [8] Previs S. F., Hazey J. W., Diraison F., Beylot M., David F. & Brunengraber H. (1996) Assay of the deuterium enrichment of water via acetylene. J. Mass Spectrom., 31(6): 639-642. [1] Price G. D., McKenzie J. E., Pilcher J. R. & Hoper W. T. (1997) Carbon-isotope variation in Sphagnum from hummock-hollow complexes: implications for Holocene climate reconstruction. The Holocene, 7: 229-233. [24] Prombo C. A. & Clayton R. N. (1993) Nitrogen isotopic compositions of iron meteorites. Geochim. Cosmochim. Acta, 57: 3749-3761. [16] Prosser S. J., Wright I. P. & Pillinger C. T. (1990) A preliminary investigation into the isotopic measurement of carbon at the picomole level using static vacuum mass spectrometry. Chem. Geol., 83: 71-88. [13] Prosser S. J. & Scrimgeour C. M. (1995) High-precision determination of 2H / 1H in H2 and H20 by continuous-flow isotope ratio mass spectrometry. Anal. Chem., 67: 1992-1997. [1, 3, 8, 38] Pruden G. & Bloomfield C. (1968) The determination of iron(II) sulphide in soil in the presence of iron(III) oxide. Analyst, 93: 532-534. [26] Puchelt H., Sabels B. R. & Hoering T. C. (1971) Preparation of sulfur hexafluoride for isotope geochemical analysis. Geochim. Cosmochim. Acta, 35: 625-628. [20, 26, 36] Pugh L. A. & Rao K. N. (1976) Intensities from infrared spectra. In: Molecular spectroscopy: Modern research. K. N. Rao (ed.). Academic Press, New York Vol. II. [34] Pupin A. M., Dennis M. J., Parker I., Kelly S., Bigwood T. & Toledo M. C. F. (1998) Use of isotopic analyses to determine the authenticity of brazilian orange juice (Citrus sinensis). J. Agric. Food Chem., 46: 1369-1373. [5] Purnell A. L. & Doolan K. J. (1983) Rapid determination of forms of sulfur in coal. Fuel, 62: 1107-1110. [26] Qi H. P., Taylor P. D. P., Berglund M. & De Bi6vre P. (1997) Calibrated measurement of the isotopic composition and atomic weight of the natural Li isotopic reference material IRMM-016. Int. J. Mass Spectrom. Ion Proc., 171: 263-268. [6] Qi H. P. & Coplen T. B. (2003) Evaluation of the 34S / 32S ratio of Soufre de Lacq elemental sulfur isotopic reference material by continuous flow isotope-ratio mass spectrometry. Chem. Geol., 199: 183-187. [40] Qi-Lu & Masuda A. (1994) The isotopic composition and atomic weight of molybdenum. Int. J. Mass Spectrom. Ion Proc., 130: 65-72. [31] Quay P. D., Emerson S. R., Quay B. M. & Devol A. H. (1986) The carbon cycle for Lake WashingtonA stable isotope study. Limnol. Oceanogr., 31: 596-611. [10] Quay P. D., King S., Wilbur D. O. & Wofsy S. (1989) 13C/12C of atmospheric CO2 in the Amazon basin: forest and river sources. J. Geophys. Res., 94: 18327-18336. [9] Quay P. D., King S. L., Stutsman J., Wilbur D. O., Steele L. P., Fung I., Gammon R. H., Brown T. A., Farwell G. W., Grootes P. M. & Schmidt F. H. (1991) Carbon isotopic composition of atmospheric CH4: Fossil and biomass burning source strengths. Global Biogeochem. Cycles, 5: 25-47. [14] Quay P. D., Wilbur D. O., Stuiver M., Hedges J. I., Devol A. H. & Richey J. E. (1992) Carbon cycling in the Amazon River: Implications from the 13C composition of particulate and dissolved carbon. Limnol. Oceanogr., 37: 857-870. [9] Quay P., Stutsman J., Wilbur D., Snorer A., Dlugokencky E. & Brown T. (1999) The isotopic composition of atmospheric methane. Global Biogeochem. Cycles, 13(2): 445-461. [14] Quemerais B., Mabon F., Naulet N. & Martin G. J. (1995) Site-specific isotope fractionation of hydrogen in the biosynthesis of plant fatty acids. Plant Cell Environ., 18: 989-998. [5] Qu6tel C. R., Thomas B., Donard O. F. X. & Grousset F. E. (1997) Factorial optimization of data acquisition factors for lead isotope ratio determination by inductively coupled plasma mass spectrometry. Spectrochim. Acta, B 52: 177-187. [32] Qu6tel C. R., Prohaska T., Hamester M., Kerl W. & Taylor P. D. P. (2000a) Examination of the performance exhibited by a single detector double focusing magnetic sector ICP-MS instrument for uranium isotope abundance ratio measurements over almost three orders of magnitude and down to pg g-1 concentration levels. J. Anal. At. Spectrom., 15: 353-358. [32]

1144

References

Qu6tel C. R., Vogl J., Prohaska T., Nelms S., Taylor P. D. P. & De Bi6vre P. (2000b) Comparative performance study of ICP mass spectrometers by means of U "isotopic measurements". Fresenius J. Anal. Chem., 368: 148-155. [32] Qu6tel C., Prohaska T., Nelms S., Diemer J & Taylor P. D. P. (2001) ICP-MS applied to isotope abundance ratio measurements: performance study and development of a method for combining uncertainty contributions from measurement correction factors. In: Plasma source mass spectrometry: the new millennium. G. Holland & S. D. Tanner (eds.). The RSC Publishers, UK.Proc of the 7th Durham Conf.: 257-269. [32] Quinn P. K. & Bates T. S. (1989) Collection efficiencies of a tandem sampling system for atmospheric aerosol particles and gaseous ammonia and sulfur dioxide. Environ. Sci. Technol., 23: 736-739. [26] Rabalais N. N., Turner R. E., Justic D., Dortch Q. & Wiseman W. (1999) Characterization of hypoxia: topic I report for the integrated assessment on hypoxia in the Gulf of Mexico. NOAA Coastal Ocean Program, NOAA Coastal Ocean Program Decision Analysis, Silver Spring, Md., Series No. 15:167 p. [15] Rabo J. A. (1976) Zeolite chemistry and catalysis. American Chemical Society, Washington DC, ACS Monograph 171:796 p. [36] Rahn T. A. & Wahlen M. (1997) Stable isotope enrichment in stratospheric nitrous oxide. Science, 278: 1776-1778. [19] Rahn T. A. & Wahlen M. (2000) A reassesment of the global isotopic budget of atmospheric nitrous oxide. Glob. Biogeochem. Cycles, 14: 537-543. [34] Rajan R. S., Brown L., Tera F. & Whitford D. J. (1980) Lithium isotopic composition in some stone meteorites. Earth Planet. Sci. Lett., 51: 41-44. [6] Ramakumar K. L., Parab A. R., Khodade P. S., Almaula A. I., Chitambar S. A. & Jain H. C. (1985) Determination of isotopic composition of boron. J. Radioanal. Nucl. Chem., 94: 53-62. [7] Rameb~ick H., Berglund M., Vendelbo D., Wellum R. & Taylor P. D. P. (2001) On the determination of the true dead-time of a pulse-counting system in isotope ratio mass spectrometry. J. Anal. At. Spectrom., 16: 1271-1274. [37] Ramesh R., Bhattacharya S. K. & Pant G. B. (1989) Climatic significance of 5D variations in a tropical tree species from India. Nature, 337: 149-150. [46] Rashid K., Krouse H. R. & McCready R. G. L. (1978) Selenium isotope fractionation during bacterial selenite reduction. In: Short papers of the fourth international conference, geochronology, cosmochronology, isotope geology, 1978. R. E. Zartman (ed.). U. S. Geological Survey: 347-348. [29] Rashid K. & Krouse H. R. (1985) Selenium isotopic fractionation during SeO3 reduction to Se0 and H2Se. Can. J. Chem. 63: 3195-3199. [29] Rau G. (1978) Carbon-13 depletion in a subalpine lake: carbon flow implications. Science, 210: 901-902. [10] Rautenschlein M., Habfast K. & Brand W. (1990). In: Stable isotopes in pediatric nutritional and metabolic research. T. E. Chapman, R. Berger, D. J. Reijngoud & A. Okken (eds.). Intercept Ltd, Andover, MA: 133-148. [8] Raven A. M., van Bergen P. F., Stott A. W., Dudd S. N. & Evershed R. P. (1997) Formation of long-chain ketones in archaeological pottery vessels by pyrolysis of acyl lipids. J. Anal. Appl. Pyrol., 40-1: 267-

285. [8]

Rayleigh Lord R.J. (1939) Nitrogen, argon and neon in the earth's crust with applications to cosmology. Proc. Roy. Soc. Lond., Ser. A 170: 451. [16] Read J. J., Johnson D. A., Asay K. H. & Tieszen L. L. (1992) Carbon isotope discrimination: Relationship to yield, gas exchange, and water-use efficiency in field-grown crested wheatgrass. Crop Science, 32: 168-175. [9] Reardon E. J. & Fritz P. (1978) Computer modeling of groundwater 13C and 14C isotope composition. J. Hydrol., 36: 201-224. [10] Reardon E. J., Allison G. B. & Fritz P. (1979) Seasonal chemical and isotopic variations of soil CO2 at Trout Creek, Ontario. J. Hydrol., 43: 355-371. [10] Rees C. E. & Thode H. G. (1966) Selenium isotope effects in the reduction of sodium selenite and of sodium selenate. Can. J. Chem., 44: 419-427. [29, 46]

References

1145

Rees C. E. & Thode H. G. (1977) A 33S anomaly in the Allende meteorite. Geochim. Cosmochim. Acta, 41: 1679-1682. [20, 26, 45] Rees C. E. (1978) Sulphur isotope measurements using SO2 and SF6. Geochim. Cosmochim. Acta, 42: 383389. [20, 26, 44, 45] Rees C. E.& Holt H. D. (1991) The isotopic analysis of sulphur and oxygen. In: Stable isotopes: Natural and anthropogenic sulphur in the environment; Scope 43. H. R. Krouse & V. A. Grinenko (eds.). John Wiley & Sons, Chichester: 43-64. [26] Rehk~imper M. & Halliday A. N. (1998) Accuracy and long-term reproducibility of lead isotopic measurements by multiple-collector inductively coupled plasma mass spectrometry using an external method for the correction of mass discrimination. Int. J. Mass Spectrom., 181: 123-133. [31] Rehktimper M. & Halliday A. N. (1999) The precise measurement of T1 isotopic compositions by MCICPMS: Application to the analysis of geological materials and meteorites. Geochim. Cosmochim. Acta, 63: 935-944. [31] Rehk~imper M. & Mezger K. (2000) Investigation of matrix effects for Pb isotope ratio measurements by multiple collector ICP-MS: verification and application of optimized analytical protocols. J. Anal. At. Spectrom., 15: 1451-1460. [31] Rehk~imper M., Frank M., Halliday A. N. & Hein J. R. (2001a) Thallium isotope variations in ferromanganese crusts. J. Conf. Abstr., 6:689 (Abstract). [31] Rehk~imper M., SchSnbachler M. & Stirling C. H. (2001b) Multiple collector ICP-MS: Introduction to instrumentation, measurement techniques and analytical capabilities. Geostand. Newslett., 25: 23-40. [31] Rein J. E. & Abernathey R. M. (1972) Rapid mass-spectrometric determination of boron isotopic distribution in boron carbide. Talanta, 19: 857-862. [7] Remaud G., Guillou C., Vallet C. & Martin G. J. (1992) A coupled NMR and MS isotopic method for the authentification of natural vinegars. Fresenius J. Anal. Chem., 342: 457-461. [5] Remaud G., Debon A. A., Martin Y. L., Martin G. G. & Martin G. J. (1997a) Authentication of bitter almond oil and cinnamon oil: application of the SNIF-NMR method to benzaldehyde. J. Agric. Food Chem., 45: 4042-4048. [5] Remaud G. S., Martin Y. L., Martin G. G. & Martin G. J. (1997b) Detection of sophisticated adulterations of natural vanilla flavours and extracts: application of the SNIF-NMR method to vanillin and p-hydroxybenzaldehyde. J. Agric. Food Chem., 45: 859-866. [5] Remaud G. S., Martin Y. L., Martin G. G., Naulet N. & Martin G. J. (1997c) Authentication of mustard oils by combined stable isotope analysis (SNIF-NMR and IRMS). J. Agric. Food Chem., 45: 1844-1848. [26, 5] Rendigs R. R. & Commeau J. A. (1987) Effects of disaggregation on a fine-grained marine mud by two ultrasonic devices. J. Sediment. Petrol., 57: 786-787. [2] Rennie M. J., Meier-Augenstein W., Watt P. W., Patel A., Begley I. S. & Scrimgeour C. M. (1996) Use of continuous-flow combustion MS in studies of human metabolism. Biochem. Soc. Trans., 28: 927932. [8] Reuter J. H., Epstein S. & Taylor H. P. Jr. (1965) O18 / O16 ratios of some chondritic meteorites and terrestrial ultramafic rocks. Geochim. Cosmochim. Acta, 29: 481-488. [20] Reynolds J. H. & Verhoogen J. (1953) Natural variations in the isotopic constitution of silicon. Geochim. Cosmochim. Acta, 3: 224-234. [20, 25] R6v6sz K. & Woods P. H. (1990) A method to extract soil water for isotopic analysis. J. Hydrol., 115: 397406. [1, 2, 21] R6v6sz K., BShlke J. K. & Yoshinari T. (1997) Determination of 5180 and 615N in nitrate. Anal. Chem., 69: 4375-4380. [15] R6v6sz K. M., BShlke J. K., Smith R. L. & Yoshinari T. (1999a) Stable isotope composition of dissolved 02 undergoing respiration in a ground-water contamination gradient. In: U. S. Geological Survey Toxic Substances Hydrology Program; proceedings of the technical meeting. D. W. Morganwalp & H. T. Buxton (eds.). U. S. Geological Survey. Water-Resources Investigations, Report WRI99-4018-C: 323-328. [15]

1146

References

R6v6sz K. M., B/Shlke J. K., Smith R. L. & Yoshinari T. (1999b) Delta 18-O composition of dissolved 02 undergoing respiration in a contaminated ground water. IAEA, Vienna. IAEA International Symposium on isotope techniques in water resources development and management, Vienna, Austria. 10-14 May, 1999. (on CD-ROM). [15] R4v4sz K. & BiShlke J. K. (2002) Comparison of 6180 measurements in nitrate by different combustion techniques. Anal. Chem., 74: 5410-5413. [15] Rhee S. K., Reed R. G. & Brenna J. T. (1997) Fatty acid carbon isotope ratios in humans on controlled diets. Lipids, 32: 1257-1263. [9] Rhee Y.-J., Yi J.-H., Kim J. T., Park H., Han J.-M., Lee J. & Baik M.-G. (1998) Monitoring of Gd photoionization process by detection of fluorescence charateristics. Resoannce Ionization Spectroscopy, Ninth International Symposium J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AIP Press Woodbury, NY. AIP Conference Proceedings 454: 261-265. [33] Rhine E. D., Sims G. K., Mulvaney R. L. & Pratt E. J. (1998) Improving the Berthelot reaction for determining ammonium in soil extracts and water. Soil Sci. Soc. Am. J., 62: 473-480. [15] Rhodes A. L. & Oreskes N. (1999) Oxygen isotope composition of magnetite deposits at E1Laco, Chile: Evidence of formation from isotopically heavy fluids. In: Geology and ore deposits of the central Andes. B. J. Skinner (ed.). Soc. Economic Geologists Inc., Spec. Publ. No. 7, Chap. 11: 333-351. [20] Ricci M. P. & Douthitt C. B. (1998) 180/16 0 in terrestrial material: Consistent and exact calculation of the 170 "ion correction" for O-containing molecules used in gas isotope ratio mass spectrometry. 5th Canadian Continuous-Flow Isotope Ratio Mass Spectrometry, Ottawa, Canada. Poster Presentation (Unpublished). [23] Rice C. A., Tuttle M. L. & Reynolds R. L. (1993) The analysis of forms of sulphur in ancient and sedimentary rocks: comments and cautions. Chem. Geol., 107: 83-95. [26] Richard L. (2001) Calculation of the standard molal thermodynamic properties as a function of temperature and pressure of some geochemically important organic sulfur compounds. Geochim. Cosmochim. Acta, 65(21): 3827-3877. [26] Richer P., Bottinga Y. & Javoy M. (1977) A review of hydrogen, carbon, nitrogen, oxygen, sulphur, and chlorine stable isotope fractionation among gaseous molecules. Ann. Rev. Earth Planet. Sci., 5: 65-110. [16,46] Richet P., Roux J. & Pineau F. (1986) Hydrogen isotope fractionation in the system H20-liquid NaA1Si308: new data and comments on D / H fractionation in hydrothermal experiments. Earth Planet. Sci. Lett., 78: 115-120. [46] Richter F. M., Davis A. M., Ebel D. S. & Hashimoto A. (2002) Elemental and isotopic fractionation of Type B calcium-, aluminum-rich inclusions: experiments, theoretical consideration, and constraints on their thermal evolution. Geochim. Cosmochim. Acta, 66: 521-540. [46] Richter R. & Hoerness S. (1988) The application of the increment method in comparison with experimentally derived and calculated O-isotope fractionations. Chemie der Erde, 48: 1-18. [20] Riciputi L. R. & Paterson B. A. (1994) High spatial-resolution measurements of O isotope ratios in silicates and carbonates by ion microprobe. Am. Mineral., 79: 1227-1230. [20] Riciputi L. R., Duckworth D. C., Barshick C. M. & Smith D. H. (1995) Isotope ratio measurements using glow discharge mass spectrometry. Int. J. Mass Spectrom. Ion Proc., 146/147: 55-64. [35] Riciputi L. R. (1996) A comparison of extreme energy filtering and high mass esolution techniques for measurement of 34S/32S ratios by ion microprobe. Rapid Commun. Mass Spectrom., 10: 282-286. [30] Riciputi L. R., Cole D. R. & Machel H. G. (1996) Sulfide formation in reservoir carbonates of the Devonian Nisku Formation, Alberta, Canada: An ion microprobe study. Geochim. Cosmochim. Acta, 60: 325-336. [26] Riciputi L. R. & Greenwood J. P. (1998) Analysis of sulfur and carbon isotope ratios in mixed matrices by secondary ion mass spectrometry: Implications for mass bias corrections. Int. J. Mass Spectrom., 178: 65-71. [30] Riebesell U., Revill A. T., Holdsworth D. G. & Volkman J. K. (2000) The effects of varying CO2 concentration on lipid composition and carbon isotope fractionation in Emiliania huxleyi. Geochim. Cosmochim. Acta, 64: 4179-4192. [46]

References

1147

Rieley G., Collier R. J., Jones D. M., Eglington G., Eakin P. A. & Fallick A. E. (1991) Sources of sedimentary lipids deduced from stable carbon-isotope analyses of individual compounds. Nature, 342: 425-427. [9] Rieley G. (1994) Derivatization of organic compounds prior to gas chromatographic - combustion isotope ratio mass spectrometric analysis: identification of isotope fractionation processes. Analyst, 119: 915-919. [8] Rieley G., Van Dover C. L., Hedrick D. B. & Eglinton G. (1999) Trophic ecology of Rimicaris exoculota: a combined lipid abundance/stable isotope approach. Marine Biol., 133: 495-499. [9] Rigby D. & Batts B. D. (1986) The isotopic composition of nitrogen in Australian coals and oil shales. Chem Geol., 58: 273-282. [16] Righi D. & Jadault P. (1988) Improving soil clay minerals studies by high-gradient magnetic separation. Clay Minerals, 23: 225-232. [2] Riley J. T., Ruba G. M. & Lee C. C. (1990) Direct determination of total organic sulfur in coal. In: Geochemistry of Sulphur in Fossil Fuels. W. L. Orr & C. M. White (eds.). American Chemical Society, Washington. ACS Symposium Series 429:429-438. [26] Rinsland C. P., Malathy Devi V., Flaud J.-M., Camy-Peyret C., Smith M. A. & Stokes G. M. (1985) Identification of 180-isotope lines of ozone in infrared ground-based solar absorption spectra. J. Geophys. Res., 90" 10719-10725. [18] Rittenberg D., Keston A. S., Rosebury E & Schoenheimer R. (1939) Studies in protein metabolism. II. The determination of nitrogen isotopes in organic compounds. J. Biol. Chem., 127: 291-299. [15] Rittenberg D. (1948) The preparation of gas samples for mass spectrographic isotope analysis. In: Preparation and measurement of isotopic tracers. D. W. Wilson, A. O. C. Nier & S. P. Reimann (eds.). J.W. Edwards, Ann Arbor, Michigan: 31-42. [15] Rittenberg D. & Ponticorvo L. (1956) A method for the determination of the 0 TM concentration of the oxygen of organic compounds. Int. J. Appl. Radiat. Isot., 1" 208-214. [23] Robert F. & Epstein S. (1982) The concentration and isotopic composition of hydrogen, carbon and nitrogen in carbonaceous meteorites. Geochim. Cosmochim. Acta, 46: 81-95. [11] Robert F. & Halbout J. (1990) A numerical model for isotopic patterns from thermal-extraction experiments. Meteoritics, 25: 291-299. [13] Robert F., Rejou-Michel A. & Javoy M. (1992) Oxygen isotopic homogeneity of the earth: New evidence. Earth Planet. Sci. Lett., 108: 1-9. [45] Roberts T. L. & Bettany J. R. (1985) The influence of topography on the nature and distribution of soil sulfur across a narrow environmental gradient. Can. J. Soil. Sci., 65: 419-434. [26] Robinson B. W. (1973) Sulphur isotope equilibrium during sulphur hydrolysis at high temperatures. Earth Planet. Sci. Lett., 18: 443-450. [46] Robinson B. W. (1974) The origin of mineralization at the Tui Mine, Te Aroha, New Zealand, in the light of stable isotope studies. Econ. Geol., 69: 910-925. [3] Robinson B. W. & Kusakabe M. (1975) Quantitative preparation of sulfur dioxide, for 34S/32S analysis, from sulfides by combustion with cuprous oxide. Anal. Chem., 47: 1179-1181. [26, 44, 45] Robinson B. W. (1995) Sulphur isotope standards. In: Reference and intercomparison materials for stable isotopes of light elements. Proceedings of a Consultants Meeting held in Vienna, 1-3 Dec 1993. IAEA-TECDOC-825, IAEA, Vienna: 39-45. [40, 43] Roedder E., Ingram B. & Hall W. E. (1963) Studies of fluid inclusions, III: extraction and quantitative analysis of inclusions in the milligram range. Econ. Geol., 58: 353-374. [1, 3] Roehl R. & Gomez J. (1995) Correction of mass bias drift in inductively coupled plasma mass spectrometry measurements of zinc isotope ratios using gallium as an isotope ratio internal standard. J. Anal. Atom. Spectrom., 10: 15-23. [31] Roether W. (1970) Water-CO2 exchange set-up for the routine oxygen-18 assay of natural waters. Int. J. Appl. Radiat. Isot., 21" 379-387. [1] Roller C. B., Weidmann D., Kosterev A., Tittel E K. & Uehara K. (2003) Quantum cascade laser spectrometer for carbonyl sulfide isotopomer analysis. 2nd International Symposium on Isotopomers, 4-6 November, Stresa, Italy. [34] Rolston J. H., den Hartog J. & Butler J. P. (1976) The deuterium isotope separation factor between hydrogen and liquid water. J. Phys. Chem., 80: 1064-1067. [1]

1148

References

Romanek C. S., Grossman E. L. & Morse J. W. (1992a) Carbon isotopic fractionation in synthetic

aragonite and calcite: Effects of temperature and precipitation rate. Geochim. Cosmochim. Acta, 56: 419-430. [46] Romanini D., Kachanov A. A., Sedeghi N. & Stoeckel F. (1997a) CW cavity ringdown spectroscopy. Chem. Phys. Lett., 264: 316-322. [34] Romanini D., Kachanov A. A. & Stoeckel F. (1997b) Diode laser cavity ringdown spectroscopy. Chem. Phys. Lett., 270: 538-545. [34] Romanini D., Kachanov A. A. & Stoeckel F. (1997c) Cavity ringdown spectroscopy: broad band absolute absorption measurements. Chem. Phys. Lett., 270: 546-550. [34] Romanini D., Kachanov A. A., Morville J. & Chenevier M. (1999) Measurement of trace gases by diode laser cavity ringdown spectroscopy. Proceedings SPIE 3821 (Environmental Sensing and Applications): 94-104. [34] Romanini D. (2003) Cavity ring down and cavity enhanced absorption spectroscopy, and trace detection, with diode lasers. Proceedings of the 4 th International Conference on Tunable Diode Laser Spectroscopy, July 14-18, 2003, Zermatt, Switzerland. [34] Rose S. & Long A. (1989) Dissolved inorganic carbon in the Tucson Basin aquifer. Ground Water, 27: 4349.[10] Rosenbaum J. M. (1993) Room temperature oxygen isotope exchange between liquid CO2 and H20. Geochim. Cosmochim. Acta, 57: 3195-3198. [46] Rosenbaum J. M. (1994) Stable isotope fractionation between dioxide and calcite at 900~ Geochim. Cosmochim. Acta, 58(17): 3747-3753. [46] Rosenbaum J., Kyser T. K. & Walker D. (1994) High temperature oxygen isotope fractionation in the enstatite-olivine-BaCO3 system. Geochim. Cosmochim. Acta, 58: 2653-2660. [46] Rosenbaum J. & Mattey D. (1995) Equilibrium garnet-calcite oxygen isotope fractionation. Geochim. Cosmochim. Acta, 59: 2839-2842. [46] Rosenbaum J. & Slagel M. M. (1995) C-O-H speciation in piston-cylinder experiments. Am. Mineralogist 80: 109-114. [46] Rosenbaum J. M., Cliff R. A. & Coleman M. L. (2000) Chlorine stable isotopes: A comparison of dual inlet and thermal ionization mass spectrometric measurements. Anal. Chem., 72: 2261-2264. [28] Rosenfeld W. D. & Silverman S. R. (1959) Carbon isotope fractionation in bacterial production of methane. Science, 130: 1658-1659. [46] Rosman K. J. R. (1972) A survey of the isotopic and elemental abundance of zinc. Geochim. Cosmochim. Acta, 36: 801-819. [31] Rosman K. J. R., De Laeter J. R. & Gorton M. P. (1980) Cadmium isotope fractionation in fractions of two H3 chondrites. Earth Planet. Sci. Lett., 48: 166-170. [31] Rosman K. J. R., Lycke W., Damen R., Werz R., Hendrickx F., Traas L. & De Bi6vre P. (1987) The preparation and use of synthetic isotope mixtures for testing mass spectrometers. Int. J. Mass Spectrom. Ion Proc., 79: 61-71. [32] Rosman K. J. R. & De Laeter J. R. (1988) Cadmium mass fractionation in unequilibrated ordinary chondrites. Earth Planet. Sci. Lett., 89: 163-169. [31] Rosman K. J. R. & Taylor P. D. P. (1998) Isotopic compositions of the elements 1997. Pure & Appl. Chem., 70(1): 217-235. [28, 32, 37] Ross P. J. & Martin E. A. (1970) A rapid procedure for preparing gas samples for nitrogen-15 determination. Analyst, 95: 817-822. [15] Ross G. J. (1969) Acid dissolution of chlorites: release of magnesium, iron and aluminium and mode of acid attack. Clays Clay Minerals, 17: 347-354. [2] Rotaru M., Birck J. L. & All6gre C. J. (1992) Clues to early solar system history from chromium isotopes in carbonaceous chondrites. Nature, 385(No. 6386): 465-470. [29] Roth E. (1997) Critical evaluation of the use and analysis of stable isotopes. Pure & Appl. Chem., 69(8): 1753-1828. [37, 40]

References

1149

Rothman L. S., Rinsland C. P., Goldman A., Massie S. T., Edwards D. P., Flaud J.-M., Perrin A., CamyPeyret C., Dana V., Mandin J.-Y., Schroeder J., McCann A., Gamache R. R., Wattson R. B., Yoshino K., Chance K. V., Jucks K. W., Brown L. R., Nemtchinov V. & Varanasi P. (1998) The HITRAN molecular spectroscopic database and HAWKS (HITRAN atmospheric Workstation): 1996 edition. J. Quant. Spectrosc. Rad. Trans., 60: 665-710. [34] Rouxel O., Ludden J., Fouquet Y. & Carignan J. (2000) Natural variations of selenium isotopes determined by multicollector plasma source mass spectrometry: Application to seafloor hydrothermal systems. J. Conf. Abstr. (Goldschmidt 2000 Conference), 5: 858-859. [31] Rouxel O., Fouquet Y. & Ludden J. (2001) Copper, selenium and sulphur isotope systematics of hydrothermal systems. J. Conf. Abstr., 6: 689. [31] Rouxel O., Ludden J., Carginan, J., Martin, L. & Fouquet Y. (2002) Natural variations of Se isotopic composition determined by hydride generation multiple collector inductively coupled plasma mass spectrometry. Geochim. Cosmochim. Acta, 66(18): 3191-3199. [20, 29] Rouxel O., Ludden J. & Fouquet Y. (2003) Antimony isotope variations in natural systems and implications for their use as geochemical tracers. Chem. Geol., 200: 25-40. [46] Rozanski K., Dulinski M. & Rzepska J. (1987) D and 180 analysis of milligram water samples. Appl. Radioat. Isot., 38: 713-716. [1] Rozanski K. & Rzepka J. (1991) Modification of D and O-18 content of water samples due to evaporation from leaky containers. Appl. Radiat. Isot., 42: 129-133. [1] RiJckmann T., Brenninkmeijer C. A. M., Neeb P. & Crutzen P. J. (1998a) Ozonolysis of non-methane hydrocarbons as a source of the observed mass independent oxygen isotope enrichment in tropospheric CO. J. Geophys. Res., 103: 1463-1470. [18] RiJckmann T., Brenninkmeijer C. A. M., Saueressig G., Bergamaschi P., Crowley J. N., Fisher H. & Crutzen P. J. (1998b) Mass-independent oxygen isotope fractionation in atmospheric CO as a result of the reaction CO + OH. Science, 281: 544-546. [18] R6ckmann T., Kaiser J., Brenninkmeijer C. A. M., Crowley J. N., Borchers R., Brand W. A. & Crutzen P. J. (2001) The isotopic enrichment of nitrous oxide (15N14NO, 14N15NO and 14N14N180) in the stratosphere and in the laboratory. J. Geophys. Res., 106: 10403-10410. [19, 34] Rudolph J., Lowe D. C., Martin R. J. & Clarkson T. S. (1997) A novel method for compound specific determination of 613C in volatile organic compounds at ppt levels in ambient air. Geophys. Res. Lett., 24: 659-662. [8] Rumble D., Palin J. M. & Hoering T. C. (1991) Laser fluorination of sulfide minerals with F2 gas. Carnegie Institution of Washington, Geophysical Laboratory. Annual Report of the Director, 1990-1991, No. 2250: 30-34. [20] Rumble D., Hoering T. C. & Palin J. M. (1993) Preparation of SF6 for sulfur isotope analysis by laser heating sulfide minerals in the presence of F2 gas. Geochim. Cosmochim. Acta, 57: 4499-4512. [20, 26, 45] Rumble D. III& Hoering T. C. (1994) Analysis of oxygen and sulfur isotope ratios in oxide and sulfide minerals by spot heating with a carbon dioxide laser in a fluorine atmosphere. Acc. Chem. Res., 27: 237-241. [20] Rumble D. III, Farquhar J., Young E. D. & Christensen C. P. (1997) In situ oxygen isotope analysis with an excimer laser using F2 and BrF5 reagents and 02 gas as analyte. Geochim. Cosmochim. Acta, 61(19): 4229-4234. [20] Rumble D. & Sharp Z. D. (1998) Laser microanalysis of silicates for 180/170 /160 and of carbonates for 180/160 and 13C/12C ratios. In" Applications of microanalytical techniques to understanding mineralizing processes. M. A. McKibben, W. C. III Shanks & W. I. Ridley (eds.). Society of Economic Geologists, Littleton, CO, USA. Vol. 7: 99-119. [20, 46] Rundel P. W., Ehleringer J. R. & Nagy K. A. (1989) Stable isotopes in ecological research. In: Ecological studies: Analysis and Synthesis, 68. W. D. Billings, F. Golley, O. L. Lange, J. Olson & H. Remmert (eds.). Springer Verlag, New York. [9] Rusak D. A., Anderson J. E., Keller R. A., Wilkerson C. W. Jr. & Kunhardt E. E. (1999) Signal dependence on light modulation frequency in optogalvanic spectroscopy. Appl. Spectrosc., 53: 1651-1653. [34]

1150

References

Russ G. P. & Bazan J. M. (1987) Isotopic ratio measurements with an inductively coupled plasma source

mass spectrometer. Spectrochim. Acta, 42B: 49-62. [31, 35] Russell J. D., Birnie A. & Fraser A. R. (1984) High-gradient magnetic separation (HGMS) in soil clay

mineral studies. Clay Minerals, 19: 771-778. [2] Russell R. D. (1971) The systematics of double spiking. J. Geophys. Res., 76(20): 4949-4955. [29] Russell R. D. (1972) Reply to discussion of "The systematics of double spiking". ]. Geophys. Res., 77(2):

381. [29]

Russell S. S., Arden J. W. & Pillinger C. T. (1991) Evidence for multiple sources of diamond from primi-

tive chondrites. Science, 54: 1188-1191. [13] Russell S. S. (1992) A carbon and nitrogen isotope study of chondritic diamond and silicon carbide. PhD

Thesis, Open University, Milton Keynes, England. [13] Russell W. A. & Papanastassiou D. A. (1978) Calcium isotope fractionation in ion-exchange chromato-

graphy. Anal. Chem., 50: 1151-1154. [29] Russell W. A., Papanastassiou D. A. & Tombrello T. A. (1978) Ca isotope fractionation on the Earth and

other solar system materials. Geochim. Cosmochim. Acta, 42: 1075-1090. [29, 31, 46] Russo R. E., Mao X. & Mao, S. S. (2002) The physics of laser ablation in microchemical analysis. Anal.

Chem., 74: 71A-77A. [20] Ryan J. N. & Gschwend P. M. (1991) Extraction of iron oxides from sediments using reductive dissolution by titanium(III). Clays Clay Minerals, 39: 509-518. [2] Ryan K. R. (1972) Ionic collision processes in gaseous N20. J. Chem. Phys., 57: 271-276. [19] Ryan P. C., Conrad M. E., Brown K., Chamberlain C. P. & Reynolds R. C. Jr. (1998) Oxygen isotope compositions of mixed-layer serpentine-chlorite and illite-smectite in hte Tuscaloosa formation (US Gulf Coast): Implications for pore fluids and mineralogic reactions. Clays Clay Minerals, 46: 357-368. [2] Rye R. O. & O'Neil J. R. (1968) The 018 content of water in primary fluid inclusions from Providencia, North-Central Mexico. Econ. Geol., 63: 232-238. [3, 20] Rye R. O., Schuiling R. D., Rye D. M. & Jansen J. B. H. (1976) Carbon, hydrogen, and oxygen isotope studies of the regional metamorphic complex at Naxos. Geochim. Cosmochim. Acta, 40: 1031-1049. [3] Rye R. O., Back W., Hanshaw B. B., Rightmare C. T. & Pearson E J. Jr. (1981) The origin and isotopic composition of dissolved sulphide in groundwaters from carbonate aquifers in Florida and Texas. Geochim. Cosmochim. Acta, 45: 1841-1950. [26] Rye R. O. (1993) The evolution of magmatic fluids in the epithermal environment. Econ. Geol., 63: 232238. [20] Saccocia P. J., Seewald J. S. & Shanks W. C. III (1998) Hydrogen and oxygen isotope fractionation between brucite and aqueous NaC1 solutions from 250 to 450~ Geochim. Cosmochim. Acta, 62: 485492. [46] Sachs J. P., Repeta D. J. & Goericke R. (1999) Nitrogen and carbon isotopic ratios of chlorophyll from marine phytoplankton. Geochim. Cosmochim. Acta, 63: 1431-1441. [46] Sadofsky S. J. & Bebout G. E. (2000) Ammonium partitioning and nitrogen isotope fractionation among coexisting micas during high temperature fluid-rock interactions: Examples from Vermont and Maine, U.S.A. Geochim. Cosmochim. Acta, 64: 2835-2849. [16] Sahoo S. K. & Masuda A. (1995) Simultaneous measurement of lithium and boron isotopes as lithium tetraborate ion by thermal ionization mass spectrometry. Analyst, 120: 335-339. [6, 7] Saito M. (1995) A contribution to the study of matrix effects in the analysis of solid samples by DC glow discharge mass spectrometry. Spectrochim. Acta, B 50: 171-178. [35] Sajjad M. I. & Tasneem M. A. (1983) An efficient hydrogen preparation system. The Nucleus, 20: 31-34.

[1]

Sakai H. & Honma H. (1966) Isotopic analysis of oxygen in rocks and minerals. I. On the conversion of

oxygen gas into carbon dioxide. Mass Spectroscopy, 14:183-186 (in Japanese; with English abstract).

[20]

Sakai H. & Yamamoto M. (1966) Fractionation of sulfur isotopes in the preparation of sulfur dioxide.

An improved technique for the precision analysis of stable sulfur isotopes. Geochem. J., 1: 35-42. [26] Sakai H., Smith J. W., Kaplan I. R. & Petrowski C. (1976) Micro-determinations of C, N, S, H, He, metallic Fe, 613C, 615N and ~)34S in geologic samples. Geochem. J., 10: 85-96. [13]

References

1151

Sakai H. & Dickson F. W. (1978) Experimental determination of the rate and disequilibrium fractionation factors of sulfur isotope exchange between sulfate and sulfide in slightly acid solutions at 300~ and 1000 bars. Earth Planet. Sci. Lett., 39: 151-161. [46] Sakai H., DesMarias D. J., Ueda A. & Moore J. G. (1984) Concentrations and isotope ratios of carbon, nitrogen and sulfur in ocean-floor basalts. Geochim. Cosmochim. Acta, 48: 2433-2441. [3] Salata G. G., Roelke L. A. & Cifuentes L. A. (2000) A rapid and precise method for measuring stable carbon isotope ratios of dissolved inorganic carbon. Marine Chem., 69: 153-161. [10] Saloman E. B. (1990) RIS data tables. Spectrochim. Acta, 45B: 37-83. [33] Saloman E. B. (1991) RIS data tables. Spectrochim. Acta, 46B: 319-378. [33] Saloman E. B. (1992) RIS data tables. Spectrochim. Acta, 47B: 517-543. [33] Saloman E. B. (1993) RIS data tables. Spectrochim. Acta, 48B: 1139-1203. [33] Saltzman E. S., Brass G. W. & Price D. A. (1983) The mechanism of sulfate aerosol formation: chemical and sulfur isotopic evidence. Geophys. Res. Lett., 10: 513-516. [26] Sanchez Chillon B., Alberdi M. T., Leone G., Bonadonna F., Stenni B. & Longinelli A. (1994) Oxygen isotopic composition of fossil equid tooth and bone phosphate: an archive of difficult interpretation. Palaeogeogr., Palaeoclimatol., Palaeoecol., 107: 317-328. [22] Sano Y., Takahata N., Nishio Y. & Marty B. (1998) Nitrogen recycling in subduction zones. Geophys. Res. Lett., 25: 2289-2292. [16] Sansone F. J., Popp B. N. & Rust T. M. (1997) Stable carbon isotope analysis of low-level methane in water and gas. Anal. Chem., 69: 40-44. [14] Santos R. V. & Clayton R. N. (1995) Variations of oxygen and carbon isotopes in carbonatites: A study of Brazilian alkaline complexes. Geochim. Cosmochim. Acta, 59: 1339-1352. [22] Santrock J., Studley S. A. & Hayes J. M. (1985) Isotopic analysis based on the mass spectrum of carbon dioxide. Anal. Chem., 57: 1444-1448. [1, 15, 44, 45] Santrock J. & Hayes J. M. (1987) Adaptation of the Unterzaucher procedure for determination of oxygen-18 in organic substances. Anal. Chem., 59: 119-127. [23] Sanyal A., Hemming N. G., Hanson G. N. & Broecker W. S. (1995) Evidence for a higher pH in the glacial ocean from boron isotopes in foraminifera. Nature, 373: 234-236. [7] Sanyal A., Hemming N. G., Broecker W. S., Lea D. W., Spero H. J. & Hanson G. N. (1996) Oceanic pH control on the boron isotopic composition of foraminifera: evidence from culture experiments. Paleoceanography, 11: 513-517. [46] Sasaki A., Arikawa Y. & Folinsbee R. E. (1979) Kiba reagent method of sulphur extraction applied to isotopic work. Bull. Geol. Surv. Japan, 30: 241-245. [26, 27] Sato R. K. & McMillan P. F. (1987) An infrared and Raman study of the isotopic species of a-quartz. J. Phys. Chem., 91: 3494-3498. [46] Sauer P. E. & Sternberg L. d. S. L. (1994) Improved method for the determination of the oxygen isotopic composition of cellulose. Anal. Chem., 66: 2409-2411. [23] Sauke T. B. & Becker J. F. (1998) Stable isotope laser spectrometer for exploration of Mars. Planet. Space Sci., 46: 805-812. [34] Saurer M., Borella S. & Leuenberger M. (1997a) 6180 of tree rings of beech (Fagus sylvatica) as a record of 6180 of growing season precipitation. Tellus, 49B: 80-92. [23] Saurer M., Aellen K. & Siegwolf R. (1997b) Correlating 613C and 6180 in cellulose of trees. Plant Cell Environm., 20: 1543-1550. [46] Saurer M., Robertson I., Siegwolf R. & Leuenberger M. (1998) Oxygen isotope analysis of cellulose: An interlaboratory comparison. Anal. Chem., 70(10): 2074-2080. [23] Savarino J. & Thiemens M. H. (1999a) Analytical procedure to determine both 6180 and 6170 of H202 in natural water and first measurements. Atmos. Environ., 33: 3683-3690. [18] Savarino J. & Thiemens M. H. (1999b) Mass independent oxygen isotope (160, 170, 180) fractionation found in the H+O2+M reaction. J. Chem. Phys., A 103: 9221-9229. [18] Savin S. M. (1967) Oxygen and hydrogen isotope ratios in sedimentary rocks and minerals. PhD Thesis (unpublished), California Institute of Technology. [2, 20] Savin S. M. & Epstein S. (1970) The oxygen and hydrogen isotope geochemistry of clay minerals. Geochim. Cosmochim. Acta, 34: 25-42. [2, 46]

1152

References

Saxton J. M., Lyon I. C., Chatzitheodoridis E., Perera I. K., Van Lierde P., Freedman P. & Turner G. (1996) The Manchester Isolab 54 ion microprobe. Int. J. Mass Spectrom. Ion Proc., 154: 99-131. [30] Scalan R. S. (1958) The isotopic composition, concentration, and chemical state of the nitrogen in igneous rocks. PhD Thesis, The University of Arkansas: 79 p. [16] Schadewaldt P., Schommartz B., Wienrich G., BriJsicke H., Piolot R. & Ziegler D. (1997) Application of isotope-selective nondispersive infrared spectroscopy (IRIS) for evaluation of [13C]octanoic acid gastric-emptying breath tests: comparison with isotope ratio-mass spectrometry (IRMS). Ctin. Chem., 43(3): 518-522. [34] Scheele N. & Hoers J. (1992) Carbon isotope fractionation between calcite, graphite and CO2: an experimental study. Contrib. Mineral. Petrol., 112: 35-45. [46] Schell D. M. (1978) Chemical and isotopic methods in nitrification studies. In: Microbiology-1978. D. Schlesinger (ed.). Am. Soc. Microbiol., Washington: 292-295. [15] Schelles W., De Gendt S., Muller V. & Van Grieken R. (1995) Evaluation of secondary cathodes for glow discharge mass spectrometry analysis of different nonconducting sample types. Appl. Spectrosc., 49: 939-944. [35] Schelles W., De Gendt S. & Van Grieken R. (1996) Optimization of secondary cathode thickness for direct current glow discharge mass spectrometric analysis of glass. J. Anal. At. Spectrom., 11: 937-941.

[35]

Schelles W. & Van Grieken R. (1996) Direct current glow discharge mass spectrometric analysis of a Macor ceramic using a secondary cathode. Anal. Chem., 68: 3570-3574. [35] Scherer J. J., Paul J. B., O'Keefe A. & Saykally R. J. (1997) Cavity ringdown laser absorption spectroscopy. History, development, and application to molecular beams. Chem. Rev., 97: 25-51. [34] Schidlowski M., Hayes J. M. & Kaplan I. R. (1983) Isotopic interferences of ancient biochemistry of carbon, sulfur, hydrogen and nitrogen. In: Earth's earliest biosphere: Its origin and evolution. J. W. Schopf (ed.). Princeton University Press, Princeton: 149-186. [27] Schiegl W. E. & Vogel J. C. (1970) Deuterium content of organic matter. Earth Planet. Sci. Lett., 7: 307313. [1] Schiffries C. M. & Rye D. M. (1990) Stable isotopic systematics of the Bushveld complex: II. Constraints on hydrothermal processes in layered intrusions. Am. J. Sci., 290: 209-245. [3] Schimmelmann A. & DeNiro M. J. (1985) Determination of oxygen stable isotope ratios in organic matter containing carbon, hydrogen, oxygen, and nitrogen. Anat. Chem., 57: 2644-2646. [23] Schimmelmann A. (1991) Determination of the concentration and stable isotopic composition of nonexchangeable hydrogen in organic matter. Anal. Chem., 63: 2456-2459. [24] Schimmelmann A. & DeNiro M. J. (1993) Preparation of organic and water hydrogen for stable isotope analysis: Effects due to reaction vessels and zinc reagents. Anal. Chem., 65: 789-792. [1, App C] Schleser G. H. (1990) Investigations of the 613C pattern in leaves of Fagus sylvatica L. J. Experim. Botanic, 41: 565-572. [24] Schlesinger W. H. (1997) Biogeochemistry - An analysis of global change. Academic, San Diego, California: 588 p. [15] Schmidt H.-L., Kempe K., Demmelmair J. & Medina R. (1992) On-line gas chromatography and isotope ratio mass spectrometry of organic compounds in the range of natural abundances. Kyoto '92 Int. Conf. Biol. Mass Spectrom. Osaka University, Japan, Proceedings: 382-383. [23] Schmidt M., Botz R., Stoffers P., Anders T. & Bohrmann G. (1997) Oxygen isotopes in marine diatoms: a comparative study of analytical techniques and new results on the isotope composition of recent marine diatoms. Geochim. Cosmochim. Acta, 61: 2275-2280. [46] Schmutz J. A. & Hobson K. A. (1998) Geographic, temporal, and age-specific variation in diets of glaucous gulls in western Alaska. The Condor, 100: 119-130. [9] Schoell M., Faber E. & Coleman M. (1983) Carbon and hydrogen isotopic compositions of the NBS-22 and NBS-21 stable isotope reference materials: An Inter-laboratory comparison. Org. Geochem., 5: 36. [40] Schoenau J. J. & Bettany J. R. (1988) A method for determining sulfur-34 abundance in soil and plant sulfur fractions. Soil Sci. Soc. Am. J., 52: 297-300. [26] Schoenau J. J. & Bettany J. R. (1989) 34S natural abundance variations in prairie and boreal forest soils. J. Soil Sci., 40: 397-413. [26]

References

115 3

Schoeninger M. J., Iwaniec U. T. & Nash L. T. (1998) Ecological attributes recorded in stable isotope ratios of aboreal prosimian hair. Oecology, 113: 222-230. [9] Schopf J. W. (1983) Earth's earliest biosphere: Its origins and evolution. Princeton University Press, Princeton: 543 p. [13]

Schopf J. W. & Klein C. (1992) The Proterozoic biosphere. Cambridge University Press, Cambridge, England: 1348 p. [13]

Schueler B., Morton J. & Mauersberger K. (1990) Measurements of isotopic abundances in collected stratospheric ozone samples. Geophys. Res. Lett., 17: 1295-1298. [18] Schumacher K., Turgeon H. & Mosandl A. (1995) Sample preparation for gas-chromatography isotope ratio mass-spectrometry - an investigation with volatile components from strawberries. Phytochem. Anal., 6: 258-261. [8] Schupp M., Bergamaschi P., Harris G. W. & Crutzen P. J. (1993) Development of a tunable diode laser absorption spectrometer for measurements of the 13C/12C ratio in methane. J. Chemosphere, 26: 1322. [34] Schiitze M. (1939) Ein neues Verfahren zur direkten Sauerstoffbestimmung in organischen Sub-stanzen. Fresenius Z. Anal. Chem., 118: 245-258. [23] Schwander H. (1953) Bestimmung des relativen Sauerstoffisotopen-Verhatnisses in Silikatgesteinen und Mineralen. Geochim. Cosmochim. Acta, 4: 261-291. [20] Schwartz H. P., Agyei F. K. & McMullen C. C. (1969) Boron isotopic fractionation during clay absorption from sea-water. Earth Planet. Sci. Lett., 6: 1-5. [7, 46] Schwarcz H. P., Harmon R. S., Thompson P. & Ford D. C. (1976) Stable isotope studies of fluid inclusions in speleothems and their paleoclimatic significance. Geochim. Cosmochim. Acta, 40: 657-665. [3] Schwarcz H. P., Melbye J., Katzenberg M. A. & Knyf M. (1985) Stable isotopes in human skeletons of southern Ontario: reconstructing palaeodiet. J. Archaeol. Sci., 12: 187-206. [22] Schwarz G., Ohm K., Yamanouchi T., Furukawa T., Kowski P. & Gernandt H. (1996) Stable isotopic composition of Antarctic air moisture and precipitation. Nankyoku Shiry6 (Antarctic Record), 40: 169178. [36] Schwarz G., Kowski P. & Gernandt H. (1998a) Saisonale Variation von Deuterium und Sauerstoff-18 in Luftfeuchte und Niederschlag tiber Antarktika. Isot. Environ. Health Stud., 34: 159-168. [36] Schwarz G., Kowski P., Kaneto S. & Gernandt H. (1998b) Meridional distributions of deuterium in atmospheric water vapour between tropical and southern polar latitudes. Mem. Natl. Inst. Polar Res., 52 (Special Issue): 102-110. [36] Scrimgeour C. M., Rollo M. M., Mudambo S. M. K. T., Handley L. L. & Prosser S. J. (1993) A simplified method for deuterium/hydrogen isotope ratio measurements on water samples of biological origin. J. Biol. Mass Spectrom., 22: 383-387. [1] Scrimgeour C. M. (1995) Measurement of plant and soil water isotope composition by direct equilibration methods. J. Hydrol., 172: 261-274. [1, 21] Scrimgeour C. M., Begley I. S. & Thomason M. L. (1999) Measurement of deuterium incorporation into fatty acids by gas chromatography isotope ratio mass spectrometry. Rapid Commun. Mass Spectrom., 13: 271-274. [8] Seaton N. A., Friedman S. P., MacElroy J. M. D. & Murphy A. (1997) The molecular sieving mechanism in carbon molecular sieves: a molecular dynamics and critical path analysis. Langmuir, 13:1199-1204. [36] Seboldt W., Linger S., Hoernes S., Grimmeisen W., Lekies R., Herkelmann R. & Burr D. (1993) Lunar oxygen extraction using fluorine. In: Resources of near-Earth space. J. S. Lewis, M. S. Matthews & M. L. Guerrieri (eds.). Univ. Arizona Press, Tucson: 129-147. [20] Segschneider H.-J., Hofmann D., Schmidt G. & Russow R. (1995) Incorporation of 15NO2 nitrogen into individual amino acids by suflowers using GC-C-IRM. Isot. Environ. Health Stud., 31: 315-325. [8] Seiler R. L. (1998) Prediction of lands susceptible to irrigation-induced selenium contamination of water. In: Environmental chemistry of selenium. W. T. Jr. Frankenberger & R. A. Engberg (eds.). Marcel Dekker, Inc., Vo|. 64: 397-418. [29]

1154

References

Selby D. & Nesbitt B. E. (1996) British Columbia; implications for P-V-T-X history of ore and post ore fluids. Biennial Pan-American Conference on Research on Fluid Inclusions, 6. PACROFI VI; Sixth biennial Pan-American Conference on Research on Fluid Inclusions; Program and Abstracts: 116117. [3] Sennaroglu A., Konca A. O. & Pollock C. R. (2000) Continuous-wave power performance of a 2.47 ~m Cr2+:ZnSe laser: Experiment and modeling. IEEE J. Quantum El., 36: 1199-1205. [34] Sephton M. A., Pillinger C. T. & Gilmour I. (1998) 613C of free and macromolecular aromatic structures in the Murchison meteorite. Geochim. Cosmochim. Acta, 62: 1821-1828. [11] Sephton M. A., Pillinger C. T. & Gilmour I. (1999) Small-scale hydrous pyrolysis of macromolecular material in meteorites. Planet. Space Sci., 47: 181-187. [11] Sephton M. A., Pillinger C. T. & Gilmour I. (2001) Normal alkanes in meteorites: molecular 613C values indicate an origin by terrestrial contamination. Precambrian Res., 106: 45-56. [11] Sephton M. A. & Gilmour I. (2001) Pyrolysis-gas chromatography-isotope ratio monitoring-mass spectrometry of macromolecular material in meteorites. Planet. Space Sci., 49: 465-471. [11] Seyfried W. E. Jr., Janecky D. R. & Brendt M. E. (1987) Rocking autoclaves for hydrothermal experiments II: the flexible cell system. In: Experimental Hydrothermal Techniques. G. Ulmer & H. L. Barnes (eds.). Wiley-Interscience, New York: 216-240. [46] Shackleton C. H. L., Roitman E., Phillips A. & Chang T. (1997) Androstanediol and 5-androstenediol profiling for detecting exogenously administered dihydrotestosterone, epitestosterone, and dehydroepiandrosterone: Potential use in gas chromatography isotope ratio mass spectrometry. Steroids, 62: 665-673. [8] Shackleton N. J. (1974) Attainment of isotopic equilibrium between ocean water and the benthonic foraminifera genus Uvigerina: Isotopic changes in the ocean during the last glacial. In: Variation du climat au cours du Pleistoc6ne. J. Labeyrie (ed.). Colloq. Intern. CNRS, Paris, Vol. 219: 203-209. [46] Shahack-Gross R., Shemesh A., Yakir D. & Weiner S. (1996) Oxygen isotopic composition of opaline phytoliths: potential for terrestrial climatic reconstruction. Geochim. Cosmochim. Acta, 60: 3949-3953. [46] Shanks W. C. III, Boehlke J. K. & Seal R. R. II (1995) Stable isotopes in mid-ocean ridge hydrothermal systems: Interaction between fluids, minerals and organisms. In: Seafloor hydrothermal systems: Physical. chemical, biological, and geological interactions. S. E. Humphris. et al. AGU, Monograph 91: 194-221. [1] Shanks W. C., Crowe D. E. & Johnson C. (1998) Sulfur isotope analyses using the laser microprobe. In: Applications of microanalytical techniques to understanding mineralizing processes. M. A. McKibben, W. C. III Shanks & W. I. Ridley (eds.). Society of Economic Geologists, Littleton, CO, USA, Vol. 7: 141-153. [46] Sharp Z. D. & O'Neil J. R. (1988) A laser-based carbon reduction technique for oxygen isotope analysis of microsamples of silicates and oxides. Geol. Soc. Am. Abstr. w. Prog., 20:A280 (Abstract). [20] Sharp Z. D. & O'Neil J. R. (1989) A laser-based carbon reduction technique for oxygen isotope analysis of silicates and oxides. In: Annual Report of the Director, Geophysical Laboratory: 1988-1989. Carnegie Institution, Washington, Vol. 89: 72-78. [20] Sharp Z. D. (1990) A laser-based microanalytical method for the in situ determination of oxygen isotope ratios of silicates and oxides. Geochim. Cosmochim. Acta, 54: 1353-1357. [2, 20] Sharp Z. D., Giletti B. J. & Yoder H. S. Jr. (1991) Oxygen diffusion rates in quartz exchanged with CO2. Earth Planet. Sci. Lett., 107: 339-348. [46] Sharp Z. D. (1992) In situ laser microprobe techniques for stable isotope analysis. Chem. Geol. (Isot. Geosci. Sect.), 101: 3-19. [26, 22] Sharp Z. D. & Jenkin G. R. T. (1994) An empirical estimate of the diffusion rate of oxygen in diopside. J. Metamorphic Geol., 12: 89-97. [46] Sharp Z. D. & Kirschner D. L. (1994) Quartz-calcite oxygen isotope thermometry: A calibration based on natural isotopic variations. Geochim. Cosmochim. Acta, 58: 4491-4501. [46] Sharp Z. D. & Moecher D. (1994) O-isotope variations in a porphyroclastic metaanorthosite: Diffusion effects and false isotherms. Am. Mineralogist, 79: 951-959. [20] Sharp Z. D. (1995) Oxygen isotope geochemistry of the A12SiO5 polymorphs. Am. J. Sci., 295: 1058-1076. [20]

References

1155

Sharp Z. D. & Cerling T. E. (1996) A laser GC-IRMS technique for in situ stable isotope analyses of carbonates and phosphates. Geochim. Cosmochim. Acta, 60(15): 2909-2916. [22, 45] Sharp Z. D. & Cerling T. E. (1998) Fossil isotope records of seasonal climate and ecology. Straight from the horse's mouth. Geology, 26: 219-222. [22] Sharp Z. D., Atudorei V. & Furrer H. (2000) The effect of diagenesis on oxygen isotope ratios of biogenic phosphates. Am. J. Sci., 300: 222-237. [22] Sharp Z. D., Atudorei V. & Durakiewicz T. (2001) A rapid method for the determination of hydrogen and oxygen isotope ratios of water from water and solid hydrous substances. Chem. Geol., 178: 197210. [1, 3] Shaw S. Y. & Lue J. T. (1980) An integrated circuit based vacuum ionization gauge meter. J. Phys. E.: Sci. Instrum., 13: 1150-1152. [39] Shearer G. & Kohl H. (1993) Natural abundance of 15N: Fractional Contribution of two sources to a common sink and use of isotope discrimination. In: Nitrogen isotope Techniques. R. Knowles & T. H. Blackburn (eds.). Academic Press Inc.: 89-125. [15] Sheets R. W., Nesbitt B. E. & Muehlenbachs K. (1996) Meteoritic water component in magmatic fluids from porphyry copper mineralization, Babine Lake area, British Columbia. Geology, 24: 1091-1094.

[3]

Shelton K. L. (1983) Composition and origin of ore-forming fluids in a carbonate-hosted porphyry copper and skarn deposit: a fluid inclusion and stable isotope study of Mines Gasp6, Quebec. Econ. Geol., 78: 387-421. [3] Shelton K. L., Taylor R. P. & So C. S. (1987) Stable isotope studies of the Dae Hwa Tungsten-Molybdenum Mine, Republic of Korea: evidence of progressive meteoric water interaction in Tungstenbearing hydrothermal system. Econ. Geol., 82: 471-481. [3] Shemesh A., Kolodny Y. & Luz B. (1988) Isotope geochemistry of oxygen and carbon in phosphate and carbonate of phosphorite francolite. Geochim. Cosmochim. Acta, 52: 2565-2572. [22] Shemesh A. (1990) Crystallinity and diagenesis of sedimentary apatites. Geochim. Cosmochim. Acta, 54: 2433-2438. [22] Shemesh A. (1992) Response to R. N. Clayton on "Silica-carbonate isotopic temperature calibration". Science, 258: 1163. [46] Shemesh A., Charles C. D. & Fairbanks R. G. (1992a) Oxygen isotopes in biogenic silica: global changes in ocean temperature and isotopic composition. Science, 256: 1434-1436. [46] Shemesh A., Ron H., Erel Y., Kolodny Y. & Nur A. (1992b) Isotopic composition of vein calcite and its fluid inclusions: implication to paleohydrologica systems, tectonic events and vein formation processes. Chem. Geol., 94: 307-314. [3] Sheppard S. M. F., Nielsen R. L. & Taylor H. P. Jr. (1969) Oxygen and hydrogen isotope ratio of clay minerals from porphyry copper deposits. Econ. Geol., 64: 755-777. [2, 46] Sheppard S. M. F. & Schwarcz H. P. (1970) Fractionation of carbon and oxygen isotopes and magnesium between coexisting metamorphic calcite and dolomite. Contrib. Mineral. Petrol., 26: 161-198. [46] Sheppard S. M. F., Nielsen R. L. & Taylor H. P. Jr. (1971) Hydrogen and oxygen isotope ratios in minerals from porphyry copper deposits. Econ. Geol., 66: 515-542. [2] Sheppard S. M. F. (1980) Isotopic evidence for the origins of water during metamorphic processes in oceanic crust and ophiolite complexes. In: Associations mafiques ultramafiques dans les orog6nes, CNRS, Colloq. Intern. No. 272: 135-147. [46] Sheppard S. M. F. (1981) Stable isotope geochemistry of fluids. In: Chemistry and geochemistry of solutions at high temperatures and pressures. D. T. Rickard & F. E. Wickman (eds.). Physics and Chemistry of the Earth, Pergamon Press, Vol. 13 & 14: 419-445. [46] Sheppard S. M. F. & Gilg H. A. (1996) Stable isotope geochemistry of clay minerals. Clay Minerals, 31: 1-24. [2, 46] Sherwood B., Fritz P., Frape S. K., Macko S. A., Weise S. M. & Whelhan J. A. (1988) Methane occurences in the Canadian Shield. Chem. Geol., 71: 223-236. [16] Sheu D. D. & Chiu C. H. (1995) Evaluation of cellulose extraction procedures for stable carbon isotope measurement in tree ring research. Int. J. Environ. Anal. Chem., 59: 59-67. [24] Shields W. R., Murphy T. J., Garner E. L. & Dibeter V. H. (1962) Absolute isotopic abundance ratios and the isotopic weight of chlorine. J. Am. Chem. Soc., 84: 1519-1522. [28, 40]

1156

References

Shields W. R., Goldich S. S., Garner E. L. & Murphy T. J. (1965) Natural variations in the abundance ratio and the atomic weight of copper. J. Geophys. Res., 70: 479-491. [31] Shields W. R., Murphy T. J., Catanzaro E. J. & Garner E. L. (1966) Absolute isotopic abundance ratios and the atomic weight of a reference sample of chromium. J. Res. Nat. Bur. Stand., 70A: 193-197. [29] Shima M. (1962) Boron in meteorites. J. Geophys. Res., 67: 4521-4523. [7] Shima M. & Honda M. (1963) Isotopic abundance of meteoritic lithium. J. Geophys. Res., 68: 2849-2854. [6] Shimamura T., Takahashi T., Honda M. & Nagai H. (1993) Multi-element and isotopic analyses of iron meteorites using a glow discharge mass spectrometer. J. Anal. At. Spectrom., 8: 453-460. [35] Shimizu N. & Hart S. R. (1982a) Applications of the ion microprobe to geochemistry and cosmochemistry. Ann. Rev. Earth Planet. Sci., 10: 483-526. [6, 30] Shimizu N. & Hart S. R. (1982b) Isotope fractionation in secondary ion mass spectrometry. J Appl. Phys., 53(3): 1303-1311. [30] Shmulovich K., Landwehr D., Simon K. & Heinrich W. (1999) Stable isotope fractionation between liquid and vapour in water-salt systems up to 600~ Chem. Geol., 157: 343-354. [46] Shouakar-Stash O., Drimmie R., Morrison J., Frape S. K., Heemskerk A. R. & Mark W. A. (2000) Online D / H analysis for water, natural gas and organic solvents by manganese reduction. Anal. Chem., 72: 2664-2666. [1] Shukolyukov A. & Lugmair G. W. (1998) Isotopic evidence for the Cretaceous-Tertiary impactor and its type. Science, 282(No. 5390): 927-929. [29] Sickman J. O., Leydecker A., Chang C. C. Y., Kendall C., Melack J. M., Lucero D. M. & Schimel J. (2003) Mechanisms underlying export of N from high-elevation catachments during seasonal transitions. Biogeochem., 64: 1-24. [15] Siebert C., N/igler T. F. & Kramers J. D. (2001) Determination of molybdenum isotope fractionation by double-spike multicollector inductively coupled plasma mass spectrometry. Geochem. Geophys. Geosys., 2, 2000GC000124. [31] Siegfriedt R. K., Wiberley J. S. & Moore R. W. (1951) Determination of sulfur after combustion in a small oxygen bomb. Rapid titrimetric method. Anal. Chem., 23: 1008-1011. [26] Sigman D. M., Altabet M. A., Michener R., McCorckle D. C., Fry B. & Holmes R. M. (1997) Natural abundance-level measurement of the nitrogen isotopic composition of oceanic nitrate: an adaptation of the ammonia diffusion method. Marine Chem., 57: 227-242. [15] Sigman D. M., Casciotti K. L., Andreani M., Barford C., Galanter M. & Bi~hlke J. K. (2001) A bacterial method for the nitrogen isotopic analysis of nitrate in seawater and freshwater. Anal. Chem., 73: 41454153.[15] Silfer J. A., Engel M. H., Macko S. A. & Jumeau E. J. (1991) Stable carbon isotope analysis of amino acid enantiomers by conventional isotope ratio mass spectrometry and combined gas chromatography/isotope ratio mass spectrometry. Anal. Chem., 63: 370-374. [9] Silva S. R., Kendall C., Wilkison D. H., Ziegler A. C., Chang C. C. Y. & Avanzino R. J. (2000) A new method for collection of nitrate from fresh water and the analysis of nitrogen and oxygen isotope ratios. J. Hydrol., 228: 22-36. [15] Silver J. A. (1992) Frequency-modulation spectroscopy for trace gas detection: theory and comparison among experimental methods. Appl. Opt., 31: 707-717. [34] Silver L. A. & Stolper E. M. (1989) Water in albitic glasses. J. Petrol., 30: 667-709. [46] Silverman S. R. (1951) The isotope geology of oxygen. Geochim. Cosmochim. Acta, 2: 26-42. [20] Silverman S. R. (1964) Investigations of petroleum origin and evolution mechanisms by carbon isotope studies. In: Isotopic cosmic chemistry. H. Craig, G. J. Miller & G. J. Wasserburg (eds.). North Holland: 92-102. [40] Simon K. (2001) Does 6D from fluid inclusion in quartz reflect the original hydrothermal fluid? Chem. Geol., 177: 483-495. [3] Simpson I. A., Evershed R. P., Bull I. D. & Dockrill S. J. (1997) Identification of manuring practices associated with early anthropogenic soil formation in Orkney, Scotland. Abstr. of Papers Am. Chem. Soc., 213: 54-GEOC. [8]

References

1157

Skorupa J. P. (1998) Selenium poisening of fish and wildlife in nature: Lessons from twelve real-worldexamples. In: Environmental chemistry of selenium. W. T. Jr. Frankenberger & R. A. Engberg (eds.). Marcel Dekker Inc., Vol. 64: 315-354. [29] Skulan J., DePaolo D. J. & Owens T. L. (1997) Biological control of calcium isotopic abundances in the global calcium cycle. Geochim. Cosmochim. Acta, 61(12): 2505-2510. [29] Slodzian G., Chaintreau M. & Dennebouy R. (1986) The emission objective lens working as an electron mirror: self regulated potential at the surface of an insulating sample. In: Secondary ion mass spectrometry. A. Benninghoven, D. S. Colton, D. S. Simons & H. W. Werner (eds.). Springer, SIMS V: 158-160. [30] Slodzian G., Chaintreau M., Dennebuoy R. & Rousse A. (2001) Precise in situ measurements of isotopic abundances with pulse counting of sputtered ions. Eur. Phys. J. Appl. Phys., 14: 199-231. [30] Smith B. F. L. (1994) Characterisation of poorly ordered minerals by selective chemical methods. In: Clay mineralogy: Spectroscopy and chemical determinative methods. M. J. Wilson (ed.). Chapman & Hall, London: 333-357. [2] Smith B. N. & Epstein S. (1971) Two categories of 13C/12C ratios for higher plants. Plant Physiol., 47: 380-384. [46] Smith D., Adams N. G. & Henchman M. J. (1980) Studies of the binary reactions of H30+.(H20)0,1,2 ions and their deuterated analogues with D20, H20 and NH3. J. Chem. Phys., 72: 4951-4957. [4] Smith D. & Adams N. G. (1988) The selected ion flow tube (SIFT). Studies of ion-neutral reactions. Adv. Atom. Mol. Phys., 24: 1-49. [4] Smith D. & Spanel P. (1996a) Application of ion chemistry and the SIFT technique to the quantitative analysis of trace gases in air and on breath. Int. Reviews Phys. Chem., 15: 231-271. [4] Smith D. & Spanel P. (1996b) The novel selected-ion flow tube approach to trace gas analysis of air and breath. Rapid Commun. Mass Spectrom., 10:1183-1198 [4]. Smith D., Wang T. S. & Spanel P. (2002a) Kinetics and isotope patterns of ethanol and acetaldehyde emissions from yeast fermentations of glucose and glucose-6,6-d2 using selected ion flow tube mass spectrometry: a case study. Rapid Commun. Mass Spectrom., 16: 69-76. [4] Smith D., Wang T. S. & Spanel P. (2002b) On-line, simultaneous quantification of ethanol, some metabolites and water vapour in breath following the ingestion of alcohol. Physiol. Meas., 23, 477-489. [4] Smith J. W., Young N. B. & Lawlor D. L. (1964) Direct determination of sulphur forms in Green River oil shale. Anal. Chem., 36:618-622. [26] Smith J. W. & Kaplan I. R. (1970) Endogenous carbon in carbonaceous meteorites. Science, 167: 13671370.[11] Smith J. W. &Batts B. D. (1974) The distribution and isotopic composition of sulphur in coal. Geochim. Cosmochim. Acta, 38: 121-133. [26] Smith J. W. & Croxford R. J. W. (1975) An isotopic investigation of the environment of deposition of the McArthur mineralisation. Miner. Deposita, 10: 269-276. [2] Smith M.-A. H., Rinsland C. P., Fridovich B. & Rao K. N. (1985) Intensities and collision broadening parameters from infrared spectra. In: Molecular spectroscopy: Modern research. K. N. Rao (ed.). Academic Press, Orlando, Vol. III. [34] Smith M. P. & Yardley B. W. D. (1996) The boron isotopic composition of tourmaline as a guide to fluid processes in the southwestern England orefield: An ion microprobe study. Geochim. Cosmochim. Acta, 60: 1415-1427. [30] Smithwick R. W., Lynch D. W. & Franklin J. C. (1993) Relative ion yields measured with a high-resolution glow discharge mass spectrometer operated with an Argon/Hydrogen mixture. J. Am. Soc. Mass Spectrom., 4: 278-285. [35] Smykatz-Kloss W. (1974) Differential thermal analysis. Springer, Berlin: 185 p. [13, 16] So C.-S., Rye D. M. & Shelton K. L. (1983) Carbon, hydrogen, oxygen, and sulphur isotope and fluid inclusion study of the Weolog tungsten-molybdenum deposit, republic of Korea: Fluid histories of metamorphic and ore-forming events. Econ. Geol., 78: 1551-1573. [3] Socki R. A., Karlsson H. R. & Gibson E. K. Jr. (1992) Extraction technique for the determination of oxygen-18 in water using preevacuated glass vials. Anal. Chem., 64: 829-831. [1] Socki R. A., Romanek C. S. & Gibson E. K. Jr. (1999) On-line technique for measuring stable oxygen and hydrogen isotopes from microliter quantities of water. Anal. Chem., 71: 2250-2253. [1, 20]

1158

References

Sofer Z. & Gat J. R. (1972) Activities and concentrations of oxygen-18 in concentrated aqueous salt solutions: Analytical and geophysical implications. Earth Planet. Sci. Lett., 15: 232-238. [46] Sofer Z. & Gat J. R. (1975) The isotope composition of evaporating brines: Effect of the isotopic activity ratio in saline solutions. Earth Planet. Sci. Lett., 26: 179-186. [46] Sohns E., Gerling P. & Faber E. (1994) Improved stable nitrogen isotope ratio measurements of natural gases. Anal. Chem., 66: 2614-2620. [8, 43] Solarzano L. (1969) Determination of NH4 + in natural water by the phenolhypochlorite method. Limnol. Oceanogr., 14: 799-801. [9] Song T. R. & Ding T. P. (1990) A new probe on application of silicon isotopes composition (630Si) in siliceous rocks to sedimentary facies analysis. Chinese Sci. Bull., 35: 761-766. [25] Sorokina I. T., Sorokin E., DiLieto A., Tonelli M., Page R. H. & Schaffers K. I. (2001) Efficient broadly tunable continuous-wave Cr2+:ZnSe laser. J. Opt. Soc. Am., B 18: 926-930. [34] Sorensen P. & Jensen E. S. (1991) Sequential diffusion of ammonium and nitrate from soil extracts to a polytetrafluoroethylene trap for 15N determination. Anal. Chim. Acta, 252: 201-203. [15] Si~derholm S., Meinander N. & Kauppinen J. (1998) The 13C/12C isotopic ratio determined by FTIR spectroscopy. AIP Conf. Proc., No. 430: 223-226. [34] Si~derlund R. & Svensson B. H. (1976) The global nitrogen cycle. Ecol. Geol. (Stockholm), 22(SCOPE Report 7), 23073. [15] Spanel P. & Smith D. (1996) The selected ion flow tube, a novel technique for the quantitative trace gases analysis of air and human breath. Med. Biol. Eng. Comp., 34: 409-419. [4] Spanel P., Davies S. & Smith D. (1998) Quantification of ammonia in human breath by the selected ion flow tube analytical method using H30 + and 02 + precursor ions. Rapid Commun. Mass Spectrom., 12: 763-766. [4] Spanel P., Smith D., Holland T. A., A1 Singary W. & Elder J. B. (1999) Analysis of formaldehyde in the headspace of urine from bladder and prostate cancer patients using selected ion flow tube mass spectrometry. Rapid Commun. Mass Spectrom., 13: 1354-1359. [4] Spanel P. & Smith D. (2000) Selected flow tube mass spectrometry analyses of stable isotopes in water: Isotopic composition of H30 + and H30+(H20)3 ions in exchange reactions with water vapor. J. Am. Soc. Mass Spectrom., 11: 866-875. [4] Spanel P. & Smith D. (2001) Accuracy and precision of flowing afterglow mass spectrometry for the determination of the deuterium abundance in hte headspace of aqueous liquids and exhaled breath water. Rapid Commun. Mass Spectrom., 15: 867-872. [4] Sparks J. P. & Ehleringer J. R. (1997) Leaf carbon isotope discrimination and nitrogen content for riparian trees along elevational transects. Oecologia, 109: 362-367. [9] Spicuzza M. J., Valley J. w., Kohn M. J., Girard J. P. & Fouillac A. M. (1998a) The rapid heating, defocused beam technique: a CO2-1aser-based method for highly precise and accurate determination of 6180 values of quartz. Chem. Geol., 144: 195-203. [20] Spicuzza M. J., Valley J. W. & McConnell V. S. (1998b) Oxygen isotope analysis of whole rock via laser fluorination: An air-lock approach. Geol. Soc. Am. Abstr. w. Prog., 30(7): 80 (Abstract). [2, 20] Spivack A. J. & Edmond J. M. (1986) Determination of boron isotope ratios by thermal ionization mass spectrometry of the dicesium metaborate cation. Anal. Chem., 58: 31-35. [7, 40] Spivack A. J. & Edmond J. M. (1987) Boron isotope exchange between seawater and the oceanic crust. Geochim. Cosmochim. Acta, 51: 1033-1043. [40] Spivack A. J., Palmer M. R. & Edmond J. M. (1987) The sedimentary cycle of the boron isotopes. Geochim. Cosmochim. Acta, 51: 1939-1949. [7] Spivack A. J., You C.-F. & Smith H. J. (1993) Foraminiferal boron isotope ratios as a proxy for surface ocean pH over the past 21 Myr. Nature, 363: 149-151. [7] Spivack A. J. & You C.-F. (1997) Boron isotopic geochemistry of carbonates and pore waters, Ocean Drilling Program Site 851. Earth Planet. Sci. Lett., 152: 113-122. [7] Spooner N., Rieley G., Collister J. W., Lander M., Cranwell P. A. & Maxwell J. R. (1994) Stable carbon isotopic correlation of individual biolipids in aquatic organisms and a lake bottom sediment. Org. Geochem., 21: 823-827. [9]

References

1159

Spratt H. G. & Morgan M. D. (1990) Sulfur cycling over an annual cycle in a cedar dominated freshwater wetland. Limnol. Oceanogr., 35: 1586-1593. [26] Sprinson D. B. & Rittenberg D. (1948) Preparation of gas samples for mass spectrometric analysis of isotope abundance. U. S. Naval Med. Bull., Supplement 115 (March-April): 82-93. [15] Sprinson D. B. & Rittenberg D. (1949) The rate of utilization of ammonia for protein synthesis. J. Biol. Chem., 180: 707-714. [15] Sridhar K., Jackson M. L. & Clayton R. N. (1975) Quartz oxygen isotopic stability in relation to isolation from sediments and diversity of source. Soil Sci. Soc. Am. Proc., 39: 1209-1213. [2] Stahl W. J. (1977) Carbon and nitrogen isotopes in hydrocarbon research and exploration. Chem. Geol., 20: 121-149. [26] Stanley K., Lyon G. L. & Stewart M. K. (1984) Hot shot reduction of water to hydrogen for isotopic analysis. Institute of Nuclear Sciences, DSIR, Lower Hutt, New Zealand. Rept. INS-R--322, February, 1984: llp. [1] Stapp P., Polis G. A. & Pinero S. (1999) Stable isotopes reveal strong marine and E1 Nifio effects on island food webs. Nature 401, 467-469 [9]. Stark J. M. & Hart S. C. (1996) Diffusion technique for preparing salt solutions, Kjeldahl digests, and persulfate digests for nitrogen-15 analysis. Soil Sci. Soc. Am. J., 60: 1846-1855. [15] Staudt W. J. & Schoonen M. A. A. (1994) Sulfate incorporation into sedimentary carbonates. In: Geochemical transformations of sedimentary sulfur. M. A. Vairavamurthy, M. A. A. Schoonen, T. I. Eglinton, G. W. Luther, and B. Manowitz (eds.). American Chemical Society, Washington. ACS Symposium Series 612: 332-345. [26] Steele L. P., Dlugokencky E. J., Lang P. M., Tans P. P., Martin R. C. & Masarie K. A. (1992) Slowing down of the global accumulation of atmospheric methane during the 1980's. Nature, 358(6384): 313316. [14] Stehr J., Krankowsky D. & Mauersberger K. (1996) Collection and analysis of atmospheric ozone samples. J. Atmos. Chem., 24: 317-325. [18] Steiner R. E., Lewis C. L. & King F. L. (1997) Time-of-flight mass spectrometry with a pulsed glow discharge ionization source. Anal. Chem., 69: 1715-1721. [35] Steinfeld J. I., Adler-Golden S. M. & Gallagher J. W. (1987) Chem. Ref. Data, 16(4): 911. [18] Stephan E. (2000) Oxygen isotope analysis of animal bone phosphate: method refinement, influence of consolidants, and reconstruction of paleotemperatures for Holocene sites. J. Archaeol. Sci., 27: 523535. [22] Stern L. A., Chamberlain C. P., Reynolds R. C. & Johnson G. D. (1997) Oxygen isotope evidence of climate change from pedogenic clay minerals in the Himalayan molasse. Geochim. Cosmochim. Acta, 61: 731-744. [2] Sternberg L. d. S. L., DeNiro M. J. & Savidge R. A. (1986) Oxygen isotope exchange between metabolites and water during biochemical reactions leading to cellulose synthesis. Plant Physiol., 82: 423427. [23] Stevens C. M. & Rust F. E. (1982) The carbon isotopic composition of atmospheric methane. J. Geophys. Res., 87(C7): 4879-4882. [14] Stevens C. M. (1995) Carbon-13 isotopic abundance and concentration of atmospheric methane for background air in the southern and northern hemispheres from 1978 to 1989. Environmental Science Division, U. S. Dept. of Energy, Publ. 4388. [14] Stevens W. B., Mulvaney R. L., Khan S. A. & Hoeft R. G. (2000) Improved diffusions methods for nitrogen and 15 nitrogen analysis of kjeldahl digests. J. Assoc. Off. Anal. Chem. Int., 38: 1039-1046. [15] Stevens W. H. (1971) Catalyst for hydrogen-deuterium exchange. Ger. Often. Rept. 2104019: 22p. [1] Stevenson E J. (1982) Nitrogen - Organic forms. In: Methods of soil analysis, Part 2. Chemical and microbiological properties. A. L. Page, R. H. Miller & D. R. Keeney (eds.). Am. Assoc. of Agronomy Inc- Soil Sci. Soc. of Amer. Inc., Madison, Wisconsin. Agronomy Series, No.9, 2nd edition: 625-641. [15] Stevenson F. J. (1996) Nitrogen - Organic forms. In: Methods of soil analysis. D. L. Sparks (ed.). Soil Sci. Soc. of Amer. Inc., Madison, Wisconsin. Ser. 5, Part 3: 1185-1200. [15] Stewart M. K. & Friedman I. (1975) Deuterium fractionation between aqueous salt solutions and water vapor. J. Geophys. Res., 80: 3812-3818. [46]

1160

References

Stewart M. K. & James M. R. (1981) Conversion of water to hydrogen for deuterium assay using uranium at 700~ Inst. Nucl. Sci., DSIR, Lower Hutt, New Zealand. Rept. INS-R-292. [1, 33] Stewart M. K. (1981) 180 and D enrichment by evaporation from sample containers. Int. J. Appl. Radiat. Isot., 32: 159-163. [1] Stichler W. (1995) Interlaboratory comparison of new materials for carbon and oxygen isotope measurements. In: Refrence and intercomparison materials for stable isotopes of light elements. IAEA-TECDOC-825, IAEA, Vienna: 67-74. [43] Stickrod R. D. & Marshall J. D. (2000) On-line nitrate-615N extracted from groundwater determined by continuous-flow elemental analyzer/isotope ratio mass spectrometry. Rapid Commun. Mass Spectrom., 14: 1266-1268. [15] St-Jean G. (2003) Automated quantitative and isotopic (C-13) analysis of dissolved inorganic carbon and dissolved organic carbon in continuous-flow using a total organic carbon analyzer. Rapid Commun. Mass. Spectrom., 17: 419-428. [10] Stoffregen R. E., Rye R. O. & Wasserman M. D. (1994a) Experimental studies of alunite: I. 180-160 and D-H fractionation factors between alunite and water at 250-450~ Geochim. Cosmochim. Acta, 58(2): 903-916. [46] Stoffregen R. E., Rye R. O. & Wasserman M. D. (1994b) Experimental studies of alunite: II. Rates of alunite-water alkali and isotopic exchange. Geochim. Cosmochim. Acta, 58: 917-929. [46] Stoffregen R. (1996) Numerical simulation of mineral-water isotope exchange via Ostwald ripening. Am. J. Sci., 296: 908-931. [46] Stoffyn-Egli P. & Mackenzie F. T. (1984) Mass balance of dissolved lithium in the oceans. Geochim. Cosmochim. Acta, 48: 859-872. [6] Stolper E. (1982) The speciation of water in silicate melts. Geochim. Cosmochim. Acta, 46: 2609-2620. [46] Stolper E. & Epstein S. (1991) An experimental study of oxygen isotope partitioning between silica glass and CO2 vapor. In: Stable isotope geochemistry: A tribute to Samuel Epstein. H. P. Jr. Taylor, J. R. O'Neil & I. R. Kaplan (eds.). The Geochemical Society, Special Publication No. 3: 35-51. [46] Stone J., Hutcheon I. D., Epstein S. & Wasserburg G. J. (1991) Silicon, carbon and nitrogen isotope studies of silicon carbide in carbonaceous ad enstatite chondrites. In: Stable isotope geochemistry: A tribute to Samuel Epstein. H. P. Jr. Taylor, J. R. O'Neil & I. R. Kaplan (eds.). The Geochemical Society, Special Publication No. 3: 487-504. [25] Stott A. W., Davies E., Evershed R. P. & Tuross N. (1997a) Monitoring the routing of dietary and biosynthesised lipids through compound-specific stable isotope (613C) measurements at natural abundance. Naturwissenschaften, 84: 82-86. [9] Stott A. W., Evershed R. P. & Tuross N. (1997b) Compound-specific approach to the ~)13C analysis of cholesterol in fossil bones. Org. Geochim., 26: 99-103. [9] Strauss H. (1997) The isotopic composition of sedimentary sulfur through time. Palaeogeogr., Palaeo-clim., Palaeoecol., 132: 97-118. [27] Strebel O., BSttcher J. & Fritz P. (1990) Use of isotope fractionation of sulfate-sulfur and sulfate-oxygen to assess bacterial desulfurication in a sandy aquifer. J. Hydrol., 121(1-4): 155-172. [29] Strelow, F. W. E. (1973) Application of ion exchange to element separation and analysis. In: Ion exchange and solvent extraction. J. A. Marinsky & Y. Marcus (eds.). Marcel Dekker. [29] Strelow F. W. E., Weinert C. H. S. W. & van der Walt T. N. (1974) Separation of lithium from sodium, beryllium and other elements by cation-exchange chromatography in nitric acid-methanol. Anal. Chim. Acta, 71: 123-132. [6] Strickland J. D. H. & Parsons T. R. (1972) A practical handbook of seawater analysis. Bull. Fish. Res. Bd. Can., 167: 1-130. [9] Stuart-Williams H. L. Q. & Schwarcz H. P. (1995) Oxygen isotopic analysis of silver orthophosphate using a reaction with bromine. Geochim. Cosmochim. Acta, 58: 3837-3841. [12, 22] Stuart-Williams H. L. Q. & Schwarcz H. P. (1997) Oxygen isotopic determination of climatic variation using phosphate from beaver bone, tooth enamel, and dentine. Geochim. Cosmochim. Acta, 61: 25392550. [22] Studley S. A., Ripley E. M., Elswick E. R., Dorais M. J., Fong J., Finkelstein D. & Pratt L. M. (2002) Analysis of sulfides in whole rock matrices by elemental analyzer - continuous flow isotope ratio mass spectrometry. Chem. Geol., 192(1-2): 141-148. [26, 45]

References

1161

Stuiver M. & Braziunas T. F. (1987) Tree cellulose 13C/12C isotope ratios and climatic change. Nature, 328: 58-60. [9] Stuiver M., Grootes P. M. & Braziunas T. F. (1995) The GISP2 6180 climate record of the past 16,500 years and the role of the sun, ocean, and volcanoes. Quater. Research, 44: 341-354. [1] Stull D. R. (1947) Vapour pressure of pure substances. Inorganic compounds. Ind. Eng. Chem., 39: 540-

550. [28]

Sturchio N. C. & C h a n L. H. (2003) Lithium isotope geochemistry of the Yellowstone hydrothermal system. Soc. Econ. Geol., Spec. Publ., 10: 171-180. [6] Stiirup S., Hansen M. & Molgaard C. (1997) Measurement of 44Ca:43Ca and 42Ca:43Ca isotopic ratios in urine using high resolution inductively coupled plasma mass spectrometry. J. Anal. Atom. Spectrom., 12: 919-923. [31] Sudzuki N. (1987) A water conversion method for D / H ratio analyes and its accuracy. Geochem. J., 21: 29-33. [1] Sugiura N., Hashizume K., Miura Y., Kiyota K. & Zashu S. (1995) Cosmogenic nitrogen isotopic disequilibrium observed in meteorites. Meteoritics, 30: 646-653. [17] Summons R. E., Jahnke L. L. & Roksandic Z. (1994) Carbon isotopic fractionation in lipids from methanotrophic bacteria: relevance for interpretation of the geochemical record of biomarkers. Geochim. Cosmochim. Acta, 58: 2853-2863. [46] Summons R. E., Franzmann P. D. & Nichols P. D. (1998) Carbon isotopic fractionation associated with methylotrophic methanogenesis. Org. Geochem., 28: 465-475. [9] Supplee J. M., Whittaker E. A. & Lenth W. (1994) Theoretical description of frequency modulation and wavelength modulation spectroscopy. Appl. Opt., 33: 6294-6302. [34] Suryanarayana M. V., Sankari M. & Gangadhara S. (1998) Determination of Li-6/Li-7 isotope ratio using two photon resonance three photon resonance ionization mass spectrometry. Int. J. Mass Spectrom., 173: 177-189. [33] Sutka R. L., Ostrom N. E., Ostrom P. H., Gandhi H. & Breznak J. A. (2003) Nitrogen isotopomer site preference of N20 produced by Nitrosomonas europaea and Methylococcus capsulatus Bath. Rapid Commun. Mass Spectrom., 17: 738-745. [19] Suvorova V. A. & Dubinina Y. O. (1996) Determinating oxygen isotope compositions on small amounts of water. Geochem. Int., 33: 50-54. [1] Suzuoki T. & Epstein S. (1976) Hydrogen isotope fractionation between OH-bearing minerals and water. Geochim. Cosmochim. Acta, 40: 1229-1240. [46] Svec H. J. & Anderson A. R. Jr. (1965) The absolute abundance of the lithium isotopes in natural sources. Geochim. Cosmochim. Acta, 29: 633-641. [6] Swart P. K. (1983) Carbon and oxygen isotope fractionation in scleractinian corals: A review. Earth Sci. Rev., 19: 51-80. [3] Swart P. K., Burns S. J. & Leder J. J. (1991.) Fractionation of the stable isotopes of oxygen and carbon in carbon dioxide during the reaction of calcite with phosphoric acid as a function of temperature and technique. Chem. Geol., 86: 89-96. [12] Swart P. K., Leder J. J., Szmant A. M. & Dodge R. E. (1996) The origin of variations in the isotopic record of scleractinian corals: II. Carbon. Geochim. Cosmochim. Acta, 60: 2871-2885. [46] Swihart G. H. (1996) Instrumental techniques for boron isotope analysis. In: Boron. Mineralogy, petrology and geochemistry. E. S. Grew & L. M. Anovitz (eds.). Mineralogical Society of America, Reviews in Mineralogy, Vol. 33: 845-862. [7] Switsur R., Waterhouse J. S., Field E. M., Carter T. & Loader N. (1995) Stable isotope studies in tree rings from o a k - technique and some preliminary results. In: Problems of stable isotopes in treerings, lake sediments and peat-bogs as climatic evidence for the Holocene. B. Frenzel (ed.). Gustav Fischer Verlag, Stuttgart, Vol. 15: 129-140. [24] Syers J. K., Chapman S. L., Jackson M. L., Rex R. W. & Clayton R. N. (1968) Quartz isolation from rocks, sediments and soils for determination of oxygen isotopes composition. Geochim. Cosmochim. Acta, 32: 1022-1025. [2] Szaran J. (1996) Experimental investigation of sulphur isotopic fractionation between dissolved and gaseous H2S. Chem. Geol., 127: 223-228. [26]

1162

References

Szostak R. (1989) Molecular sieves: Principles of synthesis and identification. Van Nostrand Reinhold, New York: 524 p. [36] Tabatabai M. A. & Bremner J. M. (1970) An alkaline oxidation method for determination of total sulfur in soils. Soil Sci. Soc. Am. Proc., 34: 62-65. [26] Tabatabai M. A. (1982) Sulfur. In: Methods of soil analysis, part 2. Chemical and microbiological properties. A. L. Page, R. H. Miller & D. R. Keeney (eds.). American Association of Agronomy Inc. - Soil Science Society of America Inc., Agronomy Series, No. 9: 501-538. [26] Tabatabai M. A. (1992) Methods of measuring of sulphur in soils, plant materials, and waters. In: Sulphur cycling on the continents: Wetlands, terrestrial ecosystems and associated water bodies. R. W. Howarth, J. W. B. Stewart & M. V. Ivanov (eds.). John Wiley & Sons, New York: 307-344. [26] Taieb R. (1990) Les isotopes de l'hydrogene, du carbone et de l'oxygene dans les s6diments argileux et les eaux de formation. Th6se, Institut national Polytechnique de Lorraine, Nancy, France: 141 p. [2] Takahata N., Nishio Y., Yoshida N. & Sano Y. (1998) Precise isotopic measurements of nitrogen at the sub-nanomole level. Anal. Sci., 14: 485-491. [17] Takala K., Olkkonen H. & Krouse H. R. (1991) Sulphur isotope composition of epiphytic and terricolous lichens and pine bark in Finland. Environ. Pollut., 69: 337-348. [26] Tan T. C., Pearson G. J. & Walker R. W. (1973) Sampling extraction and 13C/12C analysis of total dissolved CO2 in marine environments. Bedford Institute of Oceanography, Canada. Report Series, BI-R-73-16:17 p. [10] Tanaka N. & Rye D. M. (1991) Chlorine in the stratosphere. Nature, 353: 707. [28] Tanaka N., Rye D. M., Xiao u & Lasaga A. C. (1994) Use of stable sulfur isotope systematics for evaluating oxidation reaction pathways and in-cloud-scavenging of sulfur dioxide in the atmosphere. Geophys. Res. Lett., 21(14): 1519-1522. [26] Tanaka N., Rye D. M., Rye R., Avak H. & Yoshinari T. (1995) High precision mass spectrometric analysis of isotopic abundance ratios in nitrous oxide by direct injection of N20. Int. J. Mass Spectrom. Ion Proc., 142: 163-175. [19] Tang D., Yang Q., Zhou C., Kang X., Liu D. & Huang W. (2001) Genetic relationships between swamp microenvironment and sulfur distribution of the late Paleozoic coals in North China. Science in China (Series D), 44: 555-565. [26] Tang K., Feng X. & Ettl G. J. (2000) The variation in dD of tree rings and the implications for climatic reconstruction. Geochim. Cosmochim. Acta, 64: 1663-1673. [46] Tanner S. D. (1992) Space charge in ICP-MS: Calculation and implications. Spectrochim. Acta, 47B: 809823. [6] Tans P. P., de Jong A. F. M. & Mook W. G. (1978) Chemical pretreatment and radial flow of 14C in tree rings. Nature, 271: 234-235. [24] Tans P. P., Fung I. Y. & Takahashi T. (1990) Observational constraints on the global atmospheric carbon dioxide budget. Science, 247: 1431-1438. [14] Tans P. P., Berry J. A. & Keeling R. F. (1993) Oceanic 13C/12Cobservations: A new windowq on ocean CO2 uptake. Global Biogeochem. Cycles, 7: 353-368. [14] Tanweer A., Hut G. & Burgman J. O. (1988) Optimal conditions for the reduction of water to hydrogen by zinc for mass spectrometric analysis of deuterium content. Chem. Geol. (Isot. Geosc. Sect.), 73(2): 199-203. [1] Tanweer A. (1990) Importance of clean metallic zinc for hydrogen isotope analysis. Anal. Chem., 62: 2158-2160.[1] Tanweer A. (1993a) Investigation of BDH zinc for hydrogen isotope analysis by mass spectrometry. Z. Naturforsch., 48a: 739-740. [1] Tanweer A. (1993b) Determination of deuterium in brines and in hypersaline aqueous solutions by mass spectrometry using zinc as reducing agent. Analyst, 118: 835-838. [1] Tanweer A. & H a n L.-F. (1996) Reduction of microliter amounts of water with manganese for D / H isotope ratio measurement by mass spectrometry. Isot. Environ. Health Stud., 31: 1-7. [1] Tanzer D. & Heumann K. G. (1991) Determination of dissolved selenium species in environmental water samples using isotope dilution mass spectrometry. Anal. Chem., 63: 1984-1989. [29] Taube H. (1954) Use of oxygen isotopes effects in the study of hydration of ions. J. Phys. Chem., 58: 523528. [46]

References

1163

Taylor B. E. & O'Neil J. R. (1977) Stable isotope studies of metasomatic Ca-Fe-A1-Si skarns and associated metamorphic and igneous rocks, Osgood Mountains, Nevada. Contrib. Mineral. Petrol., 63: 1-49. [20] Taylor B. E., Foord E. E. & Friedrichsen H. (1979) Stable isotope and fluid inclusion studies of gembearing granitic pegmatite-aplite dikes, San Diego Co., California. Contrib. Mineral. Petrol., 68: 187205. [3] Taylor B. E., Eichelberger J. C. & Westrich H. R. (1983) Hydrogen isotopic evidence of rhyolitic magma degassing during shallow intrusion and eruption. Nature, 306: 541-545. [46] Taylor B. E. & Beaudoin G. (1993) MILES laser microprobe. Part I: System description. Geol. Surv. Canada. Current Research, Part D, Paper 1993-ID: 191-198. [20, 26] Taylor B. E. & Beaudoin G. (2000) Sulphur stratigraphy of the Sullivan Pb-Zn-Ag deposit, B.C: Evidence for hydrothermal sulphur, and bacterial and thermochemical sulphate reduction. In: The Sullivan deposit and its geological environment. J. W. Lydon, T. H6y, J. F. Slack & M. Knapp (eds.). Mineral Deposits Division of the Geological Association of Canada. Spec. Publ. No. 1: 696-719. [20] Taylor B. E., Ding T., Halas S., Br6as O. & Robinson B. W. (2001) Accurate calibration of the V-CDT scale: Proposed ~134Svalues for calibration reference materials and methods of correction for SO2based analyses. AGU Mtgn: $426 (Abstract). [20] Taylor B. E. (2004) Biogenic and thermogenic sulfate reduction in the Sullivan Pb-Zn-Ag deposit, B.C.: Evidence from micro-isotopic analysis of carbonate and sulfide in bedded ores. Chem. Geol.: in press [20] Taylor B. E. (in press) Laser-assisted micro-analysis of the sulfur isotope reference materials using the SF6-method: Calibration of the sulfur isotope reference scale and consistent discrepancies with SO2based measurements. IAEA, Vienna. TECDOC. [20, 26] Taylor B. E., Ding T., Halas S., Br6as O. & Robinson B. W. (in press) Accurate calibration of the V-CDT scale: proposed 634S values for calibration and reference materials and methods of correction for SO2 based analyses. IAEA, Vienna. [20, 26] Taylor C. B. (1973) Measurement of oxygen-18 ratios in environmental waters using the EpsteinMayeda technique. Part I: Theory and experimental details of the equilibration technique. Inst. Nucl. Sci., DSIR, Lower Hutt, New Zealand, Rept N. 556 (INS-LN-37), April, 1973:24 p. [1] Taylor C. B. & Fox V. J. (1996) An isotopic study of dissolved inorganic carbon in the catchment of the Waimakariri River and deep ground water of North Canterbury plains, New Zealand. J. Hydrol., 186: 161-190. [10] Taylor J. W. & Grimsrud E. P. (1969) Chlorine isotope ratios by negative ion mass spectrometry. Anal. Chem., 41: 805-810. [28] Taylor J. W. & Chen I.-J. (1970) Variables in oxygen-18 isotopic analysis by mass spectrometry. Anal. Chem., 42: 224-228. [23] Taylor H. E., Huff R. A. & Montaser A. (1998) Novel applications of ICPMS. In: Inductively coupled plasma mass spectrometry. A. Montaser (ed.). Wiley-VCH: 681-807. [31] Taylor H. P. Jr. & Epstein S. (1962) Relationship between O18/O16 ratios in coexisting minerals of igneous and metamorphic rocks. Part I: Principles and experimental results. Geol. Soc. Am. Bull., 73: 461-480. [2, 20, 25] Taylor H. P. Jr. (1968) The oxygen isotope geochemistry of igneous rocks. Contrib. Mineral. Petrol., 19: 171. [46] Taylor H. P. Jr. & Epstein S. (1970) Oxygen and silicon isotope ratios of lunar rocks. Earth Planet. Sci. Lett., 9: 298-310. [20] Taylor H. P. Jr. & Epstein S. (1973) 180/160 and 30Si/28Si studies of some Apollo 15, 16 and 17 samples. 4th Lunar Science Conference, Proceedings, Vol. 2: 1657-1679. [25] Taylor P. D. P., De Bi6vre P., Walder A. J. & Entwistle A. (1995) Validation of the analytical linearity and mass discrimination correction model exhibited by a multiple collector inductively coupled plasma mass spectrometer by means of a set of synthetic uranium isotope mixtures. J. Anal. At. Spectrom., 10: 395-398. [31, 32, 35] Taylor P. D. P. (1998) IRMM Certificate: Sulfur primary isotopic gas standard IRMM-PIGS-2010. Institute for Reference Materials and Measurements, Geel, Belgium: 2 p. [40]

1164

References

Taylor T. I. 6: Urey H. C. (1938) Fractionation of the lithium and potassium isotopes by chemical exchange with zeolites. J. Chem. Phys., 6: 429-438. [6] Tebo B. M. & Obraztova A. Y. (1998) Sulfate reducing bacterium grows with Cr(VI), U(VI), Mn(IV), and Fe(III) as electron acceptors. FEMS Microbiol. Lett., 162: 193-198. [29] Teece M. A., Fogel M. L., Dollhopf M. E. & Nealson K. H. (1999) Isotopic fractionation associated with biosynthesis of fatty acids by marine bacterium under oxic and anoxic conditions. Org. Geochim., 30: 1571-1579. [9] Tellier K. E., Hluchy M. M., Walker J. R. & Reynolds R. C. Jr. (1988) Application of high gradient magnetic separation (HGMS) to structural and compositional studies of clay mineral mixtures. J. Sedim. Petrol., 58: 761-763. [2] Teng J., Barshick C. M., Duckworth D. C., Morton S. J., Smith D. H. & King F. L. (1995) Factors influencing the quantitative determination of trace elements in soils by glow discharge mass spectrometry. Appl. Spectrosc., 49: 1361-1366. [35] Tennie A., Hoffbauer R. & Hoernes S. (1998) The oxygen isotope fractionation of kyanite in experiment and nature. Contrib. Mineral. Petrol., 133: 346-355. [20] Thielecke F. & Noack R. (1997) Evaluation of an automated equilibration technique for deuterium/hydrogen isotope ratio measurements with respect to assessing total energy expenditure by the doubly labelled water method. J. Mass Spectrom., 32: 323-327. [1] Thielecke F., Brand W. & Noak R. (1998) Hydrogen isotope determination for small-size water samples using ban equilibration technique. J. Mass Spectrom., 33: 342-345. [1] Thiemens M. H. & Heidenreich J. E. III (1983) The mass-independent fractionation of oxygen: A novel isotope effect and its possible cosmochemical implications. Science, 219: 1073-1075. [15, 18, 46] Thiemens M. H. & Jackson T. (1987) Production of isotopically heavy ozone by ultra-violet light photolysis of 02. Geophys. Res. Lett., 14: 624-627. [18] Thiemens M. H. & Jackson T. (1988) New experimental evidence for the mechanism for production of isotopically heavy 03. Geophys. Res. Lett., 15: 639-643. [18] Thiemens M. H. & Jackson T. L. (1990) Pressure dependency for heavy isotope enhancement in ozone formation. Geophys. Res. Lett., 17: 717-719. [18] Thiemens M. H. & Trogler W. C. (1991) Nylon production: An unknown source of atmospheric nitrous oxide. Science, 251: 932-934. [19] Thiemens M. H. (1992) Mass-independent isotopic fractionation and their applications. In: Isotope effects in gas-phase chemistry. J. A. Kaye (ed.). American Chemical Society, Washington DC: 138-154.

[18]

Thiemens M. H. (1999) Mass-independent isotope effects in planetary atmospheres and the early solar system. Science, 283: 341-345. [18, 46] Thiemens M. H. (2001) The mass-independent ozone isotope effect. Science, 283: 341-345. [46] Thiemens M. H., Savarino J., Farquhar J. & Bao H. (2001) Mass-independent isotopic compositions in terrestrial and extraterrestrial solids and their applications. Acc. Chem. Res., 34: 645-652. [15, 18] Thirlwall M. (2001) Inappropriate tail corrections can cause large inaccuracy in isotope ratio determination by MC-ICP-MS. J. Anal. At. Spectrom., 16: 1121-1125. [32] Thode H. G., MacNamara J., Lossing F. P. & Collis C. B. (1948) Natural variations in isotopic content of boron and its chemical atomic weight. J. Am. Chem. Soc., 70: 3008-3011. [7] Thode H. G., MacNamara J. & Collins C. B. (1949) Natural variations in the isotopic content of sulfur and their significance. Can. J. Res., 27B: 361-373. [44, 46] Thode H. G., Kleerekoper H. & McElchern D. (1951) Isotopic fractionation in the bacterial reduction of sulphate. Research (London), 4: 581-582. [46] Thode H. G., Monster J. & Dunford H. B. (1958) Sulphur isotope abundances in petroleum and associated minerals. AAPG Bull., 42: 2619-2641. [26] Thode H. G., Monster J. & Dunford H. B. (1961) Sulphur isotope geochemistry. Geochim. Cosmochim. Acta, 25: 159-174. [26, 40, 44] Thode H. G. & Monster J. (1970) Sulfur isotope abundances and genetic relations of oil accumulations in Middle East basin. AAPG Bull., 54: 627-637. [26] Thode H. G., Cragg C. B., Hulston J. R. & Rees C. E. (1971) Sulphur isotope exchange between sulphur dioxide and hydrogen sulphide. Geochim. Cosmochim. Acta, 35: 35-45. [46]

References

1165

Thode H. G. & Rees C. E. (1971) Measurement of sulphur concentrations and the isotope ratios 33S/32S, 34S / 32S and 36S / 32S in Apollo samples. Earth Planet. Sci. Lett., 12: 434-438. [20, 26, 45] Thode H. G. & Rees C. E. (1979) Sulphur isotopes in lunar and meteorite samples. Proc. Lunar Planet. Sci. Conf. (10th): 1629-1636. [26] Thode H. G. (1981) Sulfur isotope ratios in petroleum research and exploration: Williston Basin. AAPG Bull., 65: 1527-1535. [26] Thompson A. M. (1992) The oxidizing capacity of the earth's atmosphere: probable past and future changes. Science, 256: 1157-1165. [14] Thompson P. & Gray J. (1977) Determination of 180/160 ratios in compounds containing C, H and O. Int. J. Appl. Rad. Isot., 28: 411-415. [23] Thomson J. J. (1910) Rays of positive electricity. Philosoph. Mag., 20: 752-767. [38] Thomson J. J. (1913) Rays of positive electricity and their application to chemical analyses. Longman, Green and Co., London. [38] Thonnard N. & Lehmann B. E. (1994) The institute for rare isotope measurements: An interdisciplinary facility for RIS-based applications. H.-J. Kluge, J.E. Parks & K. Wendt (eds.). Resonance Ionization Spectroscopy, Seventh International Symposium. AIP Press Woodbury, New York, AIP Conference Proceedings 329: 335-338. [33] Thornton V. & Condon F. E. (1950) Infrared spectrometric determination of deuterium oxide in water. Anal. Chem., 22: 680-681. [34] Thurman E. M. & Malcolm R. L. (1981) Preparative isolation of aquatic humic substances. Env. Sci. Technol., 15: 463-466. [15] Thurston W. M. (1970) Automatic mass spectrometric analysis of the deuterium to hydrogen ratio in natural water. Rev. Sci. Instrum., 41(7): 963-966. [1] Thurston W. M. (1971) Steam film sampling of water for mass spectrometric analysis of the deuterium content. Rev. Sci. Instrum., 42: 700-703. [1] Thurston W. M. & James M. W. D. (1984) Absolute determination of deuterium content of heavy water standards by distillation and mass spectrometry. Anal. Chem., 56: 386-389. [1] Tieszen L. L., Boutton T. W., Tesdahl K. G. & Slade N. A. (1983) Fractionation and turnover of stable carbon isotopes in animal tissues: Implications for 613C analysis of diet. Oecologia, 57: 32-37. [9] Tilles D. (1961a) Variations of silicon isotope ratios in a zoned pegmatite. J. Geophys. Res., 66: 3015-3020. [20] Tilles D. (1961b) Natural variations in isotopic abundances of silicon. J. Geophys. Res., 66: 3003-3014. [20] Tittel E K., Richter D. & Fried A. (2003) Mid-infrared laser applications in spectroscopy. In: Solid-state mid-infrared laser sources. I. T. Sorokina & K. L. Vodopyanov (eds.). Springer Verlag, Berlin. Topics Appl. Phys., 89: 445-516. [34] Tobias H. J., Goodman K. J., Blacken C. E. & Brenna J. T. (1995) High precision D / H measurement from hydrogen gas and water by continuous flow isotope ratio mass spectrometry. Anal. Chem., 67(14): 2486-2492. [1] Tobias H. J. & Brenna J. T. (1996) High-precision D / H measurement from organic mixtures by gas chromatography continuous flow isotope ratio mass spectrometry using a palladium filter. Anal. Chem., 68: 3002-3007. [8] Tobias H. J. & Brenna J. T. (1997) On-line pyrolysis as a limitless reduction source for high-precision isotopic analysis of organic-derived hydrogen. Anal. Chem., 69: 3148-3152. [8] Todd B. J. (1955) Outgassing of glass. J. Appl. Phys., 26: 1238-1243. [1] Todt W., Cliff R. A., Hanser A. & Hofmann A. W. (1996) Validation of the analytical linearity and mass discrimination correction model exhibited by a multiple collector inductively coupled plasma mass spectrometer by means of a set of synthetic uranium isotope mixtures. J. Anal. At. Spectrom., 10: 395398. [31] Tokunaga T. K., Zawislanski P. T., Johannis P. W., Lipton D. S. & Benson S. (1994) Field investigations of selenium speciation, transformation, and transport in soils from Kestrson Reservoir and Lahontan Valley. In: Selenium in the environment. W. T. Jr. Frankenberger & S. Benson (eds.). Marcel Dekker: 119-138. [29] Tomascak P. B., Krogstad E. J., Prestegaard K. L. & Walker R. J. (1995) Li isotope variations in groundwaters from eastern Maryland. EOS Trans. AGU, 76:$112 (Abstract). [6]

1166

References

Tomascak P. B., Carlson R. W. & Shiney S. B. (1999a) Accurate and precise determinations of Li isotopic composition by multi-collector sector ICP-MS. Chem. Geol., 158: 145-154. [6, 31] Tomascak P. B., Hauri E. H. & Lassiter J. C. (1999b) Lithium isotope constraints on Hawaiian plume components. EOS Trans. AGU, 80: $354. [31] Tomascak P. B. & Langmuir C. H. (1999) Lithium isotope variability in MORB. EOS Trans. AGU, 80: F1086-1087. [6, 31] Tomascak P. B., Tera F., Helz R. T. & Walker R. J. (1999c) The absence of lithium isotope fractionation during basalt differentiation: New measurements by multicollector sector ICP-MS. Geochim. Cosmochim. Acta, 63(6): 907-910. [6, 31] Tomascak P. B., Ryan J. G. & Defant M. J. (2000) Lithium isotope evidence for light element decoupling in the Panama subarc mantle. Geology, 28: 507-510. [6] Tomascak P. B., Widom E., Benton L. D., Goldstein S. L. & Ryan J. G. (2002) The control of lithium budgets in island arcs. Earth Planet. Sci. Lett., 196: 227-238. [6] Tomascak P. B., H e m m i n g N. G. & H u m m i n g S. R. (2003) The lithium isotope composition of waters of the Mono Basin, California. Geochim. Cosmochim. Acta, 67: 601-611. [6] Tonarini S., Pennisi M. & Leeman W. P. (1997) Precise boron isotopic analysis of complex silicate (rock) samples using alkali carbonate fusion and ion-exchange separation. Chem. Geol., 142: 129-137. [7, 40] Tonarini S., Pennisi M., Adorni-Braccesi A., Dini A., Ferrara G., Gonfiantini R., Wiedenbeck M. & Grinning M. (2003) Intercomparison of boron isotope and concentration measurements. Part I: Selecion, preparation and homogeneity tests of the intercomparison materials. Geostand. Newslett., 27(1): 21-39. [40] Tong S. L. & Harrison W. W. (1993) Glow discharge mass spectrometric analysis of non-conducting materials. Spectrochim. Acta, B 48: 1237-1245. [35] Toth R. A. (1994a) Extensive measurements of H2160 line frequencies and strengths: 5750 to 7965 cm-1. Appl. Opt., 33: 4851-4867. [34] Toth R. A. (1994b) Transition frequencies and strengths of H2170 and H2180:6600 to 7640 cm -1. Appl. Opt., 33: 4868-4879. [34] Toulemonde B., Horman I., Egli H. & Derbesy M. (1983) Food-related applications of high-resolution NMR. Part II. Differentiation between natural and synthetic vanillin samples using 2H-NMR. Helv. Chim. Acta, 66: 2342-2345. [5] Toyoda S. & Yoshida N. (1999) Determination of nitrogen isotopomers of nitrous oxide on a modified isotope ratio mass spectrometer. Anal. Chem., 71: 4711-4718. [19] Toyoda S., Yoshida N., Urabe T., Aoki S., Nakazawa T., Sugawara S. & Honda H. (2001) Fractionation of N20 isotopomers in the stratosphere. J. Geophys. Res., 106: 7515-7522. [19] Toyoda S., Yoshida N., Miwa T., Mastsui Y., Yamagishi H., Tsunogai U., Nojiri Y. & Tsurushima N. (2002) Production mechanism and global budget of N20 inferred from its isotopomers in the western North Pacific. Geophys. Res. Lett., 29(3): 7-1 - 7-4. [19] Trautmann N. (1994) Ultratrace analysis of long-lived radioisotopes in the environment. H.-J. Kluge, J.E. Parks & K. Wendt (eds.). Resonance Ionization Spectroscopy, Seventh International Symposium. AIP Press Woodbury, New York, AIP Conference Proceedings 329: 243-250. [33] Truesdell A. H. (1974) Oxygen isotope activities and concentrations in aqueous salt solutions at elevated temperatures: consequences for isotope geochemistry. Earth Planet. Sci. Lett., 23: 387-396. [46] Trust B. A. & Fry B. (1992) Stable sulphur isotopes in plants: a review. Plant Cell Environ., 15: 1105-1110. [26] Trust Hammer B. A., Fogel M. L. & Hoering T. C. (1998) Stable carbon isotope ratios of fatty acids in seagrass and redhead ducks. Chem. Geol., 152: 29-41. [9] Tse R. S., Wong S. C. & Yuen C. P. (1980) Determination of deuterium/hydrogen ratios in natural waters by Fourier transform nuclear magnetic resonance spectrometry. Anal. Chem., 52: 2445-2448. [40] Tucker M. E. (1988) Techniques in sedimentology. Blackwell Scientific Publications, Oxford. [2] Tudge A. P. (1960) A method of analysis of oxygen isotopes in orthophosphate - its use in the measurement of paleotemperatures. Geochim. Cosmochim. Acta, 18: 81-93. [2, 20, 22]

References

1167

Turatti F., Griffith D. W. T., Wilson S. R., Esler M. B., Rahn T. A., Zhang H. & Blake G. A. (2000) Positionally dependent 15N fractionation factors in the UV photolysis of N20 determined by high resolution FTIR spectroscopy. Geophys. Res. Lett., 27: 2489-2492. [34] Turekian V. C., Macko S. A. & Keene W. C. (2001) Application of stable sulfur isotopes to differentiate sources of size-resolved particulate sulfate in polluted marine air at Bermuda during spring. Geophys. Res. Lett., 28(8): 1491-1494. [26] Turner J. V. (1982) Kinetic fractionation of carbon-13 during calcium carbonate precipitation. Geochim. Cosmochim. Acta, 46: 1183-1191. [46] Turner J. V. & Gailitis V. (1988) Single-step method for hydrogen isotope ratio measurement of water in porous media. Anal. Chem., 60: 1244-1246. [1] Turner P. J., Mills D. J., SchrSder E., Lapitajas G., Jung G., Iacone L. A., Haydar D. A. & Montaser A. (1998) Instrumentation for low- and high-resolution ICPMS. In: Inductively coupled plasma mass spectrometry. A. Montaser (ed.). Wiley-VCH: 421-501. [31] Tuttle M. L., Goldhaber M. B. & Williamson D. L. (1986) An analytical scheme for determining forms of sulfur in oil shales and associated rocks. Talanta, 33: 953-961. [26] Tuttle O. F. (1949) Two pressure vessels for silicate-water studies. Geol. Soc. Am. Bull., 60: 1727-1729. [46] Tyler S. C. (1986) Stable carbon isotope ratios in atmospheric methane and some of its sources. J. Geophys. Res., 91(D12): 13232-13238. [14] Tyler S. C., Crill P. M. & Brailsford G. W. (1994) 13c/12 C fractionation of methane during oxidation in a temperate forested soil. Geochim. Cosmochim. Acta, 58: 1625-1633. [46] Ueda A. & Sakai H. (1983) Simultaneous determination of the concentration and isotope ratio of sulfate and sulfide sulfur and carbonate carbon in geological samples. Geochem. J., 17: 185-196. [26] Ueda A. & Krouse H. R. (1986) Direct conversion of sulphide and sulphate minerals to SO2 for isotope analysis. Geochem. J., 20: 209-212. [26] Uehara K., Yamamoto K., Kikugawa T. & Yoshida N. (2001) Isotope analysis of environmental substances by a new laser-spectroscopic method utilizing different pathlengths. Sensors and Actuators, B 74: 173-178. [34] Uehara K., Yamamoto K., Kikugawa T. & Yoshida N. (2002) Site-selective nitrogen isotopic ratio measurement of nitrous oxide using 2-/2m diode lasers. Spectrochim. Acta, Part A, 59: 957-962. [19] Ulmer G. C. & Barnes H. L. (1987) Hydrothermal experimental techniques. John Wiley & Sons: 523 p. [46] UnsSld A. (1968) Physik der Sternatmosph/iren. Springer Verlag, Berlin: 90-91. [39] Unterzaucher J. (1940) Die mikroanalytische Bestimmung des Sauerstoffes. Chem. Berichte, 73B: 391-404. [23] Urey H. C. & Rittenberg D. (1933) Some thermodynamic properties of the HIH2, H2H 2 molecules and compounds containing the H2 atom. J. Chem. Phys., 1(2): 137-143. [46] Urey H. C. (1947) The thermodynamic properties of isotopic substances. J. Chem. Soc., 69: 562-581. [15, 18, 20, 28] Urey H. C., Lowenstam H., Epstein S. & McKinney C. R. (1951) Measurement of paleotemperatures and temperatures of the Upper Cretaceous of England, Denmark, and the southeastern United States. Geol. Soc. Am. Bull., 62: 399-416. [40] Usdowski E. & Hoefs J. (1993) Oxygen isotope exchange between carbonic acid, bicarbonat, and water: A re-examination of the data of McCrea (1950) and an expression for the overall partitioning of oxygen isotopes between the carbonate species and water. Geochim. Cosmochim. Acta, 57: 3815-3818. [46] Usdowski E., Michaelis J., BSttcher M. E. & Hoefs. J. (1993) Factors for the oxygen isotope equilibrium fractionation between aqueous and gaseous CO2, carbonic acid, bicarbonate, carbonate, and water (19~ Z. Phys. Chem., 170: 237-249. [46] Uzaki M., Yamada K. & Ishiwatari R. (1993) Carbon-isotope evidence for oil-pollution in long-chain normal alkanes in Tokyo bay sediments. Geochem. J., 27: 385-389. [8] Valentini J. J. (1987) Mass-independent isotope fractionation in nonadiabatic molecular collisions. J. Chem. Phys., 86: 6757-6765. [18] Valentini J. J., Gerrity D. P., Phillips D. L., Nieh J. C. & Tabor K. D. (1987) CARS spectroscopy of O2(1Dg) from the Hartleyband photodissociation of 03: Dynamics of the dissociation. J. Chem. Phys., 86: 6745-6756. [18]

1168

References

Valkiers S., Aregbe Y., Taylor P. D. P. & De Bi6vre P. (1998) A primary xenon isotopic gas standard with traceable values for isotopic composition and molar mass. Int. J. Mass Spectrom., 173: 55-63. [42, 43] Valkiers S., Kipphardt H., Ding T., Damen R., De Bi6vre P. & Taylor P. D. P. (1999) A primary isotopic gas standard for sulfur in the form of SF6, with Systeme International d'Unites traceable values for isotopic composition and molar mass. Int. J. Mass Spectrom., 193: 1-8. [43] Vallet C., Arendt M. & Martin G. J. (1988) Site specific isotope fractionation of hydrogen in the oxidation of ethanol into acetic acid. Application to vinegars. Biotechnol. techniq., 2(2): 83-88. [5] Vallet C., Arendt M., Mabon F., Naulet N. & Martin G. J. (1991) Combination of mass spectrometry and site-specific NMR isotope analyses in the characterization of amino acids. J. Sci. Food. Agric., 56: 167-185. [5] Valley J. M., Chiarenzelli J. R. & McLelland J. M. (1994) Oxygen isotope geochemistry of zircon. Earth Planet. Sci. Lett., 126: 187-206. [20] Valley J. W. & O'Neil J. R. (1981) 13C/12C exchange bewteen calcite and graphite: a possible thermometer in grenville marbles. Geochim. Cosmochim. Acta, 45: 411-419. [46] Valley J. W., Kitchen N., Kohn N. J., Niendorf C. R. & Spicuzza M. J. (1995) UWG-2, a garnet standard for oxygen isotope ratios: Strategies for high precision and accuracy with laser heating. Geochim. Cosmochim. Acta, 59(24): 5223-5231. [16, 20] Valley J. W. & Graham C. M. (1996) Ion microprobe analysis of oxygen isotope ratios in quartz from Skye granite: Healed microcracks, fluid flow, and hydrothermal exchange. Contrib. Mineral. Petrol., 124: 225-234. [20] Valley J. W., Graham C. M., Harte B., Eiler J. M. & Kinny P. D. (1998) Ion microprobe analysis of oxygen, carbon and hydrogen isotope ratios. In: Applicants of microanalytical techniques to understanding mineralizing processes. M. A. McKibben, W. C. III Shanks & W. I. Ridley (eds.). Society of Economic Geologists, Rev. in Econ. Geol., Vol. 7: 73-98. [20, 46] Valyi L. (1977) Atom and ion sources. Akademiai Kiado, Budapest. [38] Van der Kamp G., Van Stempvoort D. R. & Wassenaar L. I. (1996) The radial diffusion method 1. Using intact cores to determine isotopic composition, chemistry and effective porosity for groundwater in aquitards. Water Resource Res., 32: 1815-1822. [2] Van der Meer M. T. J., Schouten S. & Sinninghe Damste J. S. (1998) The effect of the reversed tricarboxylic acid cycle on the 13C contents of bacteriallipids. Org. Geochem., 28: 527-533. [9] Van Everdingen R. O., Shakur M. A. & Krouse H. R. (1982) Isotope geochemistry of dissolved, precipitated, airborne and fallout sulphur species associated with springs near Paige Mountain, Normal Range, N.W.T. Can. J. Earth Sci., 19: 1395-1407. [26] Van Hook W. A. (1972) Vapor pressure isotope effect in aqueous systems. III. The Vapor pressure of HOD (-60 to 200~ J. Phys. Chem., 76: 3040-3043. [4] Van Kreel B. K., Van der Vegt F., Meers M., Wagenmakers T., Westerterp K. & Coward A. (1996) Determination of total-body water by a simple and rapid mass-spectrometric method. J. Mass Spectrom., 31(1): 108-111. [1] van Stempvoort D. R., Reardon E. J. & Fritz P. (1990) Fractionation of sulfur and oxygen isotopes in sulfate by soil sorption. Geochim. Cosmochim. Acta, 54: 2817-2826. [26] van Trigt R., Kerstel E. R. T., Visser G. H. & Meijer H. A. J. (2001) Stable isotope ratio measurements on highly enriched water samples by means of laser spectrometry. Anal. Chem., 73: 2445-2452. [34] van Trigt R., Meijer H. A. J., Sveinbjornsdottir A. E., Johnsen S. J. & Kerstel E. R. T. (2002a) Measuring stable isotopes of hydrogen and oxygen in ice by means of laser spectrometry: the Bolling transition in the Dye-3 (South Greenland) ice core. Ann. Glaciology, 35: 125-130. [34] van Trigt R., Kerstel E. R. T., Meijer H. A. J., McLean M. & Visser G. H. (2002b) Validation of the DLW method in Japanese quail at various water fluxes using laser and IRMS. J. Am. Phys., 93: 2147-2154. [34] van Warmerdam E. M., Frape S. K., Aravena R., Drimmie R. J., Flatt H. & Cherry J. A. (1995) Stable chlorine and carbon isotope measurement of selected chlorinated organic solvents. Appl. Geochem., 10: 547-552. [28] Vanhaecke F., Moens L., Dams R. & Taylor P. (1996) Precise measurement of isotope ratios with a double-focusing magnetic sector ICP mass spectrometer. Anal. Chem., 68: 567-569. [31, 35]

References

1169

Vanhaecke F., Moens L., Dams R., Papadakis I. & Taylor P. (1997) Applicability of high-resolution ICP-

mass spectrometry for isotope ratio measurements. Anal. Chem., 69: 268-273. [31] Vanhaecke F., Moens L., Dams R., Allen L. & Georgitis S. (1999a) Evaluation of the isotope ratio per-

formance of an axial time-of-flight ICP mass spectrometer. Anal. Chem., 71: 3297-3303. [31] Vanhaecke F., Moens L. & Taylor P. (1999b) Use of ICP-MS for isotope ratio measurements. In: Induc-

tively coupled plasma spectrometry and its applications. S. J. Hill (ed.). Sheffield Academic Press/CRC Press: 145-207. [37] Vaughn B. H., White J. W. C., Delmotte M., Trolier M., Cattani O. & Stievenard M. (1998) An automated system for hydrogen isotope analysis of water. Chem. Geol., 152: 309-319. [1] Veblen D. R. & Ferry J. M. (1983) A TEM study of the biotite-chlorite reaction and comparison with petrologic observations. Am. Mineralogist, 68: 1160-1168. [16] Velinsky D. J., Pennock J. R., Sharp J. H., Cifuentes L. A. & Fogel M. L. (1989) Determination of the isotopic composition of ammonium-nitrogen at the natural abundance level from estuarine waters. Mar. Chem., 26: 351-361. [9, 15] Vengosh A., Chivas A. R. & McCulloch M. (1989) Direct determination of boron and chlorine isotopic compositions in geological materials by negative thermal-ionization mass spectrometry. Chem. Geol., 79: 333-343. [7, 28] Vengosh A., Kolodny Y., Starinsky A., Chivas A. R. & McCulloch M. T. (1991) Coprecipitation and isotopic fractionation of boron in modern biogenic carbonates. Geochim. Cosmochim. Acta, 55: 29012910. [7] Vengosh A., Chivas A. R., Starinsky A., Kolodny Y., Zhang B. & Zhang P. (1995) Chemical and boron isotope compositions of non-marine brines from the Qaidam Basin, Qinghai, China. Chem. Geol., 120: 135-154. [46] Venkitachalem T. V. & Rao A. S. (1991) Resonance-enhanced multiple-photon ionization spectroscopy of excited sulphur atoms. Appl. Phys. B - Photophys. Laser Chem., 52: 102-107. [33] Vennemann T. W. & Smith H. S. (1990) The rate and temperature of reaction of C1F3 with silicate minerals, and their relevance to oxygen isotope analysis. Chem. Geol. (Isot. Geosci. Sect.), 86: 83-88. [20] V e n n e m a n n T. W., Kesler S. E. & O'Neil J. R. (1992) Stable isotope compositions of quartz pebbles and their fluid inclusions as tracers of sediment provenance; implications for gold- and uranium-bearing quartz pebble conglomerates. Geology, 20: 837-840. [3] V e n n e m a n n T. W. & O'Neil J. R. (1993) A simple and inexpensive method of hydrogen isotope and water analyses of minerals and rocks based on zinc reagent. Chem. Geol. (Isot. Geosci. Sect.), 103: 227234. [1, 2] V e n n e m a n n T. W. & O'Neil J. R. (1996) Hydrogen isotope exchange reactions between hydrous minerals and molecular hydrogen: I. A new approach for the determination of hydrogen isotope fractionation at moderate temperatures. Geochim. Cosmochim. Acta, 60(13): 2437-2451. [46] V e n n e m a n n T. W. & Hegner E. (1998) Oxygen, strontium, and neodynium isotope composition of fossil shark teeth as a proxy for the palaeoceanography and palaeoclimatology of the Miocene northern Alpine Paratethys. Palaeogeogr., Pataeoclimatol., Palaeoecol., 142: 107-121. [22] Vennemann T. W., Fricke H. C., Blake R. E., O'Neil J. R. & Colman A. (2002) Oxygen isotope analysis of phosphates: a comparison of techniques for analysis of Ag3PO4. Chem. Geol., 185: 321-336. [20] Verchovsky A. B., Fisenko A. V., Semjonova L. F. & Pillinger C. T. (1997) Heterogeneous distribution of Xenon-HL within presolar diamonds. Meteoritics, 32: A131-A132. [13] Verkouteren R. M. & Dorko W. D. (1989) High-accuracy gas analysis via isotope dilution mass spectrometry: carbon dioxide in air. Anat. Chem., 61: 2416-2422. [43] Verkouteren R. M., Klouda G. A. & Currie L. A. (1995) The carbon dioxide isotopic measurement process: Progress at NIST on measurements, reduction algorithms and standards. In: Reference and intercomparison materials for stable isotopes of light elements. IAEA-TECDOC-825, IAEA, Vienna: 111-129. [1, 15] Verkouteren R. M., Lagergren E. S. & Velapoldi R. A. (1998) Report of Analysis. NIST, Gaithersburg, USA. Internal Report 837-6-98. [40] Verkouteren R. M. (1999) Preparation, characterization, and value assignment of carbon dioxide isotopic reference materials: RMs 8562, 8563, and 8564. Anal. Chem., 71: 4740-4746. [40, 43]

1170

References

Verkouteren R. M. & Lee J. N. (2001) Web-based interactive data processing: Application to stable isotope metrology. Fresenius J. Anal. Chem., 370: 803-810. [43] Verkouteren R. M. & Klinedinst D. B. (2003) Value assignment and uncertainty estimation of selected light stable isotope reference materials: RMs 8543-8545, RMs 8562-8564, and RM 8566. NIST, Special Publication 260-149. [43] Vet5 I., Het6nyi M., Dem6ny A. & Hertelendi E. (1994) Hydrogen index as reflecting intensity of sulphidic diagenesis in non-bioturbated, shaly sediments. Org. Geochem., 22: 299-310. [27] Vieth W. & Huneke J. C. (1991) Relative sensitivity factors in glow discharge mass spectrometry. Spectrochim. Acta, B 46: 137-153. [35] Vilagura R. A., Alexander R. B., Castro M. S., Tilden P. M., Pearl H. W., Stacey P. E. & Turner R. E. (2001) Nitrogen loading in coastal water bodies: an atmospheric perspective. American Geophysical Union, Washington DC. Coastal and Estuarine studies, Vol. 57. [15] VIM (1993) International vocabulary of basic and general terms in metrology. BIPM, IEC, IFCC, ISO, IUPAC, IUPAP, OIML, International Organization for Standardization. ISBN 92-67-01075-1. [42, 43] Vinogradov A. P. & Dontsova E. I. (1947) Izotopnyi Sostav Kisloroda Alyuminosilikatnykh Gornykh Porod (Isotopic composition of the alumino-silicate rocks). Dokl. Akad. Nauk SSSr, 58:83-84 (in Russian). [20] Vistousek P. M., Aber J. D., Howarth R. W., Likens G. E., Matson P. A., Schindler D. W., Schlesinger W. D. & Tilman D. G. (1977) Human alterations of the global nitrogen cycle: sources and consequencaes. Ecol. Appl., 7: 737-750. [15] Vitali F., Longstaffe F. J., Bird, M. I., Gage K. L. & Caldwell W. G. E. (2001) Hydrogen-isotope fractionation in aluminium hydroxides: Synthesis products versus natural samples from bauxites. Geochim. Cosmochim. Acta, 65: 1391-1398. [46] Vityk M., Krouse H. R. & Demihov Y. (1993) Preservation of 6180 values of fluid inclusion water in quartz over geological time in an epithermal environment: Beregovo Deposit, Transcarpathia, Ukraine. Earth Planet. Sci. Lett., 119(4): 561-568. [3] Voerkelius S. (1990) Isotopendiskriminierungen bei der Nitrifikation und Denitrifikation: Grundlagen und Anwendungen der Herkunfts-Zuordnungen von Nitrat und Distickstoffmonoxid. PhD Thesis, Technical University of Mfinich, Mfinich, Germany: 119 p. [15] Vogel A. I. (1951) A textbook of quantitative inorganic analysis, theory and practice. Longmans, Green and Co., London: 918 p. [28] Vogel J. C., Talma A. S. & Heaton T. H. E. (1981) Gaseous nitrogen as evidence for denitrification in groundwater. J. Hydrol., 50: 191-200. [15] Vogl J., Becker D., Birke W., Janisch N., Klingbeil P., Ludwig H. P., Riebe W., Pritzkow W. & Wermann G. (2002) Zertifizierungsbericht, Herstellung und Zertifizierung von 3 Bor-Isotopenreferenzmaterialien mit unterschiedlicher Bor-Isotopenzusammensetzung. Bundesanstalt fiir Materialforschung und -pr~ifung, Berlin, Germany. [40] Vogler E. A. & Hayes J. M. (1980) Carbon isotopic compositions of carboxyl groups of biosynthesized fatty acids. In: Advances in organic geochemistry, 1979. A. G. Douglas & J. R. Maxwell (eds.). Pergamon Press, Oxford, Vol. 12: 697-704. [9] Volkert R., Johnson C. M. & Tamashausky A. (2000) Mesoproterozic graphite depositits, New Jersey Highlands: geologic and stable isotopic evidence for possible algal origins. Can. J. Earth Sci., 37: 16651675. [45] yon Kallman H. & Lasareff W. (1932) l]-ber die Isotopenuntersuchungen (Sauerstoff, Neon und Chlor). Z. Physik, 80: 333-343. [28] Voncken J. H. L., Van Roermund H. L. M., van der Eerden A. M. J., Jansen J. B. H. & Erd R. C. (1993) Holotype buddingtonite; an ammonium feldspar without zeolitic H20. Am. Mineralogist, 78: 204209. [16] Viilkening J. & Papanastassiou D. A. (1989) Iron isotope anomalies. Astrophys. J., 347: L43-L46. [29, 31] VSlkering J., Koppe M. & Heumann K. G. (1991) Tungsten isotope ratio determinations by negative thermal ionization mass spectrometry. Int. J. Mass Spectrom. Ion Proc., 107: 361-368. [35] VSlkl J. & Alefeld G. (1975) Hydrogen diffusion in metals. In: Diffusion in Solids: Recent Developments. A. S. Nowick & J. J. Burton (eds.). Academic Press: 231-302. [46]

References

1171

Vreca P. (2003) Carbon cycling at the sediment-water interface in a eutrophic mountain lake (Jezero na

Planini pri Jezeru, Slovenia). Org. Geochem., 34: 671-680. [10] Wachniew R. & Rozanski K. (1997) Carbon budget of a mid-latitude, groundwater-controlled lake: Isotopic evidence for the importance of dissolved inorganic carbon recycling. Geochim. Cosmochim. Acta, 61: 2453-2465. [10] Wachsmann M. & Heumann K. G. (1991) Negative thermal ionization mass spectrometry of main group elements, Part 1.5 th group: phosphorus, arsenic and antimony. Int. J. Mass Spectr. Ion Proc., 108: 75-86. [39] Wachsmann M. & Heumann K. G. (1992) Negative thermal ionization mass spectrometry of main group elements. Part 2. 6th group: sulfur, selenium, and tellurium. Int. J. Mass Spectrom. Ion. Proc., 114: 209-220. [20, 26, 29, 39] Wada H. & Siuzuki K. (1983) Carbon isotopic thermometry calibrated by dolomite-calcite solvus temperatures. Geochim. Cosmochim. Acta, 47: 697-706. [46] Wada H. (1988) Microscale isotopic zoning in calcite and graphite crystals in marble. Nature, 331: 61-63. [12] Wadleigh M. A., Schwarcz H. P. & Kramer J. R. (1994) Sulphur isotope tests of seasalt correction factors in precipitation: Nova Scotia, Canada. Water Air Soil Pollut., 77: 1-16. [26] Wadleigh M. A. & Blake D. M. (1999) Tracing sources of atmospheric sulphur using epiphytic lichens. Environ. Pollut., 106: 265-271. [26] Wahlen M. & Yoshinari T. (1985) Oxygen isotope ratios in N20 from different environments. Nature, 313: 780-782. [19, 34] Wainwright S. C., Haney J. C., Kerr C., Golovkin A. N. & Flint M. V. (1998) Utilization of nitrogen derived from seabird guano by terrestrial and marine plants at St. Paul, Pribilof Islands, Bering Sea, Alaska. Marine Biol., 131: 63-71. [9] Wakeham S. G. & Volkman J. K. (1991) Sampling and analysis of lipids in marine particulate matter. Geophys. Monographs, 63: 171-179. [9] Wakshal E. & Nielsen H. (1982) Variations of ~)34S(SO4), ~)180(H20) and C1/SO4 ratio in rainwater over northern Israel, from the Mediterranean coast to Jordan Rift Valley and Golan Heights. Earth Planet. Sci. Lett., 61: 272-282. [26] Walder A. J. & Freedman P. A. (1992) Isotopic ratio measurement using a double focusing magnetic sector mass analyzer with an inductively coupled plasma as an ion source. J. Anal. At. Spectrom., 7: 571-575. [31, 35] Walder A. J. & Furuta N. (1993) High-precision lead isotope ratio measurement by inductively coupled plasma multiple collector mass spectrometry. Anal. Sci., 9: 675-680. [31] Walder A. J., Abell I. A., Platzner I. & Freedman P. A. (1993a) Lead isotope ratio measurement of NIST 610 glass by laser ablation inductively coupled plasma mass spectrometry. Spectrochim. Acta, B 48: 397-402. [35] Walder A. J., Koller D., Reed N. M., HuRon R. C. & Freedman P. A. (1993b) Isotope ratio measurement by inductively coupled plasma multiple collector mass spectrometry incorporating a high efficiency nebulization system. J. Anal. At. Spectrom., 8: 1037-1041. [31] Walder A. J., Platzner I. & Freedman P. A. (1993c) Isotope ratio measurement of lead, neodymium and neodymium - samarium mixtures, hafnium and hafnium - lutetium mixtures with a double focusing multiple collector inductively coupled plasma mass spectrometer. J. Anal. At. Spectrom., 8: 19-23. [31, 35] Walder A. J. (1997) Advanced isotope ratio mass spectrometry. II: isotope ratio measurement by multiple collector inductively coupled plasma spectrometry. In: Modern isotope ratio mass spectrometry. I. T. Platzner (ed.). John Wiley & Sons. Chemical analysis, a series of monographs on analytical chermistry and its applications, Vol. 145: 83-108. [31] Walker G. R., Woods P. H. & Allison G. B. (1994) Interlaboratory comparison of methods to determine the stable isotope composition of soil water. Chem. Geol. (Isot. Geosci. Sect.), 111: 297-306. [1, 2] Waiters L. J. Jr., Claypool G. E. & Choquette P. W. (1972) Reaction rates and 6180 variation for the carbonate-phosphoric acid preparation method. Geochim. Cosmochim. Acta, 36: 129-140. [12] Wan D. F., Li Y. H., Song H. B. & Liu Z. J. (1997) Preparation and calibration of reference material for silicon isotope. Acta Geoscientia Sinica, 18(3): 330-335 (in Chinese, with English abstract). [25]

1172

References

Wang C.-H. & Yeh H.-W. (1985) Oxygen isotopic compositions of DSDP Site 480 diatoms: Implications and applications. Geochim. Cosmochim. Acta, 49: 1469-1478. [2] Wang J., Davis A. M., Clayton R. N. & Hashimoto A. (1999) Evaporation of single crystal forsterite: evaporation kinetics, magnesium isotope fractionation, and implications of mass-dependent isotopic fractionation of a diffusion-controlled reservoir. Geochim. Cosmochim. Acta, 63: 953-966. [46] Wang J., Davis A. M., Clayton R. N., Mayeda T. K. & Hashimoto A. (2001) Chemical and isotopic fractionation during the evaporation of the FeO-MgO-SiO2-CaO-A1203-TiO2 rare earth element system. Geochim. Cosmochim. Acta, 65: 479-494. [46] Wang X. F. & Yakir D. (1995) Temporal and spatial variations in the oxygen-18 content of leaf water in different plant species. Plant, Cell Environm., 18(12): 1377-1385. [21] Wang Y., Huntington T. G., Osher L. J., Wassenaar L. I., Trumbone S. E., Amundson R. G., Harden J. W., McKnight D. M., Schiff S. L., Aiken G. R., Lyons W. B., Aravena R. O. & Baron J. S. (1998) Carbon Cycling in Terrestrial Environments. I. In: Isotope tracers in catchment hydrology. C. Kendall & J. J. McDonnel (eds.). Elsevier Science B.V., New York: 577-610. [26] Wankel S. D., Kendall C., Silva S. R. & Bernis B. E. (2002) Estimating atmospheric nitrate deposition using nitrate 6170, 6180 and 615N. EOS, Trans. AGU. Annual Fall Meeting, San Francisco, CA. December 6-10. [15] Ward M. H., Mark S. D., Cantor K. P., Weisenburger D. D. & Correa-Villasenor A. (1996) Drinking water nitrate and the risk of non-Hodgkin's lymphomia. Epidemiology, 7: 465-471. [15] Ward S., Scantlebury M., Krol E., Thompson P. J., Sparling C. & Speakman J. R. (2000) Preparation of hydrogen from water by reduction with lithium aluminum hydride for the analysis of 62H by isotope ratio mass spectrometry. Rapid Commun. Mass Spectrom., 14: 450-453. [1] Ware R. K. & Pirooz P. P. (1967) Gas adsorption on freshly broken glass surfaces- a source of error in analysis of bubbles in glass. Glass Technol., 8: 86-87. [3] Wassenaar L. I. (1995) Evaluation of the origin and fate of nitrate in the Abbotsford Aquifer using the isotopes of 15N and 180 in NO3-. Appl. Geochem., 10: 391-405. [15] Wassenaar L. I. & Koehler G. (1999) An on-line technique for the determination of the 6180 and 6170 of gaseous and dissolved oxygen. Anal. Chem., 71: 4965-4968. [15, 36] Wasserburg G. J., Lee T. & Papanastassiou D. A. (1977) Correlated O and Mg isotopic anomalies in Allende inclusions. II: Magnesium. Geophys. Res. Lett., 4: 299-302. [31] Wasserburg G. J., Jacobsen S. B., DePaolo D. J., McCulloch M. T. & Wen T. (1981) Precise determination of S m / N d ratios, Sm and Nd isotopic abundances in standard solutions. Geochim. Cosmochim. Acta, 45: 2311-2323. [31] Watanabe S., Kazunobu M. & Matsuo S. (1983) Isotopic ratios of carbonaceous materials incorporated in olivine crystals from the Hualalai Volcano, Hawaii- An approach to mantle carbon. Geochem. J., 17: 95-104. [3] Watson L. L., Hutcheon I. D., Epstein S. & Stolper E. M. (1994) Water on Mars: Clues from deuterium/ hydrogen and water contents of hydrous phases in SNC meteorites. Science, 265: 86-90. [30] Wayne D. M., Yoshida T. M. & Vance D. E. (1996) Analysis of trace impurities in palladium metal powders by glow discharge mass spectrometry. J. Anal. At. Spectrom., 11" 861-869. [35] Wayne D. M. (1997) Direct determination of trace noble metals (palladium, platinum, and rhodium) in automobile catalysts by glow discharge mass spectrometry. J. Anal. At. Spectrom., 12: 1195-1202. [35] Wayne D. M., Schulze R. K., Maggiore C., Cooke D. W. & Havrilla G. (1999) Characterization of tantalum films on analytical surfaces: Insight into sputtering of non-conductors in a dc-glow discharge using secondary cathodes. Appl. Spectrosc., 53: 266-277. [35] Webb E. A. & Longstaffe F. J. (2000) The oxygen isotopic compositions of silica phytoliths and plant water in grasses: implications for the study of paleoclimate. Geochim. Cosmochim. Acta, 64: 767-780. [46] Webb E. A. & Longstaffe F. J. (2002) Climatic influences on the oxygen isotopic composition of biogenic silica in prairie grass. Geochim. Cosmochim. Acta, 66, 1891-1904. [46] Weber D., Kexel H. & Schmidt H.-L. (1997) 13C-pattern of natural glycerol: Origin and practicle importance. J. Agric. Food Chem., 45: 2042-2046. [8]

References

1173

Weber J. N., Deines P., Weber P. H. & Baker P. A. (1976) Depth related changes in the 13C/12C ratio of skeletal carbonate deposited by the Caribbean reefframe building coral Montastrea annularis: further implications of a model for stable isotope fractionation by scleractinian corals. Geochim. Cosmochim. Acta, 40: 31-39. [44] Weber P. K., Hutcheon I. D., McKeegan K. D. & Ingram B. L. (2002) Otolith sulfur isotope method to reconstruct salmon (Oncorhynchus tshawytscha) life history. Can. J. Fish Aquat. Sci., 59: 587-591. [26] Webster C. L. (1972) Selenium isotope analysis and geochemical applications. PhD Thesis, Colorado State University. [29] Webster C. L. Jr. & Warren C. G. (1981) Selenium isotope analysis. U. S. Geological Survey, Open-File Rept., 81-992:20 p. [20] Webster C. R. & May R. D. (1992) In situ stratospheric measurements of CH4, 13CH4, N20 and OC180 using the BLISS tunable diode laser spectrometer. Geophys. Res. Lett., 19: 45-48. [34] Webster C. R., May R. D., Owen T., Herman R. L., Rossman G. R. & Yung Y. L. (1999) Tunable laser spectrometers for in-situ gas, particulate, and mineralogical studies of Mars. Centre National d'Etudes Spatiales, International Conference. Mars Exploration Program and Sample Return Missions, P8 $3(64). [33] Webster L. A., Wahl M. H. & Urey H. C. (1935) The fractionation of the oxygen isotopes in an exchange reaction. J. Chem. Phys., 3: 129. [46] Wedepohl K. H. (1995) The composition of the continental crust. Geochim. Cosmochim. Acta, 59: 12171232. [15] Wefer G. & Berger W. H. (1991) Isotope paleontology: growth and composition of extant calcareous species. Marine Geol., 100: 207-248. [46] Weidmann D., Roller C., Tittel E K., Curl R. F., Uehara K., Oppenheimer C. & DeNatale P. (2003) Development of a 4.3 ~m quantum cascade laser based 13CO2/12CO2 isotope ratio sensor. 2nd International Symposium on Isotopomers, 4-6 November, Stresa, Italy. [34] Weilacher T., Gleixner G. & Schmidt H.-L. (1996) Carbon-isotope pattern in purine alkaloids a key to isotope discriminations in C-1 compounds. Phytochem., 41: 1073-1077. [8] Weiler R. R. & Nriagu J. O. (1978) Isotopic composition of dissolved inorganic carbon in the Great lakes. J. Fisheries Res. Board of Canada, 35: 422-430. [10] Welch S. A., Beard B. L., Johnson C. M. & Braterman P. S. (2003) Kinetic and equilibrium Fe isotope fractionation between aqueous Fe(II) and Fe(III). Geochim. Cosmochim. Acta, 67: in press. [46] Wendt K., Bushaw B. A., Bhowmick G., Bystrow V. A., Kotovski N., Kratz J. V., Lantzsch J., Miiller P., NSrtersh/iuser W., Often E. W., Seibert A., Trautmann N., Waldek A. & Yushkevich Y. (1996) 89,90Sr-determination on various environmental samples by collinear resonance ionization spectroscopy. Resonance Ionization Spectroscopy 1996, Eighth International Symposium. N. Winograd & J.E. Parks (eds.). AIP Press Woodbury, New York, AIP Conference Proceddings 388: 361-364. [33] Wendt K. (1998) Ultratrace determination of long-lived radioactive isotopes. Resonance Ionization Spectroscopy, Ninth International Symposium. J.C. Vickerman, I. Lyon, N.P. Lockyer & J.E. Parks (eds.). AIP Press Woodbury, NY. AIP Conference Proceedings 454: 39-46. [33] Wendt K., Bushaw B. A., Gruning C., Horn R., Huber G., Katz J. V., Kunz P., Miiller P., NiJrtersh~iuser W., Nunnemann M., Passler G., Schmitt A., Trautmann N. & Waldek, A. (1999) Recent developments in and applications of resonance ionization mass spectrometry. Fresenius J. Anal. Chem., 364: 471-477. [33] Wenner D. B. & Taylor H. P. Jr. (1971) Temperatures of serpentinization of ultramafic rocks based on 180 / 160 fractionation between coexisting serpentine and magnetite. Contrib. Mineral. Petrol., 32: 165185. [46] Wenzel B., L6cuyer C. & Joachimski M. M. (2000) Comparing oxygen isotope records of silirian calcite and phosphate - 6180 compositions of brachiopods and conodonts. Geochim. Cosmochim. Acta, 64: 1859-1872. [22] Werle P. (1999) Laseroptical sensors for in-situ gas analysis. Rec. Res. Developm. Opt. Eng., 2: 143-160. [34] Werne J. P., Lyons T. W., Hollander D. J., Formolo M. J & Damste J. S. S. (2003) Reduced sulfur in euxinic sediments of the Cariaco Basin: sulfur isotope constraints on organic sulfur formation. Chem. Geol., 195: 159-179. [26]

1174

References

Werner R. A., Kornexl B. E., Rossmann A. & Schmidt H.-L. (1996) On-line determination of 6180 values

of organic substances. Anal. Chim. Acta, 319: 159-164. [1, 23] Werner R. A. & Brand W. A. (2001) Referencing strategies and techniques in stable isotope ratio analysis.

Rapid Commun. Mass Spectrom., 15: 501-519. [40] Werner R. A., Br~iunlich S. & Brand W. A. (2001) On-line determination of 5180 in silicates, carbonates

and other inorganic substances: First results. Abstr. Papers Am. Chem. Soc., 221 (98-GEOC, Part I, Apr. 1, 2001). [201 Westaway K. C., Koerner T., Fang Y.-R., Rudzinski J. & Paneth P. (1998) A new method of determining chlorine kinetic isotope effects. Anal. Chem., 70: 3548-3552. [28] Westgate L. M. & Anderson T. E (1982) Extraction of various forms of sulphur from coal and shale for stable sulphur isotope analysis. Anal. Chem., 54: 2136-2139. [26] Weston R. E. (1999) Anomalous or mass-independent isotope effects. Chem. Rev., 99: 2115-2136. [18] Wetherill G. W. (1964) Isotopic composition and concentration of molybdenum in iron meteorites. 1. Geophys. Res., 69(20): 4403-4408. [29] Wetherill G. W. & Chapman C. R. (1998) Asteroids and meteorites. In: Meteorites and the early solar system. J. F. Kerridge & M. S. Mathews (eds.). Univ. Arizona Press: 35-67. [11] Wheat J. A. (1971) Isotopic analysis of lithium by atomic absorption spectrophotometry. Appl. Spectrosc., 25: 328-330. [6] Wheeler M. D., Newman S. M., Orr-Ewing A. J. & Ashfold M. N. R. (1998) Cavity ring-down spectroscopy. J. Chem. Soc. Farad. Trans., 94: 337-351. [34] White A. F. & Chuma N. J. (1987) Carbon and isotopic mass balance models of Oasis Valley-Fortymile Canyon groundwater basin, southern Nevada. Water Resour. Res., 23: 571-582. [10] White J. U. (1976) Very long optical paths in air. J. Opt. Soc. Am., 66: 411-416. [34] White J. w. & Gedzelman S. D. (1984) The isotopic composition of atmospheriic water vapor and the concurrent meteorological conditions. J. Geophys. Res., 89: 4937-4939. [1] White J. w. C., Lawrence J. R. & Broecker W. S. (1994) Modeling and interpreting D / H ratios in tree rings: a test case of White Pine in the northeastern United States. Geochim. Cosmochim. Acta, 58: 851862. [46] Whittaker M., Pollard S. J. T. & Fallick A. E. (1995) Characterization of refractory wastes at heavy oilcontaminated sites - a review of conventional and novel analytical methods. Environ. Technol., 16: 1009-1033. [8] Whittaker M., Pollard S. J. T., Fallick A. E. & Preston T. (1996) Characterization of refractory wastes at hydrocarbon-contaminated sites. 2. Screening of reference oils by stable carbon isotope fingerprinting. Environ. Pollut., 94: 195-203. [8] Whittaker P. G., Barrett J. F. R. & Williams J. G. (1992) Precise determination of iron isotope ratios in whole blood using inductively coupled plasma mass spectrometry. J. Anal. Atom. Spectrom., 7: 109113. [31] Wiechert U. & Hoefs J. (1995) An excimer laser-based micro analytical preparation technique for in-situ oxygen isotope analysis of silicate and oxide minerals. Geochim. Cosmochim. Acta, 59: 4093-4101. [20] Wiechert U., Fiebig J., Przybilla R., Xiao Y. & Hoefs J. (2002) Excimer laser isotope-ratio-monitoring mass spectrometry for in situ oxygen isotope analysis. Chem. Geol., 182: 179-194. [20] Wieder R. K., Lang G. E. & Granus V. A. (1985) An evaluation of wet chemical methods for quantifying sulfur fractions in freshwater wetland peat. Limnol. Oceanogr., 30: 1109-1115. [26] Wieder R. K. & Lang G. E. (1988) Cycling of inorganic and organic sulfur in peat from Big Run Bog, West Virginia. Biogeochem., 5: 221-242. [26] Wieser M. E. (1997) Stable isotope ratio mass spectrometry of nanigram quatities of boron and sulfur. PhD Thesis, University of Calgary, Calgary, Canada: 127 p. [26] Wigley T. M. L., Plumer L. N. & Pearson F. J. (1978) Mass transfer and carbon isotope evolution in natural waters. Geochim. Cosmochim. Acta, 42: 1117-1139. [10] Wijbrands J. R. & McDougall I. (1986) 40Ar/39Ar dating of white micas from an Alpine high-pressure metamorphic belt on Naxos (Greece): the resetting of the argon isotopic system. Contrib. Mineral. Petrol., 93: 187-194. [16] Wilbur K. M. & Saleuddin A. S. M. (1983) Shell formation. In: The Mollusca. A. S. M. Saleuddin & K. M. Wilbur (eds.). Academic Press: 236-279. [46]

References

1175

Wilde S. A., Valley J. W., Peck W. H. & Graham C. M. (2001) Evidence from detrital zircons for the existence of continental crust and oceans on the Earth 4.4 Gyr ago. Nature, 409: 175-178. [30] Wiley K. F., Arlinhaus H. F. & Whitaker T. J. (1999) Routine analysis at sub-micron resolution through the use of sputter initiated resonance ionization spectroscopy. Appl. Surf. Sci., 145: 36-40. [33] Wilkinson J. J., Jenkin G. R. T., Fallick A. E. & Foster R. P. (1995) Oxygen and hydrogen isotopic evolution of Variscan crustal fluids, south Cornwall, U.K. Chem. Geol., 123: 239-254. [3] Willey J. F. & Taylor J. W. (1978) Capacitive integration to produce high-precision isotope ratio measurements on methyl chloride and methyl bromide samples. Anal. Chem., 50: 1930-1933. [28] Williams J. D. H., Syers J. K. & Walker T. W. (1967) Fractionation of soil inorganic phosphate by a modification of Chang and Jackson's procedure. Soil Sci. Soc. Am. Proc., 31: 736-739. [2] Williams J. M., Jegou H., Langlois V., Zhang B. L. & Martin M. L. (1993) The search for D-glucose derivates suitable for the study of natural hydrogen isotope fractionation. Carbohydr. Res., 245: 57-64. [5] Williams L. B., Ferrell R. E., Hutchean I., Bakel A. J., Walsh M. M. & Krouse H. R. (1995) Nitrogen isotope geochemistry of organic matter and mineral during diagenesis and hydrocarbon migration. Geochim. Cosmochim. Acta, 59: 765-779. [16] Williams L. B. & Hervig R. L. (1997) Potential for in-situ measurements of isotopic ratios in clay minerals using secondary ion mass spectrometry. 11th International Clay Conference, Ottawa, Canada (Abstract). [2] Williams L. B., Hervig R. L., Holloway J. R. & Hutcheon I. (2001a) Boron isotope geochemistry during diagenesis. Part I. Experimental determination of fractionation during illitization of smectite. Geochim. Cosmochim. Acta, 65: 1769-1782. [2] Williams L. B., Hervig R. L., Wieser M. E. & Hutcheon I. (2001b) The influence of organic matter on the boron isotope geochemistry of the gulf coast sedimentary basin, USA. Chem. Geol., 174: 445-461. [30] Williams P. (1979) The sputtering process and sputtered ion emission. Surf Sci., 90: 588-634. [30] Williams P. (1982) On mechanisms of sputtered ion emission. Appl. Surface Sci., 13: 241-259. [30] Wilson A. T. & Grinsted M. J. (1975) Palaeotemperatures from tree rings and the D / H ratio of cellulose as a biochemical thermometer. Nature, 257: 387-388. [24] Wilson A. T. & Grinsted M. J. (1977) 13C/12 C in cellulose and lignin as palaeothermometers. Nature, 265: 133-135. [24] Wilson B. E. Jr. & Wells A. J. (1946) The experimental determination of the intensities of infra-red absorption bands. I. The theory of the method. J. Chem. Phys., 14: 578-565. [34] Wilson B. E. Jr., Decius J. C. & Cross P. C. (1980) Molecular vibrations: The theory of infrared and raman vibrational spectra. Dover Publications, New York. [34] Wilson P. F., Freeman C. G., McEwan M. J., LMiligan D. B., Allardyce R. A. & Shaw G. M. (2001) Alcohol in breath and blood: a selected ion flow tube mass spectrometric study. Rapid Commun. Mass Spectrom., 15: 413-417. [4] Wilson R. G. (1995) SIMS quantification in Si, GaAs, and diamond - an update. Int. J. Mass Spectrom. Ion Proc., 143: 43-49. [30] Winchester M. R., Lazik C. & Marcus R. K. (1991a) Characterization of a radio frequency glow discharge emission source. Spectrochim. Acta, B 46: 483-499. [35] Winchester M. R., Hayes S. M. & Marcus R. K. (1991b) Determination of platinum and rhodium in alumina catalysts by glow discharge atomic absorption spectrophotometry. Spectrochim. Acta, B 46: 615627. [35] Winchester M. R., Duckworth D. C. & Marcus R. K. (1993) Analysis of nonconducting sample types. In: Glow discharge spectroscopies. R. K. Marcus (ed.). Plenum Press, New York: 263-328. [35] Winteringham E P. (1980) Nitrogen balance and related studies: a global review. In: Soil Nitrogen as Fertilizer or Pollutant. IAEA, Vienna: 307-344. [15] Wise L. E., Murphy M. & D'Addieco A. (1946) Chlorite holocellulose, its fractionation and bearing on summative wood analysis and on studies in the hemicellulose. Paper Trade J., 122: 35. [24] w i t t G. B., Moll E. J., Beeton R. J. S. & Murray P. J. (1998) Isotopes, wool, and rangeland monitoring: Let the sheep do the sampling. Environ. Management, 22: 145-152. [9] Witter A. E. & Jones A. D. (1999) Chemical characterization of organic constituents from sulfide-rich produced water using gas chromatography/mass spectrometry. Environ. Toxicol. Chem., 18: 19201926. [26]

1176

References

Wolfe R. R. (1992) Radioactive and stable isotope tracers in biomedicine. Wiley-Liss Inc., New York. [8] Wombacher F., Rehk~imper M., Miinker C. & Mezger K. (2000) Cadmium isotope analysis by MCICPMS. EOS Trans. AGU, 81: 1362. [31] Wombacher F., Rehk~imper M., Mezger K., Miinker C. & Bischoff A. (2001) Mass-dependent cadmium isotope fractionation in meteorites and experiments. J. Conf. Abstr., 6: 689. [31] Wong W. W., Caberra M. P. & Klein P. D. (1984) Evaluation of a dual mass spectrometer system for rapid simultaneous determination of hydrogen-2/hydrogen-1 and oxygen-18/oxygen-16 ratios in aqueous samples. Anal. Chem., 56: 1852-1858. [1] Wong W. W. (1985) Comparison of infrared and mass-spectrometric measurements of carbon-13/carbon-12 ratios. Int. J. Appl. Radiat. Isot., 36: 997-999. [34] Wong W. W. & Klein P. D. (1986) A review of techniques for the preparation of biological samples for mass-spectrometric measurements of hydrogen-2/hydrogen-1 and oxygen-18/oxygen-16 isotope ratios. Mass Spectrom. Rev., 5: 313-342. [1, 23] Wong W. W., Lee L. S. & Klein P. D. (1987a) Oxygen isotope ratio measurements on carbon dioxide generated by reaction of microliter quantities of biological fluids with guanidine hydrochloride. Anal. Chem., 59: 690-693. [1] Wong W. W., Lee L. S. & Klein P. D. (1987b) Deuterium and oxygen-18 measurements on microliter samples of urine, plasma, saliva and human milk. Am. J. Clin. Nutr., 45: 905-913. [1] Wong W. W., Clarke L. L., Llaurador M. & Klein P. D. (1992a) A new zinc product for the reduction of water in physiological fluids to hydrogen gas for 2H/1H isotope ratio measurements. Eur. J. Clin. Nutr., 46: 69-71. [1] Wong W. W., Clarke L. L., Johnson G. A., Llaurador M. & Klein P. D. (1992b) Comparison of two elemental-analyzer gas-isotope-ratio mass spectrometer system in the simultaneous measurement of 13C/12C ratios and carbon content in organic samples. Anal. Chem., 64: 354-358. [9] Wong W. W., Klein P. D., Parr R. M. & Clements S. A. (1993) Interlaboratory analysis of reference water samples enriched in deuterium and oxygen-18. Int. J. Appl. Radiat. Isot., 44: 561-566. [1] Woo J. C., Jakubowski N. & Stuewer D. (1993) Analysis of aluminium oxide powder by glow discharge mass spectrometry with low mass resolution. J. Anal. At. Spectrom., 8: 881-889. [35] Woodbury S. E., Evershed R. P., Rossell J. B., Griffith R. E. & Farnell P. (1995) Detection of vegetable oil adulteration using gas-chromatography combustion isotope ratio mass-spectrometry. Anal. Chem., 67: 2685-2690. [8] Woodbury S. E., Evershed R. P. & Rossell J. B. (1998a) 613C analyses of vegetable oil fatty acid components, determined by gas chromatography combustion isotope ratio mass spectrometry, after saponification or regiospecific hydrolysis. J. Chromatogr.,, A 805: 249-257. [8] Woodbury S. E., Evershed R. P. & Rossell J. B. (1998b) Purity assessments of major vegetable oils based on 613C values of individual fatty acids. J. Am. Oil Chem. Soc., 75: 371-379. [8] Wortel T. M. (1979) Some applications of zeolites in organic chemistry. Delft University Press, Delft, Netherlands: 127 p. [36] Woutersen S., Milan J. B., Buma W. J. & deLange C. A. (1997) Resonance-enhanced multiphotonionization photoelectron spectroscopy of even-parity autoionizing Rydberg states of atomic sulphur. J. Chem. Phys., 106: 6831-6838. [33] Wright I. P., Boyd S. R., Franchi I. A. & Pillinger C. T. (1988) High-precision determination of nitrogen stable isotope ratios at the sub-nanomole level. J. Phys. E: Sci. Instrum., 21: 865-875. [13, 17] Wright I. P. & Pillinger C. T. (1989) Carbon isotopic analysis of small samples by use of stepped-heating extraction and static mass spectrometry. In: New frontiers in stable isotopic research; Laser probes, ion probes, and small-sample analysis. W. C. III Shanks & R. E. Criss (eds.). U.S. Geol. Surv., Bull. 1890: 9-34. [13] Wu S., Ding T., Meng X. & Bai L. (1997) Determination and geological implication of O-Si isotope of the sediment core in the CC area, the Pacific Ocean. Chinese Sci. Bull., 42(17): 1462-1465. [25] Xenakis D., Faucher O., Charalambidis D. & Fotakis C. (1996) Two-XUV-photon ionization of argon using a harmonic radiation source. N. Winograd & J.E. Parks (eds.). Resonance Ionization Symposium 1996, Eighth International Symposium. AIP Press Woodbury, NY, AIP Conference Proceedings 388: 365-368. [33]

References

1177

Xiao Y.-K., Beary E. S. & Fassett J. D. (1988) An improved method for the high-precision isotopic measurement of boron by thermal ionization mass spectrometry. Int. J. Mass Spectrom. Ion Proc., 85: 203213. [7] Xiao Y.-K. & Beary E. S. (1989) High-precision isotopic measurement of lithium by thermal ionization mass spectrometry. Int. J. Mass Spectrom. Ion Proc., 94: 101-114. [6] Xiao Y.-K., Jin L. & Qi H.-P. (1991) Investigation of thermal ion emission characteristics of graphite. Int. J. Mass Spectrom. Ion Proc., 107: 205-213. [7] Xiao Y.-K. & Zhang C.-G. (1992) High precision isotopic measurement of chlorine by thermal ionization mass spectrometry of the Cs2C1+ ion. Int. J. Mass Spectrom. Ion Proc., 116: 183-192. [28] Xiao Y. K., Qi H. P., Wang Y. H. & Jin L. (1992a) The comparison of loading forms in isotopic measurement of lithium by thermal ionization mass spectrometry. Chin. Sci. Bull., 37: 469-472. [6] Xiao Y., Sun Y. K., Wang Y. H., Oi H. P. & Jin L. (1992b) Boron isotope composition of brine, sediments and source water in Da Qaidam Lake, Qinghai, China. Geochim. Cosmochim. Acta, 56: 1561-1568. [46] Xiao Y.-K., Liu W.-G., Qi H.-P. & Zhang C.-G. (1993) A new method for the high-precision isotopic measurement of bromine by thermal ionization mass-spectrometry. Int. J. Mass Spectrom. Ion Proc., 123: 117-123. [28] Xiao Y. & Wang L. (2001) The effect of pH and temperature on the isotopic fractionation of boron between saline brine and sediment. Chem. Geol., 171: 253-261. [46] Xiao Y., Zhou Y., Wang Q., Wei H., Liu W. G. & Eastoe C. J. (2002a) A secondary isotopic reference material of chlorine from selected seawater. Chem. Geol., 182: 655-661. [40] Xiao Y. K., Zhou Y. M., Liu W. G., Hong A. S., Wang Q. Z., Wei H. Z. & Shirodkar P. V. (2002b) The character of chlorine isotopic compositions in ocean water. Geol. Review, 48: in press (in Chinese). [40] Xue S., Yang Y.-L., Hall G. S. & Herzog G. F. (1997) Germanium isotopic compositions in Canyon Diablo spheroids. Geochim. Cosmochim. Acta, 61: 651-655. [31] Yakir D. (1992) Variations in the natural abundance of oxygen-18 and deuterium in plant carbohydrates. Plant, Cell Environ., 15: 1005-1020. [23] Yakir D. & Wang X. F. (1996) Fluxes of CO2 and water between terrestrial vegetation and the atmosphere estimated from isotope measurements. Nature, 380: 515-517. [21] Yamamuro M. & Kayanne H. (1995) Rapid direct determination of organic carbon and nitrogen in carbonate-bearing sediments with a Yanaco MT-5 CHN analyzer. Limnol. Oceanogr., 40: 1001-1005.

[9]

Yarnamoto T., Hashizume K., Matsuda K. & Kase T. (1998) Multiple nitrogen isotopic components coexisting in ureilites. Meteorit. Planet. Sci., 33: 857-870. [17] Yamulki S., Toyoda S., Yoshida N., Veldkamp E., Grant B. & Bol R. (2001) Diurnal fluxes and the isotopomer ratios of N20 in a temperate grassland following urine amendment. Rapid Commun. Mass Spectrom., 15: 1263-1269. [19] Yang J. & Epstein S. (1983) Interstellar organic matter in meteorites. Geochim. Cosmochim. Acta, 47: 21992216. [11] Yang J. & Epstein S. (1987) The effect of the isotopic composition of oxygen on the non-mass-dependent isotopic fractionation in the formation of ozone by discharge of 02*. Geochim. Cosmochim. Acta, 51: 2011-2017. [18] Yang W., Spencer R. J. & Krouse H. R. (1995) Stable isotope and major element compositions of fluid inclusions in Devonian and Cambrian dolomite cements, western Canada. Geochim. Cosmochim. Acta, 59: 3159-3172. [3] Yang C., Telmer K. & Veizer J. (1996a) Chemical dynamics of the "St. Lawrence" riverine system: 6DH20, 618OH20, 613CDIC, 634Ssulfide, and dissolved 87Sr/86Sr. Geochim. Cosmochim. Acta, 60: 851-866.

[lOl

Yang W., Krouse H. R. & Spencer R. J. (1996b) Improved techniques for stable isotope analyses of microlitre quantities of water from fluid inclusions in halite and concentrated brines. Chem. Geol., 130: 139-145. [1, 3] Yanagisawa E & Sakai H. (1983) Thermal decomposition of barium sulfate - vanadium pentaoxide silica glass mixtures for preparation of sulfur dioxide in sulfur isotope ratio measurements. Anat. Chem., 55: 985-987. [26]

117 8

References

Yapp C. J. & Epstein S. (1982) A reexamination of cellulose carbon-bound hydrogen 6D measurements and some factors affecting the plant-water D /H relationships. Geochim. Cosmochim. Acta, 46" 955-965. [46] Yapp C. J. & Pedley M. D. (1985) Stable hydrogen isotopes in iron oxides - II. D / H variations among natural goethites. Geochim. Cosmochim. Acta, 49: 487-495. [2, 46] Yapp C. J. (1987) Oxygen and hydrogen isotope variations among goethites (R-FeOOH) and the determination of paleotemperatures. Geochim. Cosmochim. Acta, 51" 355-364. [2, 46] Yapp C. J. (1990) Oxygen isotopes in iron (III) oxides 1. Mineral-water fractionation factors. Chem. Geol., 85: 329-335. [461 Yapp C. J. (1991) Oxygen isotopes in an oolitic ironstone and the determination of goethite 6180 values by selective dissolution of impurities: The 5 M NaOH method. Geochim. Cosmochim. Acta, 55: 26272634. [2] Yapp C. J. & Poths H. (1991) 13C/12Cratios of the Fe(III) carbonate component in natural goethites. In" Stable isotope geochemistry: In: Stable Isotope Geochemistry: A tribute to Samuel Epstein. H. P. Jr. Taylor, J. R. O'Neil & I. R. Kaplan (eds.). The Geochemical Society, Spec. Publ. 3: 257-270. [2] Yapp C. J. & Poths H. (1993) The carbon isotope geochemistry of goethite (cz-FeOOH) in ironstone of the Upper Ordovician Neda Formation, Wisconsin, USA: Implications for early Paleozoic continental environments. Geochim. Cosmochim. Acta, 57: 2599-2611. [2] Yapp C. J. & Poths H. (1995) Stable hydrogen isotopes in iron oxides: III. Nonstoichiometric hydrogen in goethite. Geochim. Cosmochim. Acta, 59(16): 3405-3412. [2] Yapp C. J. (1998) Paleoenvironmental interpretation of oxygen isotope ratios in oolitic ironstones. Geochim. Cosmochim. Acta, 62: 2409-2420. [2] Yapp C. J. (2000) Climatic implications of surface domains in arrays of 6D and 6180 from hydroxyl minerals: goethite as an example. Geochim. Cosmochim. Acta, 64: 2009-2025. [46] Yeh H. W. (1974) Oxygen isotope studies of ocean sediments during sedimentation and burial diagenesis. PhD Thesis, Case Western Reserve University: 136 p. [2] Yeh H.-W. & Epstein S. (1978) 29Si/28Si and 30Si/28Si of meteorites and Allende inclusions. Lunar Planet. Sci., IX: 1289-1291 (Abstract). [25] Yeh H. W. (1980) D / H ratios and late stage dehydration of shales during burial. Geochim. Cosmochim. Acta, 44: 341-352. [2] Yi W., Halliday A. N., Lee D.-C. & Christensen J. N. (1995) Indium and tin in basalts, sulfides, and the mantle. Geochim. Cosmochim. Acta, 59: 5081-5090. [31] Yi W., Budd P., McGill R. A. R., Young S. M. M., Halliday A. N., Haggerty R., Scaife B. & Pollard A. M. (1999) Tin isotope studies of experimental and prehistoric bronzes. In: The beginnings of metallurgy. A. Hauptmann, E. Pernicka, T. Reheren, and U. Yalcin (eds.). Deutsches Bergbau-Museum: 285290. [31] Yildiz O. (1995) Ist der Isotopengehalt des Holzes aus Jahrringen von B~iumen als Klimaindikator geeignet? Diploma Thesis, Fachhochschule Aachen. [24] Yong J. W. H., Wong S. C., Letham D. S., Hocart C. H. & Farquhar G. D. (2000) Effects of elevated [CO2] and nitrogen on cytokinins in he xylem sap and leaves of cotton. Plant Physiol., 124: 767-779. [21] Yonge C. J. (1982) Stable isotope studies of water extracted from speleothems. PhD Thesis, McMaster University: 298 p. [3] Yoo R. K. & Gellene G. I. (1995) Symmetry induced kinetic isotope effects in the formation of (CO2)2+. J. Chem. Phys., 102: 3227-3237. [18] Yorozu M., Okada Y. & Endo A. (1999) Detection of hydrogen isotope atoms by mass spectrometry combined with resonant ionization. Appl. Phys. B - Lasers and Optics, 69" 129-132. [33] Yoshida N. & Matsuo S. (1983) Nitrogen isotope ratio of atmospheric N20 as a key to the global cycle of N20. Geochem. J., 17: 231-239. [19] Yoshida N. & Mizutani Y. (1986) Preparation of carbon dioxide for oxygen-18 determination of water by use of a plastic syringe. Anal. Chem., 58: 1273-1275. [1] Yoshida N. & Toyoda S. (2000) Constraining the atmospheric N20 budget from intramolecular site preference in N20 isotopomers. Nature, 405- 330-334. [15, 19, 34]

References

1179

Yoshinari T. (1990) Emissions of N20 from various environments The use of stable isotope composition of N20 as tracer for the studies of N20 biogeochemical cycling. In: Denitrification in soil and sediment. N. P. Revsbech & J. Sorensen (eds.). Plenum Press, New York: 129-150. [19] Yoshinari T., Altabet M. A., Naqvi S. W. A., Cosdispoti L., Jayakumar A., Kuhland M. & Devol A. (1997) Nitrogen and oxygen isotopic composition of N20 from suboxic waters of the eastern tropical North Pacific and the Arabian Sea - measurement by continuous-flow isotope-ratio monitoring. Marine Chem., 56: 253-264. [19] You C. F., Spivack A. J., Smith J. H. & Gieskes J. M. (1993a) Mobilization of boron in convergent margins: Implications for boron geochemical cycle. Geology, 21: 207-210. [7] You C.-F., Gieskes J. M., Chen R. & Spivack A. J. (1993b) I, Br, dissolved organic carbon, B, 611B and Mn in pore waters of Site 808, Nankai Accretionary Prosm. ODP, Scientific Results, 131: 165-174. [7] You C.-F., Butterfield D., Spivack A. J., Gieskes J. M., Gamo T. & Campbell A. (1994) Boron and halide systematics in submarine hydrothermal vent fluids: effects of phase separation and sediment contribution. Earth Planet. Sci. Lett., 123: 227-238. [7] You C.-F., Chan L.-H., Spivack A. J. & Gieskes J. M. (1995a) Lithium, boron and their isotopes in sediments and pore waters of Ocean Drilling Program Site 808, Nankai Trough: Implications for fluid expulsion in accretionary prisms. Geology, 23: 37-40. [6, 7] You C.-F., Spivack A. J., Gieskes J. M., Rosenbauer R. & Bischoff J. M. (1995b) Experimental study of boron geochemistry implication for fluid processes in convergent margin. Geochim. Cosmochim. Acta, 59: 2435-2442. [7] You C. F. & Chan L. H. (1996) Precise determination of lithium isotopic composition in low concentration natural samples. Geochim. Cosmochim. Acta, 60(5): 909-915. [6, 31, 40] You C.-F., Spivack A. J., Gieskes J. M., Martin J. B. & Davisson L. (1996a) Boron contents and isotopic compositions in pore waters: new approach to evaluate the temperature squeezing artifacts-geochemical implications. Marine Geol., 129: 351-361. [7] You C.-F., Castillo P., Gieskes J. M., Chan L. & Spivack A. J. (1996b) The role of slab derived fluids in arc magmatism. Earth Planet. Sci. Lett., 140: 41-52. [7] You C.-F., Shen J.-S. & Lee T. (1999) Precise boron isotopic determination by negative thermal ionization mass spectrometry. J. Nat. Cheng Kung Univ., Sci. Eng. & Med. Sect., 34: 1-11. [7] Young E. D., Fogel M. L., Rumble D. I. & Hoering T. C. (1998a) Isotope-ratio-monitoring of 02 for microanalysis of 180/160 and 170/160 in geological materials. Geochim. Cosmochim. Acta, 62(18): 3087-3094. [20] Young E. D., Coutts D. W. & Kapitan D. (1998b) UV laser ablation and irm-GCMS microanalysis of 180 / 160 and 170/160 with applicatiopn to a calcium-aluminium-rich inclusion from the Allende meteorite. Geochim. Cosmochim. Acta, 62(18): 3161-3168. [20] Young E. D., Nagahara H., Mysen B. O. & Audet D. M. (1998c) Non-Rayleigh oxygen isotope fractionation by mineral evaporation: Theory and experiments in the system SiO2. Geochim. Cosmochim. Acta, 62: 3109-3116. [46] Young J. P., Barshick C. M., Shaw R. W. & Ramsey J. M. (1994) Application of diode lasers to the isotopically selective determination of uranium in oxides by optogalvanic spectroscopy. H.-J. Kluge, J.E. Parks & K. Wendt (eds.). Resonance Ionization Spectroscopy 1994, Seventh International Symposium. AIP Press Woodbury, NY. AIP Conference Proceddings 329: 111-114. [33] Yu F. J. (1991) The measurement of oxygen isotope composition of exceeding micro-water; capillary CO2-H20 room temperature equilibrium method. Kexue Tongboa (Chinese Sci. Bull.), 36(2): 140-144.

[3]

Yuen G., Blair N., DesMarais D. J. & Chang S. (1984) Carbon isotope composition of low molecular weight hydrocarbons and monocarboxylic acids from the Murchison meteorite. Nature, 307: 252-254.

[11]

Yun M., Wadleigh M. A. & Pye A. (in press) Direct measurement of sulphur isotopic composition in lichens by continuous flow isotope ratio mass spectrometry (CF-IRMS). Chem. Geol. [26] Yund R. A. & Anderson T. F. (1978) The effect of pressure on oxygen isotope exchange between feldspar and water. Geochim. Cosmochim. Acta, 42: 235-239. [46]

1180

References

Yung Y. L., Lee A. Y. T., Irion F. W., DeMore W. B. & Wen J. (1997) Carbon dioxide in the atmosphere: Isotope exchange with ozone and its use as a tracer in the middle atmosphere. J. Geophys. Res., 102: 10857-10866. [18] Yung Y. L. & Miller C. E. (1997) Isotopic fractionation of stratospheric nitrous oxide. Science, 278: 17781780. [19] Zaback D. A. & Pratt L. M. (1992) Isotopic composition and speciation of sulfur in the Miocene Monterey Formation; reevaluation of sulfur reactions during early diagenesis in marine environments. Geochim. Cosmochim. Acta, 56(2): 763-774. [26, 29] Zaback D. A., Pratt L. M. & Hayes J. M. (1993) Transport and reduction of sulfate and immobilization of sulfide in marine black shales. Geology, 21(2): 141-144. [29] Zadnik M. G., Specht S. & Begemann F. (1989) Revised isotopic composition of terrestrial mercury. Int. J. Mass Spectrom. Ion Proc., 89: 103-110. [31] Zaluski G., Nesbitt B. & Muehlenbachs K. (1994) Hydrothermal alteration and stable isotope systematics of the Babine porphyry Cu deposits, British Columbia: implications for fluid evolution of porphyry system. Econ. Geol., 89: 1518-1541. [3] Zawislanski P. T. & Zavarin M. (1996) Nature and rates of selenium transformations; a laboratory study of Kesteron Reservoir soils. Soil Sci. Soc. Amer. J., 60(3): 791-800. [29] Zeebe R. E. (1999) An explanation of the effect of seawater carbonate concentration on foraminiferal oxygen isotopes. Geochim. Cosmochim. Acta, 63: 2001-2007. [46] Zehr J. P. & Oremland R. S. (1987) Reduction of selenate to selenide by sulfate-respiring bacteria: Experiments with cell suspensions and estuarine sediments. Appl. Env. Microbiol., 53(6): 1365-1369. [29] Zeng Y., Mukai H., Bandow H. & Nojiri Y. (1994) Application of gas chromatography-combustionisotope ratio mass spectrometry to carbon isotopic analysis of methane and carbon monoxide in environmental samples. Anal. Chim. Acta, 289: 195-204. [14] Zhabina N. N. & Volkov I. I. (1978) A method of determination of various sulfur compounds in sea sediments and rocks. In: Environmental biogeochemistry and geomicrobiology. W. E. Krumbein (ed.). Ann Arbor Science, Ann Arbor, MI: 735-746. [26, 27] Zhai M., Nakamura E., Shaw D. M. & Nakano T. (1996) Boron isotope ratios in meteorites and lunar rocks. Geochim. Cosmochim. Acta, 60: 4877-4881. [7] Zhang B. L., Vallet C., Blanchard F., Martin G. J. & Combaut G. (1994a) Charaterization of galactose from Carrageenans by natural abundance site-specific isotope parameters. Phytochem. Analysis, 5: 291-296. [5] Zhang B. L., Quemerais B., Martin M. L., Martin G. J. & Williams J. M. (1994b) Determination of the natural deuterium distribution in glucose from plants having different photosynthetic pathways. Phytochem. Analysis, 5: 105-110. [5] Zhang B. L., Buddrus S., Trierweiler M. & Martin G. J. (1998a) Charaterization of glycerol from different origins by 2H- and 13C-NMR studies of site-specific natural isotope fractionation. J. Agric. Food Chem., 46(4): 1374-1380. [5] Zhang D. (1988) Nitrogen concentrations and isotopic compositions of some terrestrial rocks. PhD Thesis, The University of Chicago: 157 p. [16] Zhang J., Quay P. D. & Wilbur D. O. (1995) Carbon isotope fractionation during gas-water exchange and dissolution of CO2. Geochim. Cosmochim. Acta, 59(1): 107-114. [46] Zhang L., Chan L. H. & Gieskes J. M. (1998b) Lithium isotope geochemistry of pore waters from Ocean Drilling Program Sites 918/919, Irminger Basin. Geochim. Cosmochim. Acta, 62: 2437-2450. [6] Zhang Q. & Ding T. (1989) Analysis of the reference material NBS-123 and the atomic weight of sulphur. Kexue Tongbao (Chinese Sci, Bull.), 34: 1086-1089. [44] Zhang Q.-L. & Li W.-J. (1990) A calibrated measurement of the atomic weight of carbon. Chinese Sci. Bull., 35(4): 290-296. [40, 43] Zhang T. & Krooss B. H. (2001) Experimental investigation on the carbon isotope fractionation of methane during gas migration by diffusion through sedimentary rocks at elevated temperatures and pressure. Geochim. Cosmochim. Acta, 65: 2723-2742. [46] Zhang X., Nesbitt B. & Muehlenbachs K. (1989) Gold mineralization in the Okanangan Valley, southern British Columbia: fluid inclusion and stable isotope studies. Econ. Geol., 84: 410-424. [3]

References

1181

Zhang Y., Stolper E. & Wasserburg G. J. (1991a) Diffusion of a multi-species component and its role in oxygen and water transport in silicates. Earth Planet. Sci. Lett., 103: 228-240. [46] Zhang Y., Stolper E. & Wasserburg G. J. (1991b) Diffusion of water in rhyolitic glasses. Geochim. Cosmochim. Acla, 55: 441-456. [46] Zheng H. W. (1993) Calculation of oxygen isotope fractionation in hydroxyl-bearing silicates. Earth Planet. Sci. Left., 120: 247-263. [2] Zheng Y.-F., Metz P., Satir M. & Sharp Z. D. (1994) An experimental calibration of oxygen isotope fractionation between calcite and forsterite in the presence of a CO2-H20 fluid. Chem. Geol., 116: 1727. [46] Zheng Y.-F., Satir M., Metz P. & Sharp Z. D. (1999) Oxygen isotope exchange processes and disequilibrium between calcite and forsterite in an experimental C-O-H fluid. Geochim. Cosmochim. Acta, 63: 1781-1786. [46] Zhu X. K., O'Nions R. K., Guo Y., Belshaw N. S. & Rickard D. (2000a) Determination of natural Cuisotope variation by plasma-source mass spectrometry: implications for use as geochemical tracers. Chem. Geol., 163: 139-149. [31] Zhu X.-K., O'Nions R. K., Guo Y. & Reynolds B. C. (2000b) Secular variations of iron isotopes in North Atlantic deep water. Science, 287: 2000-2002. [31] Zhu X.-K., Guo Y., O'Nions R. K., Young E. D. & Ash R. D. (2001) Isotopic homogeneity of iron in the early solar system. Nature, 412: 311-313. [31] Zhu X. K., Guo Y., Williams R. K., O'Nions R. K., Matthews A., Belshaw N. S., Canters G. W., de Waal E. C., Weser U., Burgess B. K. & Salvato B. (2002) Mass fractionation processes of transition metal isotopes. Earth Planet. Sci. Lett., 200: 47-62. [46] Zinner E., McKeegan K. D. & Walker R. M. (1983) Laboratory measurements of D / H ratios in interplanetary dust. Nature, 305: 119-121. [30] Zinner E. & Crozaz G. (1986) A method for the quantitative measurement of rare earth elements in the ion microprobe. Int. ]. Mass Spectrom. Ion Proc., 69: 17-38. [30] Zinner E., Tang M. & Anders E. (1987) Large isotopic anomalies of Si, C, N and noble gases in interstellar silicon carbide from the Murray meteorite. Nature, 330: 730-732. [20, 25, 30] Zinner E. (1989) Isotopic measurements with the ion microprobe. In: New frontiers in stable isotopic research: Laser probes, ion probes, and small-sample analysis. W. C. III Shanks & R. E. Criss (eds.). U.S. Geol. Surv., Bull. 1890: 145-162. [30] Zinner E. (1998) Stellar nucleosynthesis and the isotopic composition of presolar grains from primitive meteorites. Annu. Rev. Earth Planet. Sci., 26: 147-188. [17] Zuleger E., Gieskes J. M. & You C.-F. (1996) Interstitial water chemistry of sediments of the Costa Rica Accretionary Complex off the Nicoya Peninsula. Geophys. Res. Lett., 23: 899-902. [7]


SUBJECT INDEX


1183

Subject Index

A absence of air absolute calibration absolute isotope abundance absolute isotope composition absolute isotope ratio

1006 506, 777, 890, 909 136, 902, 905, 964, 977, 978 706

165, 526, 533, 828, 874878,885,901,902,945 absolute measurement 706, 776 absolute sensitivity 853 absorption 13,15,104,156,282,378, 381, 405, 433, 446, 455471, 748-758, 760-783, 812,1040 absorption (on resin) 336 absorption lines 378 absorption spectrometry 47, 123, 907 absorption spectroscopy 380, 396, 399, 768-783 absorption spectrum 466, 776-779,880 abundance of artifacts 945 abundance sensitivity 96, 666, 694, 695, 697, 728,742,743,851,854 accelerating beams 642,655,657,664 accelerating voltage 123, 534, 656, 659, 727, 734, 735, 798, 799, 844, 847,860,862,864-866 acceleration of electrons 667, 791, 792, 859, 860, 864-866 acceleration of ions 661, 694, 791, 839, 841 acceleration potential 661, 685, 709, 804, 839, 843, 844 accelerator mass spec- 204, 609 trometer (AMS) acetic acid 40, 43, 48, 120, 125, 201, 231, 320, 512, 514, 517 acetic anhydride 165 acetone 16, 117, 339, 365, 427, 448, 451, 457, 458, 528, 529, 531, 572, 576, 577, 589, 613, 616 acid chemistry 240 acid digestion 147, 318, 319, 427, 450, 487, 571, 720, 895 acid environment 249 acid lake water 11 acid matrix 636, 637, 708 acid soil 590, 592, 1013 acid treatment 43, 56, 257, 266 acid volatile sulfur 570, 571, 574-577, 580, 603 acrylic fibers 490 acrylic matter 490

activated clay activation energies active fluorine / fluorination compounds active submarine chimney acetylene adiabatic energies adjacent peaks adjacent standard adsorbed air adsorbed moisture adsorbed water

581 581, 994 451, 526, 531 723

4,17 387 168, 848 133, 703 365 461 8, 9, 13, 47, 49, 53, 59, 72, 74, 298, 444, 493, 502 adsorption 4, 6, 8, 12, 14, 36, 72, 74, 123, 158, 189, 190, 214, 215, 318, 327-330, 338, 378, 430, 465, 471, 559, 564, 585, 591, 649, 809814, 819, 911, 926, 1016, 1019, 1034 adsorption (on resin) 330, 338 adsorption peak 811 aerated soil 586, 587, 590, 592 aerodynamic sizing 561 aerosol 132, 560, 561 aerosol (dry) 720, 725 aerosol nitrate 336, 376 aerosol particles 550, 561, 562 aerosol sulfate 376, 560, 561 Ag + exchange membrane337 Ag3PO4 crystal 482, 486-494 agar 200 AgPd alloy 1007, 1008 agricultural 2-5, 17, 19, 25, 29, 32, 180, 181, 305, 313, 320, 340 agriculture 813 air 3, 10, 18, 19, 25, 45, 95, 157, 174, 183, 189, 214223, 227, 274-280, 283, 287, 292-303, 307, 317, 341, 354, 369, 371, 380, 383, 458, 465, 475, 486, 488, 502, 557-561, 565, 566, 570, 611, 613, 616, 657, 815, 817, 858, 863, 870, 879, 895, 898, 899, 910, 945, 1013, 1014, 1018, 1024 air actuated pistons 284 air bath 20, 22, 25 air bubbling 207, 557 air conditioning 1039

1184 air extraction air flow rate air intake air leak / leaking air moisture air of breath air pollution air pressure air purification system air reference air sample air sample (modern) air sampler air standard air stream air temperature air toxics airborne emission airlock module airtight conditions airtight vial / container airway reactivity alcohol aldehydes algae algae strains algal assimilation algal culture algal population algal proteins algal species algal uptake algorithm alignment of magnet alignment of primary beam alignment / centering check aliphatic hydrocarbons alkali basalt alkali salts alkaline earth elements alkaline oxidation allanite crystal Allende chondrite altered basalt alumina tube aluminium foil boat

Subject Index 275 558,815 213 341,534 502,815,892 89,9Z 98,158 191 100,293 8O8 279,293,294,300,372 179, 272-297, 302, 303, 761,815,818 269 558 274, 280, 289, 302, 363, 364,367-372 281 473 951 381 50 215 215,220,224,328,347 378 4, 19, 25, 32, 159, 160, 163-166, 195, 196, 198, 221,858,896 159,163 650,1013 1014 1033 649 179 200 179 649 115, 168, 169, 439, 504, 634, 646, 779, 944-956, 974,979-990 851 675 458 233,234 85 559 31,867 586,588,593 66 721 1033 175 242

aluminum boat aluminosilicates alunite Amberjet (ion exch. resin) Amberlite (Na resin) Amberlyst ambient air amino acid

amino acid standard amino (acid) sugar ammonium carbamate ammonium citrate reagent ammonium contamination amorphous silica amphibole amplifier gain drift analyte matrix analytical artifact analytical chemistry analytical contamination analytical science analytical sensitivity analytical setup analytical statistics ancient (early) atmosphere ancient materials ancient organic matter ancient sediments Anderson pyrite andesite angular momentum anhydrite anhydrous environment anhydrous phosphoric acid anhydrous process aniline animal animal animal animal animal animal animal

cellulose feeding glue protein teeth tissue

190 40, 43 444, 445, 572, 1003 486-488, 492, 493 41, 333, 486-489, 493, 564 499 9Z 159,211,293,300 120, 158, 161-167, 173, 189-201, 230-236, 555, 565, 585, 588, 592, 595, 1019 199 320,344,346 24 485 190,320 43,56,1011 570,687,1017 634 708,723 123,687 154,155,639,794,935 274 820 130,139,149 204, 684 288 269,448 229 185 597 687 127 103, 748-750, 752, 754 553,568,572,573,599 255 27 246 338,339 179-183, 200, 378, 594, 595,1015 1030 177 529 200 183 182,186,191

Subject Index animal waste anion exchange chemistry anisotopic standard annual plants anode dark space Antarctic

1185 305 715

136 50L 516 791, 792 7, 80, 84, 292, 305, 310, 311, 376, 476, 681, 775, 815,885,889,945 564 Antarctic firn cores Antarctic nitrate aerosol 336 377 anthropogenic hydrocarbons anthropogenic sources 556, 557, 559 58, 68, 407, 446, 483-486, apatite 495, 496, 605, 619, 687, 688, 1013 apparent isotopic frac- 447 tionation 1 applied science 16,457 aqua regia 643 Aquadag paste 187 aquatic ecology 178,179,181,190,329 aquatic ecosystem 1014 aquatic plants 743 aquisition cycle Ar ion/atom interaction 792, 793 46 aragonite 269 Archean age 268,269 Archean chert Archean sedimentary 685,687 sulfides 716, 724 archeological artifact archeological samples 716, 724 716 archeological studies archeology 907 231,233-235,581 aromatic compounds aromatic deuterium 120 nuclei aromatic entities 235 aromatic formulation 118 (alkyl) aromatic frag233 ments aromatic hydrocarbons 163,198 aromatic positions 118 aromatic ring 118, 120 aromatic signals 117 aromatics 581,582,584 Arrhenius diagram 263 Arrhenius equation 266 Arrhenius expression 263 arsenopyrite 569 945,952 artifact scale Ascarite trap 334

Ascarite II ascorbic acid aspartic acid asphaltenes asteroid asymmetric stretching band Atlantic seawater atmosphere (the)

213,223,293 649 161,191,231,232,1019 581,582,584 229,236 779

137, 712 49, 54, 177, 180, 220, 257, 307, 309, 375-379, 389, 390, 417, 501, 538, 541, 552-559, 596, 607, 889 atmosphere monitoring 273, 290 atmospheric air 307, 364, 372, 501, 782, 894 272-304, 775, 778, 818, atmospheric CH4 951-956 atmospheric chemistry 390 376 atmospheric CO 180, 204, 213, 215, 21Z atmospheric CO2 220, 224, 272-304, 376, 779,894 atmospheric components 378, 383 349 atmospheric contamination 38O atmospheric data 189,379,380, 794 atmospheric gases 376 atmospheric H202 305 atmospheric load 291,775,778,951 atmospheric methane 9,10,416,457,502 atmospheric moisture 199, 357, 361, 392, 877, atmospheric N2 879,898 376,394,775 atmospheric N20 390 atmospheric N20 lifetime 309,329,336,343,376 atmospheric nitrate 335, 336, 376, 380, 455, atmospheric 02 1013,1015 309,376-381,384,388 atmospheric ozone 46, 98, 694, 789, 950, atmospheric pressure 1005-1008, 1012, 10181025 375,384 atmospheric species 343,380,899 atmospheric studies 3,25,47 atmospheric vapor 3,203,443,491,496,688 atmospheric water 106 atomic environment 703, 706,718,835 atomic weight 1008,1020 Au capsule 178 Australian tropical stream 1025-1027 autoclaves automated combustion 322,341

1186 automated device automated instrumentation automated Rittenberg analysis (ARA-MS) autosampler

Subject Index 178, 237, 238, 838 339 341

6, 7, 36, 165, 321, 341, 503, 570 Avogadro constant 924 921 Avogadro project Avogadro spectrometer 925 azeotropic distillation 3, 4, 31, 54,473-475, 479, 480 R-cellulose 507-521

B Ba(OH)2 reagent 211, 213 backbone procedure 419 back-diffusion 805 background concentra- 290 tion background condition 852 background contamina- 286 tion background contribution 537 background correction 743, 744, 852 background currents 669, 978 background determina- 852, 854 tion background drift 299 background emission 536 background estimation 743 background impurities 861 background information 314, 948 background interference 101 background level 287, 393, 503 background measurement852 background monitoring 470 background nitrogen 312 background noise 287, 677 background peaks 612 background pressure 873 background problem 500 background reduction 468 background samples 161 background scan 611 background science 89 background signal 742, 778 background water 287 backup filter 561 BaC12-NaOH reagent 211 bacterial experiment 650 bacterial methane 291 bacterial strains 335, 650, 994, 1012

bactericide bacteriological reduction bacterioplankton BaF2 window Balmat pyrite barite basalt basalt glass basalt standard basaltic magma basaltic melt baseline baseline baseline baseline baseline baseline baseline

conditions correction counts oscilations peaks separation

206, 211, 562 994

184 468,532 686,687 66, 43Z 465, 552, 570, 572, 573, 576, 600,602,1007 68, 125-127, 137, 350, 648, 710, 712, 904,1023 70,443,689 133,134,648 1027 1027 158, 161, 299, 614, 852,911 300 115 646 777 168,169 161, 168, 170, 173, 193 773 pro- 715

568, 579, 139, 903,

769,

174,

baseline slope baseline subtraction cedure baseline values 173, 181, 184 basic isotopic fractiona- 624 tion systematics batch reactor 647 batch reduction 9, 11, 22, 35, 36 beam defining slits 860, 861 Beer's law 768, 771, 774, 776, 780 beet root sugar 116 beet sucrose standard 476 bellows pump 457, 951 bellows reservoir (IRMS)242, 276, 297, 299, 300, 350, 357, 358, 838, 851, 879 275,410,416,950,951 bellows valves 120,231 benzaldehyde 17, 19, 44, 117, 118, 231, benzene 516,572,582,584 118,120,121,891,897 benzoic acid 315 Berthelot reaction 40, 43, 48, 61, 564, 896, bicarbonate 1018,1023 108, 161, 165, 175, 177, biochemical 181-192 103,820,907 biochemistry 202 biochemists 407,437,496,1030,1035 biogenic apatite

Subject Index biogenic mineral biogenic opal biogenic pyrite biogenic shell biogeochemistry biological activity Biological and Environmental Reference Materials (BERM) biological (re)cycling biological fluid biological fossils biological harmful radiation biological marker biological method biological process

1187 1028 52 578 1029 179,358,891,89G 899 684,717 948 306, 496, 628 3-5, 10, 17, 19, 25, 29, 32, 481 684

377 187 118 177, 181, 183, 450, 482, 625, 653, 1028 biological sample/ 31, 177, 179, 185, 186, matter 191, 194, 197, 201, 306, 313, 341, 344, 541, 555, 593, 594, 716, 1011, 1023, 1024 biological science 340 biological studies 474 biological tissue 188 biological tracer 2 biology 1, 1043 biomedical applications 756, 891, 898 biomedical fluid 11 biomedical science 122, 172 biomedical study 6 biomedicine 772 biominerals 1011, 1028 BioRad 328, 329, 331, 564, 640, 642 biosphere 177, 200, 272, 538, 552, 596 biosynthesis 116, 175, 182, 192, 196, 198-200 biosynthetic 108 biotite 348, 351-354, 357-359, 429, 441, 442, 445, 570, 687, 884, 890, 910, 1021, 1031 BiPO4 crystal 485, 486, 491 birch leaves 519 bird contamination 313 bird droppings 312 birds 172, 178, 181, 200, 312, 313, 594 bitumen 173, 578-583 bituminous organic sulfur572

black shale bladder stones blank

45, 46, 597-603 595 8, 9, 59, 127, 246, 251, 324, 328, 336, 348-350, 356-374, 410, 423-440, 452, 453, 459-461, 465, 471, 496, 497, 500, 501, 505, 550, 558, 569, 578, 640, 647, 648, 709, 729, 731, 736, 743, 748, 756, 800, 817, 818, 833, 852, 1050 blank contribution 328, 373, 374, 469, 822 blank correction 8 blank run 303 blank solution 165 bleaching 512,516,517 blend samples 729, 736, 738-740 block heater 11 blood (serum) 4, 29, 88, 98, 102, 182, 191, 197, 206, 594, 595 Boltzmann constant 105, 669, 867, 924 Boltzmann equilibrium 105, 114 Boltzmann populations 770 Boltzmann relation 105 bond stretching 385 bone 180, 186, 188, 1029, 1030 bone collagen 199, 1032 bone-dry 299 boric acid 125, 127, 128, 201, 314, 317, 322, 327, 345, 641, 645, 876, 903, 904 boric acid standard 142, 147, 148, 150 boron reagent 128 boron spike 146 borosilicate 12, 14, 15, 237, 240, 242, 532, 533, 612, 615, 805, 806 botanical origin 112, 116, 117, 121 botanical source 115 Boudouard equilibrium 498, 499 Bouguer-Lamber-Beer 771 law brachiopods 1011,1035 brackish water 639 breath (human) 2, 88, 89, 93, 95-99, 101, 102, 158, 159, 761-763, 778, 779,816 brines 4, 5, 7, 16, 1Z 2Z 29-33, 74, 130, 139-141, 269, 452,1018,1019,1034 broad peak 268,839 broad-leaved trees 515,520 bromination 490,491

1188

Subject Index

bromine pentafluoride/ 24, 30, 49, 50, 57, 400BrF5 472, 482, 483, 490-493, 524, 526-531, 542, 546, 818 bromine trifluoride/ 408, 409, 412, 417, 424, BrF3 442, 446, 447, 451, 490, 528, 531, 542, 545, 546, 924 bromoform 42 bromomethane 606, 619-621 bronze alloy 724 bronze artifact 724 bulk air 295 bulk meteorite 371 bulk tissue 190, 201 butane 231

C caffeine calcite lattice calcium carbide calcsilicate calculated equilibrium calculated isotope fractionation calibration

334, 897 573 17 78, 83 62 359

22, 30, 34, 111, 132, 274, 279, 280, 298, 300, 311, 337, 357, 398, 425, 439, 471, 476, 478, 498, 503, 599, 675, 681, 687, 745, 765, 779, 780, 784, 795, 796, 874-906, 909, 916920, 993, 994, 1000, 1006, 1028, 1032 calibration (one point) 336 calibration (two point) 980,981 calibration (mult. point) 336 478,599,795,982,983 calibration curve 906 calibration cycle 358 calibration line 881, 889, 890, 892, 893, calibration material 901, 917 35,392,505,506,794 calibration method 398,425 calibration standard 200 Calvin cycle 116 cane sugar 183 canopy air(forest) 864,865 capacitance-resistance filter 11, 18, 20, 25, 28, 63, 87, capillary (-ies) 97, 98,153, 155, 157, 168173, 276-299, 451, 502, 543, 546, 598, 611, 836,

838,853,953,1025,1040 342 290 273,274,304 59,174, 180, 248, 272, 321, 499, 602, 843, 894, 944, 947, 948, 956, 96Z 1040 carbon isotope labeled 896,1020 162,377,499,1040 carbon monoxide 489 carbon oxidation 8,29,36,343 carbon pyrolysis 7, 23, 7Z 342, 401-404, carbon reduction 408,429 carbon rod reactor 410 carbonaceous chondrite 229, 233, 720, 1023 carbonaceous meteorite 256, 260 carbonate 27, 31, 40-48, 56, 58-61, 64, 69, 74, 81, 125, 129, 132, 133, 136, 139, 180, 183, 188, 203, 211-223, 237-255, 258, 261-263, 280, 402, 405, 427, 444, 461, 483-485, 488, 495, 496, 497, 524, 529, 563, 572, 573, 579, 586, 590, 591, 610, 614, 618, 619, 625, 657, 681, 684, 707, 713, 723, 838, 852, 876883, 892-895, 902, 903, 910, 912, 1000, 1005, 1015, 1020, 1025-1030, 1035 carbonatite 64, 83, 241, 243, 618, 619, 876, 884, 892 carbonic anhydrase 26 carboniferous chomdrite 536 503 carbosorb CO2 trap 161, 164, 166, 230, 231, carboxylic acid 233, 234 514, 582, 584, 625 carcinogen 178 carniforous fish 255, 882, 894 Carrara marble 8, 9, 18, 29, 45, 46, 89, 92carrier gas 102, 132, 157-175, 225, 286, 296, 297, 303, 334, 341, 358, 419, 430, 433, 459-471, 501, 503, 546, 550, 551, 569, 571, 612, 618, 637, 638, 725, 782, 790-794, 817, 818, 853, 854, 879, 882, 1040 560, 561 cascade impactor

capsule crimping carbon cycle (the) Carbon Cycle Group carbon dioxide

Subject Index catalyst (a, the)

18, 19, 20, 88, 160, 298, 318, 319, 402, 501, 552, 598, 805, 941, 1016, 1021 catalyst poisoning 19 catalytic activity/action 17, 19, 22, 430 catalyze(d)/catalysis 7, 19, 26, 29, 75, 164, 298, 335, 344, 417, 499, 625, 1019 catalyzed graphite 333, 334, 340, 410 / carbon cathode dark space 791, 792 cathodic sputtering 789, 792 cathodoluminescence 674 causal law 699 cavity enhanced absorp- 777 tion spectroscopy (CEAS) cavity ring down 776 spectroscopy (CRDS) celestite 599 cell membrane 20O cellulose 200, 201, 474-480, 497522,883,896,1014,1030 cellulose filters 556,558,560,562 cellulose nitrate / nitrat. 508,509,1033 cellulose oxidation 499 cellulose standard 503,511,520, centrifugation 3, 4, 40-42, 46, 53, 56, 201, 335, 338, 486, 590, 617, 639 centrifugation-agitation 98 cycle centrifuge tube 188, 318, 550 ceramic frit 582 ceramic insulator 858, 862 ceramic material 500, 803 ceramic parts 863, 864 ceramic rods 864 ceramic spacers 858, 864 ceramic tube 165, 298, 334, 501, 1023 ceramics 788, 794, 859 certified isotopic compo- 648, 706 sition certified reference mate- 112, 828, 830, 881, 890, rial 891, 896, 897, 903, 907927, 935 certified site specific 112 ratio certified spike 820,827,829 certified values 733, 734, 740-742, 907, 945 chalcopyrite 548,599,716,966-968 changeover valve 836-838,843,851-853 channeltrons 670,695,697

1189 charcoal charcoal (activated) charge neutralization charge-balanced lattice charging cycle chemical bonding environment chemical conversion chemical environment chemical impurity chemical lifetime chemical memory chamical oxidation chemical reactivity chemical reactor chemical reagent chemical separation resin chemical thermometer chemical tracer chemical treated filters chemically homogene-

6, 23, 185, 338, 363 19, 338, 439, 462, 819 654, 664, 675 813 872 601 338, 541, 598 106 920 379 843 117, 197, 586, 593 547 907 4, 257 339. 342, 493 1031 557 557,558, 562, 565 1016

OUS

chemically scrubbed air 578 chemisorption 366, 813, 814, 818 chemistry 1, 152, 202, 464, 564, 673, 748, 808, 882, 929, 939, 995, 998, 1010, 1016, 1017, 1043 chemistry laboratory 496 chemistry of the phases 1006 chert 268, 269, 406, 1032 Chilean perchlorate 604 chitin 200, 1015 chlorinated solvent 7, 615 chlorine radical 291 chlorine trifluoride 24, 30, 49, 400, 404, 408C1F3 412, 422, 423, 431, 433436, 442, 446, 452, 455, 467, 468, 471 chlorite 41-43, 48, 49, 51, 351, 429, 445, 514, 1001, 1003, 1004, 1031 chloroform 44, 165, 197, 517 chloromethane 607-610, 613, 615, 616, 620, 621 chlorophyll 200, 710, 712 cholesterol 186, 191, 196-198 chondritic meteorite 724 chromatographic 74 analysis chromatographic column193, 195, 284, 494, 495, 546, 598, 602, 818, 854, 891

1190 chromatographic chromatographic ference chromatographic tope effect chromatographic

Subject Index filter inter-

416 855

iso-

168,171

peak

156, 168, 169, 297, 853, 854 171

chromatographic peak distortion chromatographic performance chromatographic separation chromatographic technique chromatography chromium

169 156, 195, 293, 335, 547, 557,621,1021 113, 125, 54Z 853, 856, 880 160, 169, 230, 232, 287, 836,853,854,856 Z 15, 623, 625, 627, 640, 644, 648, 650, 783, 829, 918 648 783 626,650

chromium chloride chromium doped... chromium isotopic fractionation chromium powder chromium reduced sulfur570-580 chromium reduction 571, 572, 576, 579, 588 Chromosorb 157, 441, 564 cinnabar 447, 720 circulating/pumping air 3, 275, 288, 815 citric acid 120, 610, 611, 618 classical fluorination 407 classical isotope frac309 tionation clay (mineral/sedim.) 15, 38-61, 422, 461, 590, 795, 813, 904, 993, 1016, 1019, 1032, 1034 clay fraction 39, 41 clay membrane 53 clay rich 22, 38, 53-56, 60, 139 clay separation 38-40 clay size 41, 42 clay standard 45, 47, 50 clean air site 274, 277 cleaning procedure 13, 320, 365, 709, 863 clear glass window 467 climate change 180, 183, 306 climate conditions 36, 497 climate monitoring 273, 274, 304 climate reconstruction 497 climate sensitive ar507 chive climate, affect on.. 272 climate, disrupt the.. 378

climate, influence on.. clinical applications clinical diagnosis clinical environment clinical investigation clinical setting clinical tracer study clinopyroxene Co catalyst CO gas laser CO standard CO2 contaminant CO2 fluorination technique CO2 laser

377 758, 784 88,98 88,779 98 779 6 68 1021 778,780 503,504 205,220,22Z 286 337

361, 362, 364, 436, 458, 459, 463, 464, 466-468, 472, 495, 532, 760, 764, 780 CO2-CH4 equilibration 1021 211-213,218,221,223 CO2-free atmosphere CO2-dissolved bicarbo- 1018 nate equilibration CO2-plant water equi- 473-481 libration CO2-water equilibration 3, 4, 17-35, 55, 66, 226, 408, 451, 476, 1018 coal 291, 559, 560, 568, 578583, 894 coal carbon 46 coal-fired power plant 559 coca leaves 172 cocaine 172 collector arrangment 534 collector array 697 collector assembly 858, 870 collector block 695, 697 collector ceramics 858 collector feed throughs 858 collector plate 847 collector slits 393, 664, 665, 675, 857, 859, 861 collector system 390, 393, 857, 859, 990 collision cell 695, 706, 722 collision gas 695, 708, 722 collisional interaction 793 colloidal graphite 643 colloidal Pt powder 494 colloidal silica 645, 646 colloidal solution 617 colloidal suspension 640, 645 color center crystals 782 color center laser 764, 772, 782 combined uncertainty 731, 733, 734, 738, 740, 745, 911, 916, 922

Subject Index 154, 157, 162, 165, 170, 182, 189, 192, 193, 225, 233, 256, 258, 293, 296, 298, 332, 333, 338-340, 342, 350, 356, 357, 366, 403, 505, 544, 545, 559, 597, 598, 602, 615, 817, 971-973, 982, 985, 988, 1051 combustion / oxidation 593 combustion / pyrolysis 258, 817, 818 combustion apparatus 294 578 combustion boat 162, 166 combustion catalyst 162 combustion chamber 182, 298 combustion efficiency combustion experiment 261, 266, 366 combustion furnace 224, 225, 294, 297-299 combustion gases 168, 365, 366 combustion interface 154, 165, 171 combustion mode 362, 366 combustion of sample on 326, 327 resin combustion procedure 601,818 171 combustion process 168, 174, 321, 341, 578, combustion products 615 162,165,173 combustion reactor 330 combustion reagent 162,302,358,544 combustion system 544,547,578,817,895 combustion technique / method combustion tube/vessel 70, 165, 298, 324, 330, 332, 336, 340, 342, 346, 544, 593, 618 common air 380 complete electrolysis 23 (in)complete fluorination 420, 443, 450, 461 complete oxidation 44 complex ecosystem 202 complex interactions 178, 180, 184 complex matrix 112, 154, 197 complex organic mixtures232, 233 compound peaks 17, 126, 129, 144, 146, 149, 152, 156, 168, 171, 284-286, 288, 290, 297299, 303, 341, 367-369, 371, 537, 593, 612, 621, 639, 641, 657, 799 compound specific 153, 156, 179, 182, 185188, 191-193, 202, 229236, 547, 596 242,1040, 1041 compressed air 492 computed statistics

combustion

1191 computerized environ- 121 ment condenser 189, 317, 467 conducting matrices 788 conduction heating 366 conodont 483, 495 conservation of momen- 749 tum constant combustion con-544 ditions 836 constant sensitivity 55, 215, 216 constant temperature (thermostated) water bath 944 consumer protection 24, 38, 58, 235, 236, 257, contaminants 285, 286, 306, 343, 373, 410, 412, 447, 448, 450, 465, 469, 496, 529, 531, 543, 546, 657, 743, 794, 817, 818, 863, 966, 982, 1040 3, 4, 8-10, 25, 28, 29, 36, contamination 47, 113, 185, 186, 190199, 211, 242, 250, 256, 257, 268, 278, 285-287, 303, 305, 312, 315, 320, 327, 331, 365, 412, 443, 450, 456-503, 524, 558, 565, 612, 625, 626, 649, 660, 707, 737, 743, 746, 779, 780, 794, 832, 847, 911, 936, 949, 950, 965, 969, 982, 983, 988, 1015, 1040, 1050 continental crust 307, 1034 continental rocks 139, 140 continental scale 183 continental tourmalines 1034 continental USA 292 continents 556 continuous flow MS (CF) 1, 7-9, 17, 22, 24, 29, 36, 168, 287, 335, 341, 343, 474, 497, 543, 544, 546, 564, 578, 595, 817, 944, 882, 972 399 CF analysis 637 CF apparatus 990 CF applications 505 CF approaches 972, 974 CF autosampler technique 973 CF combustion CF hydride generation 637

1192 CF loop CF measurement CF method CF mode CF preparation procedure CF purification method CF pyrolysis technique CF spectrophotometry CF sulfur isotope technology CF system

Subject Index 42 882, 971, 972 637, 971, 973 8, 17, 494, 501, 878 403 637, 638 476 333, 334 972, 974

676, 168, 190, 498, 593, 64Z883,891,972,973 77, 338, 390, 391, 394, CF technique / ology 459,544,626,879 CF ultracentrifugation 41 66 CF water reduction 333,468 CF with He 182,460 CF-GC-IRMS continuous fluorination 533 technique continuous isotopic spec-510 trum continuous monitoring 107, 192, 872 continuous sputtering 674,799 controlled climate boxes 1014 controlled environment 1014 conventional fluorination50, 410, 412, 413, 433, 443, 446, 447, 453, 458, 459, 461, 466, 681 conventional methodo- 471 logy 5, 8, 24, 30, 33, 120, 161, conversion (method) 165, 221, 223, 297, 313, 321, 329, 334, 335, 337, 340, 392, 410, 411, 417419, 423, 427, 430, 433, 436, 442, 446, 455, 474, 498-501, 508, 524, 541, 547, 558, 561, 556, 562, 563, 569, 570, 573, 574, 586, 590, 591, 595, 614, 775, 838, 847, 936, 1012 335, 339, 430, 501, 506, conversion (complete) 525, 586 conversion (incomplete) 9, 571, 847 conversion dynode/ 668, 679 channel plates conversion efficiency 303, 419 conversion equation 855, 894 (formula) conversion factor 940 conversion interface 1041 conversion relationship 136

conversion system 5 conversion temperature 13 cool plasma 713 Cooperative Air Samp- 292 ling Network 49, 69, 257, 258 copper oxide furnace 344 corn leaves 861 correction plate 706, 715, 717 correction technique cosmochemical fractio- 718 nation process cosmochemical implica- 122 tions cosmochemical signifi- 268 cance cosmochemistry 268, 709, 725, 912 cosmogenic mass-depen- 632 dent fractionation 391 cosmogenic nitrogen 448 cosmogenic studies 475 cotton plants 597-600 coulometric method 600 coulometric measurement 599-601 coulometry 677 counting artifact 373, 374, 658, 659, 667, counting statistics 671, 731, 734 7 Cr reactor 7, 23, 34, 36, 571, 572, Cr reduction 575, 576, 579, 588, 626, 650 181 crabs 33, 311 Craig correction 578 Cretaceous 879, 892 Cretaceous belemnite 482 critical nutrient 181 crop plants 50, 320, 322, 407, 408, cross contamination 439, 461, 539, 647, 784, 941, 953 503 cross-calibration 173, 578, 579, 584 crude oil 349 crust-mantle cycling 348 crust-mantle transfer 293, 294, 296, 297, 325, cryo-focusing 430, 433, 462, 465, 469, 470 427, 447, 452, 972 cryogenic distillation 444 crystal face 485 crystal growth 458, 484, 485 crystal lattice 349 crystal rocks environment

Subject Index crystal structure crystalline KBr crystalline ozone crystalline pyrite crystalline pyrrhotite crystalline residue crystalline rock crystalline solids crystallinity crystallization crystallographic orientation crystallographic site crystallographic structure Cu/CuO reagent Cu20 cubic pyrite CuO/copper oxide

CuO/NiO system CuO/NiO/Pt'catalyst' CuO / Pt' catalyst' CuO-Cu20 mixture CuO oxidation CuO oxygen reservoir CuO wires cw laser cyclic compound cyclic hydrocarbons cyclic structure cyclic sulfide cyclone pre-separator cyclosilicates cysteine cystine

1193 68, 70, 259, 805, 806 418 379, 383 582, 600 571, 574 165 140, 268, 349 259, 805 47, 51, 445 349, 1003 1004 810, 817, 1000, 1004 442 352, 358 75, 544, 545 549 58, 59, 68, 70, 175, 190, 258, 321, 339, 350, 352, 362, 363, 366, 401, 544, 616, 818, 973 162 162 162 59, 257, 407 494 59 170, 258, 350, 615 777 553 636 553 554 560 358 554, 555, 585, 595 554, 555, 571, 585, 595

D Daly collector/ion coun- 130, 144, 627, 695 ter data reduction 33, 36, 150, 169, 262, 311, 343, 548, 628-634, 644, 854, 944, 945, 947, 956, 958 data normalization 944, 956 dead biological material 188 Dead Sea 27, 31-33 dead time 95, 364, 667-669, 677, 679, 681, 728-743, 832, 923, 926

dead volume decripitation experiments deep earth fluids deep-sea spherule deer bones degassing degradative enzymes dehydration dehydroxylation dendrochronology dendroclimatology dendroecology denitrification denitrifier method denitrifying bacteria densimetry derivatization derivatization agents derivatization method derivatization procedure derivatization reactions derivatization scheme derivatization step derivatization technique derivatized derivatized compounds desiccator desolvating nebulizer desorption effects detectable contamination detection limit detrital deuterium labeled glucose Devarda's alloy Devonian diagenesis diagenetic diagentic opal

3, 9, 208, 275, 285, 290, 313, 455, 843 66, 358 142 632 1032 12-14, 23, 28, 32, 47, 49, 53, 57, 60, 204, 207, 349, 357,491,565,566,870 186 56,57,239,420,491,806, 1031 46, 49, 51-53, 422, 423, 445 508 508 507 308,335,336,344 306, 333, 335, 336, 343, 344 335 40,42,48,61 155, 160-166, 193, 19519G 233 15~ 160 162,164,166,165 197 160 194 198, 201 198 196, 199 160, 162, 163, 195, 199 215, 220, 328, 345, 485 698, 708-710, 713, 714, 716, 719-722 926 287 11, 65, 155, 755, 780 39, 187 101 314-324, 327, 328, 330, 332,346,347 354,579 139,140,199,201,271 38, 39, 42, 56, 484, 488, 602,689 56

1194 diamond

Subject Index 30, 257, 259, 265-267, 270, 271, 349, 350, 433, 467, 684 27O

diamondiferous diatreme diatary protein 200 dietary sources 181,198 diatom 534,535,1014,1015 diatom opal 56 diatom silica 52 diatomaceous silica 1015 diatomic species 608 dichloromethane 197-199,201,572,576 diethyl ether 158,618 differences in reactivity 444 different matrices 903 differential sputtering 659,793 diffusion 3, 16, 51, 189, 190, 264, 306, 314, 320328, 330, 345, 346, 375, 605, 809, 812, 813, 821, 828, 995, 998, 1001, 1004-1006, 1008, 1018, 1020, 1023, 1025,1026,1034 diffusion coefficient 604, 1004, 1020, 1023, 1034 diffusion constants 1004 diffusion data 1005,1024 diffusion disks 322 diffusion experiments 1004 diffusion method 189,190,330,344 diffusion packets 346 diffusion period 323,346 diffusion process 324 diffusion rates 810,812,1020,1031 diffusion technique 318,324,330,332,342 diffusional loss 72 digestion catalyst 319 digestion decomposition 124 digestion problem 832 digestion procedure 831 digestion rate 319 digestion time 319 digital conversion 115 diode laser 386, 388, 761, 772-775, 77Z 780,782,783 777 diode laser light injection direct combustion 230, 232, 322 direct diffusion 322 direct injection 11, 25, 650 direct sputtering 794 direct-sampling ion 789 source

dis-(non-)equilibrium fractionation dispersed graphite dispersion of ions dispersion range (IRMS) disposible syringe dissociation / ed

1002,1028

403 842 848, 849 213 26,45,362,369,378,384, 385, 389, 397, 408, 499, 553,860,1003 dissociation constants 158,239 dissociation process 370,384,463 dissociation reaction 26 dissolved inorganic car- 203-228 bon (DIC) dissolved organic car565 bon (DOC) dissolved organic mat- 23, 31, 314, 331 ter (DOM) dissolved salt 4, 5, 22, 23, 32, 33, 342 distillation 113, 116, 156, 189, 190, 313-327, 330, 332, 340, 345, 349, 356, 412, 414, 451, 455, 499, 531, 619, 620 distillation apparatus 316, 345, 567, 570-575, 586, 588, 591, 619 distillation equipment 316, 317, 620 distillation flask 189, 316, 317, 345 distillation method 54, 315, 320, 323, 328 distillation procedure 316 distillation processes 53, 54, 430 distillation products 116 distillation technique 54, 55, 60, 583 distillation time 315, 317, 320, 321 distillation unit 316,319 distillation yield 55, 116 divinylbenzene 125, 327 Dole effect 481 double filament 130, 138, 550, 643 double focussing 132, 549, 550, 677, 726, technique 738, 740-742, 788, 798, 804, 844 double focussing ion op- 697 tical system double spike calibration 626, 633, 650 double(isotope)spike 450, 623-651, 701, 709, 711, 713, 716, 719, 720, 724, 1012 dried atmosphere 49 Drierite 322, 345 drift correction 171, 300, 301 drift line 302 drinking water 305, 729, 736 dry (atmospheric) air 278-281,283, 287

1195

Subject Index 49,408,423 20, 25, 55, 75, 185, 208, 221, 448, 427, 451, 528, 529, 531, 533, 609, 613, 616,811,814,816 dry leaf samples 19Z 476,477,480 dry weight 200 drying agent 157,158,410 drying of air 280 drying (biol.) tissue 186,188 627 dual collector TIMS dual inlet (unspecified) 1, 6, 272, 282, 290, 300, 348,350 dual inlet (MS/IRMS) 1,4,21,22,157,168,259, 272-274, 276, 282, 299, 334, 441, 460, 539, 543, 544, 551, 564, 572, 578, 590, 875, 882, 883, 957, 972 dual inlet analysis 282, 288, 391, 971 dual inlet configuration 879 dual inlet instrument 299, 972 dual inlet measurement 273, 288, 882 dual inlet method 282,971 dual inlet mode 283,595,878,953 dual inlet system 272, 390, 391, 396, 498, 543,836-838,853,854 dual inlet technique 283,290 dual inlet valves 543 dual path gas cell 772 Dumas combustion 318,319,321,339 Dumas method 321 Dumas process 330,340 duoplasmatron 656,657,659,680 dye laser 758, 764,776 dynamic conversion 5 dynamic extraction 208 dynamic headspace 158 dynamic mass spectro- 1,262,270,348,349,362, metry 468 dynamic methods 5 dynamic measurements 697 dynamic mode 16,350 dynamic pumping 495 dynamic range 95, 363, 669, 670, 782, 795 dynamic reduction 5,16,35 dynamic systems 8,36 dynamic vacuum 528 dynamic zoom lens 69Z 727 system dry box dry ice

E Earth (the) earth materials earth science earth scientists earth's atmosphere earth's crust earth's early atmosphere earth's early history earth's surface

1,122,139,141,142,229, 235,875 626 122, 142, 152, 349, 471, 476,496,874 40O 375-377 307,625 448 540 38, 143, 236, 37Z 378, 384,625,648 85 188 17L 183 17L 183,189,191,306

East Pacific rise ecological laboratories ecological processes ecological samples / material ecological studies / 177-179, 183, 191, 198 research / investigation ecological systems 175, 178 ecologists 177, 179, 202 ecology / ecological 172, 202, 181, 596, 899 science economic geology 684 ecosystem 177-184,201,202,307 ecosystem analysis 180 ecosystem components 183 ecosystem cycles 329 ecosystem dynamics 183 ecosystem location 184 ecosystem processes 202 ecosystem seasonal va- 184 riation ecosystem studies 177, 182, 184 ecosystem survey 184 ecosystem temporal va- 183 riation ectotherms 483 effective filtering area 560 egg shell 1031 eigenstate 749 Elba tourmaline 903 electron emission current 841, 859, 864-866 electron gun 661-663, 675, 678-680 electron shell 654 electron trap 394, 398, 859, 860, 866 electronic environment 106 electrostatic interac729 tion

1196 elemental analyzer (EA) 29, 190, 202, 340, 341, 40Z 429, 542, 546, 54Z 569, 593, 595, 598, 602, 817,891,971,973 342 EA (automated) 341 EA (automated C/N) EA (high temperature) 8,340 305,319,341,971 EA-IRMS EA-CF-IRMS 358 476 elemental standard 24 electrolysis method 21,23,30 electrolysis of water 611,863 electrolyte electrolytical recoverage 401 electron bombardment 123 electron collector 840 electron impact frag396 mentation electron impact ion 390, 839, 840, 859, 907 source electron injection scheme654 electron multipliercol- 95, 132, 364, 367-369, lector/ion counter 378, 549, 550, 667-670, 677, 681, 684, 695-697, 832, 926 electronegative element 652, 653, 679 654 electronegative ions electronegative species 655 675 electronics drift electropositive element 653,654 electropositive species 655 7,695,727,728 electrostatic filter 538,597 element cycles 125, 127, 195, 284, 287, elution peak 299,854,855 859,865 emission controller 982,983,1027 empirical calibration 703, 704, 716, 717, 719, empirical correction 973 980-982 empirical derived law 699 empirical derived con- 504 stants 1036 empirical experiments 629 empirical fit empirical fractionation 484, 1035, 1036 empirical fractionation 671 law empirical increment me- 1031 thod empirical isotope correc- 979 tion empirical isotope ratio 150 empirical isotope ther- 1031 mometer

Subject Index empirical method empirical normalization empirical procedure energy dispersion energy filter

992, 993 706, 723 702 661, 694, 697, 742, 844 17, 662, 666, 676, 677, 679, 682, 742-744, 854 energy filtering 662, 666, 676, 677, 679, 682 energy window 662, 666, 677 enstatite chondrite 536 environment 38, 52, 143, 177, 186, 198, 383, 391, 541, 597, 605, 650, 746, 1008, 1027, 1032, 1038, 1039 environmental chemistry 907 environmental condi196, 625, 1014, 1035 tions environmental extremes 412 environmental factors 418 environmental geoche- 685 mistry environmental health 558 viewpoint environmental level 879, 880, 1048 environmental molecule 759, 779, 782 environmental monito- 928, 944 ring environmental N20 390 environmental para191, 202, 1028 meters environmental policy 920 issues environmental questions 175 environmental recon507 struction environmental research 343, 497, 782 environmental samples/ 149, 738, 739, 757 sampling environmental science 154,239 environmental seawater 1028 environmental study 178, 507, 778 environmental tracing 757 environmental variables 1035 environmental variation 508 environmental water 189 enzyme 26, 185, 595, 625 enzyme-mediated reac- 196, 199 tion equal size peaks 664 equilibration conditions 408 equilibration effect 504 equilibration method 5, 31, 55, 60 equilibration rate 25 equilibration system 17

Subject Index equilibration time equilibration vessels equilibrium bracketing equilibrium conditions equilibrium constant (K) equilibrium distribution equilibrium fractionation factor (c,) equilibrium isotope / ic fractionation equilibrium isotope effects equilibrium (chemical) processes equilibrium thermodynamics equilibrium time error propagation Eschka method Eschka mixture essential nutrient essential oils estuarine water ethane ethanol

1197 27 18 996 675, 867 91, 543, 994, 995 92 28, 34, 850, 993-1009, 1023,1025,1027 34, 375, 993-1000, 1004, 1006, 1009, 1018, 1021, 1024-1030,1035 2,375 375, 994 224 21,27 633,634,673,833,983 563, 565, 574-57G 587, 593,594 568,569,578,586 624 172 316-318 4,231 44, 55, 75, 101, 108, 109, 112-116, 127, 185, 189, 195, 275, 277, 320, 345, 412, 451, 457, 458, 516, 517, 529, 571, 582, 614, 816,891,896,897,908 16 175, 231 163-166

ether ethylbenzene ethyl-chloroformates (ECF) Etna basalt 903 Europa mission 758 eutrophication 305 evacuating air 20, 25, 206, 216, 217, 221 evaporation filament 131 evaporitic environment 1034 exchange equilibrium 1016 excimer laser 458,459,466,469 exhalation-inhalation 98 cycle exothermic dissolution 514 exothermic reaction 44,413 expansion bellows 7,949,950 experimental artifacts 27Z 385 experimental background237 experimental blank 337 experimental calibration 1001 experimental capsule 1001

experimental comparison932, 934, 935 experimental conditions 47, 121, 398, 491, 727, 738, 769, 1005, 1006 experimental data 16, 263, 264, 266, 267, 270,992,993,1004,1025, 1026,1036 experimental design 184,658,1010,1025 experimental details 780 experimental determi- 671,744,1010,1035 nations experimental evidence 464, 1009 experimental fractiona- 484 tion equations experimental geoche1006 mistry experimental isotope 992-1037 fractionation measurements experimental method 309, 337, 392, 992, 1005 experimental observa- 385, 386 tions experimental parameters 395 experimental petrology 1006, 1010 experimental procedure 20, 112, 923, 950 experimental results 266, 267, 727, 729, 741 experimental section 245 experimental set-up 668, 774,778, 994, 995, 1012, 1018, 1020, 1023, 1024 experimental settings 729 experimental study 32, 47, 52, 261, 309, 350, 756, 796, 992, 1001, 1003, 1009-1011, 1023, 1028, 1036 experimental system 193, 1017, 1020 experimental technique 95, 391 experimental time 114, 115 experimental work 376 experiments 13, 14, 52, 55, 74, 107, 208, 211, 218, 233, 237, 238, 240, 263, 264, 271, 322, 324, 353, 356, 358, 384, 385, 393, 398, 491, 511, 526, 568, 582, 604, 624, 625, 650, 708, 727, 731, 744, 816, 985, 9921037 explosive reagent 30 exponential equation 733, 743 exponential law 629, 631, 671, 699-704, 706, 716, 733, 788, 796, 799-802 exponential model 629, 730, 733, 735, 738, 743

1198 external external external external

Subject Index calibration contamination environment normalization

731, 735 320 1011 701, 702, 704, 706, 708, 711, 715-725 external standard 134, 138, 279, 598, 671, 701-703, 708-717, 723 extracting plate 860, 861 extraction efficiency ex- 217 periment extraction line 52, 66, 67, 70, 71, 76, 208, 357, 358, 452, 453, 491, 526, 1041 extraction plate voltage 395 extrapolation 995, 998, 1002 extraterrestrial bodies 747 extraterrestrial environ- 229, 230 ment extraterrestrial feature 235 extraterrestrial material 230, 361, 362, 374, 406, 535, 653 extraterrestrial O-iso680 tope analysis extraterrestrial organic 230, 235 matter extraterrestrial property 235 extraterrestrial samples 257, 267, 361, 371, 374 extraterrestrial source 230 region extraterrestrial water 25 extremely poisonous Cd 566 acetate

Faraday collector / cup

fat-free tissue fatty acid fatty acid standard Fe chemistry Fe isotope standard Fe oxidation feathers Fe-Cr-Ni alloy

71, 129, 132, 133, 276, 284, 364, 367, 369, 393, 550, 614, 626, 644, 645, 667, 669, 670, 679, 681, 685, 68Z 690, 694, 695, 69Z 698, 726, 728, 729, 739, 743, 832, 841, 847849, 853, 854, 857, 858, 870 191 120, 162-166, 170, 172, 175, 193, 195, 196, 198, 1014,1051 196,199 625 627,648 650,1013 181,186,594 795

feldspar feldspar crystal ferromanganese crusts fertilizer field environment field experiments filament filament filament filament

(metal) (Ni) (Pt) (Re)

filament (Ta) filament (thoriated Ir) filament (W) filament amendment proportions filament assembly filament centre filament characteristics filament conditioning filament contamination filament corrosion filament current filament electron density filament graphite technique filament length filament loading filament material filament protection filament resistance stabilizer filament safity filament stabilization filament status filament surface filament temperature filament voltage filament work function filter capsule filter discs filter glass tube filter material filter packs

56, 264, 265, 349, 356, 358, 403, 406, 429, 442, 444,445,461,530 256 714, 721 181, 305, 312, 328, 329, 344 1027 540,992,993,1005,1011, 1027,1028,1035,1036 148 621 494,868,870 130, 131, 549, 641, 643646,839,868 129,614,868 839 363,839,868 627 131 867 869 843 647 22 124, 129-131, 363, 614, 645, 646, 864 643 643 867 130, 149, 550, 561, 614, 62Z 639, 640, 643, 645, 650 147,646 864 868 866 868,869,871 870 645,843 130, 131, 368, 62Z 628, 647,857,859,867-870 865,869,871 494 331 346 521 186 346

Subject Index filter paper

215, 322, 476, 529, 558562,565-567,578 filter sheets 560 filter strips 327,328 filter system 214,560 fine-grained matrices 465 fish 178, 181, 200, 305, 483, 625,1035 fish bones 482,1035 fish flesh 490 fish otoliths 594 fish skin 594 fish tooth 484 flame sealing 215 flash combustion 601 flask air 293 flat-topped peaks 694,697,706,799 Floridian aquifer 567 flow-through reactor 421 fluid inclusion 2, 4, 8, 36, 62-8Z 269, 349,358,359,452,1030 fluor-apatite 425 fluoroscence techniques 747, 756 fluorescent instruments 202 fluoric acid 365 fluoride 124, 162, 442, 444, 454, 456, 457, 464, 467, 468, 614,617 fluoride(-rich) atmo461, 972 sphere fluoride compound 451, 453, 455, 471 fluoride reagent 420 fluorinated gases 946 fluorinated plastic 412 fluorinated polymer 168 fluorinated substance 107 fluorinating compounds 546 reactivity fluorination / ed 49, 58, 66, 400-472, 482, 484, 488, 490-492, 524534,626,816,1031 fluorination / ing agent 49,58,409,418,456,546 fluorination chemistry 471 fluorination cycles 413 fluorination line 412, 416, 418, 443, 453, 455,462 fluorination reaction 49, 405, 413, 418-421, 443, 453, 455, 460, 464, 468,471,491 fluorination reagent 403-409, 414-416, 419, 420, 423, 450, 457, 459, 461, 466, 468, 471, 524, 526 fluorination stripping 443,453

1199 fluorine

58, 107, 162, 400, 408, 411-413, 414, 418, 420, 421, 443, 445, 446, 452454, 461, 524, 532, 534, 545, 957, 972 400, 409, 421, 448

fluorine(-bearing) compounds fluorine frequency 114 fluorine gas 404, 409, 626 fluorine generator 413, 532 fluorine molecules 199 fluorine nuclei 113 fluorine oxidation 49, 401 fluorine pump 413, 418, 452 fluorocarbon contami448 nant fluorosilicic acid 445, 525 fly dilution 820-822 focus position drift 147 foil boat (small/tiny) 242, 337, 341 Fomblin 46 food quality control 172 food science 1, 103, 121 foraminifera 130, 1011, 1015, 1029 foraminifera shell 132, 141, 152 foraminiferal tests 144 formaldehyde 195 formic acid 588, 591, 638 forrest ecosystem 183 forrest soil 1013 forsterite crystal 1023 fossil fuel combustion 557, 559, 562 fossil hydrocarbons 235 fossil sediments 601 Fourier transform IR 779 spectroscopy Fourier transformation 47, 105, 114, 115, 380, (FTIR) 382, 763-765, 779, 782 fractional approach to 1000 equilibrium fractionated distillation 157 fractionation standard 325 fractionation factor 19, 20, 28, 34, 52, 55, 133, 141, 147, 243, 246, 247, 359, 360, 396, 408, 673, 796, 800, 802, 850, 975, 976, 992-1037 fragment ions 154, 390-397, 957, 958, 962, 986, 987 fragment pattern 396 fragment ratio 394, 395 framboidal pyrite 549 freeze drying 187, 188, 194, 314, 330, 339, 576, 1014

1200 frequency modulation spectroscopy (FMS) frequency window fresh leaf fresh water

Subject Index 773

114 474-481 1Z 33,314,324,331,335, 345,346,911,1013,1028, 1030,1035 1012 freshwater bacteria 1015,1030 freshwater diatoms freshwater ecosystem 306 freshwater environment 1012 freshwater fish 181 fruit juice 116, 481, 896, 913 fulvic acid 563, 588 fundamental standard 929 fusel oil 118

G 66, 441, 447, 548, 549, 552,568,570,1025,1030 359,429,441-443,466 garnet 171 gas calibration 8, 74, 76, 77, 88, 153, gas chromatograph(y) 154, 160, 161, 168, 169, (GC) 173, 182, 190-199, 223225, 233, 272, 282-284, 292, 321, 333, 350, 395, 430, 447, 448, 469, 471, 474, 476, 478, 494, 495, 508, 544, 546, 557, 609, 612, 613, 620, 621, 818, 853,854,880,891,971 GC technology 153-176 GC-C-IRMS 153-155, 157, 160-166, 168, 170-173, 175, 182, 192-194, 196, 197, 199, 201,350 gas equilibration 1Z 25,32,250,941,948 gas evolution technique 205, 206, 208, 210, 220, 226 29 gas permeable membrane 417, 418, 455, 81Z 1039, gas storage 1041 62-87 gaseous inclusion 55,224,225,615 gas-tight syringe 510 Gaussian distribution 168 Gaussian function 284,301,467 Gaussian peak Gaussian shaped signals 168 Gaussian source of noise 681 gelatinous mass 525 gelatinous silica 525 galena

genesis geochemical cycle geochemical tracer geochemistry

geographical origin geographical latitude geographical site geological analysis geological application geological community geological context geological environment geological fluid geological implications geological investigation geological material geological origin geological phenomenon geological processes geological recycling geological sample

geological settings geological studies geological time geology geoscience geostandard shale geothermal mudpot geothermal system geothermal well geothermal water geothermometer/metry

38,56,140,141,445 1036 152 103, 141, 154, 400, 496, 523, 596, 604, 605, 619, 622, 680, 685, 788, 789, 803, 805, 907, 992, 993, 1009,1036,103Z 1043 103,108,115,116,121 108 172 654 123,138,139,496 789 199 607 2 143 65 122, 123, 127, 132, 138, 139, 569, 570, 573, 614, 652, 664, 692, 717, 794 892 690 122,141,653 224 4, 8, 122, 128, 130, 132134, 139, 141, 256, 568, 569, 572-574, 609, 619, 707,709,713-725 78 124,141 62,820,1043 1 652, 788 597 1029 1028,1030 1029,1030 2 51, 52, 406, 444, 445, 993, 1000, 1031

geothermometric calibration glass ampoule glass boat glass fiber filter glass vial glass standard glass wool glass-ceramic standard

1000 3,7,35,894 215,216 186, 187, 189, 190, 214, 322, 323, 327, 341, 558, 560-562,610,617 1Z 26,161,198,214,215, 32G 1019 688, 788,798,802-804 293,294,564 80O

Subject Index glassy carbon Global Meteoric Water Line (GMWL) globalisation glove box glow discharge ion source glow discharge mass spectrometry (GDMS) glucose glutamic acid glycine glycolipids gold wire Gorleben aquifer grain fragments grain size

grain size distribution grain size fraction grain size range grain size segregation grain size separation granite graphite

graphite capsule graphite container graphite crucible graphite powder graphite reagent graphite reduction graphite resistance furnace graphite rod graphite slurry graphite standard grass-like plants gravimetric calibration

1201 76, 77, 334, 340, 494, 501, 503 15, 65 928 214, 215, 218, 223, 558, 568 789, 790, 791, 798, 803, 8O4 143, 788-804 101,200,201,896,897 232,235,884,89Z 1019 199,200,231,232 2O0 559 565 464 11-13, 23, 39, 54, 55, 68, 72, 240, 245, 253, 262, 265-267, 373, 418, 421, 434, 442, 463, 812, 996, 998, 1003, 1005, 1016, 1017,1020 39,41,42 68 41 41 39,41,42 39, 81-83, 85, 86, 264, 265, 395, 402, 431, 530, 568,569 24, 25, 30, 49, 271, 334, 340, 401-403, 430, 482, 483, 488, 489, 492, 493, 495, 529, 614, 636, 643, 64G 789, 808, 876, 883, 884, 895-89G 910, 967969,102G 1030 402 402 529 402,492,493 492,493 401,482,489 401 417, 423, 457 147, 614 494 517 631, 633, 635

gravimetric determination gravimetric (isotope) mixtures gravimetric preparation gravimetric techniques greenhouse effect greenhouse gas greenhouse gas lifetime greenhouse warming Greenland Ice Sheet Precipitation (GISP) Grimm-type discharge cell Grimm-type ion source groundwater

566, 567, 571, 574, 827 539, 540, 889, 920, 921, 924, 932, 941 627 570 291 272, 280, 377, 390 378 291 476, 887 798

789, 791 2, 36, 134, 142, 180, 181, 203, 204, 306, 312, 541, 567, 624-626, 649, 650, 757, 903, 904, 1034 groundwater chemistry 55 guanidine (hydro24, 30, 31, 33, 66, 76, 333, chloride) 334, 340, 408 gypsum 43, 552, 568, 572, 573 H H3+ contribution H3+ correction H3+ effect H3+ factor H3 + formation H3+ ion H3+ production H3+ signal H3BO3 reagent H202 oxidation H2S evolution peaks H2-water equilibration Haarhof / Van-derLinde function habitat hair halite crystal Hall probe halloysite halogen fluoride

21, 22, 33, 850, 853 842, 851, 1016 33, 34 848, 850 847, 848 841, 843, 850 22 21 345 565 581-583 3, 4, 7, 17-35, 55, 889 168

108, 177 182, 594 65 364, 861, 862 47-49, 53, 60 410-412, 418-420, 446, 451, 453, 455, 461, 471 halogen waste products 448 halogen 341, 410, 412, 604, 619, 621, 622, 655, 939 hard coal 46 hardwood 514 Harriott-type gas cell 772

1202 Hay(e)sep H-bearing contaminants headspace air headspace equilibration health and waste issues heavily contaminated heavy liquid heavy oil heavy water Heliobacter pylori hemicellulose hemoglobin herbaceous plants herbivore teeth herbivores heterogeneous pathway heterogeneous reaction heterogeneous sulfide heterogeneous system heterogeneous target heteronuclear 02 molecules hexane hexanediethyl ether hexapole collision cell Hg catalyst HgC12 poisoning HgC12 pyrolysis high detection sensitivity high dispersion IRMS high(er) mass resolution

Subject Index 157, 294 38 1013 226 319 313 42, 570 173, 581, 583 1,17 779 200, 201, 512-517 625 517 437 180 557 995, 996, 998, 1001 660 996,998, 1000 652 387 54, 191, 198, 502, 1051 198 696, 723 499 2 497, 499, 505 782 175 364, 551, 666, 676, 677, 693, 697, 790 334, 349

high temperature combustion high temperature oxi405, 529, 600 dation high-sensitivity analy- 795 tical technique high-sensitivity MS 259, 362 high-voltage discharge 499 reactor HNO3-Br2 digestion 574-576 HNO3-Br2 oxidation 594 Hokko beads 19, 22 holocellulose 508-520 homeotherms 483 homogeneous compound881 homogeneous crystal 690 homogeneous elemental 877 reservoir homogeneous gas 5 homogeneous magnetic 840, 844 field

homogeneous natural material homogeneous pathway homogeneous reaction homogeneous reference material homogeneous sample homogeneous seawater homogeneous solution homogeneous standard homogeneous surface homogeneously heating homogenized resin homonuclear diatomics horizontal momentum host lattice hostile environment HPLC HPLC peak HR-ICP-MS human activity human arbitrariness human error human hair human hands human health human inhalation human milk human nose human origin human population human samples human subjects human waste humans humic acid humid air humid locations humidity humidity conditions humidity control humidity levels humidity of air hydration hydration limited pro-

691 557 995, 999-1001 881, 882, 896, 900, 909 508, 512, 519, 586, 725 136 481 259, 511, 690 12 364 328 387 664 813 756 6, 7, 153, 1019 158 692, 693, 697, 710, 713, 715, 718, 720 390 914, 926 17 594 194 378 561 2, 10 19 2 291 11 173, 816 305 175, 177, 594 330, 563, 565 89, 280, 502 280, 281 185, 186, 280, 456, 473, 675, 1014, 1024, 1032, 1033, 1035 9, 1032 1039 172 3, 502, 558, 1014, 1032 26, 27, 31, 491, 806 26

cess

hydration reaction hydride contribution hydride formation hydride generation hydride interferences

26 681 743 636-639,647, 651, 744 536

1203

Subject Index hydride process hydride technique hydrides hydriodic acid hydrobromic acid hydrocarbon

hydrocarbon contaminants hydrocarbon oil hydrocarbon skeletons hydrochloric acid hydrofluorocarbons (HFCs) hydrogen atmosphere hydrogen background hydrogen diffusion hydrogen gas hydrogen peroxide

637,638 637,638 657,666,707 569,570,588,591 450 4, 44, 46, 54, 58, 59, 158, 159, 185, 193, 195-198, 230, 231, 361, 366-369, 401, 657, 794, 805, 814, 894,951,952 245 46 233, 236 195, 317, 332, 1018 377 559,1023 856 1006,1008 4,9,68,71,819,843,847, 850,1016,1051 44, 45, 48, 484, 489, 490, 559,579,618 236,582 307 204

hydrogenation hydrological data hydrological evolution of water 2,3,25 hydrological fluid hydrological processes 142 142 hydrological reservoir 35, 36, 313 hydrological sample 203 hydrological systems hydrological tracer study 2 hydrologists 306 1, 103, 189, 907 hydrology 35 hydrology laboratory 17, 19, 21 hydrophobic (Pt) catalyst 323 hydrophobic tape 538, 552, 562, 569, 683 hydrosphere 141 hydrothermal activity 38, 83, 443, 444, 1026 hydrothermal alteration hydrothermal apatite 407 hydrothermal aragonite 46 hydrothermal deposits 349, 1004 hydrothermal exchange 1011 hydrothermal fields 723 130, 139, 142 hydrothermal fluid 719 hydrothermal molybdenite hydrothermal quartz 68 hydrothermal reaction 140

hydrothermal rocking bombs hydrothermal system hydrothermal water /fluid hydrous iron oxides hydrous mineral hydrous phases hydrous pyrolysis hydrous silicates hydrous whole-rock samples hydroxybenzalacetone hydroxyl (OH) hydroxyl group hydroxyl hydrogen hydroxyl ions hydroxyl oxygen hydroxyl radical hydroxyl site hydroxyl water hydroxylamine hygroscopic hygroscopic fluoride hygroscopic salts hypersaline brine hypersaline water hypobromite oxidation hypochlorite hypothetical contamination

1007, 1025, 1026 56, 1028 32, 130, 139, 142, 149 56,625 2, 4, 8, 11, 53, 69, 687, 1020,1021 69 233 423,444 453 117 164, 377, 625 47, 116, 160, 265 46, 508 265 5O 291, 377, 484 112, 495 49 428, 429, 450 413, 425, 442, 444, 457, 491, 502, 506, 895, 923 451 631 32 11 317, 319, 321, 340 44, 45, 573 735

IAEA (International 34, 35, 111, 280, 300, 305 Atomic andEnergy 307, 310, 311, 504-506, Agency) 539, 599, 739-741, 772, 784, 853, 874-906, 910912, 917, 930, 931, 934, 943-946, 948, 951, 953, 981, 984, 1048 IAEA Advisory Group 892 Meeting IAEA Consultants' 886 Meeting IAEA experts meeting 883, 902 IAEA intercomparison 946 885 IAEA interlaboratory comparison IAEA laboratory 13 IAEA panel group 887 IAEA Panel Meeting 885

1204 IAEA reports IAEA technical contract ice (frozen water) ice bath ice core ice crystals ice slush ice trap Icelandic basalt ice-water bath ICP ion source ICP technique ICP-MS

ICP-MS experiment ICP-MS standard ICP-TOF-MS ideal standard igneous rock illite impermeable glass impregnated filters impurity interference impurity ions impurity isotopes

impurity of'.

BrF3 lignin (in cellulose) oxygen SeF4 sodium transition metals 80Se (in double spike) incomplete combustion incomplete diffusion incremental heating incubated soil incubation experiments incubation time incubation temperature independent isotope fractionation Indiana zinc industrial standard inert gas infrared (IR) IR (near) IR (mid)

Subject Index 887 903 3,186,187,313,75Z 875 621 29,290,336,772 275,277 208,215,216 207 648 615 132, 651, 692, 726, 803, 907 47 123, 127, 132, 138, 551, 621, 692-745, 788, 789, 794-800, 804, 820, 828, 923,926,937 729 648 692,693 259 39,64,404,648,716,721 39-51, 60, 80, 429, 445, 1030,1031 250 558 138 102 633,634 409 511 461 429 128 549 630 161, 171 325 349, 356, 870 649 1013 325, 346 1012 389 11-13 168 74, 157, 188, 410, 414, 455, 503, 545, 778 104, 753, 759-787, 915, 1036 396, 466, 783 766, 767, 780, 783

IR (far) IR absorption

382 156, 396, 399, 767, 880, 907 IR beam 756 IR detector 772, 778 IR emission 380 IR lamp 609 IRlaser 458-460, 463, 466, 467, 471, 772, 972 IR photon 766 IR radiation 377, 390, 458, 756, 779, 780, 783 IR spectrometer / try 380, 382, 759-787, 907 IR spectroscopy/ic 156, 391, 396, 399 IR spectrum 760, 769, 880 IR wavelength 458 inhomogeneous dust par-555 ticles inhomogeneous reference538, 900, 901, 911 material inhomogeneous sample 519 series in-house standard 598, 599, 934, 982 initial experiment 356, 384 injection current 773 injection flow rate 288 injection nebulizers 709, 722 injection valve 6, 7, 293 in-line combustion 333 in-line configuration 456 in-line gas purifier 286 in-line reaction traps 433, 461 in-line reactor 410 in-line sampling 1021 inner leaf section 481 inorganic acid 164 inorganic carbon 27, 189, 602, 892, 893 inorganic carbonates 1023, 1035 inorganic cation com812 plexes inorganic cation ex125 change inorganic chemistry 831 inorganic components 307, 579 inorganic compound 313, 315, 538, 541, 552, 553, 563, 568, 570, 574, 575, 578, 585-587, 590, 592, 595 inorganic conditions 496 inorganic dust particles 555 inorganic experiments 1018, 1023 inorganic forms 315, 585 inorganic fraction 321, 576, 592

1205

Subject Index inorganic isotopic reservoir inorganic material inorganic matter inorganic minerals inorganic nitrogen

186

189 8,177 579 179, 305, 312-316, 322, 323,325,330 inorganic phosphate 490,1013,1035 inorganic processes 482,624 inorganic residue 576 inorganic samples 605 inorganic sequestration 314 inorganic species 2Z 329,562,565 inorganic standard 547 inorganic substances 334,501 inorganic sulfates 547, 581, 585, 588, 590592,594 inorganic sulfides 547, 581 inorganic sulfur 552, 553, 562, 563, 565, 568, 570, 573-576, 579, 581, 585-588, 590, 592, 595 inorganic system 1007,1010,1024 insectivorous bats 198 insects 186,555 in-situ analysis 123, 405, 421, 444, 459, 462-466, 469, 471, 495, 535, 537, 548, 652, 661, 684, 688-692, 715, 725, 747,972 in-situ analytical capa- 652,691 bility in-situ chemical destruc- 378 tion in-situ chemistry 748 in-situ CO source 291 in-situ fluorination 464,465 in-situ laser sampling 459,482,483 in-situ measurement 50,433,652,723 in-situ method / techni- 1,380,483, 754 que / system in-situ micro sampling 467 in-situ photochemical 378 production in-situ Rb / Sr dating 757 in-situ reaction 465 in-situ reaction volume 405 in-situ sampling 56,466 in-situ stable isotope 725 studies in-situ 40K decay 269 instrument / ation 96-98, 107, 110, 115, 135, 154, 165, 175, 182, 188, 192, 202, 273, 274, 288,

instrumental instrumental instrumental instrumental instrumental

341, 376, 380, 383, 391, 395, 398, 450, 453, 535, 548-550, 594, 596, 614, 626, 627, 642, 644-646, 650, 653, 65G 663, 664, 674-679, 692-699, 706, 708, 710, 712, 714, 725740, 743-745, 747, 757, 758, 766-768, 772, 775784, 788-790, 794-798, 814, 824, 830, 835, 838, 843, 844, 847, 851, 856, 863, 870-873, 878, 880882, 907, 908, 912, 923, 926, 940, 946, 947, 950, 953, 954, 971, 972-975, 991,1038-1041 analysis 122,143,196 aspects 654,853 artifact 850,852 background726, 731, 736, 740, 743 (mass)bias 698, 701, 706, 708, 718, 725, 923 blank 709 calibration 779 characte- 728

instrumental instrumental instrumental ristics instrumental conditions 108, 123, 129, 397, 709, 710 instrumental configura- 93 tion instrumental corrections 33, 311 instrumental design 663, 803, 856 instrumental detection 65 limits instrumental (mass)dis-623-635, 644, 647, 650, crimination 651, 698, 701-703, 708, 716, 720, 721, 724 instrumental drift 124,284, 778,784,838 instrumental effects 675,833,848 instrumental factors 123,133,670 instrumental features 678 instrumental (isotope / 50, 141, 149, 152, 536, mass) fractionation 549, 652, 653, 655, 658, 659, 662, 666, 670, 672,677, 680, 687, 690, 699, 701, 706, 709-713 instrumental installa1038 tion instrumental knowledge 835 instrumental limitations 687 instrumental magnetic 108 field

1206 instrumental manufacturers instrumental memory instrumental methods instrumental offset instrumental parameters instrumental performance instrumental procedures instrumental resolution instrumental response instrumental sensitivity instrumental setting instrumental setup instrumental stability instrumental step instrumental sources instrumental techniques insulated ice chest insulating materials insulators (samples)

Subject Index 155, 165, 954 709 404, 450 427 671, 674, 776 697 123 771, 779 842 725, 729, 824 738, 832 674 778, 784 743 726 141, 143, 152, 460 3, 313 789, 794, 795 654, 675

interaction between:

adsorbent-nitrogen 158 analyte-stationay 171 phase environment-orga177 nisms fluid-mineral 687 fluid-rock 38, 349 gas-filament surface 843 ions-He or H2 651 laser-mineral 459 laser-sample 433, 436, 458, 461, 464 light-atoms 749 organisms 177, 191 plasma-organic matter 46 rotation-vibration 770 seawater-ocean crust 139, 1033 sediment-water 141 solute-water 31, 1018 solute-stationary 160, 168 phase surface water-ground- 204 water various carbon species 203 water-rock 142 interaction with high 863 speed ions intercalibration 132, 273, 290, 723, 904 intercomparison 280, 952-956 intercomparison material506, 539, 853, 881, 945, 949, 951, 956 intercomparison program280, 303 intercomparison results 505

intercomparison studies 290, 850, 898 intercrystalline fractio- 50, 51 nation interhalogen fluorides 400, 404, 409, 410, 412, 415, 442, 454, 455 interhalogens 401 inter-laboratory cali34, 273, 290, 311 bration inter-laboratory stan646 dard interlayer water 47, 49, 50, 52, 53, 55, 60 internal laboratory 408, 671, 731, 900, 1027 calibration internal normalization 149, 629, 699-703, 706, 718 157, 193, 195, 210, 241, internal standard 278, 408, 476, 478, 481, 512, 646, 794, 796, 882, 898, 900, 906, 1051 international standard 34, 138, 200, 259, 311, 360, 476, 478, 488, 599, 605, 606, 647, 759, 772, 784, 850, 855, 880, 903, 929, 935, 944, 953 international system of 907, 909, 912 measurement international system of 914 units 1043 internet technology interplanetary dust par- 661 ticles interstellar environment 229, 236 interstellar space 230, 236 interstitial water 53, 55, 139, 140 intracellular structure 682 intra-crystalline 423, 443 intra-mineral isotope / ic 444, 445 fractionation inverse linear relation 325 iodine pentoxide 500 iodomethane 609, 611, 613, 616, 618, 621 ion beam dispersion 550 ion beam double focus- 726 sing ion beam drift 857 ion beam monitoring 990 ion chemistry 91, 97, 856 ion (exchange)chromato-125, 127, 129, 158, 195, graphy 198, 337, 339, 552, 614, 716, 720, 721, 724 ion collection statistics 288 ion exchange resin 125, 129, 155, 313, 317, 325-338, 482-492, 499,

1207

Subject Index

ion microscope ion ion ion ion ion ion ion

optica crossover optical elements optica properties optica system peak repeller source

ion source compound ion source conditions ion source design ion source geometry ion source material ion source memory ion source parameters ion source plates ion source pressure ion source schematics ion source setting ion source technology ion source vacuum ionic contaminants ionic strength ionization cage ionization chamber ionization efficiency

ionization filament ionization sources ion-laser pump irradiation of protons irreproducible values isobaric contamination isobaric interference

564, 565, 614, 617, 637641, 646, 714, 1019 652, 656, 657, 663, 664, 675, 677, 678, 680 664 657 844, 851 694, 697 96 860, 964 95, 123, 127, 135, 137, 156, 162, 165, 168, 171, 175, 199, 283, 287, 289, 299, 302, 312, 344, 364, 368, 371, 374, 378, 390, 391, 394, 395, 398, 407, 466, 501, 539, 543, 546, 547, 593, 613, 642, 643, 657, 694, 766, 790, 796, 798, 803, 836, 838-843, 847, 848, 851, 857, 858865, 924, 926, 953, 957, 971, 990 124, 137 22, 368, 374, 395, 843, 859 842 873 125, 127, 131, 138 408, 446 173, 174, 395, 398 642 798, 841, 854 839 361, 873 789 636 487, 965 226, 323, 326, 336, 346, 610,617 859, 860, 863 22, 393-395, 397, 398 131, 138, 144, 146-149, 152, 303, 494, 535, 643, 645, 650, 654, 692, 694, 725, 746-748, 859, 860 130,131, 550, 643, 861 144, 926 782 114 123 149 101, 145-148, 156, 162, 168, 337, 343, 446, 537,

isobaric peaks Isogeochem (list) iso-propanol isothermal equilibrium isotope change monitoring isotope compositional bracketing isotope dilution / IDMS

545, 561, 635, 644, 664, 666, 695, 707, 710-719, 723, 731, 798, 800, 822, 832, 847, 853, 926, 958 536 19, 34, 1043, 1044 199 216 1014

645, 708, 743, 841,

996

88, 125, 540, 632-634, 637, 727, 729, 736, 738740, 745, 820-834, 907, 918, 928-930, 936, 937, 939 isotope (ratio) drift 133, 300, 478, 647, 947, 948, 952 isotope exchange kinetics496 isotope equilibration 20, 21, 27, 973 isotope/ic fractionation 2, 25, 34, 51, 54, 70, 74, 108, 122, 124, 128-130, 138, 141, 146, 147, 156158, 160, 171, 182, 189191, 193, 196, 199-202, 245, 249, 307, 311, 317, 320, 327, 329, 375, 384, 396, 402, 406, 407, 430, 436, 444, 447, 460, 463466, 471, 473, 483, 484, 497, 515, 525, 540, 541, 546, 555, 566, 586, 595, 603, 611, 6 1 3 , 624629,634, 649, 650, 655, 672, 701, 706, 716, 718, 719, 722, 724, 772, 813, 815, 889, 894, 895, 948950, 966, 996, 998, 1006, 1009, 1015, 1017, 1019, 1028, 1029, 1033, 1039, 1045, 1047 isotope fractionation 1045-1047 calculator isotope/ic fractionation 113, 375, 454, 781, 895 effect isotope/ic fractionation 28, 359, 992, 993, 1004, factors 1005, 1036 isotope hydrology group 622 isotope memory 455 isotope monitoring 707 isotope ratio measure909 ment science

1208 isotope ratio monitoring 350, 441, 460, 468, 495, 597, 598, 713, 854, 856, 944, 953 isotope salt effect 4, 31-33 isotopic calibration 34, 124, 154, 155, 171, 311, 503, 506, 526, 874906, 1025, 1029, 1036, 1051 isotopic diffusion 836 isotopic disequilibrium 1003 isotopic enriched tracer 190 isotopic equilibrium 21, 22, 29, 31, 62, 89, 371, 372, 471, 541, 559, 996, 1002, 1011, 1035 isotopic fingerprint 108, 121, 329, 723 isotopic fractionation 150, 708 artifact isotopic fractionation 1036 interaction isotopic fractionation 376 processes isotopic geothermome- 406,993,1031 try isotopic labeled drugs 682 isotopic labeling 540 isotopic lifetime 379 isotopic mass balance 445,477 isotopic mass fractiona- 677 tion drift isotopic peaks 77,670,742 isotopic standard 35, 123, 124, 132, 133136, 155, 380,476, 533, 604, 606, 627, 701, 708, 710, 714, 935, 945, 946, 949, 951, 967, 977-984, 1051,1052 isotopic thermometer 1030,1031 isotopic tracer 1,140,142,152,161,182, 306, 314, 316, 338, 341, 343, 507, 540, 621, 623, 649,856,947 isotopically depleted 2, 5, 51, 56,64, 68, 72, 158, 191, 199, 200, 245, 249, 251, 254, 273, 280, 387, 389, 540, 803, 889, 891, 892, 899, 900, 990, 1035 isotopically enriched 2, 6, 17, 35, 49, 72, 92, 107, 125, 128, 133, 158, 172, 178, 183, 188, 199, 201, 241, 243, 249, 251254, 260, 264, 308, 323, 328, 380, 382, 387, 397, 473, 474, 477, 481, 491,

Subject Index 540, 543, 566, 579, 621, 631, 634, 655, 714, 719, 723, 729, 75Z 780, 820834, 880, 888, 889, 891, 896-900, 903, 921-924, 930, 932, 935, 939, 940, 977, 989, 990, 992, 1001, 1004,1012,1020 182

isotopically heterogeneous isotopically homogene- 549,648,898,904,982 OUS

isotopically inhomogeneous isotopically labeled isotopomer

isotopomer bands isotopomer calibration technique isotopomer detector isotopomer determination isotopomer distribution isotopomer ion mass spectra isotopomer overlap isotopomer ratio isotopomer triplets

566 52, 682, 1004, 1049 89-95, 108, 110, 111, 168, 343, 344, 380, 387, 391394, 398, 399, 759, 766, 768-770, 772, 775, 778780, 782, 784, 786, 854 779 392 778 396 386 95 92 392, 396-399 849

jarosite 572 Johnson-Nishita reduct. 573, 586, 588, 591, 593, 594 juices 2, 4, 29, 35

K Kalrez kaolinite KBr KBr coarse scrap KBr furnace KBr reactor Kel-F kerogen

468 15, 39-51, 56, 57, 60, 61, 423, 429, 443-445, 461, 1021, 1029, 1031 410, 414, 418, 419, 421, 423, 435, 436, 439, 461, 470, 620 418 456 414, 418, 419, 456 411-416, 455, 462 58, 574, 579-581

Subject Index kerogen sulfur kerosene K-factor

579, 580 54 730, 740, 828, 829, 919, 920, 924, 940-942 K-factor drift 832 Kiba (reducing)reagent 581, 594, 595, 597, 598, 600 573,576,579 Kiba reduction 595 kidney stones kinetic analysis of data 258 810,814 kinetic diameter 239,264,813,994,1035 kinetic effect 550, 661,676, 681, 792, kinetic energy 793,844,859 925,926 kinetic gas theory 994 kinetic information 2, 94, 156, 160, 175, 234, kinetic isotope effect 235, 243, 247, 252, 339, 375, 387, 430, 450, 541, 548, 993, 1003, 1019, 1023 kinetic isotope / ic frac- 51, 244, 291, 309, 375, 649, 994, 1001, 1018, tionation 1021-1025,1030 337,994,1003 kinetic process 403 kinetic reaction 388,1026 kinetic studies 248 kinetically induced fractionation 249 kinetically related effects 316 Kjeldahl apparatus Kjeldahl determination 350 306, 315, 316, 318, 319, Kjeldahl digestion 340 Kjeldahl distillation 306,314-316,318,350 306,321,339 Kjeldahl method 319,321 Kjeldahl procedure Kjeldahl technique 315,350 659,674 Kohler illumination 659,674 Kohler method 20O Krebs cycle

L labeled compound 20, 103, 168, 178, 1049 labeled fatty acids 175 labeled food products 115 labeled gas 393, 398 labeled inorganic nitro- 322 gen labeled position 392, 396, 397 labeled storage room 418 labeled water 1002

1209 labeled working standard labeling studies laboratory atmosphere laboratory environment laboratory experiment laboratory intercomparison laboratory standard

laboratory tissue lamproite landfills large-scale ecosystem laser ablation

laser analysis laser applications laser bandwidth laser based fluid inclusion extraction laser beam

laser laser laser laser laser

chamber characteristics desorption determination diffraction particle sizer laser drilling laser emission laser extraction laser energy laser firings laser fluorination laser frequency laser heating laser laser laser laser

intensity interaction ionization light

393 540 285 779 376, 389, 647 290, 850 22, 28, 34, 243, 311, 357, 408, 476, 478, 648, 719, 880, 882, 898, 900, 905, 906, 929, 949, 966, 967970 457 269-271 209,210,212,775 179 1, 8, 350, 408, 409, 466, 495, 542, 547, 551, 570, 692, 711, 715, 723, 725, 756,788, 789 50,464,482,495 459,468 749, 752, 753 76 364, 374, 405, 459, 461, 462, 466, 467, 532, 533, 548, 715, 748, 751, 752, 755,772, 789 362,366,367,373 458,466 757 5O 42 76 362, 780 361, 364, 453, 458, 459, 532 458, 459, 463, 464, 466, 468,471 548 50, 405, 407, 412, 447, 458-461, 464, 465, 471, 495,537 777 50, 350, 358, 362, 400, 403, 407, 421, 495, 532, 972 772,777 436,458,459,461 747, 755, 793 754,772,776,777,914

1210

Subject Index

laser methods laser methodologies laser microanalysis laser noise laser peak laser photoacoustic spectroscopy laser pit laser power laser (micro-)probe laser pulse laser pyrolysis laser radiation laser laser laser laser laser laser laser laser

reaction craters resonance sampling selection source spectrometry spectroscopy system

laser technique laser tuning laser wavelength lateritic soil lattice atoms lattice binding energy leading peak lead-salt (diode) laser leaf/leaves leaf discs leaf length leaf litter leaf margin leaf material leaf mesophyll water leaf petiole leaf punch leaf pyrolysis leaf samples leaf segment leaf surface leaf temperature

44Z 461,537 403 430, 453, 461, 465, 469, 471 772,773,776 467 780 496 414, 465, 495, 748, 753, 780 359, 532, 782, 971, 972, 979,982,985,993 495,752 408 76, 405, 458, 467, 776, 777,780 464 746, 747,756 40Z 459 465,471 772,773,782,783 784 772,780,782 50,7Z 364,386,458,459, 466, 469, 495, 496, 725, 783 50, 361, 400, 403, 458, 483,495 775 458, 459, 463, 466, 467, 748,773,775 42,43 792 792, 793 168,169 760-763, 765, 772, 774, 783 180, 182, 183, 473-475, 478, 479, 514, 515, 518, 1029,1032,1033,1035 475, 479,481 481 188 481 514 474 474 474, 475, 479, 481 479-481 474, 476-478 475 481 473

182,183,473 1014 273, 473-481, 497, 1014, 1029,1032 164 Lewis acid 712 Li isotope standard 129,133,136 Li2CO3 standard LiBO2-V205 flux mixture 350 496 life science 169 lifetime of GC column 124, 132, 654, 693, 695, light elements 698, 702, 72Z 853, 911, 944,992,993 508, 509, 511, 512, 514, lignin 516,517,520,521 196,197 lignin-phenols 241,876,884,893,904 limestone 771 linear absorption 107 linear amplifiers 629 linear approximation 842 linear behavior linear/non-linear beha- 549 vior linear calibration curve 795 647 linear change 294 linear column 15,628,699,851 linear correction 363, 368, 478, 518, 703, linear correlation 704,707 671,868,960 linear dependence 795 linear dynamic range 629,731 linear equations 628,629 linear fractionation 369,848 linear function 299,300,302 linear interpolation 671, 699, 700, 702, 730, linear law 733 941 linear measurements 734, 735 linear model linear molecule density 780 344,385,390 linear molecules 479,515,518,539 linear regression 264, 396, 420, 702, 842, linear relationship 851 665 linear representation 782,853 linear response 903 linear shift 582 linear temperature increase 22 linear time-correction 302, 395, 398, 551, 669, linearity 774, 843, 851, 854, 873, 875, 931, 933, 936, 942 leaf tissue leaf transpiration leaf water

1211

Subject Index linearity limits of the IRMS Lipari obsidian lipids lipopolysaccharide liquid air liquid chromatograph liquid chromatography liquid reagent liquid-solid chromatography lithium aluminium hydride (LAH) LAH reduction lithogenic material lithogenic S compounds lithology lithosphere living material living organism living tissue liver tissue local ecosystem local standard long-term performance drift long-term storage Lorentzian rule low ionization sensitivity low-pass filter low-temperature interaction low-temperature environment lunar basalt lunar fines lunar grains lunar materials lunar nitrogen lunar regolith lunar rocks lunar samples lunar soil lunar spectra Luer ground joint Luer lock

477 903 182, 189-198, 201, 511, 517,1012-1014 200 402,421,429,450 193 153,199,201,1019 412 584 16, 572 23, 572 568 569 62 538,552, 568, 596 1011,1036 1 197 594 306 784,905,906 906 187,211,224,227,228 844 144 774 683 236 625 443 361,366,373,426 977 373 361 526,648 526,977 267,688 382 206,208,215 207

M Madan galena magnesia reagent magnetic interactions

65 485 104, 106

magneto-optical trap Maiella limestone main peaks mais germ oil major earth reservoirs major element chemistry major peak mammals Manhattan Project mannitol mantle (Earth) mantle epithelium mantle materials mantle petrologic work mantle (derived) rock mantle source marble standard Mariana trough marine air marine diatom marine ecosystem marine limestone Mars (planet) Mars mission Mars Pathfinder Mars Polar Lander Martian atmosphere Martian crustal fluid mass balance mass balance approach mass balance calculations mass balance consideration mass balance correction mass balance equation mass balance studies mass bias drift mass bias correction mass centered peak mass discrimination

mass dispersion mass filter mass filtered ions mass fragmentation analysis mass interference

759 903 608 172 1028,1036 795 199,583 200,483,1030,1035 1 147,200,722 140,141,683,1034,1036 1011 349 349 140,349,648,682,688 183 241 85,137 561 1015 172 892 688 758 913 758 688 687 224,335,432,477,574 325,474 53,56,196 55, 59 852 8, 371, 484 624 702, 710, 713, 723 699, 701, 704, 709, 715 636 94, 96, 122, 123, 132, 133, 135, 156, 158, 160, 171, 551, 698-704, 706, 708, 710, 713, 714, 716-718, 720, 721, 723-725, 737, 804, 923, 924, 978 132, 175, 664, 665, 679, 697, 851 97, 657, 658 95 394, 399 38, 58, 101, 158, 335, 361, 367, 409, 450, 508

1212 mass peak

90, 128, 149, 150, 644, 645, 728 mass resolution 367, 657, 664-666, 675, 677, 679-682, 685, 693, 697, 706, 727, 739, 798, 800 mass resolution capabi- 679 lity mass resolution require- 667, 675 ment 109, 367, 536, 664-666, mass resolving power 766, 771, 788 832 mass spectral interferences 795, 804, 835 mass spectrograph 34 mass spectrometer tuning 89, 95, 98, 146, 173, 607, mass spectrum / tra 636, 639, 641, 646, 664, 735, 839, 847, 873, 958, 972, 975, 986, 987 540, 541 mass-(in)dependent conversion mass-dependent isotope 25, 148, 623, 624, 710 fractionation mass-independent iso- 394, 534, 540, 992, 994 tope fractionation 378, 383, 804, 844 Mattauch-Herzog instrument (MS) Mattauch-Herzog geo- 679 metry 115 mathematical window functions 2, 36, 197, 261, 343, 492, matrix (a, the) 655, 673, 716, 721, 724, 729, 935, 946 681 matrix correction 788, 793, 828 matrix dependence 831 matrix difference 125, 1 3 8 , 395, 5 3 5 , matrix effect 570,657,672, 673, 6 8 8 , 690, 708, 710, 723, 726, 793, 794, 903, 947 133, 768, 939 matrix element 158 matrix interaction 789 matrix-matched standard 795 matrix matching 17 matrix materials 556 maximum conversion 465 measurable isotopic fractionation 644, 647, 743, 744, 985, measurement cycle 990

Subject Index measurement statistics measurement transparency medical applications medical field medical research medical science medical study medicine membrane desolvation membrane filtration membrane valves memory (effect)

memory issue memory signal memory size mercaptan mercuric cyanide mercury (native) mercury(II) chloride mercury pistons mesh size of exchange resin mesquite leaves metabasalt metabolism metal boat metal capsule metal catalyst metal fluoride metal hydrides metal oxide catalyst metal reactor metal reagent metal-fluoride melts metasediments metasomatic alteration metasomatic products metastable species lifetime metastable state lifetime metavolcanic meteoric sources meteoric water

690 822,833 944 896 176 1,122 98,176, 716 103,907 698, 713 556,563-56Z 572,590 836 4-9, 22, 23, 25, 29, 98, 14Z 159, 285, 286, 288, 294, 300, 313, 321, 452, 497, 498, 500-505, 543, 561, 585, 598, 60Z 608, 709, 722, 729, 743, 772, 784,816,842,852,941 114 852 114 553,554,585 333 319,421,461,499,720 312, 330, 497, 499, 500, 504,505 838 332 191 79,351 177,191,200 190 1006,1026 24 416,420,442 798 973 8 4,5 464 78-81, 84, 256, 269, 271, 351,354,356,357,360 348,359 351 386 386, 387 79, 80 63 15, 36, 682, 1016, 1032

Subject Index meteorite

meteorite (Allende) meteorite (Brownfield) meteorite (Cation Diablo) meteorite (iron)

1213 2,122-124,229,233,235, 236, 261, 266, 349, 362, 378, 404, 429, 443, 523, 526, 534, 625-628, 632, 688, 713-715, 718, 720, 723, 724, 788, 789, 803, 816,818,900,977 422,540,710,721 724 538,717

26~ 268, 350, 715, 717, 718,879,900,977 685,687,688 meteorite (Martian) meteorite (Murchison) 229,230-235 724 meteorite (Orgueil) 362 meteorite (primitive) meteorite (Semarkona) 724 meteorite (Sikhote Alin) 538, 900 meteorite impact 632 meteoritic amino acids 236 meteoritic components 632 meteoritic compounds 230,232 meteoritic diamond 266 meteoritic organic fea- 235 ture meteoritic organic frac- 233 tion meteoritic organic matter 229, 230, 236 meteoritic refractory 653, 689 inclusion meteoritic samples 123, 361, 816 meteoritic SiC 261 meteoritic sulfur 879 meteorological inviron- 508 mental variations 384, 508 meteorological parameters 4, 7, 67, 75, 76, 169, 231, methane 272, 291-303, 377, 766, 774, 775, 818, 847, 892, 894,1012,1023 methane combustion 298, 300 methane oxidation 1013 methanol 44, 127, 164, 166, 197, 198, 201, 223, 317, 318, 322, 323, 338, 516, 561, 584,609,1012 methanosulfonate (MSA)559, 561, 562 methanotropic bacteria 1013 methemoglobinemia 305 methionine 554, 555, 585, 595, 899, 902

149, 193, 195, 418, 445, 502, 541, 661, 668, 671, 680,687 537 methodology study 608,614,615 methyl chloride 163,166 methylates 164,166 methylation 822,928,929,935 metrology 133 Mg contaminants Mg isotope standard 710 Mg perchlorate 503,593,598 16 Mg-Pt reagent 57, 261, 264, 265, 348, mica 349, 351-354, 356, 357, 429,431,530,1001,1004, 1017 351 mica crystal mica sheets 240 351,356 micaceous rock 306,314 micro diffusion Micro Isotope Laser Ex- 462 traction System (MILES) micro reduction appa550 ratus micro scale fluorination 458, 460 microalgae 1014 microbe cultures 650 microbeam technique 459, 1017 microbes 291, 308, 344, 649 microbial conversion 591 microbial denitrifica330 tion microbial denitrifier 306 method (non-)microbial reduc- 343 tase microbial respiration 211 1011 microbiology microdiamond 270, 271 micro-IR 1017 micrometeorite 681 microporosity 843 microwave plasma 1022 Mid Atlantic 85 migration of air 410 milk 2,31 milk casein 595 millipore filter 215, 488 mineral composition 38, 407, 408, 690 mineral contamination 38, 43 mineral deposition 62 mineral grains 361, 422, 461, 462 mineral horizon 592

methodology

1214 mineral identification technique mineral oil mineral phases mineral separates mineral specimen mineral standard mineral surface mineral synthesis mineralization mineralogy

Subject Index 44

895 39,62,68,72,535,673 351-353,681 524,548,551 673 4,458,541,549 994 185 46,673,1005,1010,1016, 1017 miniaturization 451,458,461,46Z 471 minimal contamination 555 minor contamination 252, 657 minor peak 96, 144, 583 mixed matrices 673 Mn reactor 16 Mn reduction 7, 23, 36 model seawater 1015 moisture 20, 28, 206-208, 281, 457, 461, 474, 502, 685, 815, 816, 1003, 1040, 1041 moisture content 53, 1013 moisture traces 170 molar mass 526, 920, 924, 936, 939 molecular beam 378, 388, 657 molecular flow 838, 924, 925 molecular fragmentation 390, 393, 394, 655, 860, 987, 990 molecular interference 666, 714 molecular ion beam 393, 657 molecular radical 839 molecular sieve 3, 76, 158, 165, 174, 175, 189, 296, 297, 318, 337, 357, 383, 403, 430, 441, 465, 469, 470, 501, 503, 557, 559, 805-819 mollusk/ca 1011, 1028, 1029 molybdenite 719, 720 momentum transfer 655 Monel alloy 414, 429, 453, 462 monochromatic radiation758 monoisotopic 123, 128, 957 monosaccharide 201 Monte Carlo simulation 964 Montelupo clay 903 MORB 85,139,1034 mosses 507,516,517 mudstone 43 multi(ple)-collector 1,129,132,134,344,393, (MC) 450, 533, 546, 550, 645, 649, 650, 670, 674, 679, 681, 68Z 690, 692-725,

MC-ICP-MS multiple interactions multiple-pass gas cell multiple-reflection gas cell muscle muscle tissue muscovite mustard plants m / z peaks

733, 745, 800, 802-804, 90Z 926,992,1036 132-139, 450, 649-651, 692-725,992,993,1019 182 774, 775 774 191,594 181, 182, 191, 19Z 198, 594 351, 353, 356, 429, 445, 1003,1021,1031 594 97,29Z 299,358,598

N N contamination N20 contamination N20 cycle Na2CO3 standard Nation Nation dryer / drying Nation membrane Nation tube nails Nakhla pyrrhotite nano-diamond narcotic drugs narrow band-pass filter narrow peaks National Metrology Institution (NMI) national standard natural abundance

318,320,330 282 390 209 168,28Z 469 283,284,28Z 593 287,469 168,469 186,594 685 256,260,266 172 764, 778 853 822,824,935

935 1, 6, 103, 104, 108, 109, 115, 121, 155, 161, 172, 175, 178, 180, 306, 312, 314-316, 320, 323, 328, 378, 391, 392, 49Z 499501, 503, 550, 606, 632, 639, 754, 760, 761, 763, 769, 772, 775, 84Z 852, 898,947,972,986 natural abundance stan- 328 dard natural air 760-763,778 natural brines 32,1019 natural compound 108, 175, 597, 599, 881, 892 natural concentrations 323 natural contribution 632 natural cycling 273 natural distribution 947

Subject Index natural natural natural natural natural natural

ecosystem 200 element 628 environment 399, 523, 1027, 1034 exchange process 650 fluids 149 gas 7,157,173,159,169,291, 349, 579, 585, 894, 900, 951,1041 natural gas standards 948,951,956 natural groundwater 204 natural grain size 41 natural hydrocarbon gas 884, 893 natural isotope compo- 629-632, 633, 635, 648, sition 825, 829-831, 833, 889, 897, 923, 935 natural isotope / ic 122, 624, 631, 701, 702, fractionation 708, 718 natural isotopic ratios 16, 629, 630, 634, 635, 647 natural levels 975 natural material 122, 139, 147, 523, 538, 604, 605, 648, 691, 715, 717, 720, 875, 877, 880, 883, 900, 903, 911, 941, 947 natural matrices 903 natural mixture 123 natural origin 103, 112, 118, 120, 121, 128 natural phases 992 natural precursors 118, 120 natural products 118, 1016 natural profile 267 natural reference 137, 538 natural samples 36, 44, 119, 128, 135, 143, 152, 159, 193, 206, 209, 211, 222, 226, 266, 267, 315, 319, 320, 323, 328, 373, 376, 392, 406, 450, 585, 598, 600-602, 604, 606, 621, 625-628, 646, 709, 720, 721, 723, 756, 767, 772, 941, 985, 1002 natural sources 119, 121, 122, 556, 557 natural species 389 natural stable isotopes 112, 142, 146, 378, 387, 609, 623, 631, 635, 715, 717, 718, 720, 724, 835, 897, 899 natural standard 646, 647 natural state 233 natural settings 624, 626 natural substances 540 natural system 52, 993

1215 natural natural natural natural

uranium values variability variation

natural water

natural wetlands naturally occuring processes NBS22 oil NBS120c phosphorite standard Nd glass laser Nd-YAG laser nebulization negative peak negative TIMS

735, 738, 739 632 116, 209, 211, 713 122, 406, 549, 604, 605, 621, 623, 627, 630, 693, 713, 717, 718, 725, 831, 875 2, 4, 5, 17, 19, 21, 25, 36, 125, 133, 142, 203, 204, 206, 209, 211, 212, 219, 222, 224, 227, 312-315, 320, 329, 482, 483, 639, 767, 772, 780, 885, 888 775 391 896, 884, 897, 910 488-493

403 403, 465-467, 544, 548 138, 550, 711 859 143-145, 148, 149, 151, 152, 450, 451, 494, 550, 637, 642, 643, 747, 1015 neutralization 332, 337, 415, 418, 419, 421, 442, 451, 455, 488, 655, 661 neutralization processes 655 neutralization reactor 421 NH4+-free environment 320 Ni alloy wire 578 Ni bomb/tube pyrolysis 8, 23, 30, 497, 498 Ni catalyst 502, 1021 Ni-Cr-Co alloy 1025 Ni-Thoria catalyst 1021 nickelized graphite 334, 404 nickelized graphite 342 reactor nicotinic acid 319 Nier Johnson geometry 697, 727 Nier type ion source 364, 859, 860, 872 ninhydrin oxidation 320 NIST (National Insti34, 129, 133, 142, 305, 310, 311, 487, 533, 646, tute of Standards 706, 707, 712, 715, 716, and Technology) 719, 721, 750, 798-803, 874-906, 910, 918, 919, 922, 944948, 951, 953, 954, 956, 1049 nitrate contamination 329 nitrate radical 377

1216 nitrate reduction

320, 324, 330, 343, 344, 624, 649 nitrate standard 309, 335 nitric acid 10, 12, 13, 127, 320, 365, 366, 572, 579, 610, 611, 617-619, 637, 643, 739 nitrification 312, 1013, 1033 nitrogen atmosphere 215, 218, 570, 574, 579, 586, 595, 877, 1023 nitrogen isotope labeled 320, 345, 392, 396, 398 nitrogen standard 259, 357, 360, 368, 369, 371, 898 181 nitrogenous nutrient nitrous oxide reductase 335 899, 902 N-methyl anthranilic ester 835 Nobel Prize 1008 noble metal alloys noble metal ion-exchan- 813 ged zeolite non-absolute (isotope) 828 amount ratio non-adiabatic collision 385 non-aqueous environ240 ment non-baseline conditions 278 non-condensable gas 55, 59,76, 208, 216, 223, 379, 412, 451, 533, 818 non-dispersive technique 778 non-electrolytes 32 non-fluorination techni- 468 que non-homogeneous leaf 481 water non-hygroscopic 425, 486, 633 non-linear discrimina- 647 tion drift non-linear equations 629 non-linear expression 629 non-linear fractionation 687 non-linear scale contrac- 784 tion non-linearity 284, 288, 302, 312, 370, 647, 825, 838, 842, 848, 851, 852, 868, 875, 923 non-methane hydrocar- 377 bons (NMHC) non-quantitative con160, 162 version non-radiogenic isotope 623 non-reactive window 468 non-seasalt sulfate 557, 561, 562 non-silicate fraction 41 non-silicate standard 358

Subject Index 514 517 20, 150, 152, 699, 782, 784, 886, 887, 891, 895, 898, 900, 903, 954, 955, 1051 normalization correction 15, 150, 151 normalization laws 671 normalization procedure 144, 149, 151 normalization process 668 North Atlantic 82, 562, 714 North Pacific 137 Northrop-Clayton 998, 1000, 1001 method n-pentane 582, 584, 613 NSS aerosol 559 nuclear activities 738, 739 nuclear application 122 nuclear charges 110 nuclear chemistry 907 nuclear contamination 738 nuclear magnetic reso- 103-121,896 nance (NMR) NMR interactions 104, 105, 110 NMR magnetic inter104, 106 actions NMR power decoupling 109, 114 nuclear magnetogyric 104, 108 ratio nuclear mass spectro831 metry nuclear material 919 nuclear Overhauser en- 106, 110 hancement (NOE) 230 nuclear processing 124,540 nuclear reactions 1,122,903 nuclear reactor 729,738,919 nuclear safeguards 122 nuclear science 104, 749,753, 754 nuclear spin 124 nuclear technique nuclear test ban treaties 273 nuclear weapons 920 nucleation process 1003 nuclei 103-114, 118, 120, 766, 835 nucleopore filter 187 nucleosynthesis 2 nucleosynthetic effects 626 nucleosynthetic process 977 nuclidic mass 376 nutrient autoanalyzer 202 nutrient concentration 179, 181, 184, 187 nutrient distribution 179

non-wood sample non-woody plants normalization

1217

Subject Index nutrient loading nutrient source nutrient supply nutrients nutritional tracer study

179 179 1014, 1024 180, 181, 189, 306 6

O 180 thermometer 908, 1031 02 contaminant 430 02 oxidation 464 02 / CO2 reactor 430 observed isotope / ic 158, 375 fractionation oceanic basalt 688, 1033 oceanic chert 1032 oceanic rock 130 off-line (method/system)6, 157, 204, 239, 240, 255, 335-337, 350, 498-500, 502, 505, 511, 520, 598600, 602, 896, 1051 off-line attachments 192 off-line batch mode 16 off-line combustion 199, 201, 333, 336, 337, 340, 891, 896 off-line determination 603 off-line preparation 7, 8, 430, 432, 542, 595, 597-600, 883 off-line process 239 off-line pyrolysis 233, 497, 498 off-line studies 337 off-line technique 157, 191, 300, 336, 497, 498, 511, 519, 520, 891 oil 3, 4, 46, 172, 578, 579, 581-584, 883, 894 oil field 1012 oil free compressed air 1040 oil free pump 951 oil mist filter 1041 oil polution 173 oil sand 583 oil sand bitumen 581 oil shale 575, 599 oil seep 1033 oil spills 173 oil wells 159 oil-fired power plant 559, 562 olive oil 172 olivine 68, 70, 125, 402, 406, 429, 441-443, 464, 466, 471, 530, 673, 1031 olivine crystal 70 one-line sample combus- 164 tion

on-line (method/system)2, 6, 7, 11, 17, 23, 25, 29, 30, 74, 88, 96, 102, 157, 173-175, 159, 186, 202, 333, 335, 336, 403, 429, 497, 498, 501-504, 506, 508, 510, 511, 520, 547, 593, 598, 817 on-line analysis 156, 391, 505 on-line collection 418 on-line combustion 162, 319, 333, 334, 543, 891, 896 on-line coupling 174 on-line determination 603 on-line equilibration 18, 25 on-line furnace 258 on-line GC 853 on-line heating 350 on-line inlet 433 on-line IRMS 156, 1041, 1051 on-line measurement 8, 782, 851, 853, 1041 on-line neutralization 418, 419 on-line preparation 6, 290, 319, 542, 599, 879 on-line pyrolysis 174, 195, 233, 497, 500, 501, 847, 1041 on-line reactor 401 on-line reduction 7, 25 on-line technique 6, 497, 500, 511, 519, 520, 547, 598, 879, 891, 896, 971 on-line thermal conver- 1041 sion opal 56, 420 opaline frustules 1015 optical absorption tech- 282 nique optical components 467, 777, 842 optical detection 759, 766 optical excitation 780 optical feedback 777 optical geometry 735 optical grade window 421 optical isolator 774 optical isotope labora782 tory optical isotope ratio 784 instrumentation optical isotope ratio 760, 766, 780, 782 measurement optical microscopy 39, 659, 679, 1017 optical parametric 783 oscillation optical path length 768, 769, 771, 774-776 optical prism 844 optical properties 123

1218 optical pumping optical pyrometer optical set-up optical shielding optical spectroscopy optical spectrum optical system optical technique / method optically active 03 optogalvanic effect optogalvanic signal optogalvanic spectroscopy ore composition ore samples organ tissue organic acid organic assemblage organic carbon

Subject Index 754, 755 402 777 384 125,380,803 766 467 756, 759, 766, 779, 781, 782 377 780,880 780 780

716 707, 716 1024 163,164,366 236 47, 184, 185, 188, 200, 225, 345, 484, 529, 54Z 565, 600, 602, 895-89Z 1014,1023,1031 346 organic carryover organic cation exchanger 125 578 organic coatings 30Z 579 organic components 4, 7, 46, 112, 153, 155, organic compound 156, 158, 168, 172, 174, 175, 184, 189, 191, 194, 230, 132-234, 278, 313, 315, 334, 338, 343, 344, 362, 373, 407, 450, 490, 496, 506, 538,541, 546, 54Z 554, 569, 571, 574, 575, 578, 581, 588, 590, 594-596, 605, 615, 624, 635, 637, 639, 646, 895, 896,1041,1051 229,232,478 organic constituents organic contamination 70, 257, 270, 336, 338, 365, 489, 492, 565, 639, 646,1040 44,149 organic contents organic cosmochemist 229 245 organic dust 19Z 198 organic extract(ion) 235 organic feature 546 organic fluoride compounds 233 organic fraction organic fragment peaks 641 organic hydride peaks 639 organic hydrogen 46,1014, 1051

organic impurities organic interference organic isotope ratio organic inventory organic mass spectrometry organic matter/material

101, 895 315, 346, 636, 646, 647 173 230 154, 175, 844

3, 8, 11, 22, 24, 40,43-46, 58-60, 64, 69, 125, 127, 138, 149, 177, 178, 181, 182, 184, 186, 189-191, 200, 201, 203, 229, 230, 235, 236, 256, 25Z 268, 269, 271, 319-321, 338, 365, 366, 448, 473, 476, 482, 484, 489, 490, 498, 501, 506, 508, 54Z 573, 595, 613, 624, 682, 891, 895,1014 organic molecules 103, 339, 489, 624, 635, 798 organic network 233 organic nitrogen 256, 268, 305, 306, 312321, 330, 339, 344, 346, 899 organic oxygen 973 organic phosphate 490 organic precursors 236 organic reducing agents 582 organic reference mate- 895, 89Z 902 rial 173,642 organic residue 30Z 310, 341, 342, 49Z organic samples 507, 515, 546, 580, 615, 81G 899,973 322,345 organic solutions 7, 44, 195, 196, 208, 229, organic solvents 516,517,555,572,581 329,639 organic species 498,505 organic standard organic structure 553 45,334,501 organic substances 1013 organic substrate 552-554, 562, 565, 568organic sulfur 57G 579, 581,, 585, 586, 588,590-596,603 349,1007 organic systems 44,59,191,360,579,580, organic-rich 1012 41, 43, 44, 46, 53, 57, 58, organics 565, 636, 639-641, 646, 1021 organisms environment 196 organochlorine 615, 616 origin of life 229

1219

Subject Index 346 1003,1017 243-245, 249, 250, 278, 300, 515, 519, 644, 681, 685 883 outlier exclusion 277, 278 outlier rejection 883 outlier test 335 overnight incubation 897,899,1026,1027 oxalic acid oxic surface environment 329 59, 64, 75, 118, 188, 203, oxidation 262, 265, 267, 407, 427, 430, 490, 552, 565, 567, 571, 572, 586, 594, 595, 637, 639, 649, 967, 968 162 oxidation catalyst 199, 303, 342 oxidation furnace 307 oxidation numbers 967 oxidation procedure 319 oxidation promotion 75 oxidation rate 45, 46, 430, 624 oxidation reaction 973 oxidation reactor 377, 552, 553, 601, 624, oxidation state 625, 641, 941 637, 638 oxidation trap oxidation-reduction be- 619 havior oxidative pyrolysis 68, 70 oxidizing environment 625 oxidizing reagent 489 oxyfluorides 617 oxygen atmosphere 59, 256, 258, 265, 362, 366, 403, 489, 548, 972 oxygen sink 162 oxygen-free atmosphere 500 ozone/03 291, 309, 375-390, 1022, 1024 ozone chemistry 376, 382 ozone cycling 385 osmosis Ostwald ripening outlier (results)

P Pacific ocean Pacific seawater palaeoenvironment paleoclimate paleo-dietary habits paleohumidity palladium foil parabola spectrograph

137, 292, 550, 559, 561, 1012 137, 712 198, 482, 483, 597 38, 483 173 482 8 835

para-hydroxybenzaldehyde parent-daughter comparison parent-daughter correction parent-daughter fragments parent-daughter method parrafin wax Parr bomb Parr bomb combustion Parr bomb oxidation Parr bomb technique partial electrolysis partial exchange experiments partial exchange method partial exchange reactions partial combustion partial fluorination partition function passive fluorination Pd filter Pd foil membrane Pd-Pt catalyst peak area peak broadening peak data peak definition peak detection peak distortion peak flanks peak height peak hopping peak integration peak intensity peak interference peak maximum peak overlap peak profile peak reproducibility peak resolution peak schematics peak separation peak shape

117, 120 989 987, 990 987 988, 989 4, 224 578-580, 582, 584, 594, 615 574, 579 593 581, 595, 615 30 995,998 998, 999, 1001 64 515 51, 52, 420, 423, 427, 443-445, 464, 471 384, 770 5O 8, 175 175 1021 115, 162, 226, 297, 302, 303, 503, 840 169, 170, 855 168 170 168 170, 171 843 115, 297, 364, 366-368, 372, 664, 799, 832 634, 646, 989 168, 286 96, 113, 115, 768, 769, 773, 795 113, 794 664 113, 115, 156, 161, 168170 367 285 664, 665 665 666 96, 170, 665, 680, 693, 694, 844, 845, 851, 873

1220 peak shift peak signal peak size peak slope peak symmetry peak tailing peak to tail signal peak width peak-jump(ing) peat pectin pedosphere PeeDee belemnite PeeDee formation pegmatite perchlorates perchloric acid periodic table permeability permeable tape permeable to H2 permeable to liquid permeable to water persulfate oxidation petroleum petrologic persuits petrologic problems petrologic study petrologic system phenyl group phenylacetaldehyde phenylacetic acid phenylethanol phloem phosphate chemistry phosphate contamination phosphate radical phosphatic tissue phospholipids phosphoric acid

Subject Index 125,126 743 199,312 665 115 169,171,299,742,744 666 226, 296, 664, 69Z 799, 855 29,76,465,549,728,847 568,585-591 200 638,541,552,596 879 879,892 351,406 124,125,60G 605 124,426,524 653, 694, 789, 790, 793, 835,939 41, 53, 378, 418, 752, 1006,1008 323 1024,1027 513 469 6O8 44, 185, 235, 578, 579, 583,951 358 349,358 349,359 360 553 118 119,120 118-120 473,481 239 644

482,484-486 483,490 196 24, 58, 12Z 206-209, 211, 215, 216, 218, 220-223, 237-239, 243, 245, 248, 250, 253-255, 569, 588, 591,645,895 496 phosphorus cycle photoacoustic detection 767, 778, 779 photoacoustic spectro- 762-764,780 scopy photochemical control 384

photochemical equili384 brium model photochemical forma- 377 tion photochemical lifetime 379 photochemical produc- 378 tion photochemical smog 378 photodissociation 377, 385 photolysis lifetime 379 photon flux 750, 752 photon stop 727 photorespiratory process 200 photosynthesis 177, 200, 206, 211, 272, 1030 photosynthetic fixation 183, 200, 1013 photosynthetic fluxes 290 photosynthetic metabo- 200 lism photosynthetic pathway 116 photosynthetic process 177 photosynthetic produc- 481 tivity photosynthetic uptake 272 phototoxic effect 378 phyllosilicates 56 physical-chemical 267 background physiological changes 183, 191 physiological effects 108 physiological invariant 182 source physiological processes 182, 507 physiological questions 175 physiological variation 182 phytoplankton 178, 179, 182, 184, 187, 200, 1014 pioneering application 993 pioneering technique 138 pioneering work 349, 622 piston-cylinder capsule 1017 plagioclase 68, 351, 352 Planck's constant 752, 924 planetary cores 625 planetary differentia632 tion planetary system 229 planet-forming processes 2 229, 361 planets plant appearance 185 473, 1033 plant environment 174, 191, 197, 201, 473plant matter / material 475, 479, 497, 507, 517, 593, 594

Subject Index plant mesophyll plant parts plant physiology plant press plant samples plant plant plant plant plant plant plant plant

sap sources species structure tissue transpiration rate type water

plant wax esters plant-derived hydrocarbons plants

plasma (blood) plasma (ionized gas) plasma amino acids plasma analysis plasma ashing plasma conditions plasma discharge plasma flicker plasma furnace plasma interference plasma ionization plasma lactate plasma power plasma source plasma structure plasma torch plastic syringe platinized platinum black Pleistocene Pliocene pneumatic air supply pneumatic control pneumatic cylinder pneumatic valves

1221 182 473 497,507 185 3,181,185,197,474,476, 593,626 474 192 517 518 180,197,592,593 473 517 2, 473, 474, 481, 1014, 1032 196 196 3, 22, 29, 108, 11Z 17Z 179, 180, 181, 185, 195, 197-200, 272, 378, 473, 497, 516, 592-594, 649, 1011, 1013, 1014, 1024, 1028, 1030, 1032, 1033, 1035,1036 2,10,31,32 45, 132, 458, 489, 550, 650, 651, 656, 694, 70Z 720, 791-794,796,926 173 139 40,44,45,48,58 713 748 694, 726 45,46,489 696,698 693,698,700,701,706 161 46 138, 139, 551, 651, 692, 694, 697-699, 701, 709, 713, 725,74Z 796 790, 792 550 26,20Z 212-214,225 5,9,23,423 16 495 495 950 474,476 475 284,462,950

Poisson statistics 667, 855 polar hydrocarbons 230 polar ice 564 polar organic compounds236 Polar Science Program 889 polarization 749, 753, 754 polarized light 749 pollen 555 polyatomic interference 651 polyatomic ion 642, 651, 666, 707, 715 polycarbonate 186, 506 polycyclic sulfide 554 polyethylene 883, 884, 895, 897 polyethylene beaker 525 polyethylene bottle 3, 212, 213, 313, 618 polyethylene flask 1013 polyethylene foil 910 polyethylene glycol 163, 170 polyethylene screen 326 polymerization 239, 895 polythionate 553 poorly defined matrices 673 Porapak 285, 286, 612 Poraplot 296 pore water 15, 22, 38, 53-56, 60, 132, 204, 572, 590, 603 porous membrane 54 porous resin beads 19 porous Teflon frit 640 porphyrins 200 possible interaction 792 possible isotope/ic frac- 36, 343, 350 tionation positive peak 859 positive TIMS 143-150, 152, 614, 549 post-depositional envi- 483 ronment potential contamination 41 potential health hazard 578 potential isotopic frac- 5, 34, 311 tionation potential mass fractio- 144, 146 nation artifact potential sensitivity 988 powdered wood grain 507, 521 size power density 463,548,751 power generator (UPS) 1039 power law 671, 699, 700, 702, 704, 706,707,730,788,796 Precambrian 268 Precambrian cherts 268 precious / noble metal 1006,1025 capsule

1222 precipitation (rain, ..) precipitation/rainfall collectors precipitation procedure precipitation process precipitation reactions precipitation reagent precipitation technique

Subject Index 2, 36, 306, 497, 557, 564, 1033 312, 556

212 221 203 212 211-214, 217-220, 222, 228 pre-combustion 25Z 365,366 pre-concentration 293-29Z 303,556,565 pre-concentration device 295, 303 predominant sink 377 prefluorination 50, 57, 432, 435, 438, 440, 456, 467, 496 preliminary experiment 50, 819 presolar grains 362 pre-sputtered 798 pressure-equilibration 283, 285 (-volume) pressure-filtered 590 pretreatment 4, Z 16,1Z 22,23,38,44, 45, 318, 319, 422, 424, 426-428, 432, 435, 438, 440, 556, 563, 568, 759, 766,775,781,895 primary artifact 945,954 primary beam 123, 535, 536, 549, 652, 655, 656, 658, 660-662, 66Z 670, 674, 675, 679, 680,685,688 primary ion 535,655,658,668 primary ion beam 53Z 652, 655, 65Z 659, 674,680,682 primary ion delivery 658 primary ion spot 660 primary reference mate- 892, 898, 900, 902-905 rial primary sink 291 primary standard 133, 772, 784, 881, 929933, 944, 949, 954, 978, 979 primates 181 principles of statistics 872 pristane / phytane ratio 578 probability functions 975 pronounced isotopic 123 fractionation propanol 4, 165, 533 protein 4, 32, 161, 182, 195, 198200, 489, 592, 624, 1014 Pt boat 617,618 Pt capsule 1008, 1020, 1026

Pt catalyst

17, 19, 20, 21, 55, 403, 473, 850 Pt-A1203 catalyst 21, 474 PTFE membrane 323, 572 pulse counting 364, 667-669, 739, 740 pulsed laser 777 pump oil 46, 456, 457, 850 purines 200 Pyrex 10, 13, 14, 16, 24, 28, 65, 72, 208, 215, 220-222, 274, 275, 77, 287, 337, 410, 414, 427, 467, 495, 525, 528, 583, 611, 1021 pyridineacetic anhydride 201 pyrite 441, 447, 547-549, 552, 568-576, 579, 581, 582, 599, 602, 603, 686, 723 pyrite fraction 574, 576 pyrite grains 465, 685 pyrohydrolysis method 617 Pyrolysis/pyrolyzed 22, 65, 70, 87, 174, 175, 233, 256, 258, 262, 265, 268, 334, 342, 362, 366, 373, 401, 404, 408, 474, 474-481, 497-506, 582, 818, 847, 1041, 1051 pyrolysis fragments 175 pyrolysis interface 174, 175 pyrolysis reactor 8, 494, 501, 1051 pyrolysis system 174, 175, 342, 476, 478 pyrolysis technique 24, 476, 479, 481, 494, 497-506 pyrolysis temperature 8, 499 pyrolytic products 175, 497, 500 pyrolytic reaction 174 pyroxene 70, 431, 433, 441, 442, 530, 570 pyroxene glass 443 pyrrhotite 441, 447, 465, 568, 570

Q Q (quadrupole)-ICP-MS 132, 692, 693, 698, 702, 708-717, 720, 727, 734 Q-switched laser 753 quadrupole field disper- 734 sion quadrupole mass filter 93, 95 quality control 2, 36, 290, 293, 301, 343, 739, 1051 quality control material 881, 882, 904 quality control purpose 905 quality control tank 278, 279 quantitative conversion 4, 5, 923

Subject Index quantitative peak collection quantum cascade (QC) laser quantum interaction quartz (mineral)

1223

158 760,765,783

748 46, 56, 5~ 66, 68, 69, 72, 74, 359, 401-403, 406, 420, 421, 429, 433, 436, 442-445, 451, 457, 458, 461, 463, 465-467, 530, 532-534, 658, 878, 996, 1003-1005, 1009, 1011, 1020,1030,1031 quartz arenite 465 quartz furnace 858 quartz insulator 791 quartz phenocryst 443,444 quartz sand 533-535 quartz siltite 465 quartz standard 453 quartz vein 71,358,359 quartz wool 5,59 quartz wool filter 559 quartz xenolith 68 quartzite 535 quasi simultaneous arri- 668 val (QSA) effect quaternary science 154

R Rabi frequency RC filter radial diffusion cell method radiation radiation intensity radiation wavelength radiative forcing radio frequency radio frequency heating radio frequency induction furnace radio frequency radiation radiogenic radiogenic contents radiogenic decay radiogenic effects radiogenic isotope rain rainfall

750, 752 859 55,60

rainstorms rainwater rainy season Raman Raman microprobe Raman spectroscopy random interleaved sampling Raney-Nickel alloy rapid irradiation rare bird species rare earth elements raspberry ketone Rayleigh distillation reacted sample storage reaction bottle incubation reaction equilibrium reaction kinetics reaction pathways reaction products reaction products reactivity reactive compound reactivity of zinc reactor (unspecified) reactor (high temperature) reactor tube reactor surface reactor temperature reactor walls reagent

366,777,783 790 458 272,291 45,114,115,384,789 404 402 650 523,632 632 671 632 623, 624, 627, 628, 632, 693,699, 701, 715 3, 36, 308, 317, 325-327, 331,332,556 108,116,312

reagent blank reagent contribution reagent cost reagent cylinder reagent mixture reagent neutralization reagent poisoning reagent pressure reagent reactivity

306 108,313,559 172 101Z 1036 77 65,68,77 777 588 463 820 666,806,1023 114,117-119 430,923,1033 14 346 91,92 2Z 91,998 203,309,653,839 39, 252, 356, 402, 419, 421,451 442 1

11 5, 7, 23, 165, 418, 421, 500, 502, 583 578 583 5OO 494 5O0 15, 23-25, 30, 36, 50, 51, 138, 148, 160, 164, 165, 184, 185, 217, 220, 227, 238, 241, 250, 251, 318, 320, 330, 338-340, 343, 352-356, 404, 408-413, 417, 418, 420, 422, 424, 426, 428, 431-437, 440, 441, 443, 451 ,452, 455, 456, 516, 543, 572, 578, 637, 638, 644, 729, 743 127, 452 23 412 412, 415, 417 321 416, 418 4 466 16

1224 220, 410, 415-419, 455, 462,527,528,895 415 reagent transfer 213 reagent water sample 387 recombination theory 1004,101Z 1027 recrystallization 30Z 625 redox chemistry 624 redox conditions 624 redox indicator 624,625 redox process 30Z 493,624,1026 redox reaction 625 redox transformations reduced inorganic sulfur 571, 587, 588, 592 reducing agent 7, 333, 483, 553, 581, 582, 650 reducing environment 562,625 reducing reagent 5 550,650 reductants injection 7 reduction / pyrolysis 55O reduction flask 973 reduction reactor 342 reduction system 52,239,240,250 re-equilibration reference gas injection 284,285,880,882 4Z 112, 131, 133, 136, reference material 241, 243, 261, 26Z 271, 30Z 310, 311, 33Z 341, 392, 506, 526, 533, 538, 539, 544, 546, 551, 596, 598, 599, 606, 648, 706, 729, 735, 736, 738, 765, 768, 769, 772, 774, 778, 784, 786, 828, 833, 855, 874-927, 932-936, 944, 949, 954, 959, 974, 978, 981,985,1048,1049 reference material sto- 888,896,953 rage 171,297,855 reference peaks 15,33,135,311,336,604, reference standard 605, 648, 798, 803, 882, 1052 413 re-fluorination 1017 refractive index oil 70Z 708 refractory oxides 173 refractory waste 339,500,703-707 regression line 475,502,1014 relative humidity 104, 109, 494, 670, 779, relative sensitivity 795 498 reliable standard 690 reliable statistics 298 repeated oxidation 916 replicate experiment reagent storage

Subject Index reproducible analysis / measurement reproducible apparant mass discrimination reproducible data reproducible DIC concentration reproducible experiment reproducible extraction technique reproducible filament loading reproducible isotope mass-discrimination reproducible isotope values

495, 507, 675, 688, 690, 726 734 69, 548, 690 209 1011 508, 516, 520 627 675

45, 193, 195, 209, 217, 310, 311, 358, 447, 488, 495, 507, 546, 726, 1050 74 reproducible method reproducible procedure 222 reproducible reaction 243 reproducible repeata729 bility reproducible results 209, 219, 674, 966 reproducible yield 488 residual artifact 673 resin/diffusion method 328 resin exchange capacity 327, 639 resin exchange column 125, 323, 326, 328, 331, 338, 564, 640, 642 resin from bitumen 582 resin volume (in column) 125, 127, 326, 564 resonance excitation 755 resonance frequency 104, 105, 107, 114 resonance ionization 746-756 resonance signal 106 respiration 206 respiratory fluxes 273, 290 respiratory releases 272 retardation lens 695, 696 reverse electrolysis 863 Rh catalyst 1021 rhyolite 127, 137, 140, 903, 904 risk assessment 928 Rittenberg analysis 322 Rittenberg method 340, 341 Rittenberg oxidation 321 Rittenberg technique 321 Rittenberg Y-tube 340 river water 132, 139, 140, 1033 rock matrix 570, 647 rock matrix standard 647 rock standard 132, 137, 138, 647, 795 rodent teeth 437 root respiration 203, 1014

Subject Index rose oil rotary evaporator/ion rotational constant rotation-vibration bands rotation-vibration motions rotation-vibration interaction rotation-vibration transition routine chemistry routine normalization rubber septum RuMP cycle

34S enriched spike salicilic acid saline brine saline fluid saline groundwater saline water

1225 118 197,201, 564 770 766 759 770 770 487 902 55, 221, 224, 571 1013

540 318 130, 1018 6 134 11, 22, 324, 639, 1034, 1035 saliva 2, 10, 31, 88 salt 2, 27, 32, 33, 40, 43, 146, 185, 310, 318, 319, 330, 337, 342, 346, 410, 421, 490, 553, 610, 620, 998, 1016, 1018, 1027 salt lakes 140 salt solution 31, 32, 1019 Salton Sea 1029 sample blend 821 sample capsule 615, 1008 sample combustion 298, 341 sample digestion proto- 314 col sample filament 550,643,645 sample impurity 138 sample incubation 188,345 sample injection 6,7,9,156,162,169,171, 206, 211, 283, 285, 295, 297.470 sample introduction 285, 293, 458, 468, 708710, 713-715, 717, 725 sample loop 6,283-285,288,293,296, 298,302,303,470 sample loop equilibra- 285 tion sample matrix / ces 178, 188, 450, 596, 635, 637, 651, 706-708, 731, 735, 738, 796, 973 sample peak 168, 289, 299

sample purity sample storage

38, 138, 710 2, 3, 157, 179, 184-187, 193, 200, 205, 206, 211, 212, 214, 215, 220, 228, 280, 312, 475, 479, 562, 593 sampling protocol 178, 184 sandstone 39, 42, 79 satellite observations 385 scale contraction 539, 784, 900 scale conversion 887, 888 scale shrinkage 7, 14, 15, 34, 35, 311 scale stretching 34, 35, 311 Schumann-Runge bands 378 Schtitze-Unterzaucher 500,501 method science 928 Se contamination 643, 647 Se oxidation 649 Se reduction 626, 636, 637, 649, 1018 sea-air interface 559 seafloor basalt 139 sealed-tube 12, 14, 59, 189, 191, 201, 333, 353, 404, 406, 611, 616, 1024 sealed-tube combustion 326, 327, 332, 339, 340, 351 sealed-tube decripitation 358 sealed-tube extraction 348,358 sealed-tube experiment 348, 1020 seasalt sulfate 556, 557 seasalt tracer 557 seawater 32, 125, 132-142, 316, 323, 335, 336, 490, 606, 614, 710, 712, 721, 722, 900, 904, 905, 1015, 1018, 1019, 1028, 1029, 1032, 1034, 1036 seawater chloride 904 seawater compartments 878 seawater derived 559 seawater matrix 149 seawater pH 1035 seawater samples 25, 129, 132, 324, 723, 903,904 seawater source 63 seawater sulfate 556,900 seawater temperature 1035 seawater-rich oceanic 689 crust secondary ion 1, 535-537, 549, 652-691 secondary ion beam 653, 655-657, 659, 661664, 669, 675, 677-679

1226 secondary ion characte- 658 ristics secondary ion composi- 755 tion secondary ion collection 668 secondary ion current 549 secondary ion emission 658 secondary ion extraction 656, 658, 674, 675, 679 secondary ion generation 549 secondary ion mass spec-123, 405, 444, 523, 535, trometry/SIMS 537, 543, 548, 549, 551, 652-691, 788, 789, 799, 800, 802, 992, 993, 1015, 1018, 1019 secondary ion optical 663, 679 axes secondary ion signal 659, 670 secondary ion speciation 657 secondary ion spectra 664 secondary ion system 674 secondary ion transmis- 536 sion secondary ion tuning 660 conditions secondary ion yield 658 secondary ionization 654, 655, 746-748 sedimentary apatites 486 sedimentary basin 140, 349 sedimentary deposits 483, 813 sedimentary phosphate 483 sedimentary processes 38 sedimentary rock 43, 81, 88, 404, 568, 574, 575 sedimentary samples 547 sedimentary sulfides 685, 687 sedimentary sulfur cycle 572 sedimentary system 547 sediments 38, 39, 41, 45, 56, 62, 64, 124, 132, 139, 140, 141, 180, 184, 185, 188, 191, 196-198, 204, 227, 228, 270, 271, 336, 450, 568, 572, 576-579, 597, 602, 603, 632, 639, 649, 650, 689, 715-720, 972, 1012, 1019, 1033 seeds 555, 978, 1023 seismic monitoring 273 selected ion monitoring 153, 154 selective fluorination 404 seleniferous shale 624 self-diffusion 1004, 1020, 1023 self-neutralization 661 Semarkona chondrite 688

Subject Index semi-conductor laser semipermeable membrane semipermeable Pt semisynthetic sensitivity

775 168, 1006

1008 121 103, 108, 109, 130, 132, 144-146, 155, 15Z 284, 29Z 362, 364, 375, 395, 398, 399, 401, 453, 460, 468, 500, 508, 535, 550, 632, 634, 658, 666, 679, 680, 691, 69Z 698, 713, 723, 725, 74Z 773, 782, 832, 838, 839, 841, 84Z 848, 851, 854, 880, 926, 944,957,959,964 sensitivity estimates 482 sensitivity limitation 108, 109 sensitivity optimization 154 sensitivity specification 842 septum 11,98, 208, 216, 217, 225, 550, 616 septum coring 206 septum tubes 206-209, 211, 215-217 septum-sealed 17, 29, 98 sequential diffusion 324 Se-rich environment 626 seriously contaminated 281 serpentine 52 serpentinized peridotite 123, 889 224-226 serum bottles shaking cycles 26 shale 39, 43-46, 53, 54, 56, 58, 140, 354, 580, 597, 624, 1023,1025 shell 6,69,1011,1028,1035 22,50Z 516,517 shrubs Si isotope standard 533 Si metal matrix 654 594 Siberian mammoth side filament 861 signal dispersion (NMR) 120 significant contamination49, 320, 566 significant isotope / ic 591, 625, 636 fractionation 637 silane coating 3,186,188,322,357,430, silica gel 433, 439, 469, 549, 550, 617,643-645,809 silica gel matrix 643 silica glass capsule 1006 silica window 76 silicate carrier 223 silicate detritus 125

Subject Index silicate materials silicate matrix silicate melt silicate rock silicate standard silicate studies silicate system silicates (minerals) silicic acid silver capsule silver cyanide silver filters single collector silicon tetrafluoride silver boat single filament ion source site specific characterization site specific deuterium content site specific isotope ratio site specific nitrogen isotope analysis site specific parameters skimmer (cone) smectite SMOB SMOC Sn catalyst snow

snowmelt SO2 gas lifetime SO3 radical sodium-spiked zinc software (computer)

softwood soil

soil amino acid

1227 350 465 253 124,348,349 35Z 360 350 349,358,359 24,39,41,50,52,56,124, 348,350,359 1015,1032 332,503 333,334,340 312 139, 427, 645, 684, 692, 693, 697, 788, 790, 798, 800,803,804,926 405 337,341,342 643,644,867

soil carbonates soil digest soil extracts soil fraction soil gas soil horizon soil hydrolysates soil material soil microbes soil phosphate soil properties soil samples soil scientists soil solutions soil sulfate soil sulfer compounds soil water

103 896 110, 111 390-399 113 133, 550, 551, 728, 730, 796 41, 44-49, 52, 53, 60, 1017,1032 606 606,609,877,904,905 493 3,308,331,536,775 306,885 552 238 14 33, 115, 168, 169, 284, 286, 294, 299, 311, 599, 731, 798, 833, 852, 856, 954,964 514-516 3,10,15,38-61,172,173, 179-181, 184, 188, 197, 203, 312, 315, 322, 329, 341, 344, 428, 450, 559, 585, 586, 588, 590, 591, 649, 745, 788, 789, 803, 1013,1033 173

soil zone soil-respired CO2 solar absorption solar furnace solar material solar nebula solar neutrino experiment solar radiation solar system solar system (early) solar system environment solar system oxygen solar wind Solenhofen limestone solid hydrous substance solid standard soluable nitrate reductace (NaR) solute interaction effects solvent solvent extraction / extracts solvent impurities solvent storage solvent system solvent-insoluable sonification Soufre de Lacq

180,183 636,638 314, 315, 320-323, 344, 345 588 55,817,900 541,586,590 344 586,590 291,649 58 59O 3, 10, 15, 22, 41, 44, 54, 55, 344, 586-591, 626, 795,1013 42,306,315,341 322,345 590,591 586,590 2, 15, 22, 29, 54, 55, 273, 49Z 1014,1035 312 183 38O 1023 1023 229 772 377 236, 361, 628 229, 627, 632, 689, 710, 715 229 977 688 892 481 209, 218, 476, 478 344 1016 107, 109, 110, 113, 117, 165, 185, 188, 191, 19Z 198,572,576,70Z 708 113, 191, 230, 572, 576, 579,583,584,717 155 1041 118 233 41 884,900,901,910,981

1228 source environment source filament source tuning sources and sinks South Pacific Soxhlet apparatus Soxhlet distillation Soxhlet extractor / tion Soxhlet system space shuttle mission space technology spark discharge conversion spark source of MS spatial resolution

spectral interference spectrographic spectroscopic interference sphalerite

Subject Index 232 368 279 179, 272, 273, 375, 378, 379, 390, 399 559 44, 513, 517, 572 513 191, 513, 516, 517, 579, 580 197 382 866 404

sputtering fractionation sputtering ion source sputtering procedures sputtering process sputtering region square peak SrC12-NH4OH reagent stable isotope methodology stainless steel filter standard

790, 795 8, 405, 407, 439, 444, 453, 459, 464, 471, 482, 483, 495 ,496, 547, 551, 652, 659, 682, 691, 747, 754, 755, 758, 788, 972 706-708, 717, 727, 832 11 726

441, 447, 465, 466, 549, 568, 570, 599, 884, 900, 901, 910, 981, 1025, 1030 spike compounds 738 spike material 736, 738, 821, 830, 930 spike solution 628, 633, 634, 640, 641, 736 spiked standard 634, 636 spin-lattice (NMR) 105, 113 splitless injection 171, 159 spring water 142, 209, 649 sputter cleaning 688, 794 sputter rate 535, 536, 655, 673, 793 sputter site 652, 657, 661 sputter yield 793 sputtered area 789 sputtered atoms 535, 536, 672, 790, 792795 sputtered material 655 sputtered neutrals 793 sputtered pit / crater 549, 799 sputtering / sputtered 405, 532, 537, 549, 652, 653, 655, 664, 671, 675, 678, 680, 754, 755, 792794, 843, 862 sputtering behavior 793 sputtering characteristics 795 sputtering effects 848 sputtering event 674

standard standard standard standard standard standard standard standard

standard

670, 674 123, 789 664 670 661 299 211-214,218, 220-222 178, 324

54 6, 15, 16, 21, 34, 35, 44, 132, 133, 136-138, 155, 157, 171, 191, 196, 204, 209, 218, 240, 241, 255, 274, 279, 282, 297, 301, 303, 308, 311, 312, 325, 335, 392, 502-506, 523, 547, 599, 604, 606, 623, 627, 631-638, 646-653, 659, 668-676, 681, 682, 685, 687, 690, 702-723, 768, 772, 776, 798, 799, 830, 832, 842, 850, 873906, 908, 913, 929, 932, 935, 941, 944, 945, 953, 966-990, 994, 1045, 1051, 1052 additions 125, 325 analysis 300, 647, 681 artifacts 945 compound 193, 195, 201 conversion 337 data 681, 954 definition of 6 307 deviation 7, 45, 50, 54, 129, 130, 132, 141, 145, 146, 174, 175, 199, 207, 219, 234, 243, 278, 279, 298, 290, 301, 302, 324, 334, 337, 383, 384, 396, 478, 488, 493, 494, 506, 511, 515, 520, 534, 601, 615, 621, 646, 685, 726, 735, 738, 765, 799, 801, 882, 887, 890, 954, 955, 977, 978, 989 error 100, 303, 325, 328, 480, 481, 505, 520, 648, 744, 889, 965 fluorination 413

standard line standard gas

35, 165, 300-302, 310, 311, 364, 392, 393, 408, 501, 503, 543, 855, 898,

1229

Subject Index

standard heterogeneity standard isotope abundance standard isotopic scale standard manifold standard material

978, 1051 672 101

953 274 259, 261, 311, 392, 547, 648, 681, 682, 759, 768, 784, 795, 798, 803, 944, 946,968 651 standard matrix / ces standard measurement 202, 364, 702, 716, 720, 724,873 5, 15, 70, 164, 31G 328, standard method 348,615 101 standard mixture 60, 319, 357, 364, 374, standard procedure 524,610,646,852 Standard Reference Ma- 946 terial (SRM) 44, 96, 710, 712 standard samples 208, 219, 223, 317, 324, standard solutions 633, 634, 648, 704, 708721 556 standard state conditions 5O standard system 275,280,294,302,368 standard tank 35, 39, 49, 53, 351, 357, standard technique 703,716,852,853 154 standard tool 729, 733, 734, 744, 834, standard uncertainty 876, 883, 887, 893, 89Z 899, 901, 904-906, 916, 937,945 149,491,969 standard value 6, 20-22, 28, 34, 35, 92, standard water 280, 408, 476-478, 481, 850,885 2, 20, 32, 35, 112, 310, standardization 476, 668, 701-703, 708716, 723, 854, 879, 880, 886,896,968,973 16, 133, 243, 249, 282, standardized 309, 610, 632, 894, 945, 952-954 standard-sample brack- 284, 288, 702, 703, 710, eting 711, 713-716, 723 starch 200 stars 229, 755 static batch method 5, 9 static charge 549, 845 static dilution 392

static electricity discharges static gas cell static mass discrimination effect static mass spectrometer static measurement static mode static multicollection static reduction static vacuum stationary phase stationary phase wax statistical normalization statistical outlier steam distillation steel SRB1098 standard steel SRM368 standard steel standard stepped combustion stepped heating stepped pyrolysis step-wise fluorination steroids (a)stigmatic focussing storage conditions stratosphere stratospheric 03 stretching motions stretching transitions strong acid resin strong base resin structural anhydrous feldspar styrene divinyl benzene sublimation sublimation rate subsequent conversion substantial isotope fractionation Sudan-1 sugar sugar fermentation sugar-beet suite of standards sulfanilamide

1041 770 699 1, 70, 259, 265, 267, 268, 348-350, 358, 362, 460, 468,817,818 697 363,652 847 9 528 155, 157, 160-163, 168171 163 309 954,955 189, 312, 315, 316, 318320,326,327 350 259 795,804 257, 258, 260, 266-271, 362,817 49,256-271,349,350 262,264 56,443,445,453 163,170 664,846,847 208,228 291,377-386,390,399 377, 378, 380, 382, 385, 386,389 766,767 780 639 489,617,641 265 17, 19, 125, 327 12, 13, 548, 1020 1023 575 359 338,339 4, 32, 116, 332, 883, 896, 897,913 117 112 279 599

1230 591 585, 559, 569, 573, 575, 585, 96Z 993, 994, 1012, 1018,1033 sulfate standard 967 568,619 sulfide ores 723,985 sulfide standard 125, 231, 232, 234, 236, sulfonic acid 32Z 553,554 450 sulfosalt minerals 552 sulfur chemistry sulfur compound oxida- 543, 559, 568, 570, 574, tion 595, 1018 sulfur cycle 538, 572, 596 sulfur hexafluoride 545 sulfur isotope methodo- 540 logy sulfur isotope standard 978, 981, 984 sulfur oxyfluoride 546 sulfur trioxide 967 sulfuric acid 32, 305, 312, 314, 317, 318, 342, 346, 347, 553, 607-609 sulfuric acid catalyst 609 sulfurous acid 553 sulphamic acid 237-255, 315, 318, 330 sulphamic acid (dry) 237 summary of experiments 345 suprasil window 433 surface contaminants 593, 688, 793 surface water 2, 32, 156, 203, 204, 306, 541, 678, 680, 950 surficial environment 180, 625 (a)symmetric isotopomer 380, 382, 387-389 synthetic 121, 599 synthetic Ag2S 461 synthetic air 1040, 1041 synthetic carbonates 241, 253-255, 1016, 1023 synthetic compound 881 synthetic data 633, 634 synthetic fluid 2 synthetic gases 947 synthetic glass 903 synthetic goethite 56, 57 synthetic material 120, 506, 880, 947, 1023 synthetic mineral 1001 synthetic mixed matrices 673 synthetic (isotope)mix- 526, 922-924, 926, 930, ture 932, 933, 940, 941, 947 synthetic molecular 806, 808 sieves synthetic nitrates 336 synthetic origin 103, 116, 118 synthetic polymer 469 sulfate desorption sulfate reduction

Subject Index synthetic precursors 118 synthetic process 235 synthetic product 175 synthetic pyrite 571 synthetic reagent 599 synthetic sample 119, 721 synthetic SiC 260 synthetic solids 805 synthetic solution 721 synthetic source 119 synthetic spectra 779 synthetic standard 123, 135, 599, 647 synthetic sulfanilamide 599 synthetic sulfate 599,601 synthetic sulfur com59Z 599 pounds 173 synthetic testosterone 891 synthetic wines 805,806,810,813 synthetic zeolite synthetically produced 883 substances 9-11, 26, 54, 55, 206, syringe 212-214, 221, 225, 571, 1013 213 syringe caps 207 syringe filter 213,214,21Z 219 syringe precipitation technique T table sugar tail contribution tandem filters tank air tap water tap water vapour taphonomy tar tarry products target chemistry TCA cycle technology Teflon Teflon bag Teflon cap Teflon ferrules Teflon filters Teflon holders Teflon liners Teflon matrix material Teflon membrane Teflon O-rings

332 33 562 276,279,300 34,90,454 95 483 4,813 5OO 690 200 928 323,345,469 19,21 207 295, 414, 415, 448, 454, 457 556,560-562 513,514,522 3,313 17 346,347 293

1231

Subject Index Teflon plug Teflon reaction tube Teflon seal Teflon sheets Teflon tubing Teflon-lined cap tektite temperate soil temperature drift temperature monitoring terrestrial terrestrial apatite standard terrestrial carbon terrestrial contamination terrestrial ecosystem terrestrial environment terrestrial standard terrestrial / land plants Tertiary Tesla coil testing environment tetramethyl urea tetrathionate thawing-pumpingfreezing cycles theoretical peak therapeutic drug thermal conversion thermal conversion EA (TC/EA)) thermal decomposition

10 446 54,454 543 316,550,583 19L 198 534,535 1013 1039 418,448 2,25,123,139 687 178 235,236,362 180,290 235 68G 720,724 180,200,273,1014 578 1022 747 111, 112, 117 553 55 270 122 1041 334

25G 33G 343, 398, 542547 thermal desorption 559 thermal diffusion 548,836,1024 thermal ion source 130,657,857,907 thermal ionization 123, 124, 128, 138, 147, 626, 629, 635, 712, 71Z 720, 747,789,793,857 thermalionizationmass 123, 12Z 129, 133, 135, spectrometry/TIMS 13G 138, 139, 141-152, 450, 540, 543, 549-551, 55Z 559, 561, 562, 614, 623-651, 669, 689, 692694, 700-702, 709, 710, 712, 713, 715-718, 724, 725, 739, 741, 795, 796, 798-800, 802, 820, 828, 839, 861, 867, 923, 1012, 1019 TIMS analyses 701, 796 TIMS data 699,700,712,802

TIMS filaments TIMS instrument TIMS measurement TIMS reference values TIMS results TIMS source TIMS studies TIMS values time-of-flight (TOF) Titan mission thermodynamic (TD) data TD fractionation TD framework TD function TD isotope effect TD principles TD properties TD relationship thermodynamically stable thermogenic methane thermogravimetric analysis thermometric applications thermometry thermostatic bath thin layer chromatography thiourea Thode reduction Ti-alloy autoclaves time window tin boat tin capsule tissue tissue water Toepler pump toluene toluol-ethanol tomato plants tooth / teeth tooth apatite tooth enamel topaz

623 626, 627, 645, 646, 650, 692 133, 138, 698, 709, 710, 713,716, 720,739 701 740 644,645 713 133,802 664,692,693,789,804 758 402 309 444 994 156,158 203 379, 401 603,994 566 291 49 484 1005 485 195 899,902 560,573 1026 168 328,339,341,347 318, 321, 328, 332, 341, 342, 346, 476, 493, 503, 593,598 181, 185, 186, 18Z 191, 199,594,595 474 5, 6, 23, 49, 430, 814, 1051 44,54,231,475,516,582, 584 512-517 1013 180-183, 186, 439, 482, 595,1029,1030 180,183,1029 407,43Z 483,495,496 467,530

1232 total impurity content total inorganic carbon (TIC) total organic load total system drift tourmaline trace contaminants traceability

Subject Index 940 600

331 300 441,688,904,1031 11,12,965,984 833, 883, 908, 909, 911915, 919, 920, 926, 928944,949 875,914,929-937 traceability chain 540 tracer compounds 160,161 tracer dilution 6,161,316,338,540,621, tracer study 649,852 168,171 trailing peak transbasalt 85 882 transfer standard transient sputtering loci 674 1021 transition metal catalyst 813 transition metal ionexchanged zeolite 312 treated waste 528,531 treatment of waste 500,539 tree ring cellulose 497 tree ring record 516 tree ring research 180,508 tree ring study 180, 512, 517, 838, 1033, tree rings 1036 triethanolamine (TEA) 558 199 trifluoroacetic anhydride 195 trimethylsilyl ether 424, 429, 447, 538, 539, troilite 599, 959, 967 200 trophic interactions 177, 180, 181, 191, 200, trophic level 201 181, 187 trophic relationship 191 trophic structure 191 trophic web analysis 186 tropical regions 43 tropical soil 291, 377-383 troposphere 183 tropospheric CO2 377, 378, 382, 383, 389 tropospheric 03 390, 394, 399 tropospheric N20 892, 910 TS limestone 820-822 tsetse flies 24, 59, 241, 612, 616, 949 tube cracker tube cracker (automated) 321 tuna oil 175

tunable IR laser tunable laser tunable radiation Tuttle-type vessel two stage reactor 1,1,1 trichloroethane

767 386, 388, 751, 758, 760, 761, 763, 765, 774-776, 780, 782 778 1025 175 615

U ultrasonic agitator / ion ultrasonic bath ultrasonic cleaning ultrasonic probe ultrasonic tank ultrasonication ultraviolet (UV) UV laser UV light UV photodissociation UV photolysis UV (ir)radiation UV spectrum UV/Vis spectrophtometer uncertainty propagation undisclosed contamination unidirectional conversion universal gas constant universe unpolarized light unstable reagent uranium furnace uranium hydride uranium oxide uranium reactor uranium reduction uranium turnings urban environment urea uric acid urine US Environmental Protection Agency

208, 209, 221, 558, 572 206-208, 516, 518, 560 432 41 197 197 76, 104, 377, 431, 433, 436, 437, 753 77, 433, 436, 458, 459, 466, 469, 495, 496, 715, 723, 725, 972 125, 384, 390, 466, 496, 496, 560, 565 377 377 236, 377, 639, 641, 646 466, 471 202 916 12 541 924 2 749, 754 340 5,59 6 5 5 5, 6, 13, 22-25, 29, 33-36, 66, 766 55 383 32, 191, 329, 344, 898, 899 191 2, 10, 19, 31, 88, 98, 102, 130, 174, 322, 595 305, 378

Subject Index

1233

V

W

Vacutainer

waste Br waste disposal waste line waste products water bath

16, 17, 206, 208-210, 215, 216, 226, 228 vacuum distillation 3, 4, 29, 54, 410, 412, 474 vacuum extraction line 74, 217, 350, 458, 493 vacuum fluorination ap-406, 409, 415, 416, 453, paratus / line 455, 456 valve mixing 33, 311 Van der Waals disper168 sion forces vanadium pentoxide/ 407, 544, 545, 593, 895, V205 973, 982 V205 combustion 988 V205 reagent 598 vapor phase equilibra- 205, 223-228 tion variable combustion con-972 ditions variable dispersion optics729 (non-)vascular plants 197, 198, 200, 516, 517, 518, 1014 Venus mission 758 vibrational energy 458, 766 virtual material 912 viscous flow 348-350, 836, 838, 853 Viton 76, 274, 275, 436, 463, 468 vitreous carbon 342, 502 Voigt profile 769 volatile fluoride 420 volatile hydrides 637 volatile meteoritic com- 233 pound volatile organic com153, 157, 159, 502, 615, pound 951 volcanic air samples 761 volcanic arcs 139 volcanic ash 140, 376 volcanic gases 349 volcanic glass 70 volcanic rock 85, 132, 133 voltage window 680 volume diffusion 261, 1004 VUV synchrotron radia- 757 tion Vycor (glass) 14, 1020 Vycor flask 1021 Vycor tube/vessel 10, 12, 14, 330, 333, 336, 1020, 1021

411 528 451, 456, 543 418, 419, 442, 455-457 20, 22, 25, 44, 207, 208, 513, 514, 529-531, 578, 615, 1014, 1015 27, 326 water chemistry 680 water contaminant 36, 68, 74, 496 water desorption 53, 74, 232, 475 water extraction 4, 6, 10, 13, 15, 66, 76, water reduction 850 2, 3, 28, 36, 52, 88-102, water vapor 165, 189, 2391 255, 275, 281, 291, 293, 296, 300, 323, 330, 340, 341, 525, 559, 794, 1014, 1020 619 water-soluable salt water-soluable solvent 4 wavelength modulation 773 spectroscopy (WMS) wax paper funnel 216 waxes 163, 195, 475, 517 weathered basalt 139 weathering 38, 39, 42, 140, 141, 445, 1028, 1032 weathering environment 1028 wet air 281 wet chemistry 482-489, 525 wetland 209, 291, 559 wetland runoff 210,212 whalebones 198 whales 181 wheat plants 593, 1014 white-mica 348, 351-357 whole ecosystem 192 whole rock matrices 973 whole tissue 201 Wien filter 657, 658 wildlife 594, 625 window materials 463 wine 2, 19, 112, 116, 117 Wilson-Wells method 771 wood 183, 188, 197, 457, 502, 507-521, 1033, 1035 wood density 510 woody plant tissue 201 woody plants 516 working environment 415, 863, 1017 working lifetime 199

1234

working standard

Subject Index

21, 28, 32-35, 117, 118, 311, 396, 398, 534, 879, 882, 908, 929-931, 978

111, 113, 392, 393, 543, 855, 916, 915,

X X-ray fluorescence xylem

46 473, 474

Y yeast cells

101

Z 53, 125, 189, 314, 318, 342, 379, 805-819, 1020, 1032 817 zeolite crystal 780 zero background 32, 33, 310, 311, 952 zero-enrichment zinc artifact 13 12 zinc hydroxide 11, 13 zinc oxidation zinc reactor 5 11, 14 zinc reagent 8-10, 22, 23, 29, 31-36, zinc reduction 66, 76, 889 zinc reduction (dynamic) 8 9,10 zinc shot 409, 441, 466, 535, 681zircon 683,787 463,467,495 ZnSe window 996,997 zoisite 697 zoom lens dispersion 187,200,201 zooplankton zeolite

Handbook of Stable Isotope Analytical Techniques

Handbook of Stable Isotope Analytical Techniques, Volume I (Handbook of Stable Isotope Analytical Techniques)

Handbook of analytical techniques

Stable Isotope Geochemistry

Principles of Stable Isotope Distribution

Principles of Stable Isotope Distribution

Stable Isotope Ecology

Stable Isotope Ecology

Stable Isotope Geochemistry

Stable Isotope Geochemistry

Handbook of Instrumental Techniques for Analytical Chemistry

Handbook of instrumental techniques for analytical chemistry

Handbook Of Instrumental Techniques For Analytical Chemistry

Handbook of Analytical Techniques, 2 Volume Set

Handbook of Environmental Isotope Geochemistry

Stable Isotope Geochemistry (Reviews in Mineralogy and Geochemistry, Volume 43)

Dean's Analytical Chemistry Handbook

Dean's analytical chemistry handbook

Analytical Techniques in Combinatorial Chemistry

Analytical Techniques In DNA Sequencing

Analytical Techniques In Combinatorial Chemistry

Analytical Techniques In DNA Sequencing

Analytical Techniques In DNA Sequencing

Analytical techniques for biopharmaceutical development

Analytical Techniques in Combinatorial Chemistry

Analytical Techniques in Materials Conservation

Dean's Analytical Chemistry Handbook

Analytical instrumentation handbook

Dean's Analytical Chemistry Handbook

Analytical Techniques for Biopharmaceutical Development

Handbook of Stable Isotope Analytical Techniques