Noble Gases in Geochemistry and Cosmochemistry 0939950596

Table of contents :
Page 1
Titles
Table of Contents
Noble Gases in Geochemistry and Cosmochemistry
1 An Overview of Noble Gas Geochemistry and Cosmochemistry
Donald Porcelli, Chris J. Ballentine, Rainer Wieler
2
Noble Gases in the Solar System
ix
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Titles
3
Noble Gases in Meteorites - Trapped Components
4 Noble Gases in the Moon and Meteorites:
Radiogenic Components and Early Volatile Chronologies
x
Page 3
Titles
5
Cosmic-Ray-Produced Noble Gases in Meteorites
6
Martian Noble Gases
7
Origin of Noble Gases in the Terrestrial Planets
xi
Page 4
Titles
8
Noble Gas Isotope Geochemistry of
Characterization of Mantle Source Reservoirs
xii
Page 5
Titles
9 Noble Gases and Volatile Recycling at Subduction Zones
David R. Hilton, Tobias P. Fischer, Bernard Marty
10
Storage and Transport of Noble Gases
Tibor J. Dunai, Donald Porcelli
Page 6
Titles
11 Models for the Distribution of Terrestrial Noble Gases
and Evolution of the Atmosphere
D. Porcelli, C.J. Ballentine
xiv
Page 7
Titles
12
Production, Release and Transport of
Chris J. Ballentine, Pete G. Burnard
13 Tracing Fluid Origin, Transport and Interaction in the Crust
xv
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Titles
14
Noble Gases in Lakes and Ground Waters
Rolf Kipfer, Werner Aeschbach-Hertig, Frank Peeters, Marvin Stute
15
Noble Gases in Ocean Waters and Sediments
xvi
Page 9
Titles
16 Cosmic-Ray-Produced Noble Gases in
Terrestrial Rocks: Dating Tools for Surface Processes
Samuel Niedermann
xvii
Page 10
Titles
17
K-Ar and Ar-Ar Dating
18 (U- Thj/He Dating: Techniques, Calibrations, and Applications
xv iii
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Reviews in Mineralogy and Geochemistry Volume 47 Noble Gases in Geochemistry and Cosmochemistry FOREWORD Noble Gases in Geochemistry and Cosmochemistry is the 47th volume in a series that was initiated by the Mineralogical Society of America (MSA) in 1974 under the title “Short Course Notes.” Within a few years the series was re-named Reviews in Mineralogy, which from the mid-1980s was added to the 4000 or so scientific periodicals tracked by the Institute for Scientific Information in their Journal Citation Reports. Beginning in 2000 with Volume 39, the series was given a new title, Reviews in Mineralogy and Geochemistry (RiMG), in response to the partnership that was formed at that time between MSA and the Geochemical Society—GS in the mini-logo that now appears on the spine of all RiMG volumes: G MSA

The two societies manage their own short courses (if any) and edit their own volumes; Paul Ribbe is editor for MSA and Jodi Rosso for GS. For this volume, both editors were involved. There was no short course, although an attempt was undertaken to get the volume printed in time for the V.M. Goldschmidt conference in Davos, Switzerland (mid-August 2002) at which there was a major symposium on noble gases. Noble Gases in Geochemistry and Cosmochemistry was edited by D. Porcelli, C. J. Ballentine, and R. Wieler, all of whom were at the Institute for Isotope Geology and Mineral Resources, ETH—Zürich when this project was begun [their present affiliations are in the heading of Chapter 1, page 1]. If all proceeds according to plan, this will be the first RiMG volume to be published in electronic format as well as in paper copy. Visit the MSA website, www.minsocam.org, for breaking news. Jodi Rosso & Paul Ribbe West Richland, Washington Blacksburg, Virginia July 18, 2002 PREFACE

Noble Gases – Noble Science The scientific discoveries that have been made with noble gas geochemistry are of such a profound and fundamental nature that earth science textbooks should be full of examples. Surprisingly, this really is not so. The “first discoveries” include presolar components in our solar system, extinct radionuclides, primordial volatiles in the Earth, the degassing history of Mars, secular changes in the solar wind, reliable present day mantle degassing fluxes, the fluxes of extraterrestrial material to Earth, groundwater paleotemperatures and the ages of the oldest landscapes on Earth. Noble gas geochemistry has scored so many such “firsts” or “home runs” that it should permeate a lot of earth science thinking and teaching. Yet rather surprisingly it does not. Noble gas geochemistry also is a broader and more versatile field than almost any other area of geochemistry. It pervades cosmochemistry, Earth sciences, ocean sciences, climate studies and environmental sciences. Yet most modern Earth, planetary and environmental science departments do not consider noble gas geochemistry to be at the top of their list in

terms of hiring priorities these days. Furthermore, with the exception of Ar geochronologists, noble gas geochemists are a surprisingly rare breed. Why is the above the case? Perhaps the reasons lie in the nature of the field itself. First, although noble gas geochemists work on big problems, the context of their data is often woefully under-constrained so that it becomes hard to make progress beyond the first order fundamental discoveries. Noble gas data are often difficult to interpret. Although some concepts are straightforward and striking in their immediate implications (e.g. mantle 3He in the oceans), others are to this day shrouded in lack of clarity. The simple reason for this is that in many situations it is only the noble gases that offer any real insights at all and the context of other constraints simply does not exist. Some examples of the big issues being addressed by noble gases are as follows and I have deliberately posed these as major unresolved questions that only exist because noble gas geochemistry has opened windows through which to view large-scale issues and processes that otherwise would be obscure. (1) Is the presolar noble gas component present in a tiny fraction of submicroscopic meteoritic C or is it ubiquitously distributed? (2) How did solar noble gases get incorporated into the Earth? (3) How did solar noble gases survive the protracted accretion of the Earth via giant impacts? (4) What is the origin of the noble gas pattern in the Earth’s atmosphere? (5) Why are the Earth and Mars almost opposites in terms of the relative isotopic differences between atmosphere and mantle? (6) What is the present source of Earth’s primordial helium? Can we ignore the core? (7) What is the 20Ne/22Ne of the mantle, how was it acquired and why is it different from the atmosphere? (8) How does one reconcile the strong fractionation in terrestrial Xe with data for other noble gases? (9) How much radiogenic 40Ar should the Earth have? How well do we know K/U? (10) Are the light isotopes of Xe the same in the mantle and the atmosphere? If not, why not? (11) How are noble gases transported through the creeping solid earth? (12) How does one explain the heat – helium paradox? (13) How incompatible are the noble gases during melting? (14) How are atmospheric components incorporated into volcanic samples? (15) How are the excess air components incorporated into groundwater? (16) Why are continental noble gas paleotemperature records offset from oceanic temperature records? Noble gas data tell us that the Earth and solar system represent very complex environments. When we make our simple first order conclusions and models we are only at the tip of the iceberg of discoveries that are needed to arrive at a thorough understanding of the behavior of volatiles in the solar system. Who wants to hear that things are complicated? Who wants to hire in a field that will involve decades of data acquisition and analysis in order to sort out the solar system? Sadly, too few these days. This is the stuff of deep scientific giants and bold, technically difficult long-term research programs. It is not for those who prefer superficiality and quick, glamorous, slick answers. Noble gas geochemists work in many areas where progress is slow and difficult even though the issues are huge. This probably plays a part in the limited marketability of noble gas geochemistry to the nonspecialist. Second, noble gases is a technically difficult subject. That is, noble gas geochemists need to be adept at technique development and this has to include skills acquired through innovation in the lab. Nobody can learn this stuff merely with a book or practical guide. Reading “Zen and the Art of Motorcycle Maintenance” (by Robert Pirsig) would give you a clearer picture. This magnificent MSA-GS volume is going to be enormously useful but on its own it won’t make anybody into a noble gas geochemist. Although the mass spectrometry principles are not complex, the tricks involved in getting better data are often self taught or passed on by working with individuals who themselves are pushing the boundaries further. Furthermore, much of the exciting new science is linked with technical developments that

allow us to move beyond the current measurement capabilities. Be they better crushing devices, laser resonance time of flight, multiple collection or compressor sources - the technical issues are central to progress. Lastly, noble gas geochemists need a broad range of other skills in order to make progress. They have to be good at mass spectrometry as already stated. However, nowadays they also need to be able to understand fields as different as mantle geochemistry, stellar evolution, cosmochemistry, crustal fluids, oceanography and glaciology. They are kind of “Renaissance” individuals. Therefore, if you are thinking broadly about hiring scientists who love science and stand a good chance of making a major difference to our understanding of the solar system, earth and its environment – I would recommend you hire a really good noble gas geochemist. However, the results may take a while. If you want somebody who will crank out papers at high speed and quickly increase the publication numbers of your department then you may need to think about somebody else. The two are not mutually exclusive but think hard about what is really important. This volume is nothing short of a fantastic accomplishment. This is true in many ways. The editors are to be commended for working so efficiently to pull it off. However, there are others who also should be honored in this preface. Of course, the authors have, between them, written what is bound to be an essential handbook for all those who are interested in geochemistry. But they are building on the work of some remarkably clever and insightful scientists who had the original vision that has led to the big discoveries that I have already talked about at the start of this little section. In this case we are talking of amazing people like: Al Nier, who developed noble gas mass spectrometry and K-Ar geochronology; John Reynolds, who first discovered 129Xe excesses in the Richardton meteorite, then wrote that astonishing paper entitled “The Age of the Elements” and had the clever of idea of doing stepwise degassing of irradiated samples to establish that the radiogenic xenon was derived from iodine sites; Igor Tolstikhin and Boris Mamyrin, who first discovered primordial 3He; Brian Clarke, who developed the Clarke source and first showed 3He variations in ocean waters; Harmon Craig, who first developed our understanding of primordial 3He being released from the Earth into the oceans; Claude Allègre, who first measured Xe isotopic anomalies in MORB; Craig Merrihue and Grenville Turner, who first developed 40Ar-39Ar dating; Peter Signer, who first defined the Planetary noble gas component; Keith O’Nions and Ron Oxburgh, who formulated the heat-helium paradox; and Minoru Ozima and Frank Podosek, who established the key reference work for all those interested in noble gas geochemistry. Many more could be named but it is time for me to stop and to say I am looking forward to seeing this book launched at the Goldschmidt Conference in Davos where we should be able to toast some of the above and their remarkable accomplishments. I encourage you to jump into these pages of deep science and “learn loads” about the truly enormous array of interesting research and discovery that has followed from noble gas geochemistry. Alex N. Halliday Zürich, May 2002 .

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An Overview of Noble Gas Geochemistry and Cosmochemistry Donald Porcelli, Chris J. Ballentine, Rainer Wieler

INTRODUCTION NOBLE GAS MASS SPECTROMETRY THE ATMOSPHERIC STANDARD PRODUCTION OF NOBLE GAS ISOTOPES BEHA VIOR OF THE NOBLE GASES Water, natural gas, and oil partitioning Silicate melt solubilities Crystal-melt partitioning Iron-silicate partitioning Adsorption Diffusion Atmosphere losses NOBLE GASES IN COSMOCHEMICAL AND GEOCHEMICAL STUDIES Noble gases in cosmochemistry Noble gases in the mantle Noble gases in surface reservoirs Noble gases and geochronology ENDNOTE ACKNOWLEDGMENTS REFERENCES APPENDIX I

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1 2 2 .4 5 6 6 6 6 6 7 7 7 8 10 11 13 14 14 14 18

Noble Gases in the Solar System Rainer Wieler

INTRODUCTION ANALYTICAL TECHNIQUES THE SUN Solar noble gas abundances Helium in the Sun Noble gases in the solar corpuscular radiation THE GIANT PLANETS The He abundances The Ne-Xe abundances in Jupiter Isotopic ratios in Jupiter VENUS THE MOON Trapped solar (and other) noble gases in the lunar (and asteroidal) regoliths Nitrogen in lunar samples Cosmogenic noble gases in lunar samples THE EXOSPHERES OF THE MOON AND MERCURY Sources and sinks COMETS INTERPLANETARY DUST PARTICLES AND MICROMETEORITES

ix

21 21 23 23 26 27 36 36 39 40 42 43 .43 52 53 55 57 58 59

ELEMENTARY PARTICLES IN INTERPLANETARY Galactic cosmic rays Interstellar and interplanetary pick-up ions Anomalous cosmic rays ACKNOWLEDGMENTS REFERENCES

3

SPACE

61 61 62 63 63 63

Noble Gases in Meteorites - Trapped Components Ulrich Ott

OVERVIEW AND HISTORY ELEMENTAL ABUNDANCE PATTERNS ISOTOPIC COMPOSITIONS Helium Neon Argon Krypton Xenon ORIGINS AND HISTORY Q(PI )-gases Gases in presolar silicon carbide and graphite Gases trapped in presolar diamond (HL, P3 and P6 components) Ureilite gases Subsolar and sub-Q gases Sitings Other (lesser) components RELATIONS Relationship between Q and solar gases A more primitive component (in Xe and only in Xe?) CONCLUDING REMARKS ACKNOWLEDGMENTS REFEREN CES

4

71 73 74 75 76 77 78 80 82 83 84 87 89 89 90 90 91 91 94 96 96 96

Noble Gases in the Moon and Meteorites: Radiogenic Components and Early Volatile Chronologies Timothy D. Swindle

INTRODUCTION LONG-LIVED NUCLIDES: CHRONOLOGY OF SOLAR SYSTEM EVOLUTION Solar system impact history Extinct radionuclides: Chronology of solar system formation Iodine-xenon 244Plutonium Other fissioning nuclides? Other radiogenic noble gases ACKNOWLEDGMENTS REFERENCES

x

101 102 103 l 08 109 117 120 120 121 121

5

Cosmic-Ray-Produced Noble Gases in Meteorites Rainer Wieler

INTRODUCTION THE PRODUCTION OF COSMOGENIC NUCLIDES IN METEORITES Fundamentals Production systematics Cosmogenic noble gases produced by capture of low-energy neutrons Isotopic abundances of cosmogenic noble gases EXPOSURE AGE DISTRIBUTIONS OF METEORITES Undifferentiated meteorites Differentiated meteorites Exposure ages and dynamical models of meteorite delivery COMPLEX EXPOSURE HISTORIES THE COSMIC RA Y FLUX IN TIME TERRESTRIAL AGES COSMOGENIC NOBLE GASES PRODUCED BY SOLAR COSMIC RAYS (SCR) ACKNOWLEDGMENTS REFERENCES

6

125 126 126 128 139 141 144 146 149 153 155 159 161 162 162 163

Martian Noble Gases Timothy D. Swindle

STUDYING MARS AS ANOTHER PLANET DISCOVERY OF MARTIAN METEORITES MARTIAN ATMOSPHERE OTHER MARTIAN NOBLE GAS RESERVOIRS Martian interior (Chassigny) Other reservoirs I: Nakhlite (and ALH84001) Xe Other reservoirs II: Shergottite Ar Paleoatmosphere MARTIAN CHRONOLOGY 40 Ar- 39 Ar dating Cosmic-ray exposure ages ACKNOWLEDGMENTS REFERENCES

7

171 171 175 180 180 181 182 183 184 184 185 187 187

Origin of Noble Gases in the Terrestrial Planets Robert O. Pepin, Donald Porcelli

INTRODUCTION CHARACTERISTICS OF PLANETARY NOBLE GASES Planetary noble gas abundance patterns Terrestrial noble gases Noble gases on Venus Noble Gases on Mars ACQUISITION OF PLANETARY NOBLE GASES Early post-nebular solar wind Adsorption on pre-planetary grains or protoplanetary bodies Gravitational capture Accretion of volatile-rich planetesimals

xi

191 193 193 196 205 206 208 209 210 210 2 I3

LOSSES AND MODIFICATIONS Losses during accretion Sputtering Hydrodynam ic escape Primordial xenon OUTSTANDING ISSUES ACKNOWLEDGMENTS REFERENCES

8

OF PLANETARY NOBLE GASES

216 216 217 2 18 231 235 239 239

Noble Gas Isotope Geochemistry of Mid-Ocean Ridge and Ocean Island Basalts: Characterization of Mantle Source Reservoirs David W. Graham

INTRODUCTION BACKGROUND Noble gas chemical behavior Mantle structure and noble gases HELIUM Significance Radiogenic production Mid-ocean ridge basalts Ocean island basalts NEON Significance Nucleogenic production Mid-ocean ridge basalts Ocean island basalts ARGON Significance Radiogenic production Atmospheric contamination Mid-ocean ridge basalts Ocean island basalts KRYPTON XENON Significance Radiogenic production Mid-ocean ridge basalts Ocean island basalts COUPLED RADIOGENIC/NUCLEOGENIC PRODUCTION Neon-helium systematics Argon-helium systematics Xenon-neon-helium systematics MANTLE ABUNDANCE PATTERNS OF NOBLE GASES PRINCIPAL OBSERVATIONS SUMMARY ACKNOWLEDGMENTS REFERENCES

xii

247 249 249 252 254 254 255 255 265 279 279 279 281 283 284 284 284 284 285 288 289 289 289 290 291 292 293 293 297 .300 301 302 303 305 305

9

Noble Gases and Volatile Recycling at Subduction Zones David R. Hilton, Tobias P. Fischer, Bernard Marty

INTRODUCTION SAMPLING FOR NOBLE GASES Volcanic and geothermal fluids Hot spring waters and groundwaters Geothermal wells Natural gases Mafic phenocrysts and xenoliths Submarine glasses Problematic issues of noble gas analysis NOBLE GAS SYSTEMATICS OF ARC-RELATED VOLCANISM Helium isotope systematics of arc-related volcanism Neon and argon isotope systematics of arc-related volcanism Krypton and xenon isotope systematics of arc-related volcanism Relative noble gas abundance systematics of arc-related volcanism NOBLE GASES IN BACK-ARC BASINS Helium isotopes in back-arc basins Neon isotopes in back-arc basins Argon, krypton and xenon isotopes in back-arc basins RECYCLING OF VOLA TILES AT SUBDUCTION ZONES: A MASS BALANCE APPROACH The global volatile output at arc volcanoes: The 3He approach Volatile output at individual arcs-S02 and the power law distribution Using helium to resolve volatile provenance Problematic issues regarding estimates of volatile output at arcs Volatile output at the back-arc and fore-arc regions The volatile input via the trench Volatile mass balance at subduction zones Global implications of volatile recycling at subduction zones SUMMARY AND FUTURE WORK ACKNOWLEDGMENTS REFERENCES

10

3 19 320 320 321 322 322 322 323 323 324 325 330 333 333 334 334 337 338 340 340 341 349 35 I 3 52 353 355 358 361 362 362

Storage and Transport of Noble Gases in the Subcontinental Lithosphere Tibor J. Dunai, Donald Porcelli

INTRODUCTION SUBCONTINENTAL MANTLE AS GEOCHEMICAL RESERVOIR Thickness of the lithosphere Composition of the lithosphere Age of the lithosphere Possible xenolith noble gas components NOBLE GASES IN ULTRAMAFIC XENOLITHS AND PHENOCRYSTS Fluid inclusions as hosts for noble gases Helium isotopic variations He-Sr relationships He-C relationships The heavier noble gases Regional studies ofthe sources of xenolith mantle rare gases The mantle sources of xenolith He Introducing He into the lithosphere XU1

371 .372 372 373 374 374 377 .377 .381 382 .384 386 .388 394 396

DIAMONDS Ancient He Ancient Xe Mantle sources of diamond rare gases MANTLE VOLA TILES IN THE CONTINENTAL CRUST CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

11

397 398 .399 .399 .400 400 401 40 I

Models for the Distribution of Terrestrial Noble Gases and Evolution of the Atmosphere D. Porcelli, C.J. Ballentine

INTRODUCTION RADIOGENIC COMPONENTS IN THE BULK EARTH AND ATMOSPHERE Radiogenic He and nucleogenic Ne Radiogenic Ar Radiogenic and fissiogenic Xe MANTLE NOBLE GAS CHARACTERISTICS Helium isotopic compositions in the mantle Neon isotopic compositions in the mantle Argon isotopic compositions in the mantle Xe isotopes and a nonresidual upper mantle Mantle noble gas relative abundances Noble gas fluxes and mantle concentrations Undepleted mantle Coupled degassing of noble gases TRACE ELEMENT ISOTOPE CONSTRAINTS ON MANTLE RESERVOIRS Relationship between noble gases and radiogenic isotopes Mass and character of the depleted mantle Timing of mantle depletion and recycling PHYSICAL CONSTRAINTS ON MANTLE MODELS Mantle reservoirs Geophysical evidence for the scale of mantle convection Early Earth history NOBLE GAS MANTLE MODELS Single reservoir degassing Limited interaction box models Steady state box models Break-up of a previously layered mantle The lower boundary layer reservoir The upper boundary layer reservoir Deeper or 'abyssal' layering Heterogeneities preserved within the convecting mantle Depleted, high He/U mantle Storage of noble gases in the core Subduction of meteoritic He CONCLUSIONS Overall assessment of the models Some persistent misconceptions Important parameters that are still unknown ACKNOWLEDGMENTS REFERENCES

xiv

.411 .412 .412 414 .415 .419 .419 .420 .423 .427 .428 .430 434 .435 .436 .436 .436 .438 .438 438 .439 443 .446 446 .448 452 .457 457 .459 .459 .460 462 .463 465 466 .466 467 .468 469 469

12

Production, Release and Transport of Noble Gases in the Continental Crust Chris J. Ballentine, Pete G. Burnard

INTRODUCTION RADIOGENIC, NUCLEOGENIC, AND FISSIOGENIC NOBLE GASES The subsurface neutron flux and reaction probability Helium Neon Argon Krypton and xenon Cosmogenic noble gas production Interplanetary dust accumulation Production rates in continental crust today and over the history of the Earth RELEASE OF NOBLE GASES FROM MINERALS IN THE CRUST Recoil loss Diffusive loss from minerals Mineral breakdown/diagenesis/metamorphism/alteration TRANSPORT OF NOBLE GASES FROM THE DEEP CRUST TO SHALLOW LEVEL SySTEMS Diffusion: A viable transport mechanism to degas the continental _crust? Differential release and transport of helium and argon Accumulation and release: Fluid flow in the crust.. HEAT AND HELIUM The relationship between 4He and heat.. Transport of mantle heat and helium through the crust Transport of crustal heat and 4He MAGMATIC NOBLE GASES IN THE CRUST Tectonic control on magmatic fluid location Fossil magmatic gases ACKNOWLEDGMENTS REFERENCES

13

481 .482 .482 487 490 495 498 501 503 5 04 506 5 06 509 512 514 515 517 5 19 521 521 521 524 525 525 529 529 529

Tracing Fluid Origin, Transport and Interaction in the Crust Chris J. Ballentine, Ray Burgess, Bernard Marty

INTRODUCTION PHYSICAL CHEMISTRY OF NOBLE GASES IN CRUSTAL FLUIDS Henry's law and the assumption of ideality Non-ideality in the gas phase Non-ideality in the fluid phase Noble gas solubility in water and oil PHASE EQUILIBRIUM AND FRACTIONATION OF NOBLE GASES Liquid-gas phase partitioning of noble gases Liquid-liquid phase partitioning of noble gases Relative fractionation Rayleigh fractionation Re-solution and effervescence Multiple subsurface fluid phases Diffusion or kinetic fractionation RESOLVING DIFFERENT NOBLE GAS COMPONENTS IN CRUSTAL FLUIDS NOBLE GASES IN HYDROCARBON GAS AND OIL RESERVOIRS Identifying and quantifying groundwater/gas/oil interaction Case studies NOBLE GASES IN ANCIENT GROUNDW ATERS AND CRUSTAL DEGASSING Sources of He isotopes in groundwaters Deep aquifers and the crustal He flux

xv

539 540 540 541 544 546 548 548 549 550 551 555 558 5 59 5 62 565 5 67 569 5 81 582 584

Advective versus diffusive transfer of noble gases in basins Tentative synthesis MAGMATIC FLUIDS IN THE CRUST Mantle degassing in the continental crust The noble gas imprint.. NOBLE GASES IN MINERAL DEPOSITS AND HYDROTHERMAL FLUIDS Post-entrapment modification of He and Ar isotopes Noble gas mixtures in mineralizing fluids '" Description and analysis of multi-component noble gas mixtures in ore fluids Further developments ACKNOWLEDGMENTS REFERENCES

14

Noble Gases in Lakes and Ground Waters Rolf Kipfer, Werner Aeschbach-Hertig,

Frank Peeters, Marvin Stute

INTRODUCTION ANALYTICAL TECHNIQUES NOBLE GAS COMPONENTS IN WATER Atmospheric noble gases 1: Solubility equilibrium Atmospheric noble gases 2: Excess Air Radiogenic He (and Ar) Terrigenic He Radioactive noble gas isotopes Man made pseudo-conservative trace gases Data evaluation and interpretation Conceptual models for noble gases in water Separation of the components Interpretation APPLICATIONS IN LAKES Mixing and the distribution of dissolved substances in lakes Distribution of noble gases in lakes Application of3H-3He dating in lakes Quantification of vertical exchange rates and vertical turbulent diffusivities He flux from the continental crust and oxygen depletion Noble gases from the Earth's mantle Applications in ground water Dating of young ground waters Dating of old ground waters Noble gas recharge temperatures Excess air NOBLE GASES IN ICE Gravitational separation Thermal diffusion Helium isotopes REFERENCES

15

590 592 595 595 598 600 601 60 I 605 608 60 8

615 617 619 619 624 629 629 630 633 633 634 641 647 650 651 653 654 658 661 664 666 667 676 679 683 687 687 687 689 690

Noble Gases in Ocean Waters and Sediments Peter Schlosser, Gisela Winckler

INTRODUCTION TRACING OCEAN CIRCULATION USING 3HE AND TRITIUM Tritium/3He method Exemplary results from tritium/3He studies MANTLE 3HE General background Geochemical background 3He plumes in the ocean ;

xvi

701 702 702 706 709 709 709 710

THE FLUX OF 3HE AND 4HE FROM THE SEAFLOOR 4He 3H e EXTRATERRESTRIAL 3HEIN DEEP-SEA SEDlMENTS Del ivery of extraterrestrial 3He to the ocean sediments Applications of the lOP-derived 3He method NOBLE GASES IN DEEP-SEA BRINES The Red Sea The Eastern Mediterranean AIR/SEA GAS EXCHANGE STUDIED BY DUAL TRACER RELEASE EXPERIMENTS Background SF6lHe Method Results from oceanic dual gas tracer releases PERSPECTIVES REFERENCES

16

711 711 71 3 715 715 717 720 720 721 721 721 723 723 723 725

Cosmic-Ray-Produced Noble Gases in Terrestrial Rocks: Dating Tools for Surface Processes Samuel Niedermann

INTRODUCTION THE FUNDAMENTALS OF SURFACE EXPOSURE DATING Production mechanisms of cosmogenic nuclides in terrestrial rocks Temporal variation of cosmogenic nuclide production Cosmogenic nuclide production at depth and on eroding surfaces Derivation of exposure ages and erosion rates Error considerations PRODUCTION RATES OF COSMOGENIC NUCLIDES Experimental determinations of production rates Production rates obtained by model calculations EXPERIMENTAL ISSUES Retentivity of minerals for cosmogenic He and Ne Discrimination of cosmogenic against trapped, radiogenic, and nucleogenic components Cosmogenic noble gases as a nuisance Cosmogenic noble gases versus radionuclides APPLICATION EXAMPLES Dating of lava flows Glacier movement and ice sheet evolution Rates of erosion and soil accumulation Rates of tectonic uplift Earthquakes and landslides Further applications Future prospects ACKNOWLEDGMENTS REFERENCES

xvii

731 732 732 740 748 750 754 755 755 761 762 762 763 769 770 771 771 772 773 775 775 776 776 777 777

17

K-Ar and Ar-Ar Dating Simon Kelley

INTRODUCTION - A BIT OF HISTORY THE K-AR AND AR-AR DATING METHODS Introduction Assumptions CALCULATING K-AR AND AR-AR AGES The values of constants and estimation of errors ARGON DIFFUSION AND SOLUBILITY Argon diffusion (and its use to determine thermal histories) Argon solubility (and the causes of extraneous argon) APPLICATIONS Thermochronology Dating young volcanic eruptions High-precision ages on altered basalts Dating low-temperature processes Unique samples Acknowledgments REFERENCES

785 786 786 787 788 792 795 795 798 806 806 809 809 809 810 810 810

18 (U- Thj/He Dating: Techniques, Calibrations, and Applications Kenneth A. Farley INTRODUCTION Rationale TECHNICAL ASPECTS He ingrowth DIFFUSION BEHA VIOR Apatite Hematite Titanite Zircon Garnet The a-emission correction Analytical procedures, accuracy, precision and mineral standards INTERPRETATION OF HE AGES AND EXAMPLES He cooling ages Some case studies FUTURE PROSPECTS ACKNOWLEDGMENTS REFERENCES APPENDIX: ANALYTICAL TECHNIQUES

xv iii

819 819 820 820 821 822 825 825 826 827 82 7 832 835 835 837 839 840 841 844

1

An Overview of Noble Gas Geochemistry and Cosmochemistry D. Porcelli1, C. J. Ballentine2, and R. Wieler Institute for Isotope Geology and Mineral Resources ETH Zürich, Sonneggstrasse 5 8092 Zürich, Switzerland [email protected] 1

2

Now at: Dept. of Earth Sciences, University of Oxford Parks Road, Oxford OX1 3PR, United Kingdom

Now at: Dept. of Earth Sciences, University of Manchester, Oxford Road, Manchester M13 9PL, United Kingdom

INTRODUCTION A wealth of fundamental information regarding the Earth and solar system is based upon observations of the highly volatile elements He, Ne, Ar, Kr, and Xe. At first, this may seem surprising, considering that these elements are generally thought to reside almost entirely in the atmosphere, and so are considered strongly ‘atmophile.’ However, increasingly sophisticated analytical techniques have provided the means for precisely measuring their abundances in a wide range of geological and cosmochemical materials. Fittingly, these elements are known collectively as the rare gases, reflecting their general scarcity in geological materials. It is this feature that continues to provide challenges for analysts. These elements are also the noble gases, in tribute to their disdain for engaging in chemical consort with other species. Such behavior has been responsible for the early difficulties in their detection, and facilitates their continuing migration to the atmosphere. However, others refer to these as ‘the inert gases,’ which seems to imply that their behavior is dictated by a lack of interest in chemical reaction, a deficiency in chemical drive. Overall, the choice of appellation depends upon whether scarcity, nobility, or inertness is considered the most important characteristic. Regardless of their motivations, these noble gases can be profitably considered together, because physical and chemical properties vary systematically with atomic weight. However, much of the utility of noble gases is based on the widespread variations in their isotopic compositions. This is related to their overall depletion, which has made these elements vulnerable to isotopic modification from nuclear processes involving relatively more abundant parent elements. The wide applicability of noble gas systematics is due to the range of such processes. In cosmochemistry, fundamental contributions have been made to understanding the sources and distributions of volatiles throughout the solar system, to identifying the preservation of nucleosynthetic anomalies in meteorites, and to defining early solar system chronologies. Studies of the distribution of noble gases within the Earth are a critical component in studies of mantle geochemistry and the formation history of the atmosphere. Within the crust, noble gases have been key components in studies of crustal evolution, of flow patterns in hydrological systems and ocean basins, and in a range of dating techniques. The present volume contains a series of focused reviews of noble gases across the solar system, in the Earth’s mantle, in the crust, and as applied in geochronology. These are written by researchers closely associated with each field of research. Other books that are of interest in complementing these works include the earlier review of He Isotopes in Nature (Mamyrin and Tolstikhin 1984) and the recently published second edition of 1529-6466/02/0047-0001$05.00

DOI:10.2138/rmg.2002.47.1

2

Porcelli, Ballentine & Wieler

Noble Gas Geochemistry (Ozima and Podosek 2001), which provide different perspectives, as well as detailed discussions of some background information that is not covered in the same detail in the more application-oriented papers of the present volume. NOBLE GAS MASS SPECTROMETRY The progress of noble gas geochemistry and cosmochemistry has been paced by the rate of developments in mass spectrometry. This has been driven by the need to attain increasingly greater precision to distinguish the often subtle variations in isotopic compositions, higher sensitivity to measure the low abundances found in many materials, and lower blanks to remove interference from atmospheric gases. The earliest measurements and ages, including U-He ages, did not involve the detection of isotopes, relying only on the total abundances of the elements present. Different isotopes of a single element were first detected for a noble gas, Ne (Thomson 1914), and following this Dempster (1918) and Aston (1919) produced the first focusing machines for relative isotope abundance measurements. The Nier design mass spectrometer (Nier 1947), including what has become the standard design for the electron bombardment source, provided precise compositions of the atmosphere. Reynolds (1956) introduced the highsensitivity static mass spectrometer design that is essential to most modern noble gas studies. The high resolution required for precise 3He/4He measurements was achieved using a magnetic resonance mass spectrometer (Mamyrin et al. 1969; see Mamyrin and Tolstikhin 1984) and high-resolution sector mass spectrometer (Clarke et al. 1969). Meteorite research has made particular demands upon noble gas analysis as it was found that there are noble gas compositional variations over very fine spatial scales. Separation of noble gas carrier phases using chemical procedures were developed, leading to the isolation of essentially pure exotic ‘Ne-E’ by Junck and Eberhardt (1979), then the discovery of gas-rich acid-resistant residues by Lewis et al. (1975), and culminating in the isolation of the first pure presolar grains (Lewis et al. 1987; Zinner et al. 1987). The sensitivity required to obtain high precision isotopic compositions of small abundance samples, especially for the heavy noble gases, was achieved through advances in pulse counting and developments in ultra-low blank extraction systems (Hohenberg 1980), design of the high transmission Baur-Signer source with rotational symmetry (Baur 1980), resonance ionisation mass spectrometry (Gilmour et al. 1994), and the recent construction of a compressor source with an increase in sensitivity for He and Ne by two orders of magnitude (Baur 1999). Another development has been in designing mass spectrometers for planetary probes, which have provided compositions from other planets (see Nier and Hayden 1971). Also, space missions have made measurements of the solar wind and cosmic rays (see Wieler 2002a, this volume). Important recent examples are the solar wind Ne composition determination by the SOHO mission (Kallenbach et al. 1997) and the Ne isotopic analyses during solar energetic particle events by ACE (Leske et al. 1999). THE ATMOSPHERIC STANDARD The standard for noble gas measurements, and the reference for discussions of data, is the composition of the terrestrial atmosphere. The abundances of the noble gases are in Table 1, along with those of other major and minor gases that are more abundant than Xe. With the exception of Ar, which is dominated by radiogenic 40Ar, the noble gases are present as trace constituents. Isotopic compositions are provided in Table 2. Since air is the reference standard for laboratory analyses, measurements of other media are typically normalized to these values.

Noble Gas Geochemistry and Cosmochemistry Table 1. Composition of the terrestrial atmosphere. Constituent Dry Air N2 O2 Ar CO2 Ne He CH4 Kr H2 N2O CO Xe Rn

Volume Mixing Ratio 1 7.81 × 10-1 2.09 × 10-1 (9.34±0.01) × 10-3 3.7 × 10-4 (1.818±0.004) × 10-5 (5.24±0.05) × 10-6 1-2 × 10-6 (1.14±0.01) × 10-6 4-10 × 10-7 3 × 10-7 0.1-2 × 10-7 (8.7±0.1) × 10-8 ~6 × 10-20

Total Inventory (cm3STP) (3.961±0.006) × 1024

(3.700±0.004) × 1022 (7.202±0.016) × 1019 (2.076±0.020) × 1019 (4.516±0.040) × 1018

(3.446±0.040) × 1017 2 × 105

Based on dry tropospheric air. Water generally accounts for ≤4% of air. Other chemical constituents have mixing ratios less than Xe. Data from compilations by Lewis and Prinn (1984) and Ozima and Podosek (2001). CO2 data from Keeling and Whorf (2000).

Table 2. Noble gas isotope composition of the atmosphere. Isotope 3 4

He He

20

Relative abundances (1.399±0.013) × 10 ≡1

-6

Percent molar abundance 0.000140 100

Ne Ne 22 Ne

9.80±0.08 0.0290±0.0003 ≡1

90.50 0.268 9.23

36

≡1 0.1880±0.0004 295.5±0.5

0.3364 0.0632 99.60

0.6087±0.0020 3.9599±0.0020 20.217±0.004 20.136±0.021 ≡ 100 30.524±0.025

0.3469 2.2571 11.523 11.477 57.00 17.398

2.337±0.008 2.180±0.011 47.15±0.07 649.6±0.9 ≡ 100 521.3±0.8 660.7±0.5 256.3±0.4 217.6±0.3

0.0951 0.0887 1.919 26.44 4.070 21.22 26.89 10.430 8.857

21

Ar Ar 40 Ar 38

78

Kr Kr 82 Kr 83 Kr 84 Kr 86 Kr 80

124

Xe Xe 128 Xe 129 Xe 130 Xe 131 Xe 132 Xe 134 Xe 136 Xe 126

After compilation by Ozima and Podosek (2001). He: Mamyrin et al. (1970). Ne: Eberhardt et al. (1965). Ar: Nier (1950), Steiger and Jäger (1977). Kr, Xe: Basford et al. (1973).

3

4

Porcelli, Ballentine & Wieler PRODUCTION OF NOBLE GAS ISOTOPES

The dominant cause of the extensive isotopic variations seen in natural samples is production of noble gas nuclides by nuclear processes. The parents of simple decay schemes producing noble gases are listed in Table 3. 4He is copiously produced by the 238 U, 235U, and 232Th decay series, each of which involves decay of various intermediate nuclides before producing stable Pb. Many of the transformations in these series involve alpha decay, producing alpha particles that require only the acquisition of electrons to become 4He. Spontaneous fission, although accounting for a very small fraction of 238U decay, produces Xe isotopes that can cause appreciable isotopic variations due to the generally low concentrations of Xe. The short-lived nuclides 129I and 244Pu also produce Xe (and Kr) isotopes, although these were only present within the Earth in significant amounts early in solar system history. The fission Xe isotope yield spectra for 238U and 244 Pu decay are given in Table 4. Spontaneous fission of 238U also produces Kr, but generally does not produce appreciable isotope variations. Some of the other reactions that produce noble gas nuclides within the Earth are listed in Table 5. There are also some radioactive noble gas isotopes of interest that are listed in Table 6. A comprehensive review of radiogenic, nucleogenic and fissiogenic noble gas isotope production is given by Ballentine and Burnard (2002, this volume). The heaviest noble gas, Rn, is represented in the environment by only one isotope, Rn. It is produced by 226Ra (as part of the 238U-206Pb decay chain), and decays with a half-life of only 3.8 days. Its distribution therefore reflects that of its parent and shortterm transport processes. Therefore, it is more fruitfully considered with the U-Th series nuclides, and is not included in this volume. 222

Table 3. Half-lives of parent nuclides for noble gasesa Nuclide 3

H 238 U

235

U 232 Th 40

K Pu 129 I 244

Half-life 12.26 a 4.468 Ga

0.7038 Ga 14.01 Ga 1.251 Ga 80.0 Ma 15.7Ma

Daughter 3

He He 136 Xe 4

4

He 4 He 136 Xe 40 Ar 136 Xe 129 Xe

Yield (atoms/decay) 1 8b 3.6 × 10-8 c (4.4±0.1) × 10-8 d 7b 6b Ameas ln λ A sat − Ameas

Pstable =

(1)

Cstable Texp

(2)

Here Texp is the exposure age, λ is the decay constant of the radionuclide, and Asat and Ameas are the saturation activity and the measured activity of the radionuclide at the time of fall, respectively. Pstable and Cstable are the production rate and the measured concentration of the stable nuclide, respectively. This method requires that all meteorites considered had a simple exposure history. The production rates deduced in this way are basically the mean values over the last few half-lives of the radionuclide. We will discuss in section The Cosmic Ray Flux in Time that this approach can be used to study possible variations of the GCR flux with time. A further example of the multi-meteorite approach is the work by Eugster (1988), who presents production rates for cosmogenic 3He, 21Ne, 38Ar, 83Kr, and 126Xe as a function of 22Ne/21Ne for various chondrite classes. The data are based on noble gas analyses of a considerable number of ordinary chondrites whose exposure ages were determined by the 81Kr-Kr method, which is largely self-correcting for shielding (see the respective subsection below). The production rate equations and the correction factors F for different chemical compositions are given in Table 1. These factors are based on earlier work in which the relative production rates of each nuclide from various elements or element combinations had been determined either by analysing different minerals from the same meteorite (e.g., Bogard and Cressy 1973) or from model calculations that are discussed below (e.g., Hohenberg et al. 1978). The Eugster formalism has the virtue of yielding a unique production rate for a given value of the routinely determined shielding parameter 22Ne/21Ne. However, as the actual relationship between 22Ne/21Ne and production rate is ambiguous for relatively high shielding, for chondrites the formulas in Table 1 should only be applied for 22Ne/21Ne ratios larger than about 1.08-1.10. They become unreliable for 22Ne/21Ne ratios below about 1.07, as they imply increasing production to very large depth, which is clearly not the case (these limiting 22Ne/21Ne ratios are slightly different for meteorites with higher or lower than chondritic Mg/(Mg + Al + Si) ratios). Furthermore, for very large shielding above some 50 cm or so the trend of 22Ne/21Ne with shielding reverses, as is shown by both model calculations Table 1. Shielding-corrected production rates in chondrites (Eugster 1988) F factors for various chemical classes 3

22

21

P( He)=F(2.09-0.43× Ne/ Ne) 21

22

21

P( Ne)=1.61F/(21.77× Ne/ Ne-19.32)

CI

CM

CO

CV

H

L/LL

EH

EL

1.01

1.00

0.99

0.99

0.98

1.00

0.97

1.00

0.67

0.79

0.96

0.96

0.93

1.00

0.78

0.96

38

22

21

0.75

0.88

1.03

1.10

1.08

1.00

0.98

0.89

83

22

21

P( Kr)=0.0196F/(0.62× Ne/ Ne-0.53)

0.71

0.94

1.02

1.13

1.00

1.00

0.75

0.80

P(126Xe)=F(0.0174-0.0094×22Ne/21Ne)

0.66

0.93

1.18

1.40

1.00

1.00

0.72

0.72

P( Ar)=F(0.112-0.063× Ne/ Ne)

-8

-12

Production rates P in [10 cc STP/(g × Myr)] for He, Ne and Ar, and in [10 cc STP/(g × Myr)] for Kr and Xe (1 cc STP=2.687 × 1019 atoms). All values from Eugster (1988), except P(38Ar) for which a ~12% lower value than that given by Eugster (1988) is adopted (Schultz et al. 1991; Graf and Marti 1995). 22Ne/21Ne is the ratio of the cosmogenic component. The shielding correction proposed here is approximately valid only for 22Ne/21Ne > 1.08-1.10, and may fail completely for the rare very large meteorites (see text).

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130

and analyses of the very large Gold Basin meteorite (Masarik et al. 2001; Wieler et al. 2000b; see following subsection). However, such large meteorites are probably very rare, and the 22Ne/21Ne shielding parameter in practice is useful in a majority of cases. Note also that more recent studies indicate that P(38Ar) by Eugster (1988) appears to be about 12% too high (Schultz et al. 1991; Graf and Marti 1995). In Table 1 we adopted the lower P(38Ar) values, although the issue is not settled (Eugster et al. 1998b). Systematic studies on single meteorites and semiempirical models. An example of the second approach to determine production systematics is the study of cosmogenic He, Ne, and Ar, some radionuclides and nuclear tracks in the large chondrite Knyahinya (Graf et al. 1990b) along with follow-up studies that included cosmogenic Kr and Xe (Lavielle et al. 1997) as well as further radionuclides (Reedy et al. 1993; Jull et al. 1994). The main mass of Knyahinya was split in two almost equally sized fragments upon its fall. Fortunately the cut passed very close to the preatmospheric center of the meteorite and the slab included an edge where only a very few cm has been ablated. This happy accident allowed the determination of concentration gradients of the various cosmogenic nuclides as a function of preatmospheric depth. The data were also used to determine the parameters of a semiempiric model of cosmogenic nuclide production (Graf et al. 1990a). This model, based on earlier work by Signer and Nier (1960, 1962) and others, describes production of each nuclide by two depth- and size-dependent terms with only one free parameter each, both of which are independent of shielding. Based on the Knyahinya data and an independently determined exposure age of this meteorite, the model is quite successful in predicting cosmogenic noble gas and radionuclide production rates as a function of depth in chondrites of different sizes (see also figures in Wieler et al. 1996). Similar semiempirical models have been used to describe nuclide production in stonyiron or iron meteorites (e.g., Honda 1988). Voshage (1978, 1984) combined noble gas data with determinations of the very long-lived cosmogenic radionuclide 40K (T1/2 = 1.25 Gyr) in iron meteorites (see also The Cosmic Ray Flux in Time section). Many further comprehensive studies on single meteorites have either contributed to our understanding of cosmogenic nuclide production systematics or the exposure histories of the studied meteorites. Some examples are the investigations on Grant (Signer and Nier 1960, 1962; Graf et al. 1987), Keyes (Wright et al. 1973), St. Severin (Schultz and Signer 1976; Lavielle and Marti 1988), Jilin (Begemann et al. 1996; Heusser et al. 1996 and references therein), Chico (Garrison et al. 1992), Bur Gheluai (Vogt et al. 1993); Canyon Diablo (Heymann et al. 1966; Michlovich et al. 1994), and Gold Basin (Kring et al. 2001; Wieler et al. 2000b). Physical models. Much progress has also been made in modeling cosmogenic nuclide production in meteorites based essentially only on physical principles and without using any free parameters except for the absolute cosmic ray flux. These models simulate all the processes relevant for particle production and transport using Monte Carlo techniques (Masarik and Reedy 1994; Masarik et al. 2001; Leya et al. 2000a, Leya et al. 2001b and references therein). The basic equation underlying these models is (e.g., Leya et al. 2000a). N

Pj (R, d, M) = ∑ ci i=1

NA Ai

3 ∞

∑ ∫ σ j,i,k (E) × J k (E, R, d, M)dE

(3)

k =1 0

where Pj(R,d,M) is the production rate of nuclide j as a function of the meteoroid radius R, sample depth d, and the solar modulation parameter M (this parameter is a measure for the reduction of the GCR particle flux by the Sun's magnetic field; see also Niedermann 2002). The first sum goes over all target elements i, while the index k in the second sum represents the reaction particle type (primary or secondary proton, secondary neutron). NA is Avogadro's number, Ai the mass number (in amu) of the target element i, ci the

Cosmic-Ray-Produced Noble Gases in Meteorites

131

abundance of element i (in g/g). σj,i,k(E) is the excitation function for the production of nuclide j from element i by reactions induced by particles of type k, and Jk(E,R,d,M) is the differential flux density of particle type k, which depends on the energy E of the reacting particles (as well as on R, d, and M). Note that the primary GCR intensity is assumed to be constant in time (see also The Cosmic Ray Flux in Time section). Note also that the models by Leya and coworkers and Masarik and coworkers take into account primary and secondary alpha particles only approximatively, by multiplying production rates from Equation (3) by a factor of 1.55 (the 12% alpha particles in the GCR have 55% as much mass—and hence energy—as the 87% GCR protons). Equation (3) shows that the most critical ingredients for such models are the numerous excitation functions needed. The data base for proton-induced reactions is fairly complete for most relevant noble gas nuclides as well as for most of the important radionuclides. An important exception is 36,38Ar production from Ca, for which cross sections are still based only on theoretical nuclear models (e.g., Blann 1971). Measured cross sections for neutron-induced reactions are, however, very scarce (e.g., Sisterson et al. 2001). This is a serious problem because secondary neutrons usually dominate the cosmogenic nuclide production. Leya (1997) and Leya and Michel (1997) therefore derived excitation functions for neutron-induced reactions with the data from five thicktarget simulation experiments, where stony or iron spheres of various radii and filled with a large number of target foils were isotropically irradiated by energetic protons (Michel et al. 1986; Leya et al. 2000b and references therein). Neutron excitation functions were derived for all reactions where the proton-induced production could be reliably calculated, based on their known excitation functions, and so subtracted. Figures 2-6 present some important results from the models by Leya et al. (2000a) and Masarik et al. (2001). Figure 2 shows that both models do reproduce the measured 21 Ne depth profile in Knyahinya well, and hence can be expected to reliably predict nuclide production in meteorites of a wide range of sizes. Remarkably, secondary neutrons contribute about two thirds to the total 21Ne production at the surface and this fraction increases to 85% near the center. This illustrates the importance of reliable neutron cross section data. Figures 3 and 4 show the 3He and 21Ne production rates, respectively, in the two most abundant meteorite classes, the H and L chondrites, as a function of depth and size. As noted above, for average-sized meteorites (R < 40 cm), production rates vary within only about a factor of 1.5. On the other hand, for the Gold Basin chondrite with its preatmospheric radius of perhaps 3 m, nuclide concentrations vary by more than an order of magnitude (Kring et al. 2001; Wieler et al. 2000b). This meteorite is almost represented by the lines denoting an infinite radius (2π). Figure 5 shows P(21Ne) values versus the 22Ne/21Ne ratio for objects of various sizes and samples of various depths according to calculations by Leya et al. (2001a). Also shown is the empirical curve according to Eugster (1988) that is given in Table 1. The Figure shows that the relation between P(21Ne) and 22Ne/21Ne is more or less unique only for 22Ne/21Ne ≥ 1.13, i.e., for meteorites with radii less than some 15 cm or for nearsurface samples of somewhat larger bodies. In more heavily shielded samples, P(21Ne) may differ by up to a factor of 2 at a given 22Ne/21Ne. However, this large spread is only observed for objects with radii above ~85 cm (filled symbols), which fortunately are rare. For more common sizes and mean values of 22Ne/21Ne of 1.08-1.14, the model predictions agree quite well with the Eugster (1988) calibration curve. At lower shielding, the empirical calibration curve runs parallel to the model results, although with an offset of up to 30%. In summary, Figure 5 indicates that the shielding-corrected P(21Ne) values according to Eugster (1988) are usually correct to within some ± 15% for average

132

Wieler

Figure 2. Comparison of measured (squares) and modelled (solid lines) 21Ne concentrations in the L/LL chondrite Knyahinya. The measured data are from Graf et al. (1990b), and the two model curves from Leya et al. (2000a) and Masarik et al. (2001). Both models assume a radius of Knyahinya of 45 cm and an exposure age of 39 Myr. Also shown are the individual contributions by primary protons and secondary protons and neutrons (Leya et al. 2000a). Note that even at the surface, most of the 21Ne is produced by secondary particles.

Figure 3. GCR-induced production rates of 3He in ordinary chondrites of classes H (left ordinate) and L (right ordinate). For LL chondrites, all values are 2% larger than those of L chondrites. The curves represent meteoroids with radii of between 5 and 120 cm. Also included is a depth profile on an infinitely large flat body (2π). Model calculations are from Leya et al. (2000a) and updates (I. Leya, pers. comm. 2001).

Cosmic-Ray-Produced Noble Gases in Meteorites

Figure 4. Same as Fig. 3 but for 21Ne. For LL chondrites, the L-scale on the right ordinate has to be multiplied by 1.02. Data from Leya et al. (2000a).

Figure 5. The symbols show the GCR production rate of 21Ne as a function of the shielding parameter 22Ne/21Ne in H chondrites of various radii and samples from variable depths, calculated according to Leya et al. (2001a). Also shown is the empirical relation between P(21Ne) and 22Ne/21Ne according to Eugster (1988). The model data show that only for 22Ne/21Ne ≥ 1.13 there exists an essentially unequivocal P(21Ne) value. At lower 22 Ne/21Ne values, P(21Ne) varies by up to twofold at a given 22Ne/21Ne. Nevertheless, for meteorite radii not larger than ~50 cm, the Eugster (1988) relation is in agreement with the model results to within ±15% for 1.08 < 22Ne/21Ne < 1.14. It is unclear whether the physical model or the empirical correlation describes the data better for 22 Ne/21Ne > 1.13.

133

134

Wieler

shielding (22Ne/21Ne = 1.08-1.13). So far it is unclear whether the physical model or the empirical correlation predicts P(21Ne) more accurately for 22Ne/21Ne > 1.13, as samples with such low shielding with independently calibrated exposure ages are rare. Figure 6 shows the utility and the ambiguities of using 22Ne/21Ne as a shielding parameter in yet another way. The model by Masarik et al. (2001) reproduces this ratio very well for meteorites of the size of Knyahinya (R ~45 cm). However, in very large objects, this ratio increases with shielding, such that, e.g., a sample from the center of a R = 2 m chondrite has the same 22Ne/21Ne value as a sample from 5 to 10 cm below the surface of Knyahinya. Elemental production rates. The relationships shown in Figures 2-6 constrain production rates for meteorites with the chemical composition of ordinary chondrites, but they are not directly applicable to other meteorite classes with distinctly different concentrations of the relevant target elements. To do this, production rates from each major target element are needed, and these are also provided by the models. Table 2 is a much abbreviated version of lists of elemental production rates of 3He, 21Ne, and 22 Ne/21Ne for meteoroids of various sizes provided by Leya et al. (2000a) for Ne and I. Leya (pers. comm. 2001) for He, respectively. Also given are 38Ar production rates from Fe and Ni (Leya et al. 2001b). Production of 38Ar from Ca, the major target element in stony meteorites, has not yet been calculated with the “Hannover model” developed by R. Michel, I. Leya and coworkers, because of the lack of experimentally determined cross sections. We therefore reproduce in Table 2 the older 38Ar values from Ca (and other elements) presented by Hohenberg et al. (1978) for infinitely large bodies (2π exposure geometry), although the accuracy of these values is unclear. Elemental production rates for other noble gases are also given by Regnier et al. (1979) and Reedy (1981).

Figure 6. In the main panel, the 22Ne/21Ne ratio predicted by the model of Masarik et al. (2001) is compared with data from Knyahinya samples from known preatmospheric depths (Graf et al. 1990b). The inset shows the modelled ratios for various meteorite radii and a 2π irradiation. At depths below ~50 cm, 22Ne/21Ne increases again. Therefore, the same 22Ne/21Ne ratio corresponds to widely different depths, which illustrates the ambiguity of this ratio as a shielding index in very large meteorites.

148 166 163 190 170 202 186 232 188 246 173 210 161 138 118 146 127 83 15

O

113 124 125 144 132 158 148 188 152 205 144 180 135 130 97 113 98 64 12

Mg

3

124 130 130 139 131 139 135 149 132 148 120 119 112 78.5 90 86 66 40 7

Al

He 138 143 142 149 143 150 147 158 144 156 133 127 123 86.2 100 90 70 43 8

Si 86.6 89.0 88.2 90.8 87.3 87.2 86.8 88.2 82.8 82.7 74.5 62.0 68.1 38.7 58 49 35 20 3

Fe 53.2 69.0 70.2 98.7 80.4 121.0 102.7 159.2 111.7 184.3 108.0 171.7 104.3 120.6 59.9 93.1 83.1 52.7 9.9

Mg 32.6 37.8 37.4 45.9 39.6 49.9 44.7 59.3 45.5 65.0 41.8 53.6 39.2 36.5 24.7 34.3 28.9 17.7 3.0

Al

21

25.2 28.8 28.8 34.9 30.6 38.0 34.6 45.5 35.6 51.2 33.2 43.4 30.9 30.2 19.9 27.1 23.7 15.0 2.8

Si

Ne 8.57 8.77 8.69 8.89 8.53 8.40 8.36 8.32 7.88 7.63 7.02 5.44 6.33 3.20 5.40 4.50 3.06 1.63 0.22

Ca 2.03 1.99 1.97 1.86 1.88 1.66 1.75 1.46 1.60 1.14 1.41 0.69 1.26 0.37 1.31 0.73 0.37 0.16 0.01

Fe 1.290 1.200 1.191 1.091 1.152 1.054 1.095 1.020 1.076 1.002 1.071 0.969 1.066 1.020 1.061 0.927 0.896 0.905 0.917

Mg 1.188 1.224 1.228 1.265 1.260 1.325 1.309 1.377 1.339 1.413 1.369 1.470 1.388 1.537 1.190 1.175 1.188 1.208 1.255

Si 1.197 1.197 1.190 1.187 1.187 1.185 1.184 1.188 1.176 1.151 1.174 1.159 1.189 1.150 1.224 1.196 1.163 1.145 1.116

Ne/21Ne

Al

22

1.256 1.236 1.234 1.197 1.223 1.175 1.200 1.127 1.185 1.075 1.179 1.022 1.180 0.994 1.235 1.100 1.027 0.978 0.908

Ca

61 113 103 72 12.0

Ca

Ar

11.8 11.1 7.48 3.91 .42

Ti

38

7.47 7.43 7.39 7.23 7.19 6.66 6.78 6.17 6.33 5.32 5.60 3.63 5.08 2.00 4.42 3.14 1.96 0.97 0.12

Fe

6.54 6.58 6.50 6.41 6.33 5.89 6.01 5.52 5.60 4.79 4.95 3.22 4.47 1.80 3.96 2.83 1.75 0.78 0.10

Ni

Production rates [in 10-10 ccSTP/(g × Ma)] according to model calculations (1 ccSTP = 2.687 × 1019 atoms). Given are values for the surface (subscript s) and the center (subscipt c), respectively, of meteorites with radii between 5 and 100 cm. The lowermost 5 rows give the values for an infinitely large body (2π irradiation) at depths between 0 and 500 g/cm2. Most data are from I. Leya and coworkers (3He: I. Leya, pers. comm. 2001; 21Ne (4π): Leya et al. 2000a; Ne (2π): Leya et al. 2001; 38Ar from Fe and Ni: Leya et al. 2001). 38Ar from Ca and Ti (2π) values are from Hohenberg et al. 1978. These values should not be compared straightforwardly with the respective 38Ar data from Fe and Ni by Leya et al. 2001, because the latter values are considerably higher than P(38Ar) from Fe and Ni also reported by Hohenberg et al. 1978. 22Ne/21Ne from Fe as target not given, because P(22Ne) from Fe reported by Leya et al. (2000a, 2001) is unreliable. This table is a summary of much more extended data sets given in the original publications.

5,s 5,c 10,s 10,c 15,s 15,c 25,s 25,c 40,s 40,c 65,s 65,c 100,s 100,c 2π, 0 2π, 40 2π, 100 2π, 200 2π, 500

radius/ depth

Table 2. Elemental production rates of cosmogenic He, Ne, and Ar in meteorites

Cosmic-Ray-Produced Noble Gases in Meteorites 135

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The matrix effect. When calculating a production rate for a sample with a given chemical composition by multiplying the elemental production rates times the fractional abundance of the respective target element and summing over all relevant elements, one ignores a possible dependence of elemental production rates on the overall chemical composition of a meteorite. This so-called matrix effect can arise because particle fluxes and spectra are a function of the chemical composition of the target. For example, the matrix effect is expected to be important for silicate inclusions in stony-iron meteorites, where the high Fe concentrations lead to an enhanced flux of secondary neutrons relative to stony meteorites (Begemann and Schultz 1988; Masarik and Reedy 1994). The matrix effect will be particularly pronounced for nuclides largely produced by relatively lowenergy neutrons, e.g., 21Ne from 24Mg and 38Ar from Ca. Albrecht et al. (2000) showed that 21Ne production rates in silicates from mesosiderites on average are some 60% higher than would be expected for samples with the same composition but in a silicateonly matrix. This result is in agreement with the model calculations by Masarik and Reedy (1994). For stony meteorites of variable composition the matrix effect is generally expected to be < 10%, however, an exception being 38Ar production from Ca, which is several ten percent lower in eucrites and aubrites compared to chondrites (Masarik and Reedy 1994). Production rate ratios P(10Be)/P(21Ne) and P(26Al)/P(21Ne). Because of the ambiguity of the parameter 22Ne/21Ne, other methods for deducing shielding-corrected production rates are often useful. These methods operate with a radioactive and a stable nuclide pair that have production rate ratios that can be assumed to depend only weakly on the irradiation conditions. We already mentioned the systems 10Be-21Ne, 26Al-21Ne, and 53Mn-21Ne, which can be used not only to deduce average production rates as shown above, but in principle also to derive a shielding-corrected 21Ne production rate for an individual sample. Based on the Knyahinya data, Graf et al. (1990a) proposed the following relations for L chondrites (in atoms/atom): ⎛ P(10 Be) = (0.140 ± 0.001) + (0.02 ± 0.05) × ⎜⎜ 21 P( Ne) ⎝ ⎛ P( 26 Al ) = (0.37 ± 0.02) − (0.4 ± 0.7) × ⎜⎜ 21 P( Ne) ⎝

22 21

22 21

⎞ Ne − 1.11⎟⎟ Ne ⎠

⎞ Ne − 1.11⎟⎟ Ne ⎠

(4)

(5)

Note that the two production rate ratios P(10Be)/P(21Ne) and P(26Al)/P(21Ne) are given as functions of 22Ne/21Ne and are thus basically also subject to the reservations on the reliability of this shielding parameter. However, the dependence of P(10Be)/P(21Ne) on 22Ne/21Ne in Equation (4) is actually very weak, which reflects the fact that the ratio 10 Be/21Ne is nearly constant over the entire range of shielding represented by the Knyahinya samples. The pair 10Be-21Ne therefore offers a way to determine shieldingcorrected exposure ages in chondrites (Eqn. 4), if 10Be can be assumed to be saturated, i.e., for exposure ages above ~6 Myr. A word of caution is necessary, however. In contrast to the model of Graf et al. (1990a), that by Leya et al. (2000a) predicts P(10Be)/P(21Ne) in chondrites to depend substantially on the ratio 22Ne/21Ne. If so, 10Be21 Ne ages are correct only to within the limits imposed by the reliability of 22Ne/21Ne as shielding indicator. Leya et al. (2000a) estimate that for 22Ne/21Ne < 1.10 10Be-21Ne ages are ambiguous to within ±15%. In iron meteorites this ambiguity does not appear to exist. Both data and models show that the ratio P(10Be)/P(21Ne) is constant over a wide range of shielding. Lavielle et

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al. (1999a) derive a P(10Be)/P(21Ne) value of 0.55, in perfect agreement with the modelbased value by Leya et al. (2000a), which is essentially independent on shielding for all meteorites with radii not larger than one meter. Graf et al. (1987) had deduced a P(10Be)/P(21Ne) ratio of 0.772±0.05 with data from the Grant iron meteorite, by adopting an exposure age derived with the 40K-41K method. This ratio is 40% higher than that published by Lavielle et al. (1999a), which is based on the 36Cl-36Ar age of Grant and another iron meteorite. As Cl-Ar and K-K ages of iron meteorites differ systematically by ~40%, probably because the cosmic ray intensity has changed with time (see The Cosmic Ray Flux in Time section), the production rate ratios P(21Ne)/P(10Be) given by Graf et al. (1987), Lavielle et al. (1999a), and Leya et al. (2000a) therefore all agree very well. If we adopt the 36Cl-36Ar age scale, a 10Be-21Ne age of iron meteorites can therefore be calculated by ⎛ 21 Ne ⎞ ⎟⎟ × 0.55 Texp = ⎜⎜ 10 ⎝ P( Be) ⎠

(6)

where 21Ne is the concentration in atoms per unit of mass and P(10Be) the production rate in atoms per unit of mass per unit of time. In iron meteorites, the P(26Al)/P(21Ne) ratio also appears to be rather insensitive to shielding. Leya et al. (2000a) predict an essentially constant value of 0.45, somewhat higher that that of 0.38 proposed by Hampel and Schaeffer (1979), which is based on the 36 Cl-36Ar age scale. The mean value of 0.56 by Graf et al. (1987) (based on the 40K-41K age scale) is again 25-47% higher than the other two values. However, in contrast to the model prediction (Leya et al. 2000a), the measured P(26Al)/P(21Ne) in Grant increases somewhat with increasing shielding (Graf et al. 1987). The 36Cl-36Ar, 129I-129Xe, and 81Kr-Kr methods. These methods do not involve Ne/21Ne, and are thus self-correcting for shielding in a more strict sense than the 10Be21 Ne method in chondrites. The 36Cl-36Ar pair has the advantage that in metallic Fe-Ni samples most of the 36Ar is produced through its radioactive precursor 36Cl, and hence the production rate ratio P(36Ar)/P(36Cl) is undoubtedly essentially shielding-independent. One advantage of the 81Kr-Kr technique is that both radioactive and stable nuclide are isotopes of the same element krypton, so only an isotopic ratio measurement is needed. Note, however, that these methods also will yield erroneous ages if the assumption of a single stage exposure history of the studied sample is violated. 22

Lavielle et al. (1999a) and Graf et al. (2001) give values for P(36Cl)/[γP(36Cl) + P(36Ardirect)] of 0.84±0.04 for Fe and 0.78±0.03 for Ni. These values should be shielding-independent for production by both protons and neutrons. The coefficient γ = 0.981 is the branching ratio for the decay of 36Cl into 36Ar. With an average Ni concentration of 9% in chondritic metal, the 36Cl-36Ar age can be calculated by ⎛ 36 Ar ⎞ Texp = 0.363 * ⎜⎜ 36 cos ⎟⎟ ⎝ Cl ⎠

(7)

if [atoms per gram] are used as units for the concentrations of cosmogenic (subscript cos) 36 Ar and 36Cl, or ⎛ 36 Ar ⎞ Texp = 427 * ⎜⎜ 36 cos ⎟⎟ ⎝ Cl ⎠

using [10-8 ccSTP/g] for 36Ar and [dpm/kg] for 36Cl.

(8)

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Begemann et al. (1976), Leya et al. (2000a) and Albrecht et al. (2000) give coefficients of 425, 433 and 430, respectively, for Equation (8). Whereas 36Cl-36Ar ages are thus basically very reliable, they are quite difficult to determine, because very highpurity metal separates are required to avoid contamination with 36Cl and 36Ar from silicate impurities (e.g., Graf et al. 2001). Also, 36Cl and 36Ar have to be analyzed on two different, often small aliquots. Therefore, rather few 36Cl-36Ar exposure ages have been reported so far, although the method is becoming increasingly popular. Another potentially interesting method where the stable cosmogenic noble gas nuclide is predominantly produced through a radioactive precursor is the live-129I-129Xe method (Marti et al. 1986). All cosmogenic 129Xe produced from Te will have gone through 129I (half-life 15.7 Myr). For samples where 129Xe contributions from Ba can be corrected for (using 124,126Xe) and for which 129Xe contributions from extinct 129I from the early solar system are of no concern, this system is ideal in being self-correcting for both shielding and variable Te abundance. Marti et al. (1986) determined the exposure age of the big Cape York iron meteorite with the live-129I-129Xe couple in troilite inclusions. Otherwise, the method has not been widely applied so far, but as 129I analyses by accelerator mass spectrometry are becoming routine, it should prove very useful to study cosmic ray flux variations over timescales of a few ten million years. The 81Kr-Kr technique was proposed by Marti (1967). Cosmogenic Kr is mainly produced from the target elements Rb, Sr, Y, and Zr. 81Kr has a half-life of 229,000 yrs. Marti and Lugmair (1971) observed that the isotopic ratios of cosmogenic Kr vary systematically in Apollo 12 lunar rocks, due to variable shielding. They observed the following relations (all Kr concentrations or isotopic ratios refer to the spallation Kr component): ⎛ 80 Kr + 82 Kr ⎞ ⎟⎟ P( 81 Kr ) / P( 83 Kr ) = 0.95⎜⎜ 83 ⎝ 2× Kr ⎠

and P( 81Kr ) / P( 83 Kr) = 1.262

78 83

Kr + 0.381 Kr

(9)

(10)

These relations allow determination of a shielding-corrected exposure age based on a single Kr analysis The factor 0.95 in Equation (9) represents the isobaric fraction yield of 81 Kr (Marti 1967). Note that this value has recently been redetermined to ~0.92 for chondritic abundances of the relevant target elements (B. Lavielle, pers. comm. 2001). The second relation is insensitive to 80Kr and 82Kr from neutron capture on Br, which is present in relatively large or Br-rich meteorites (see next section). Note, that these relations are strictly valid only for samples which have the same relative abundances of the major target elements as Apollo 12 rocks. For example, Apollo 11 samples, which have ~2-3 times higher Zr/Sr ratios than Apollo 12 samples, have a P(81Kr)/P(83Kr) that is slightly higher at a given 78Kr/83Kr value than predicted by Equation (10) (Lugmair and Marti 1971). However, the difference is only about 1-3%. Moreover, Finkel et al. (1978) showed that the Kr spallation systematics derived from Apollo 12 samples are also valid for the two chondrites San Juan Capistrano and St. Severin, apparently because Apollo 12 rocks and chondrites have rather similar Zr/Sr ratios (although very different Rb/Sr ratios, Rb contributing substantially to Kr production in chondrites but not in lunar samples). Equations (9) and (10) are thus also widely used today for meteorites. Eugster (1988) presented a correlation between P(81Kr)/P(83Kr) and 22Ne/21Ne, for ordinary chondrites, but from the only available 81Kr depth profile in Knyahinya it is unclear whether this relation holds for low shielding (Lavielle et al. 1997). The 81Kr-Kr method has the

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disadvantage of often not being very precise, both because of the low abundance of 81Kr and because large corrections for non-cosmogenic Kr are often required. An important caveat is that all procedures for deducing production rates discussed above assume a single-stage exposure. In other words, it is assumed that a meteorite was ejected from its parent body from at least several meters depth, where it had been completely shielded from the GCR, and then travelled to Earth without ever changing its shape by a further collision or by space erosion (see Complex Exposure Histories section). Only in this case is the present-day activity of a radionuclide a good measure for the production rate of a stable nuclide, as the radionuclide concentration depends only on the shielding of the sample during the past few half-lives, whereas the stable nuclide has been accumulating during the entire exposure of the sample to cosmic rays. However, we will see in the Complex Exposure Histories section that for a fairly high fraction of all meteorites this requirement of a single exposure stage is probably not fulfilled. Complex exposures may well go unnoticed, e.g., if a single noble gas analysis on a single sample has been carried out only. A factor further limiting the accuracy of production rates is a possible change in the mean GCR flux over the timescales of interest. We will discuss in section The Cosmic Ray Flux in Time that such variations on a million-year scale appear to be modest but that the mean flux during the last few or few ten million years probably has been some 40 to 50% higher than the mean over the past 1-2 Gyr. The resulting additional uncertainty of exposure ages is often only a minor concern, however. This is because one is often more interested in whether exposure ages of different meteorite classes cluster at certain values rather than in the absolute position of such clusters. It needs, however, to be taken into account if ages based on production rates derived from different nuclide systems are compared with each other, particularly for iron meteorites (see the respective subsection below). In summary, production rates of cosmogenic nuclides in meteorites have quite a variable level of accuracy and reliability. In general, the more nuclides that are determined on the same sample and the more samples that are measured from the same meteorite, the higher will be our confidence in a stated exposure age. A single age determination will usually not be more accurate than to within 15-20%, with the probable exception of 36Cl-36Ar ages, which may have a precision of better than 10%. Uncertainties may even be considerably larger than 20%, e.g., if an unusually heavy shielding is not recognized. On the other hand, we will see below that the accuracy that is presently achieved generally is good enough to study the history of meteorites as small bodies in space and collisional processes on parent bodies. Cosmogenic noble gases produced by capture of low-energy neutrons So far, we have mainly discussed production of cosmogenic nuclides by neutrons or protons with energies high enough (tens of MeVs or higher) to break up a target nucleus. The resulting cosmogenic nuclide (commonly referred to as “spallation product”) usually has a lower mass than the original target. In contrast to this, target nuclei may also capture secondary cosmic ray neutrons which have already been slowed down to much lower energy without having yet interacted with a nucleus. These are so-called thermal ( 200 × 10-8 cc/g) of all iron meteorite classes. (4He/3He)cos Table 3. Isotopic composition of cosmogenic correlates with the widely used He, Ne, and Ar. shielding parameter 4He/21Ne, Corresponding 4 increasing with shielding from He/ 3He 4 He/ 21Ne range ~3 to ~4.5. However, this correlation is often disturbed in iron meteorites 3.2-4.4 200-440 meteorites that suffered 3H 1 stony meteorites 5.2±0.3 losses. 6.1±0.32 In most stony meteorites, He is overwhelmingly radiogenic, such that (4He/3He)cos is difficult to measure directly (and of rather little practical concern). The two values given in Table 3 have both been derived from a suite of meteorites with low and presumably rather constant radiogenic 4He. These are L chondrites which record a major collision on their parent body some 500 Myr ago that led to a 4

22

chondrites

Ne/ 21Ne

stony meteorites

0.88-0.98a

1.05-1.25 36

iron meteorites

Corresponding Ne/ 21Ne range

20

Ar/ 38Ar

0.60-0.665 (mean 0.63) ~0.65

Corresponding He/ 21Ne range

4

450-200

1: Heymann (1967). 2: Alexeev (1998); other values see text. a: range somewhat uncertain, see text.

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142

complete loss of the radiogenic 4He accumulated up to that time (Swindle 2002b). Within their limits of uncertainty, the L chondrite values should also hold for other stony meteorite classes. This is also indicated by some stony meteorites, in particularly diogenites, which have measured 4He/3He ratios as low as ~5 due to very low U and Th concentrations (Welten et al. 1997). Neon. We have seen above that (22Ne/21Ne)cos in stony meteorites is quite variable, which makes it useful as a shielding parameter. The range stated in Table 3 is for chondrites (see Fig. 6), and will be different for stony meteorites with rather different relative Mg, Al, and Si abundances. The differences can be estimated by means of Table 2. The (20Ne/21Ne)cos ratio in chondrites is difficult to evaluate because at least minor contributions of primordial or atmospheric 20Ne are ubiquitous. Consideration of only ordinary chondrites of high exposure ages and high petrographic types 5 and 6 (i.e., little or no primordial Ne; Schultz and Franke 2000) suggests that (20Ne/21Ne)cos increases with increasing (22Ne/21Ne)cos (Table 3). Argon. Iron meteorites again often allow direct measurement of the (36Ar/38Ar)cos ratio. Iron meteorites with relatively high 38Arcos show an inverse correlation between 36 Ar/38Ar and the shielding parameter 4He/21Ne (Table 3). (36Ar/38Ar)cos decreases by up to ~10% with higher shielding. The average value is ~0.63. Determining (36Ar/38Ar)cos in stony meteorites is again compromised by the common presence of primordial or atmospheric Ar and sometimes also by neutron-capture-produced 36Ar. Yet, the lowest measured 36Ar/38Ar ratios in achondrites poor in primordial gases and with relatively high exposure ages are ~0.63, similar to the mean value of iron meteorites (Schultz and Franke 2000). More achondrite values cluster around ~0.65, however, which is a widely adopted (36Ar/38Ar)cos value for stony meteorites. It appears rather surprising that 36Ar and 38Ar are produced in almost equal proportion from the target elements Ca and Fe, respectively. Krypton. The average isotopic composition of cosmogenic Kr in lunar soils and in chondrites is given in Table 4. These data are from compilations by Pepin et al. (1995) and Lavielle and Marti (1988). When correcting a measured Kr composition for a trapped component, the assumption is commonly made that the cosmic-ray-produced contribution on 86Kr can be neglected, so that the measured 86Kr can be assumed to be entirely trapped. Note that 80,82Kr in different meteorites may contain variable contributions from Table 4. Isotopic composition of cosmogenic Kr and Xe (relative to 83Kr, 126Xe ≡ 1) 78

Kr

80

Kr

82

Kr

84

Kr

86

Kr

average, lunar bulk soils1

0.20(2)

0.54(7)

0.72(5)

0.32(10)

≡0

chondrites2

0.18(4)

0.60(8)

0.76(8)

0.67(15)

≡0

124

average, lunar bulk samples1

Xe

0.56(3)

128

Xe

1.48(6)

Chondritic Ba/REE3

0.595(10) 1.52(10)

meteorites4

0.55(10)

1.40(15)

129

Xe

1.64(15)

1.6(2)

130

Xe

131

Xe

132

Xe

134

Xe

136

Xe

0.95(7)

5.30(42)

0.77(20)

0.05(3)

≡0

0.98(06)

3.77(16)

0.83(06)

0.044(15)

≡0

1.0(2)

2.5(5)

1.0(2)

0

≡0

Uncertainties in ( ) given in units of the last digit. 1: Pepin et al. 1995. 2: Lavielle & Marti 1988, mean and standard deviation from 11 chondrites. 3: Hohenberg et al. 1981, based on Angra dos Reis achondrite data. 4: Kim & Marti 1992.

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Table 5. Exposure age ranges and clusters of meteorite classes Class Chondrites H1

L2

Range (Myr)

Clusters (Ma)

Comments

1-80

7.6 & 33 7.0? 24 40 28 15 & 5 15 10? 28? 40? 25? & 8? & 3.5? 9? 29? 0.22

all petrographic types H5 a.m. falls only H6 only mainly L5 & L6 mainly L5 & L6 &40Ar-poor 40 Ar-poor mainly LL5 & LL6 & 40Ar-rich LL6 only mainly LL3 mainly LL4 clusters need confirmation

1-70

LL3

0.03-70

EH & EL4

0.07-66

CO, CV, CK5

0.15-63

CM & CI6 R7

0.05-7 0.2-50

Other meteorites HEDa,8

Aubrites9 Lodranitesacapulcoites10 Ureilites11 Iron meteoritesb IVA12 IIIAB13 Mesosiderites14 Lunar15 Martian15

5-76

~17-130 4-10

21 & 38 12? 50? 45-80??

CV, CK only CO, CV, CK CM only (CI poor statistics) many ages between ~7-40 Ma H&E&D Eucrites only? diogenites only? cluster unclear (see text) similar as 7 Ma peaks for H chondrites

0.1-34 10-2300 255 207 460 10-180 60 Myr as the 10-Myr-lifetime curve and this model distribution may also be consistent with the small but discernible fraction of meteorites with exposure ages 60 Myr are absent and quite few ages are below 2

Cosmic-Ray-Produced Noble Gases in Meteorites

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Myr. A notable exception is the EH3 chondrite Galim/b, which seems to be an inclusion in the LL-chondritic polymict breccia Galim/a with an exposure age of only ~70,000 years (Patzer and Schultz 2001). Rumuruti chondrites. Schultz and Weber (2001) report an age range of 0.2-50 Myr for 18 members of this class. Many ages fall between 7 and 40 Myr, but Schultz and Weber (2001) note that the remarkably low age of about 200,000 yr of the R chondrite Northwest Africa 053 demonstrates that short transfer times from the Asteroid belt are possible. The number of R chondrites is too low to expect clear-cut exposure age clusters. Carbonaceous chondrites. Carbonaceous chondrites are in many respects the most primitive meteorites. The exposure age distributions of the various chemical classes reveal a clear dichotomy (Fig. 8). The CV, CO, and CK histograms essentially look very similar to those of the other stony meteorite classes, with most ages falling in the range of between a few Myr up to 40-60 Myr (Scherer and Schultz 2000). The very low age of only 0.15 Myr of the CO chondrite Isna is remarkable. On the other hand, the chemically most primitive classes CI and CM have strikingly low-exposure ages of less than 7 Myr, the majority of them falling between a mere 50 kyr and 2 Myr (Nishiizumi et al. 1993; Caffee and Nishiizumi 1997; Eugster et al. 1998a; K. Nishiizumi, pers. comm. 2001, see Wieler and Graf 2001). CV and CK chondrites possibly show a peak in their exposure age distribution at ~9 Myr and all three classes CV, CK, and CO may display a cluster at ~29 Myr (Scherer and Schultz 2000). If these common events could be verified, this would support the idea that these classes are closely linked to each other (Kallemeyn et al. 1991). CI meteorites are much too scarce to reveal any possible clustering, but CM chondrites show a very distinct peak at 0.2-0.25 Myr. Why does the majority of CM and CI chondrites have such low-exposure ages in a range almost unoccupied by all other stony meteorite classes? We noted above that one reason may be the mechanical weakness of these meteorites and their parent bodies, which may strongly reduce their mean lifetime against collisional destruction. This would imply than CM (and CI?) chondrite meteoroids are abundantly produced. However, as also noted above, the high exposure ages of the fragile aubrites illustrate that other factors must be involved, such as mean transfer times and dynamical lifetimes, as discussed below. Perhaps the distinct 0.2- to 0.25-Myr peak is due to a collision on an immediate CM parent body with an orbit crossing that of the Earth (Caffee and Nishiizumi 1997). Besides the one distinct peak, the exposure ages of CM chondrites are distributed rather evenly, perhaps suggesting that a comparatively large fraction of them derives from relatively frequent smaller collisions. Differentiated meteorites HED meteorites. The Howardite-Eucrite-Diogenite (HED) clan is the largest class of achondrites. It is of particular importance here because it is believed that we know the parent body: the asteroid Vesta, or perhaps some of the smaller “Vestoids” thought to have been spalled off Vesta in a very large collision (Binzel and Xu 1993). The HED exposure age distribution (Fig. 9) is again strikingly similar to those of ordinary chondrites: all ages fall between 5-80 Myr, and distinct peaks emerge at 20-25 and 35-42 Myr (Eugster and Michel 1995; Welten et. al. 1997). Welten et al. (2001b) propose that there is an additional peak at ~50 Myr for the diogenites, although this is not resolved in Figure 9. Possibly as few as 5 or 6 events can explain most of the HED exposure ages (Eugster and Michel 1995; Welten et al. 1997). Based on this low number, Welten et al. (1997) and Bottke et al. (2000) favor Vesta rather than the Vestoids as the major HED source.

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Figure 9. Exposure age distributions of HED meteorites (howardites, eucrites, diogenites) and aubrites or enstatite achondrites. The HED data are from the compilation by Welten et al. (1997), with only the (shielding corrected) 81Kr-Kr ages being displayed for the eucrites. Included are 4 new diogenite ages from Welten et al. (2001b). In addition, the ~3 Myr age of the howardite Kapoeta has been added (Caffee and Nishiizumi 2001). For data sources of aubrites see text. Note the different abscissa scales between the lowermost panel and the others.

Aubrites (enstatite achondrites). This class is similar in many respects to enstatite chondrites, although the two classes are thought to derive from different parent bodies (Keil 1989). In the bottom panel of Figure 9, the exposure ages of aubrites are shown. These ages are more uncertain than those of chondrites, because production rate systematics have been less well studied for this rare class. The ages shown here rely on published cosmogenic Ne data (Schultz and Franke 2000) and the production model by Leya et al. (2000a), or, in a few cases, on 81Kr-Kr analyses (Miura et al. 2000). In addition, we include ages given for eight aubrites from Antarctica by Lorenzetti et al. (2001). In the first systematic study of the cosmic ray record of aubrites, Eberhardt et al. (1965) already noted that aubrites have unusually long exposure ages among stony meteorites. This has become even more accentuated as the production rates we adopt here are lower than those used by Eberhardt and coworkers. Note on the other hand that the exposure age of the Norton County aubrite of 240-280 Myr reported by Begemann et al. (1957) has later been revised downwards and is shown here as only 130 Myr. Yet, it is quite remarkable that this meteorite with the first ever published cosmic ray exposure age still holds the record among stones, within uncertainties together with Mayo Belwa, another aubrite (Lorenzetti et al. 2001). Eberhardt et al. (1965) noted a cluster at 40-50 Myr, but uncertainties in production rates are too large to confirm this cluster in Figure 9 (at roughly 50-60 Myr with present-day production rates). Nevertheless, exposure ages of aubrites clearly are higher than those of most enstatite chondrites, consistent with the conclusion that aubrites derive from a separate parent body. The long interplanetary journeys of the fragile aubrites also indicate a distinctly different orbital evolution compared to other meteorite classes, resulting in longer survival times against ejection from the solar system or collision with a planet, as already pointed out by Eberhardt and coworkers. Acapulcoites and lodranites. These meteorite classes, also called primitive

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achondrites, are believed to derive from the same parent body and are residues from partial melting of chondritic precursors (McCoy et al. 1997). Exposure ages of both classes show a remarkably tight clustering between 4 and 10 Myr, in the same range as the prominent 7 Myr cluster of the H chondrites (Weigel et al. 1999; Terribilini et al. 2000a; Ma et al. 2001). Terribilini et al. conclude that this may be due to one impact (with subsequent secondary break-ups of a larger meteoroid) or several impacts closely spaced in time. They point out that the coincidence of exposure age clusters of acapulcoites/lodranites and H chondrites might suggest an enhanced collisional activity in the asteroid belt some 7 Myr ago. Ureilites. Goodrich (1992) and Scherer et al. (1998) compiled exposure ages of ureilites. Values range between 0.1 and 34 Myr. Again, the majority of the ureilites, i.e., >70% of the 22 meteorites compiled by Scherer et al. (1998), have ages above 3 Myr. No exposure age clusters are observed. However, since many ureilites contain relatively large amounts of primordial Ne, a shielding correction via 22Ne/21Ne is often not possible. Therefore, exposure ages of ureilites often have a quite high uncertainty, and so clusters, even if present, might be less easily recognized than in other meteorite classes. Mean activities of 10Be and 26Al in ureilites are ~20% lower than expected, indicating that many ureilites had a small preatmospheric size (Aylmer et al. 1990). Iron meteorites and stony-irons. Reliable exposure ages for iron meteorites are often even more difficult to obtain than for stony meteorites, because quite a few iron meteorites had preatmospheric sizes of one to several meters, so that large production rate variations due to variable shielding are common. As a somewhat extreme example, 3He concentrations in fragments of the R ~15 m Canyon Diablo meteorite vary by a factor of 10,000 (Heymann et al. 1966), and even in the 13 most gas-rich samples analyzed by these authors, He, Ne, and Ar concentrations varied by ~20 to 35 times. Reliable corrections for production rates over a wide range of shielding are therefore a prerequisite in exposure age studies of iron meteorites. Considerable efforts are being spent to achieve this, because iron meteorites often have much higher exposure ages than stony meteorites and therefore are particularly important for studying possible long term variations of the flux of the galactic cosmic radiation. Three main methods are He-Ne-Ar, 40K-41K, and 36 Cl-36Ar. The relationship between noble gas production rates in iron meteorites and shielding parameters such as 4He/21Ne has been described by Signer and Nier (1960, 1962) using a semiempirical model. The 40K-41K method, involving the very long-lived radionuclide 40K (half-life 1.25 Gyr) has been developed by H. Voshage (Voshage 1978 1984, and references therein) and the 36Cl-36Ar method has recently systematically been applied to iron meteorites by Lavielle et al. (1999a, 2001). These latter studies are also discussed in section The Cosmic Ray Flux in Time. Figure 10 shows the exposure age distribution of iron meteorites. Ages >200 Myr are based on the 40K-41K method (Voshage 1978), while younger ages are based on 38Ar (Lavielle et al. 1985), because few K-ages exist for the age range 0-200 Myr. The Figure may thus well be biased somewhat due to sample selection, possibly underestimating the fraction of iron meteorites with relatively low-exposure ages. The abscissa may also be compromised by a possible GCR intensity change as discussed in section The Cosmic Ray Flux in Time. Nevertheless, it is clear that the majority of iron meteorites has exposure ages exceeding the highest known age of a stony meteorite of some 130 Myr or so. Two clusters of exposure ages have been recognized for a long time, indicating major collisions on the IIIAB and IVA iron meteorite parent bodies about 650 Myr and 375 Myr ago, respectively, using the 40K-age scale (Voshage 1978; Lavielle et al. 1985). Lavielle et al. (2001) proposed that the 375 Myr peak actually represents two events at 255 and 207 Myr ago in the 36Cl-scale, whereas the 650-Myr peak corresponds to an event at 460 Myr in this scale (Lavielle 2001, pers. comm). In contrast to other iron

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Figure 10. Exposure age distributions of iron meteorites and mesosiderite stony-irons. Iron meteorite data > 200 Myr were obtained by the 40K-41K method (Voshage 1978), ages < 200 Myr are based on 38 Ar (Lavielle et al. 1985). Iron meteorite groups IVA and IIIAB show exposure age clusters around 375 and 650 Myr (40K-scale) and are shown separately. The 375 Myr cluster appears to be composed of two independent peaks not shown here (at 207 and 255 Myr in the 36Cl-scale, see text). Mesosiderite ages (36Cl-36Ar) are from Albrecht et al. (2000), as updated from Begemann et al. (1976). Note the widely different abscissa scales of the two panels.

meteorites, the two classes IIA and IIE show exposure age ranges not unlike those of stony meteorites. Ages of IIA and IIE irons may be as low as a few Myr (Olsen et al. 1994), with the most reliable among the lowest values being around 10 Myr (e.g., Rafrüti: Terribilini et al. 2000b; Kodaikanal, Braunau, and Watson: Lavielle 2001, pers. comm.). A study by Begemann et al. (1976), updated in Albrecht et al. (2000), determined Cl-36Ar ages of 16 stony-iron meteorites, mostly mesosiderites (Fig. 10). The values fall between ~10-180 Myr (10-160 Myr for mesosiderites alone), with about a third of the values being larger than 100 Myr. The age range for stony-irons is thus broadly in between those of stones and most irons. Welten et al. (2001c) note a hint for an exposure age cluster around 60-70 Myr. 36

The comparatively very old exposure ages of iron meteorites, and perhaps also the rather high ages of stony irons, are probably partly the result of their mechanical strength and hence long survival times against collisional destructions. Long collisional lifetimes could also explain why iron meteorites probably derive from considerably more parent bodies than stony meteorites. Petrographic and chemical evidence suggests that there are 90 or so different original iron meteorite parents (Wasson 1995), and a large number of them are sampled only by relatively small meteorites (Wasson 1990). Lunar and Martian meteorites. At the time of this writing, our collections contain about 20 individual meteorites (corrected for pairings) from the Moon and 18 from Mars. A list of the Martian meteorites is given by Swindle (2002b), who also briefly discusses their cosmic ray record. Exposure ages of Martian meteorites are also compiled and

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discussed in detail by Nyquist et al. (2001). Cosmogenic noble gases and radionuclides are a major tool for deciding whether different meteorite specimens, e.g., found in a desert nearby to each other, belong to the same fall or whether different meteorites from any find location were ejected by the same event (source crater pairing). Quite as expected, most meteorites from the nearby Moon have had considerably shorter transit times than the meteorites from the asteroid belt discussed so far. Most exposure ages of lunar meteorites are below 1 Myr, and many even below 0.1 Myr. Very remarkable is the short transit time of Dhofar 026 of at most 4 Kyr (Nishiizumi and Caffee 2001). The highest age found so far is ~8 Myr (Polnau and Eugster 1998; Nishiizumi et al. 1999). Very few, if any, lunar meteorite falls are source crater paired (Warren 1994). Essentially all lunar meteorites have a complex exposure history, i.e. they also contain cosmogenic nuclides acquired on the Moon (e.g., Warren 1994; Polnau and Eugster 1998). This means that most lunar meteorites were ejected in relatively minor cratering events from within the uppermost very few meters of the lunar surface, where they had previously spent up to several hundred million years. Nishiizumi et al. (1999) note a correlation between the depth of ejection and the transit time to Earth, with meteorites from smaller events arriving earlier. Martian meteorites had travelled for considerably longer than lunar meteorites, with transit times of between 0.7 and 20 Myr (Nyquist et al. 2001). Unlike for lunar meteorites, source crater pairing is common for Martian meteorites. The number of individual ejection events on Mars needed to account for the available meteorites is controversial, however (see also Swindle 2002b). Based on crystallisation ages, four to five impacts were proposed by Nyquist et al. (1998) (see also Nyquist et al. 2001). On the other hand, exposure ages suggest 7-9 events (Nyquist et al. 2001, Nishiizumi et al. 2000; K. Nishiizumi, pers. comm. 2000). Nyquist et al. (2001) discuss scenarios involving secondary collisions to reconcile the two lines of evidence. Since, unlike lunar meteorites, none of the Martian meteorites show signs of a complex exposure (Warren 1994; Nyquist et al 2001; however see also Hidaka et al. 2001), the bodies originally launched from Mars would have to have been quite large. The apparently single-stage exposure history also indicates that all Martian meteorites were launched from at least a few meters below the Martian surface (Warren 1994). A caveat to this conclusion is that a parent body exposure stage is more difficult to detect in meteorites having had a longer subsequent exposure during their transit to Earth. However, Martian meteorites on Earth are usually considerably larger than their lunar counterparts and hence probably also had a larger preatmospheric size, which also suggests that they stem from larger events than lunar meteorites. This is indeed expected given the different escape velocities from the two planets. Micrometeorites and interplanetary dust particles (IDPs). The noble gas record in micrometeorites and IDPs is discussed by Wieler (2002). Cosmic ray exposure ages of micrometeorites from Greenland (Olinger et al. 1990) and of deep-sea spherules (Raisbeck and Yiou 1989) are between < 0.5 and 26 Myr, suggesting a cometary origin of most of these particles (Raisbeck and Yiou 1989). Cosmic ray exposure ages of IDPs are uncertain. Pepin et al. (2001) report large 3He excesses in some IDPs which, if cosmicray-produced, would imply GCR ages of up to more than one Gyr, much larger than orbital lifetimes of small particles. The authors consider that this might be due to very long exposures in the regoliths of Kuiper Belt comets, but they also discuss the possibility that the excess 3He is primordial (see Wieler 2002). Exposure ages and dynamical models of meteorite delivery Here we discuss constraints imposed by the observed distributions of exposure ages on ideas and models of how meteorites get delivered to Earth from the asteroid belt, the

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Moon and Mars (see also Wieler and Graf 2001). Stony meteorites from the asteroid belt. It has been known for a long time that ~15% of all meteorites falling today were produced in one—or more probably two— collisions manifested in the prominent 7 Myr peak in the exposure age histogram of the H chondrites. It has often been thought, however, that this was an exceptional event and that a quasi continuous production of meteorites in many smaller events is the rule. We have seen above that to some extent this discussion goes on today, as it remains open whether most of the meteorites not obviously belonging to a major peak in an exposure age distribution should be assigned to the “background” or to one of a few additional events, and how many of even these meteorites should be shifted to one of the large peaks. Furthermore, for all but the most common meteorite classes, even the quite impressive data base available today still often does not allow an unambiguous recognition of peaks. Nevertheless, exposure age peaks are identified or suspected in the histograms of many classes, which allows us to reiterate a conclusion already made by Anders (1964): a large part of the meteorites were produced in a few distinct major collisions. The quantification of this statement is somewhat difficult, but we note that more than 50% of all H chondrites derive from a few events, whereas small iron meteorites appear to be due mostly to a quasi continuous production. It also has long since been recognized that orbital resonances with planets are a major factor determining the delivery of meteorites from the asteroid belt (Wetherill and Williams 1979; Wisdom 1983; Gladman et al. 1997; Farinella et al. 2001). Until recently, a common assumption was that the velocity of meteorites ejected from their parent asteroids needs to be sufficiently high to place them into a resonance, e.g., the 3:1 Kirkwood gap in the main asteroid belt, where an object orbits three times around the Sun in the time it takes Jupiter to complete one orbit. This has led to the hypothesis that only large collisions on parent bodies close to a resonance could produce large numbers of meteorites on Earth. This would qualitatively explain the distinct peaks in exposure age histograms, but it might be difficult to account for the observed flux of meteorites with the limited number of viable parent bodies (see discussion by G.W. Wetherill in Greenberg and Nolan 1989). Recent dynamical models now indicate that objects placed into a so-called chaotic zone in a resonance will have a much shorter dynamical lifetime than has previously been thought. Within a very few million years they will either collide with the Sun or be ejected from the solar system (Gladman et al. 1997). As these times are considerably less than typical exposure ages of almost all meteorite classes, in this scenario most meteorites need to spend most of their lifetime as meter-sized objects in the main belt, slowly drifting into a resonance (Morbidelli and Gladman 1998). It has recently been realized that the asymmetric radiation of thermal energy, the so-called Yarkovsky effect, can provide enough momentum to allow meter-sized objects to drift into a resonance on timescales comparable with meteorite exposure ages (e.g., Bottke et al. 2000). The Yarkovsky effect appears to be able to explain also the paucity of meteorites with exposure ages less than a few million years (Bottke et al. 2000). If so, asteroids throughout large parts of the main belt may be potential meteorite parent bodies. This, however, is only compatible with the marked peaks in the exposure age distributions, if a very few ones among these many potential parent asteroids contribute a major part of all meteorites. These selected bodies presumably should be among the largest asteroids. Wieler and Graf (2001) discuss some evidence that this might indeed be the case. On the other hand, the original H chondrite parent body or parent bodies, responsible for a large fraction of all meteorites today, were rather small, with a radius of only about 100 km (Pellas and Fiéni 1988). Bottke et al. (2000) also discuss this problem. Clearly, exposure age distributions with improved temporal resolution will be needed to further rule on the importance of the Yarkovsky effect as meteorite delivery mechanism.

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Even though meteorites may be brought to Earth quickly from within a resonance in the main belt, it is not clear whether the 0.2-Myr peak of the CM chondrites can be explained this way. This would require a very large collision on a CM parent body in a position where the time needed to drift into a resonance would be extremely short and from which we would see today perhaps just some forerunners. It appears more likely that the 0.2 Myr CM chondrites are the products of a collision on a parent body already on an orbit crossing that of the Earth (Caffee and Nishiizumi 1997). Iron meteorites. Meter-sized iron meteorites have a slower Yarkovsky drift rate than stones of the same size (Farinella et al. 2001), which is in qualitative agreement with the much older exposure ages of iron meteorites (see also Bottke et al. 2000; Morbidelli and Gladman 1998). The old ages lead to the expectation that many more collisions are recorded in the iron meteorite histograms compared to those of stony meteorites, consistent with the rareness of peaks in the former. The comparatively large number of iron meteorite parent bodies can be explained if iron meteorites are able to drift into a resonance from even larger portions of the belt than stones, because of their longer collisional lifetimes due to their mechanical strength. Wasson (1990) proposed that small iron meteorites move faster through the belt than large ones due to their higher ejection velocity, which would account for the high number of parent bodies sampled by small iron meteorites. Lunar and Martian meteorites. Gladman et al. (1996) calculate that most lunar meteorites that end up on Earth do so within several 104 years. Fragments escaping this fate will reach orbits outside the Earth's immediate influence after about a million years. The calculations are thus essentially in agreement with the observations. Gladman (1997) calculates transfer times from Mars in agreement with the observed exposure ages for material ejected just slightly above escape velocity. Martian meteorites travel much longer than lunar ones simply because the orbits of Martian ejecta will not cross Earth's orbit initially. Given the closeness of the Moon, it appears astonishing that we do not have many more lunar than Martian meteorites. Gladman (1997) notes, however, that the higher fraction of escaped lunar meteorites that end up on Earth and the higher number of fragments per impact on the Moon may be largely offset by higher cratering rates on Mars and the fact that many Martian meteorites are source-crater-paired. It is thus not quite clear whether or not the similar numbers of lunar and Martian meteorites actually constitutes a puzzle. If it does so, this may be explained by mechanical properties of the two planetary surfaces. The low porosities of lunar meteorites suggest that compaction of loose regolith and simultaneous ejection may not work so that only the small fraction of already coherent rocks on the lunar surface may have a chance to end up as meteorites (Warren 2001). COMPLEX EXPOSURE HISTORIES The term complex exposure is commonly used to denote a situation where an entire meteorite acquired part of its cosmogenic nuclides detectable today in a larger body than the one immediately prior to atmospheric entry. A complex exposure history thus indicates a break-up of a precursor meteoroid or an ejection of a meteoroid from a nearsurface location of a parent asteroid or planet. In the first case we talk about two different so-called 4π exposure stages, in that during both stages each sample saw cosmic rays from all directions in space, while in the second case a 2π irradiation is followed by the 4π exposure. It may also occur, however, that only certain parts of a meteorite (e.g., clasts or individual mineral grains) show excesses of cosmogenic nuclides relative to other fractions of the meteorite. This is commonly called pre-compaction exposure, and the precompaction stage may have occurred early in solar system history or later in a dust and gravel layer on the parent body surface, the regolith.

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Thus, meteorites with a complex exposure history may help us to constrain collisional dynamics in the asteroid belt, the dynamics of asteroidal regoliths, and perhaps the energetic particle environment in the early solar system. On the other hand, we mentioned above that complex exposure histories often hinder our efforts to determine exposure age distributions. Sometimes, a complex exposure history is quite easily recognized, for example if cosmogenic noble gases indicate an exposure during several ten million years in an average-sized meteorite, whereas the 10Be activity is lower than the saturation value of medium-sized meteorites. Often, however, analyses of many different nuclides in different samples, if possible from well defined positions relative to each other, are required to unambiguously decide whether a meteorite had a simple or complex history. Nuclides such as 60Co with its steeply rising production rate with preatmospheric depth (e.g., Heusser et al. 1996; see Fig. 1) or nuclear tracks with their steeply declining production rate with preatmospheric depth (Bhattacharya et al. 1973) are particularly useful but such data are rarely available (and 60Co has decayed completely a few decades after the fall of a meteorite). It should thus not come as a surprise that only for few meteorites a complex exposure history has been unambiguously documented or at least found to be very likely. Vogt et al. (1993) and Herzog et al. (1997) list just 15 stony meteorites with a probable or certain complex history. It is even more difficult to exclude a complex exposure history (a prominent example is Knyahinya, where from numerous analyses a single stage lasting ~40 Myr has been deduced; Graf et al. 1990b). Therefore, the fraction of meteorites with a complex history cannot be estimated with confidence. Wieler and Graf (2001) guess that it could be around 30% for stony meteorites. Similarly, Lavielle et al. (1985, 2000) note that about one-fourth of all iron meteorites show indications of a complex history. Some meteorites with a well documented complex history are Jilin, Bur Gheluai, Tsarev, Torino, Gold Basin and QUE93021 (Begemann, et al. 1996; Vogt et al. 1993; Wieler et al. 1996; Welten et al. 2001d,e). The second stages lasted on the order of 1-15 Myr and the first stages sometimes up to a hundred or a few hundred million years. The size of the first stage body is often unconstrained, but very long first stages seem to require an asteroidal parent body. Meteorites with nominally short exposure ages appear to often show a discernible complex history (Herzog et al. 1997; Merchel et al. 1999). This is not too surprising, because a first stage contribution is easier to detect if the subsequent second stage lasted only briefly. In these examples, first stages typically lasted on the order of 10 Myr, second stages only around a few hundred thousand years. This does not mean, however, that very low 4π exposure ages are more common than it may appear from the exposure age histograms, because the first stage may already have occurred in a meteorite-sized body rather than on a parent asteroid. This is illustrated by the H-chondrite Jilin, which is a good example of how complex histories smear out peaks in exposure age histograms. Jilin was probably ejected in one of the 7-Myr events and broken up in a second collision 0.3 to 0-6 Myr ago, whereas calculated single stage ages are between 1-5 Myr only, depending on the sample analysed (Begemann et al. 1996; Heusser et al. 1996). We mentioned above that a large part of the few H chondrites with 104 years in Antarctica), and “inherited” 40Ar, radiogenic 40Ar that was incompletely degassed when the rock formed.

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Paleoatmosphere On Earth, after decades of searching, it is not clear that anyone has ever found a sample of paleoatmosphere with noble gas isotopic ratios distinct from the modern ratios. For Mars, with only 20 samples to work with, a strong case can be made that paleoatmosphere has been found. To preserve a sample of paleoatmosphere, a rock would have to have trapped some atmosphere long enough ago that the noble gas isotopes in the atmosphere had been able to evolve significantly in the interim. On Earth, that is difficult. Most rocks are not that old. Many rocks that are old enough (e.g., Archean) are not the kinds of rocks that would have trapped atmosphere within them, since virtually the only mechanisms on Earth whereby atmospheric trapping occurs are those operating during sedimentation. Finally those rocks that might have had an opportunity to trap an isotopically distinct atmosphere have usually been metamorphosed or have suffered through other processes that could have caused them to lose that atmosphere. On Mars, ALH84001 is a rock that is more than 4 Ga old (old enough), it apparently suffered several shock events early in the history of Mars (which could have implanted some atmospheric gas), but it has apparently been altered little in the last 4 Ga, since Mars lacks plate tectonics and, to a large extent, the action of liquid water. But does it have paleoatmosphere? The identification has largely been made based on expectations of what the Martian atmosphere might have been like 4 Ga ago. A definitive identification would require some demonstration that the identified component was implanted 4 Ga ago, which would be difficult. The highest 129Xe/132Xe in ALH84001 is about 2.16, 10% lower than measured in the glasses in the shergottites. In addition, the 129Xe/136Xe/132Xe systematics are not consistent with a combination of current Martian atmosphere, a Martian interior component like that identified in Chassigny, and terrestrial contamination. On this basis, the Xe has been argued to be an ancient component, rather than modern atmosphere (Gilmour et al. 1998; Mathew and Marti 2001; Murty and Mohapatra 1997). In addition, the 15N/14N and 40Ar/36Ar ratios in ALH84001 are not as high as in the modern Martian atmosphere. Since those ratios are expected to rise over time as a result of atmospheric loss and outgassing, respectively, that could mean that ALH84001 has trapped Martian paleoatmosphere that is not as evolved (Grady et al. 1998; Mathew and Marti 2001; Murty and Mohapatra 1997). Unlike the case for Xe, different authors have come up with different isotopic ratios for the “unevolved” N and Ar. For δ15N, the values obtained have been +7‰ (Mathew and Marti 2001) or ≤+200‰ (Grady et al. 1998) or ≥+46‰ and 60 Myr) and substantial (>97%) from throughout the planet.

where 136Y244 is the fission yield of 136Xe. Figure 8 shows the fraction f of retained 130Xe, plotted as a function of the time of loss, for various present values of 129*Xe/136*PuXe. It can be seen that for each value of 129*Xe/136*PuXe, there is a minimum time of loss, which corresponds to very large depletions; late losses correspond to the maximum fraction of Xe retained. For the atmosphere, 129*Xe/136*PuXe = 4.4, and so the minimum time of loss is 95 Ma, with losses >99%. Therefore, the Earth or Earth-forming materials initially contained at least 102 times the 130Xe presently seen in the atmosphere. Loss of Xe during

Figure 9. Xe isotope compositions of terrestrial precursors and present reservoirs. U-Xe, fractionated to match the light Xe isotopes (Fig. 7), provides an initial composition of the atmosphere, to which radiogenic 129Xe and fissiogenic 136Xe have been added. Similar fractionation of solar wind (SW) Xe produces Xe that is too heavy to supply the atmosphere. MORB typically have 129Xe/130Xe and 136Xe/130Xe ratios that are greater than the atmosphere due to radiogenic and fissiogenic additions and are correlated, with the range likely due to variable atmospheric contamination of samples. The most precise measurement of mantle Xe is for CO2 well gas. When uranogenic Xe is subtracted, the well gas falls substantially below the MORB data, and indicating that the ratio of 244Puderived 136Xe to I-derived 129Xe is lower than that of the atmosphere.

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a single, brief event may have been driven by a Moon-forming impact. While more complicated loss histories can be constructed, the general conclusion that very large losses of Xe occurred is inescapable. Radiogenic Xe in the mantle. Data for MORB indicate that the mantle has Xe/130Xe and 136Xe/130Xe ratios greater than those of the atmosphere (Fig. 9). There is considerable uncertainty and debate regarding how much of the 136Xe excess is due to 244 Pu, rather than 238U, decay. The most recent data suggest that ~30% is due to 244Pu decay (Kunz et al. 1998), with considerable uncertainty due to scatter in the data. This is also consistent with measurements of mantle-derived Xe in well gases that indicate m1 (constituents 2), leads to escape fluxes F2 of the minor, heavier constituents given by F2 = F1

N2 N1

⎡ m c − m2 ⎤ ⎢ ⎥ ⎣ mc − m1 ⎦

(2)

for mC ≥ m2. The critical or “crossover mass” mC in Equation (2), representing the smallest mass for which the escape flux F2 of a particular mass m2 in constituent 2 is zero, is defined (Hunten et al. 1987) as m c = m1 +

kTF1 bgX1

(3)

where k = Boltzmann’s constant, T = atmospheric temperature, g = gravitational acceleration, X1 is the mole fraction of the abundant constituent 1 (assumed to remain near unity throughout the escape episode), and b is the diffusion parameter (the product of diffusion coefficient and total number density) of mass m2 in the constituent 1 gas. Note from Equation (3) that a large F1 yields a large crossover mass mC; for F1 high enough that mC >> m2 > m1, losses described by Equation (2) are maximum (atmospheric “blowoff”) and species in constituent 2 escape close to the ratios N2/N1 of their atmospheric inventories—i.e., without fractionation. Maximum fractionation occurs when F1 drops to the level where mC approaches m2. Constituent 1 is usually assumed to be molecular hydrogen. Constituent 2 is taken to be a multicomponent noble gas mixture—excluding He, which is only weakly bound in terrestrial planet atmospheres and escapes readily by other processes. The diffusion parameter b in Equation (3), and thus the crossover mass mC, differs for different elements S in constituent 2. If mC(S) is known (or assumed) for one element, say Xe, then Equation (3) for both Xe and S yields ⎡ b( Xe ) ⎤ m c (S) = m 1 + [m c (Xe ) − m1 ] ⎢ ⎥ ⎣ b( S ) ⎦

(4)

for the S crossover mass, where S represents Kr, Ar or Ne. Values of b(S) for noble gas diffusion in H2 at various temperatures are known (Mason and Marrero 1970; Zahnle and Kasting 1986). They increase from Xe to Ne by a factor of approximately 2 that varies only slightly with temperature, and are identical or nearly so for isotopes of a given element. The energy required for escape of a particle with mass m1 from its local gravitational field, at radial distance r ≥ rS from a body of mass M and radius rS, is Gm1M/r ergs per particle. If the global mean solar EUV input at heliocentric distance R and time t is φ(R,t) ergs cm-2 sec-1, the energy-limited escape flux is F1 (R , t) =

rφ(R , t)ε φ( R, t) ε = Gm 1M / r Gm 1M

particles cm −2sec −1

(5)

where ε is the fraction of incident EUV energy flux converted to thermal escape energy of m1. Energy input φ may be expressed in terms of the current mean EUV flux at Earth φ ⊕ (tp ) , heliocentric distance R, and the ratio ΦEUV of the flux at past time t to that at present time tp. Defining φ(R, t p ) = (R ⊕ / R)2 φ⊕ (t p ) and Φ EUV (t) = φ(R, t) / φ( R, t p ) yields φ(R, t) = (R ⊕ / R) 2 φ⊕ ( t p )Φ EUV (t) . Substituting this into Equation (5) and combining Equations (5) and (3), with g = GM/r2 and M expressed in terms of rs and planetary density ρ, yields

Origin of Noble Gases in the Terrestrial Planets ⎛ 3 ⎞ ⎛ 1 ⎞ ⎛ R ⊕ ⎞ ⎛ kT ⎞ ⎛ φ ⊕ (t p ) ε ⎞ ⎛ r ⎞ ⎟ ⎜ ⎟⎜ m c ( S) = m 1 + ⎜ ⎟ ⎜ ⎟ ⎜ ⎟⎜ ⎟ ⎝ 4πGρ ⎠ ⎝ rs ⎠ ⎝ R ⎠ ⎝ X1 ⎠ ⎝ b( S)m 1 ⎠ ⎝ rs ⎠ 2

3

2

221

3

ΦEUV (t)

(6)

relating crossover masses of minor species S to the EUV flux irradiating a planetary atmosphere at heliocentric distance R. It can be seen here that once the history of the driving energy source for loss, i.e., the EUV flux, is known, and assumptions are made about the initial inventory and ongoing supply of the major volatile species, H2, the losses of each species can be calculated. Information on what ΦEUV(t) might have been in the early solar system comes from astronomical observations of radiation from young solar-type stars at various stages of pre- and early-main-sequence evolution. Since early solar EUV radiation could not have penetrated a full gaseous nebula to planetary distances, the applicable time dependence of stellar activity in the present model is that following dissipation of the dense accretion disks surrounding the classical T-Tauri stars at stellar ages of up to ~10 Ma (Walter et al. 1988; Strom et al. 1988; Walter and Barry 1991; Podosek and Cassen 1994). Among present observational data, soft (~3 to 60 Å) X-ray fluxes are most likely to be representative of at least the short-wavelength coronal component (λ < 700 Å) of the EUV spectrum. Figure 10 shows X-ray luminosities LX, relative to that of the present sun, from observations of T-Tauri stars without disks (the “Weak T-Tauri Stars”: Walter et al. 1988) and older solar-type main-sequence field stars (Simon et al. 1985), plotted against estimates of stellar ages (Walter and Barry 1991). Although considerable scatter from a single functional dependence of ΦX on age is evident, most of the data for t between ~50 and 200 Ma do indicate a decline by factors of ~5-10 from levels at ~20 Ma. In current escape models (Pepin 1991, 1994, 1997), ΦEUV is assumed to fall off exponentially, with a mean decay time τ of 90 Ma, along the dashed curve in Figure 10; the solid portion of the curve identifies the period of interest for EUV-driven loss of planetary atmospheres. An alternative, power law function for ΦX(t) of the type suggested by Feigelson and Kriss (1989) is also plotted. Models of hydrodynamic escape analogous to those for exponential flux decay have been constructed for this kind of power law dependence (Pepin 1989). Results are similar enough to indicate that the particular mathematical form of the decay in stellar flux through the period of atmospheric evolution is not centrally important. Decline of the EUV flux over the first few hundred Ma of solar history, by roughly the factor suggested by present astronomical data, is a crucial requirement of the modeling because, from Equation (6), it results in declining crossover masses to the level where fractionations described by Equation (2) and calculated below are large enough to match Xe isotopic distributions in current atmospheres. In the case where energy deposited in an atmosphere by EUV radiation or from some other source declines exponentially with time, the flux F1 of constituent 1 is given by ⎡ −(t − t o ) ⎤ F1 = F1o exp ⎢ ⎣ τ ⎥⎦

(7)

where F10 is the initial escape flux at t0 and τ is the mean decay time of the energy source. Equation (3) then requires the crossover mass mC(S) to decrease as well, from its initial value mC0(S) at time t0 and flux F10 to its final value mCf(S) at time tf when escape either terminates due to hydrogen depletion or undergoes a transition to diffusion-limited hydrogen escape through a atmosphere of increasing molecular weight; in this latter case only Ne and perhaps Ar will continue to be lost (see Zahnle et al. 1990a). If mC(S) declines through the mass m2 of a trace species S, loss of m2 ends when mC(S) = m2 but escape of lighter isotopes or elements continues for as long as the mC(S) for each specific constituent is greater than its particular mass.

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Figure 10. Observed X-ray luminosities LX relative to the present sun vs. stellar ages for "weaklined" T-Tauri stars (WTTS) and solar-type main sequence stars (MSS) (Pepin 1991). All WTTS data are from Walter et al. (1988), MSS data from Simon et al. (1985) and Walter and Barry (1990). Numbers adjacent to the plotted WTTS points are stellar mass estimates in tenths of a solar mass. Underlined symbols identify measurements that yielded only lower limits on luminosity. In the models explored in the text, both soft X-ray (ΦX) and EUV (ΦEUV) luminosity enhancements are taken to decline exponentially with a mean decay time τ = 90 Ma over the period of interest for planetary atmospheric evolution (indicated by the solid portion of the plotted curve). A ΦX ∝ (1/t)s power law decay with s = 0.6 (Feigelson and Kriss 1989) is shown for comparison.

Fractionating effects of the escape process can now be calculated analytically if specific assumptions are made about the time dependence of the major (constituent 1) inventory N1—that it is either replenished as fast as it escapes (constant inventory model), or is lost without replenishment along with the minor atmospheric species (Rayleigh fractionation model). In both cases the inventories N2 of minor species, here the noble gases, are assumed to be in the atmosphere at t0 and are lost without replenishment during the escape episode. For Rayleigh fractionation, adopted for this discussion, Equations (2), (3), and (7) and the definitions F1 = -dN1/dt and F2 = -dN2/dt may be combined and integrated, in the limit of X1 ≅ 1, mC0 > m2 > m1, and mCf ≥ m2, to yield ln

Nf2 N1f αo ⎡ m 2 − m1 ⎤ ⎡ Nf1 moc − m1 ⎤ ⎢ ⎥ ln ⎢ ⎥ o = ln o + N2 N1 1 − α o ⎣ m oc − m 1 ⎦ ⎣ N o1 m 2 − m 1 ⎦

(8)

for the fractional depletion of a particular minor species of mass m2 (Hunten et al. 1987;

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Pepin 1991). N10,N1f and N20,N2f are the initial and final inventories of the major and minor constituents. The dimensionless parameter αo is defined as αo = F10 τ/N10. The fractional retention of constituent 1 at tf is given by ⎡ m oc − m fc ⎤ Nf1 =1 − α ⎥ o⎢ o N o1 ⎣ m c − m1 ⎦

(9)

(Pepin 1991). From the definitions of mc (Eqn. 3) and αo, the initial inventory of constituent 1 is N o1 =

bgX1(m oc − m 1 ) τ particles cm − 2 . kTα o

(10)

The total duration tf – to of the escape episode is related to mean energy decay time and the initial and final crossover masses by ⎡ m f − m1 ⎤ t f − t o = −τ ln ⎢ co ⎥. ⎣ m c − m1 ⎦

(11)

Note that equations containing t involve only time differences, not their absolute values (e.g., Eqns. 7 and 11). Consequently the fractionations of initial isotopic and elemental ratios generated by hydrodynamic escape do not depend on any specific choice for t0, the time in solar evolutionary history when atmospheric escape begins. Permitted values of t0 are constrained, however, by the solar EUV flux needed to drive an escape episode. For example, as noted in the following section, EUV-driven Xe loss from Earth requires a flux that exceeds the present solar level by a factor of ~450, and thus a t0 no later than ~100 Ma if the flux history follows the τ = 90 Ma exponential in Figure 10. It is assumed that t0, whatever its value, marks the time at which dust and gas in the nebular midplane had cleared to the extent that solar EUV radiation could penetrate to planetary distances, so that EUV-driven atmospheric loss would not have occurred prior to t0 (Prinn and Fegley 1989). Isotopic and elemental fractionations generated in constituent 2 residues are calculated from separate solutions of Equation (8) for particular choices of mass m2 and a reference mass [m2]ref. Then (N2/N20)/([N2/N20] ref) = {N2/[N2] ref}/{N20/[N20] ref} is the corresponding fractionation factor. For example, if S is xenon and the inventories N2(S) and [N2(S)]ref are respectively taken to represent 124Xe and 130Xe, the isotopic fractionation is (124Xe/124Xe0)/(130Xe/130Xe0) = (124Xe/130Xe)/(124Xe/130Xe)0 where (124Xe/130Xe)0 is the original (unfractionated) ratio. Adjustable modeling parameters include the mass m1 of the abundant light constituent (taken to be H2 in most models); atmospheric temperature T; initial crossover mass mC0 for one noble gas—for example, for mC0 (Xe) as in Equation (4)—and its lower value mcf for the same species when the loss episode terminates (crossover masses for the remaining trace gases are fixed by Eqn. 4), which with m1, T, and b(S) sets the initial value of F1 via Equation (3); the decay constant τ for the thermal energy supply; the parameter α0 = F10τ/N10 which for a given τ sets the initial ratio of escape flux to inventory for H2; and the isotopic and elemental ratios of noble gases in the initial primordial atmosphere. It’s important to note that allowed parameter space is not unconstrained. In Equation (9), for example, the light constituent abundance N1 cannot be negative; N1f/N10 must therefore be ≥ 0, which limits the permitted combinations of α0 and mCf values for a particular choice of mC0. With suitable choices of parameter values, hydrodynamic escape of hydrogen-rich primary atmospheres and outgassed volatiles from the terrestrial planets (plus sputtering

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losses on Mars), operating in an astrophysical environment for the early solar system inferred from observation of young star-forming regions in the galaxy and from models of planetary accretion, can account for most of the known details of noble gas distributions in their present-day atmospheres. The question, of course, is whether the required values or ranges of values required for the free modeling parameters reflect realistic astrophysical and planetary conditions in the early history of the solar system. The conditions required on each planet for obtaining the observed noble gas characteristics are described in the following sections. Application to Earth. Current modeling requires the Earth to have acquired two isotopically primordial volatile reservoirs during accretion, one in the planet’s interior, perhaps populated by a combination of nebular gases occluded in planetary embryo materials and dissolved in molten surface materials, and the other coaccreted as a primary atmosphere degassed from impacting planetesimals or gravitationally captured from ambient nebular gases during later planetary growth. These “isotopically primordial” reservoirs are characterized by isotope ratios measured in the solar wind, with the important exception of Xe; here the U-Xe composition, which differs substantially from solar wind Xe at the two heaviest isotopes, is required (see Fig. 7 and discussion in the “Terrestrial noble gases” section). Evolutionary processing proceeds in two stages. In the first stage, substantial depletion of Xe from the primary atmosphere occurs, driven by deposition of atmospheric energy. This sets the Xe inventory and generates the extensive Xe isotopic fractionation that is presently observed in the terrestrial atmosphere. However, the other, lighter noble gases are greatly overdepleted at the end of the first stage. Subsequently, mixing of the fractionated atmospheric residue with species degassed from the second, interior reservoir compensates for the overfractionation of Kr, Ar, and Ne elemental and isotopic abundances in the first stage, and produces the presently observed distributions of these gases. Modeling must therefore aim to define not only the hydrodynamic escape parameters that generate terrestrial Xe, but also the relative elemental abundances in both the primary atmospheric and interior reservoirs required to produce the final composition of the other noble gases.

In the original formulation of the model (Hunten et al. 1987; Pepin 1991), hydrodynamic losses of primary atmospheric volatiles are driven by intense EUV radiation from the young evolving Sun. Hydrogen outflow fluxes strong enough to enable Xe escape from Earth, and fractionation to its present isotopic composition, required atmospheric H2 inventories equivalent to water abundances of up to a few wt.% of the planet’s mass, and early solar EUV fluxes up to ~450× present levels—large but, as seen in Figure 10, not unrealistic enhancements if nebular dust and gas had dissipated to levels low enough for solar EUV radiation to penetrate the midplane to planetary distances within 100 Ma or so. Energy sources other than solar EUV absorption may have powered atmospheric escape. Benz and Cameron (1990) suggested that hydrodynamic loss driven by thermal energy deposited in a giant Moon-forming impact could have generated the well-known fractionation signature in terrestrial Xe. Their model of the event calls for rapid invasion of the pre-existing primary atmosphere by extremely hot (~16,000 K) dissociated rock and iron vapor, emplacement of an orbiting rock-vapor disk with an inner edge at an altitude comparable to the atmospheric scale height at this temperature, and longer-term heating of the top of the atmosphere by reaccretion of dissipating disk material. As yet there are no detailed theoretical calculations supporting the proposition that the physical conditions and energy-decay timescale needed to implement hydrodynamic escape could have been met in the aftermath of a giant impact on Earth. If a short postimpact escape episode did in fact occur, resulting in Rayleigh fractionation of whatever

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remnant of the primary atmosphere survived direct—and presumably non-fractionating— ejection in the impact event, required atmospheric H2 inventories would be reduced by an order of magnitude or more compared to models in which losses are driven only by solar EUV radiation (Pepin 1997); this is in part a consequence of the scaling of N10 with the energy decay constant τ in Equation (10), where, as noted below, τ for post-impact energy dissipation is assumed to be small. EUV-powered escape is still needed at some time following the impact event to account for Ne isotopic distributions (see below), but flux intensities for Ne-only loss are also about an order of magnitude lower than those required to lift heavier noble gas species out of the atmosphere. With appropriate choices for free parameters, both this and the original EUV-only model (Pepin 1991) can generate close matches to noble gas elemental and isotopic distributions in the contemporary terrestrial atmosphere. The interplay of atmospheric escape fractionation and subsequent degassing required to achieve these matches is illustrated for the hybrid giant impact (GI)-EUV model in Figure 11. The mean decay time τ of energy deposited by the impact is presumably small, reflecting relatively rapid energy dissipation following the event: here τ is arbitrarily chosen to be 1000 years, and with this and other model parameter values, Equation (11) yields an escape episode that runs its course in a comparable time. Primordial atmospheric U-Xe is fractionated by GI-driven hydrodynamic escape to its present abundance and nonradiogenic isotope ratios. About 85% of the initial 130Xe inventory is lost from the planet. The resulting “nonradiogenic Earth atmosphere” is used in Figure 11 as the Xe reference composition. Post-escape isotopic evolution of atmospheric Xe is largely restricted to degassing of radiogenic 129Xe, and of 131-136Xe generated primarily by spontaneous fission of 244Pu, from the upper mantle and crust (see Fig. 13, below). This constraint on “pollution” of the nonradiogenic Xe isotopic distribution generated by hydrodynamic escape, either by subsequent degassing of solar Xe from the interior reservoir or by substantial addition of isotopically different Xe carried in later-accreting material, is central to the modeling (see discussion below in Outstand-ing Issues). Increasingly severe fractionations of the lighter noble gases from their primordial isotopic compositions are imposed by the parameter values for GI-driven Xe escape. Residual Kr and Ar are both isotopically heavy (Fig. 11) and strongly depleted relative to the present atmosphere; only 6% and 0.8% respectively of the initial 84Kr and 36Ar inventories (and 0.4% of the 20Ne) survive the event. Subsequent mixing of these residues of the primary atmosphere with solar-composition Kr and Ar degassed from the interior reservoir raises their abundances to present-day values, and yields generally good (≥ 1σ) matches to contemporary isotopic compositions except for a ~2‰ (3σ) excess of 86 Kr. Neon is a special case in that 20Ne/22Ne and 21Ne/22Ne ratios in the fractionated residual atmosphere are substantially higher than at present, and later addition of outgassed solar Ne elevates them still more (Fig. 11). Here an episode of solar EUV energy deposition driving hydrodynamic escape of Ne at some time after GI fractionation and outgassing is needed to generate the contemporary 20Ne/22Ne ratio. But, as noted above, the EUV-driven H2 escape flux must now be only intense enough to lift Ne, but not the heavier noble gases, out of the atmosphere. The waning solar EUV flux (Ayres 1997) may still have been sufficiently high (~60× present levels) to drive Ne-only escape at solar ages up to ~250 Ma (Fig. 10), with the actual timing —somewhere in the interval between Ne degassing and ~250 Ma—determined by the timescale for sufficient reduction of EUV dust-gas opacity in the nebular midplane (Prinn and Fegley 1989). Note from Figure 11 that EUV escape fractionation yielding a match to the present 20 Ne/22Ne ratio overfractionates 21Ne/22Ne by a few percent, implying a 21Ne deficit in

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Figure 11. Evolution of terrestrial Xe, Kr, Ar and Ne from primordial atmospheric distributions to present-day compositions by Giant Impact (GI)-driven fractionation, addition of outgassed (OG) solar Kr, Ar and Ne, a later stage of solar EUV-powered Ne escape fractionation, and subsequent outgassing of radiogenic and fissiogenic Xe and nucleogenic 21Ne (after Pepin 1997). The δMXe and δMKr representations are defined by δM = 1000 [(R/Rref) – 1] ‰; here R = MXe/130Xe and MKr/84Kr respectively, and the references Rref are the corresponding isotope ratios in the nonradiogenic Earth atmosphere for Xe (Pepin 1991) and the present Earth atmosphere (Basford et al. 1973) for Kr. Xe data from Table 1; solar and atmospheric Kr from Wieler (2002, Table 5, refs. 2 and 4); solar 36 Ar/38Ar = 5.80 ± 0.06, 20Ne/22Ne = 13.84 ± 0.04, and 21Ne/22Ne = 0.0334 ± 0.0003 from Palma et al. (2002).

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the modeling. The atmospheric abundance of 21Ne, however, is subsequently augmented by outgassing over geologic time of a nucleogenic 21Nen component generated primarily by 18O(α,n)21Nen reactions in the Earth’s crust and mantle (Wetherill 1954). Estimates of the fractional abundance of 21Nen in the present atmosphere are consistent with the amount needed to elevate the initially lower 21Ne/22Ne ratio in Figure 11 to its present value by post-escape degassing (see Porcelli and Ballentine 2002, this volume).

Figure 12. Modeling results for the elemental compositions of primary atmospheres on Earth and Venus, plotted relative to solar abundance ratios. The baseline model for impact-driven escape from Earth (dashed line) assumes H2 loss from a 2000-K atmosphere. Different assumptions for atmospheric temperature and the identity of the abundant escaping constituent generate the shaded spread in Kr/Ar and Xe/Ar ratios originally present on Earth (see text). The stippled region represents the factor of ~5 range in initial terrestrial 20Ne/36Ar, on either side of the Venus ratio, for which the Ne isotopic matches shown in Figure 1 can be generated by EUV-driven loss. All elemental data are from Pepin (1997).

Primordial heavy noble gas abundances, calculated from present inventories backward through the degassing and escape stages of evolution, yield the dashed-line relative abundance pattern shown in Figure 12 for the 84Kr/36Ar and 130Xe/36Ar ratios present in the pre-impact Earth atmosphere. Estimating the initial 20Ne abundance on Earth is complicated by the terminal EUV-driven Ne fractionation stage. In calculating the Ne isotopic evolution shown in Figure 11 and the dashed-line elemental pattern in Figure 12, initial 20Ne was assumed for illustration to be present in the same ratio relative to 36Ar as in Venus’ primary atmosphere (see below). Solutions yielding final Ne isotope ratios identical to those in Figure 11 exist, with differing EUV fractionation parameters and amounts of degassed 21Nen, for primordial atmospheric 20Ne/36Ar ranging from ~1/5 to 5 times the Venus ratio (stippled area in Fig. 12). These modeling results assume an atmospheric temperature T = 2000 K and H2 as the abundant hydrodynamically escaping species. Neither choice would be appropriate for escape from a very hot dissociated atmosphere, an environment likely to have characterized early stages of giant-impact-driven loss. However the results are not sensitive to either temperature or the exact identity of the carrier gas. The shaded area in

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Figure 12 represents the relatively small spread in initial elemental composition introduced by assuming atmospheric temperatures ranging from 300 K to 10,000 K and, at temperatures ≥ 4000 K, H or O rather than H2 as the dominant light atmospheric constituent. Application to Venus. Data from in situ compositional measurements of the Venus atmosphere by mass spectrometers and gas chromatographs on the Pioneer Venus and Venera spacecraft are reviewed and assessed by von Zahn et al. (1983); an updated summary is set out in Table 8 of Wieler (2002). One might suspect that planets as alike as Earth and Venus in size and heliocentric distance would have acquired compositionally similar primary atmospheres from similar sources. It is not obvious, however, from comparison of volatile mass distributions on Earth and Venus, that these two atmospheres are end products of similar evolutionary processes acting on similar primordial volatile sources. Absolute abundances on Venus exceed those on Earth by a factor >70 for 36Ar, but only by factors of ~3-6 for Kr and Xe. Consequently, as noted above, there is a pronounced solar-like signature in relative Ar:Kr:Xe abundances. This similarity does not extend to Ne: the 20Ne/36Ar ratio is low, close to terrestrial. Venusian 20Ne/22Ne, however, is significantly higher (i.e., more solar-like) than on Earth, and the nominal value of the 36Ar/38Ar ratio is somewhat above the terrestrial value. There are no measurements of Kr and Xe isotopic compositions.

Since volatile compositions on Venus are known only approximately, if at all, isotopic constraints on modeling parameter values are much weaker than in the terrestrial case, and for the heavy noble gases are missing entirely. However the EUV flux driving Ne escape in the Earth model discussed in the preceding section must also irradiate Venus at the same time—i.e., ΦEUV(t) in Equation (6) is the same for both planets—and this planetary interdependence together with the isotopic and elemental information we do have allows construction of at least a preliminary model for Venusian volatile evolution. If EUV heating efficiencies and H2 escape altitudes are assumed to be similar, it is seen from Equation (6) that the ratio of the crossover mass mC(S) on Venus to that on Earth for the same species is a function only of relative planetary radii, densities, heliocentric distances, and atmospheric temperatures. It turns out that the relatively weak solar EUV flux needed for Ne-only loss from Earth is still strong enough at Venus’ orbital position to drive outflow of Kr and lighter gases from this smaller and less dense planet. But mC0(Xe) on Venus falls below the mass of the lightest Xe isotope, and so Venusian Xe is not lost and its nonradiogenic isotopic composition, presently unknown, is predicted to be unaltered from its primordial composition. Results of this model of EUV-driven loss of an isotopically solar and elementally near-solar primordial atmosphere from Venus are sensitive to only one of the few remaining adjustable modeling parameters—αo, which fixes the initial H2 inventory N10 via Equation (10). A single value of this parameter generates matches well within their 1σ uncertainties to the very limited isotopic data we currently have for Venus’ presentday noble gases (20Ne/22Ne and 36Ar/38Ar only: see Table 8 in Wieler 2002, this volume). It is interesting that there is no discernible evidence in the data on hand that outgassing has played a significant role in establishing contemporary noble gas inventories. Fractionating loss of the primary atmosphere, governed by the parameters above, generates by itself approximate matches to observed compositions. Thus, in contrast to the case for Earth, the presence of an outgassed component on Venus is not required. This is not to say that the planet could not have degassed at some time during or after this stage of atmospheric evolution. But outgassed species would comprise only modest fractions of the large present-day Venusian atmospheric inventories even if bulk planetary concentrations were comparable to those on Earth. In fact, the atmospheric

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abundance of 40Ar suggests that outgassing rates on Venus have been substantially less than terrestrial rates (see “Noble gases on Venus” section). Elemental ratios characterizing the primary (pre-loss) Venus atmosphere in this evolutionary model are plotted in Figure 12. Although these computed primordial 84 Kr/36Ar and 130Xe/36Ar ratios inherit the large uncertainties associated with measurements of present Venusian abundances, their nominal values fall squarely within the range of estimates calculated above for Earth’s pre-impact primary atmosphere—a strong implication that noble gases on both planets could have evolved, clearly in quite different ways, from the same primordial distributions in the same types of primary planetary reservoirs. This result seems reasonably robust, provided of course that fractionating, GI-driven hydrodynamic escape actually did occur on Earth. However a central test which the Venus model must eventually confront is whether its predicted Xe and Kr isotopic compositions, respectively unfractionated and slightly fractionated with respect to primordial atmospheric compositions, are in accord with future measurements. Venus appears to represent an extreme in the range of atmospheric compositional patterns displayed by the terrestrial planets, and, as discussed below in “Outstanding Issues”, this and other central questions could be answered, or at least more quantitatively addressed, if we can eventually manage to revisit the planet for moderately accurate in situ measurements of all noble gas isotopic distributions, or, even better, return an atmospheric sample to Earth for laboratory study. Application to Mars: losses by hydrodynamic escape and sputtering. Information about volatile abundances and compositions on and in Mars come from in situ measurements by Viking spacecraft instruments and, by a great stroke of fortune, from high-precision laboratory analyses of Martian atmospheric gases trapped in the SNC meteorites (Swindle 2002, this volume). Early modeling of the history of noble gases on Mars based on these data suggested that they could have evolved from primordial to present-day distributions through two early episodes of hydrodynamic atmospheric escape (Pepin 1991). This model, however, did not address a number of processes that now appear germane to Martian atmospheric history. One, gas loss and fractionation by sputtering, was later proposed to be the dominant mechanism governing atmospheric CO2 evolution on Mars over the past ~3-4 Ga (Luhmann et al. 1992; Zhang et al. 1993). Another, atmospheric erosion (Melosh and Vickery 1989), appears increasingly important (Chyba 1990, 1991; Zahnle 1993). In the absence at that time of a plausible mechanism, the possibility of isotopic evolution of noble gases heavier than Ne after the termination of hydrodynamic escape was not considered by Pepin (1991), and only qualitative attention was paid to the eroding effects of impact on abundances of all atmophilic species prior to the end of heavy bombardment ~3.8 to 3.7 Ga ago.

The more recent models of Martian atmospheric evolution constructed by Jakosky et al. (1994) and Pepin (1994) incorporate the sputtering loss mechanism proposed by Luhmann et al. (1992) and Zhang et al. (1993), and explore its consequences for elemental and isotopic fractionation of the noble gas and nitrogen in the residual atmosphere. They divide Martian atmospheric history into early and late evolutionary periods, the first characterized by an initial episode of hydrodynamic escape that sets the Martian Xe inventory (discussed later in the “Primordial Xenon” section), followed by high CO2 pressures and a possible greenhouse, and the second by either a sudden transition to a low pressure environment similar to present-day conditions on the planet— perhaps initiated by abrupt polar CO2 condensation ~3.7 Ga ago (Gierasch and Toon 1973; Haberle et al. 1992, 1994)—or by a more uniform decline to present pressure. Jakosky et al. (1994) focused on the late evolutionary epoch, and showed that contemporary Ne, Ar and N2 abundances and isotope ratios—including the uniquely low

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Martian 36Ar/38Ar ratio (Swindle 2002, this volume)—could have been generated by sputtering losses from an atmosphere that was continuously replenished by degassing of meteoritic (CI) N2 and isotopically solar Ne and Ar, with the rate vs. time dependence of degassing chosen to be similar to estimates of volcanic flux vs. time over this period. The second of these models (Pepin 1994) included some thoughts on early evolution of volatile distribution on Mars. The EUV-powered hydrodynamic escape episode driving Ne-only loss from Earth, and loss of Kr and lighter gases from Venus, would have been intense enough on Mars to lift all the noble gases out of its primordial atmosphere. Early in this pre-3.7-Ga epoch, Xe isotopes were therefore assumed to have been hydrodynamically fractionated to their present composition, with corresponding depletions and fractionations of lighter primordial atmospheric constituents. Subsequent CO2 pressure and isotopic history was dictated by the interplay of estimated losses to impact erosion, sputtering, and carbonate precipitation, additions by outgassing and carbonate recycling, and perhaps also by feedback stabilization under greenhouse conditions. In a subsequent model of the early Martian atmosphere, Carr (1999) examined the influences of these same mechanisms in controlling CO2 pressure history, and was led to similar results and conclusions. It should be stressed, however, that since almost nothing is actually known about the values of the parameters governing these various processes, models of this epoch are no more than qualitative illustrations of how they might have driven early atmospheric behavior. Pepin’s (1994) treatment of post-3.7-Ga evolution of Martian CO2, N2, and the noble gases, although differing somewhat in detail from that of Jakosky et al. (1994) and extended to include the radiogenic isotopes 40Ar and 129Xe, generated results close to those derived by Jakosky et al. Both models assume that early and late evolutionary stages on Mars were separated by atmospheric CO2 pressure collapse near 3.7 Ga. Sputtering loss of an atmospheric species relative to that of CO2 is directly proportional to its exobase mixing ratio with CO2, and so sputtering fractionation of the atmospheric noble gas inventory is generally modest in a pre-3.7-Ga atmosphere dominated by CO2 (Jakosky et al. 1994). Pressure collapse of the major atmospheric constituent abruptly increases the mixing ratios of pre-existing Ar, Ne and N2 at the exobase, and since escape fluxes are proportional to exobase mixing ratios (see “Sputtering” section above), CO2 collapse triggers their rapid removal by sputtering. This has the interesting consequence that no isotopic memory of their earlier processing survives. Current abundances and isotopic compositions are entirely determined by the action of sputtering and photochemical escape on gases supplied by outgassing during the late evolutionary epoch, and final distributions of the light noble gases and nitrogen are therefore decoupled from whatever their elemental and isotopic inventories might have been in the pre-3.7 Ga atmosphere. The present atmospheric Kr inventory (Fig. 6) also derives almost completely from solar-like Kr degassed during this period, which overwhelms the fractionated component inherited across the collapse episode —as noted above in the “Sputtering” section, both Kr and Xe are too massive to be sensibly affected by sputtering loss and fractionation during the late evolutionary stage. Consequently, among current observables, only the Xe isotopes and δ13C survive as isotopic tracers of atmospheric history prior to its transition to low pressure. The assumption that early hydrodynamic escape fractionated the nonradiogenic Xe isotopes to at least approximately their present composition severely limits subsequent additions of unfractionated Xe to the atmospheric inventory by outgassing (consistent with the low degree of planetary degassing deduced above from radiogenic noble gas isotopes) or latestage veneer accretion. This constraint, which threads through all hydrodynamic loss models for the terrestrial planets, is perhaps their most vulnerable characteristic (see Outstanding Issues section below).

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Jakosky and Jones (1997) reviewed progress up to that time in understanding some of the aspects of Martian atmospheric evolution. There have been several developments since the work of Jakosky et al. (1994) and Pepin (1994). Hutchins and Jakosky (1996) revisited the late evolution sputtering-degassing models to investigate in more detail the parameters controlling the evolution of Ne and Ar abundances and isotopes (including radiogenic 40Ar), in particular those relating to Martian degassing history. Their interesting conclusion was that the outgassing flux of Ar and Ne attributable to degassing during epochs of volcanic activity would have been ~1-3 orders of magnitude too low to appropriately balance sputtering losses, and thus another major source of juvenile volatiles must have contributed to the atmosphere over geologic time, perhaps via input from gas-enriched hydrothermal systems. Hutchins et al. (1997) explored the important question of the extent to which a Martian paleomagnetic dipole field would have throttled sputtering losses by deflecting the solar wind around the upper atmosphere, and calculated the conditions, as functions of the time when paleomagnetic suppression of the sputtering mechanism ended, under which the combination of sputtering and degassing would still have generated present-day Ar and Ne distributions. The possible influence of an early global magnetic field on atmospheric evolution became more than a purely theoretical consideration with the discovery by Mars Global Surveyor of large-scale remnant magnetic lineations in the old Martian southern highlands (Connerney et al. 1999). The question of the existence of an active dynamo in Mars appears to be no longer if, but when. Primordial xenon

Attempts to derive the compositions of the contemporary atmospheres on Earth and Mars from primitive nebular, planetesimal, or planetary source reservoirs have had to confront, in one way or another, an apparent decoupling of the evolutionary histories of Xe from those of Kr and the lighter noble gases. On Mars, for example, Xe is regarded as a highly fractionated residue of a likely CI- or solar-like progenitor composition; Kr, in contrast, appears to be unfractionated from an isotopically solar composition (Swindle 2002, this volume). This implies spatially and/or temporally separate provenances for these two species, processing mechanisms operating at different times, or both. In the models discussed in this section, atmospheric Xe inventories on both planets are considered to be fractionated relicts of their primary atmospheric Xe, while most of the Kr and lighter noble gases are products of planetary outgassing, compositionally modified on Mars by sputtering losses during the late evolutionary epoch. Xenon therefore plays a crucial role in models of atmospheric evolution in which noble gases are fractionated from their initial compositions to isotopically heavier distributions by early hydrodynamic escape—it is the only observable among the noble gases that preserves a signature, albeit fractionated, of primordial atmospheric composition. With the assumption that nonradiogenic Xe isotope ratios in present-day atmospheres on Earth and Mars were generated in this way, backward modeling from these ratios through the fractionating process can in principle identify likely parental Xe compositions and thus the probable sources of noble gases in pre-escape atmospheres. Current results of this exercise, using the Xe component compositions listed in Table 1, are reviewed in Pepin (2000). Applied to Earth, a modeling procedure involving derivation of fractionation relationships between nonradiogenic terrestrial and meteoritic Xe compositions simultaneously identifies a composition named U-Xe as primordial Xe, and establishes the presence of an atmospheric Xe component due principally to fission of extinct 244Pu, as noted earlier in this chapter. Hydrodynamic escape of U-Xe leaves its fractionated residue on Earth. Evolution of atmospheric Xe from this early composition to its present isotopic state by subsequent degassing of fission and radiogenic

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components from the crust and mantle is shown schematically in Figure 13. To fill in the deficits at isotopes 129Xe and 131-136Xe, Pu-Xe must comprise 4.65±0.30% of atmospheric 136 Xe, and 6.8±0.5% of the present abundance of 129Xe is from decay of extinct 129I. (A more accurate re-calculation of U-Xe composition by Pepin (2000) is responsible for the increase in the Pu-Xe contribution from Pepin’s (1997) estimate of 3.9%).

Figure 13. Post-escape evolution of Xe in the Earth's atmosphere (Pepin 2000). Escapefractionated U-Xe defines an initially nonradiogenic terrestrial Xe composition (NEAXe, Table 1) to which radiogenic and fissiogenic components generated by decay of 129I and 244Pu in the crust and mantle were subsequently added by planetary outgassing.

The derived U-Xe composition is identical to that measured for solar-wind (SW) Xe except for relative underabundances of the two heaviest isotopes—an unexpected difference since the modeling otherwise points to solar wind compositions for the lighter noble gases in the primordial terrestrial atmosphere. However, as pointed out earlier (see Fig. 7 and associated discussion), SW-Xe cannot serve as primordial terrestrial Xe because its required fractionation generates a large overabundance of 136Xe in the present atmosphere. Fractionated U-Xe ratios fall below the current atmosphere at 131-136Xe, defining a nonradiogenic terrestrial Xe spectrum to which a fissiogenic component is later added (Fig. 13); this property of the U-Xe composition is intrinsic to its derivation. In contrast to Earth, Martian Xe apparently did not evolve from a U-Xe progenitor. Modeling derivation of primordial Xe composition on Mars is based on analyses of atmospheric gases trapped in glassy phases of SNC meteorites (Swindle 2002, this volume). Present ambiguities in this data base are such that two different solar-system Xe compositions, carbonaceous chondrite (CI)-Xe and SW-Xe, are possible candidates—but not U-Xe. Exclusion of U-Xe as the dominant primordial atmospheric inventory on Mars, despite the implication of the terrestrial modeling that it was a major component of the nebular gas phase, requires that accretion of CI- or SW-Xe-rich materials from sources more localized in space or time has overwhelmed the isotopic signature of its presence.

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Swindle et al. (1986) pointed out that nonradiogenic Xe trapped in the glassy lithology of the SNC meteorite EET79001 strongly resembles mass-fractionated CI-Xe. Pepin (1991, 1994) chose CI-Xe to represent the primordial Xe composition on Mars in modeling the isotopic evolution of its atmosphere, in part based on Swindle et al.’s observation but also because, as noted earlier, a meteoritic source is consistent with models of the bulk chemical composition of the planet. The fact that fractionated CI-Xe by itself provides an excellent match to the Mars atmospheric composition, in particular at the four heaviest isotopes, has the interesting and somewhat unsettling consequence that additional contributions to the Xe inventory from degassed fission Xe are either very minor or absent altogether. And yet the present atmosphere is clearly heavily enriched in 129 Xe from extinct 129I decay (Swindle 2002). The presence of radiogenic 129Xe but apparently little if any 244Pu-fission Xe, a situation quite unlike that on Earth, has been awkward to reconcile with models of Martian geochemical evolution and degassing history. Swindle and Jones (1997) considered this problem in detail, and constructed an alternative model of atmospheric evolution on Mars—using an atmospheric Xe composition close to that measured by Swindle et al. (1986)—with the specific objective of accommodating, if possible, a Pu fission Xe component. They chose SW-Xe rather than CI-Xe as primordial Xe, and demonstrated that an appropriate fractionation of this composition fell below measured Martian atmospheric 131-136Xe/130Xe ratios by amounts completely consistent within error with the presence of an additional component with PuXe fission yield ratios; moreover the calculated fissiogenic 136Xe abundance was ~5% of total 136Xe, similar to the fraction derived above for Earth. Mathew et al. (1998) revised the earlier data base for Martian atmospheric Xe by including measurements of Xe composition in the Zagami shergotite and recalculating corrections for spallation Xe produced during the space exposures of SNC meteorites to cosmic-ray irradiation. The average of the data sets they judged to be of highest quality, plotted in Figure 14 referenced to SW-Xe, differs from Swindle et al.’s (1986) composition by 206Pb+8 4He+6β235 U ----> 207Pb+7 4He+4β232 Th----> 208Pb+6 4He+4β40 K--ec-> 0.105 40Ar 238 U Fission

4,468 704 14,010 1,250 4,468

47.4 45.2 39.8 0.71

EXTINCT RADIOACTIVITY 129

Xe/130Xe 136,134 Xe/130Xe

129

I ----> 129Xe+β244 Pu Fission

16 82

NUCLEOGENIC REACTIONS FROM U AND TH DECAY 21

Ne/22Ne

18

O(α,n) ---> 21Ne 24 Mg(n,α) ---> 21Ne

One of the most significant observations is the ubiquitous presence of ‘excess’ 3He in mantle-derived rocks from ocean ridges and islands, indicating that primordial volatiles are still escaping from the Earth’s interior. The highest 3He/4He ratios, along with 20Ne/22Ne and 21Ne/22Ne ratios that approach solar values, are found at ocean islands, most notably Iceland and Hawaii. The He-Ne isotope systematics of these ocean island basalts are currently the strongest geochemical evidence that some portions of the mantle have remained relatively undegassed over geologic time. These are striking findings, because they directly conflict with some geophysical and geochemical models that argue that the mantle convects as a single system and that no primordial or undegassed material remains in the Earth’s interior. The high 40Ar/36Ar and non-atmospheric 129Xe/130Xe in MORBs also provide fundamental clues to ancient planetary outgassing. Volatile loss from the Earth’s interior over time, during ancient formation of the ocean and atmosphere and/or by continuing depletion through partial melting and magma generation, produces a range of parent/daughter ratios and is the primary cause for variable ratios of radiogenic to nonradiogenic noble gas species. Plate tectonic recycling also plays a significant role for the budget of the lithophile parental nuclides U, Th and K. In contrast, the importance of subduction in the mantle budgets of heavier noble gases such as Ar and Xe is still debatable (Staudacher and Allègre 1988; Porcelli and Wasserburg 1995a,b; Sarda et al. 1999a; Burnard 1999b). Several books (Alexander and Ozima 1978; Ozima and Podosek 2002; Mamyrin and Tolstikhin 1984; Matsuda 1994) and review articles (Craig and Lupton 1981; Lupton 1983; Farley and Neroda 1998; Ozima 1994; McDougall and Honda 1998) currently provide a comprehensive background to terrestrial noble gas geochemistry. The aim of this chapter is to provide an up-to-date overview of key observations on noble gas isotopes in ocean ridge and island basalts that bear on models for the composition and evolution of the Earth’s mantle (e.g., see the chapter by Porcelli and Ballentine 2002, this volume).

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BACKGROUND Noble gas chemical behavior As a group, the noble gases are chemically inert, exhibiting only weak van der Waals type interactions. This serves to produce systematic and predictable variations within the group as a whole, often resulting in a coherent light-to-heavy noble gas trend that reflects the physical processes at hand, such as vapor/liquid/solid partitioning, molecular diffusion, and adsorption. Igneous processes such as partial melting, crystal fractionation, and magmatic degassing of major volatile species (CO2 and H2O) should lead to systematic elemental fractionations of the noble gases that are related to their variation in atomic size and the extent to which their electron cloud may be polarized (Keevil 1940). The elemental abundances of noble gases in oceanic basalts may therefore provide clues about petrogenetic processes (e.g., Dymond and Hogan 1978; Batiza et al. 1979). Our ability to characterize noble gas reservoirs in the mantle depends to some extent on understanding vapor/liquid and solid/liquid partitioning. For example, if the amount of gas loss from a magma is known to be small (or zero), then measured noble gas concentrations in a basalt could be used in combination with vapor/melt and mineral/melt partition coefficients to estimate mantle source abundances. A meaningful determination of the relative elemental abundances in the mantle source requires a ‘correction’ of the measured concentrations in igneous samples for any fractionation or contamination that occurred during petrogenesis. This is a formidable task, and it can only be carried out effectively by using isotopic relationships among noble gas species. Argon provides a clear example. 40Ar/36Ar ratios >30,000 have been measured in some MORBs. In the case of a MORB sample with 40Ar/36Ar = 5,000 (a seemingly high value because it is more than an order of magnitude above the air ratio of 296) it is still possible that >80% of the measured Ar is derived from atmospheric contamination, and it would be erroneous to draw conclusions about noble gas abundance ratios in the mantle source region based on the concentration data alone. Fortunately, much of this problem can be circumvented through the ratios of certain radiogenic species (e.g., 4He/40Ar*). This approach is discussed in the section on Coupled Radiogenic/Nucleogenic Production. Much less uncertainty and fewer assumptions are involved when isotopic ratios (e.g., He/ He vs. 40Ar/36Ar) are compared directly. Therefore, wherever possible, this review focuses on isotopic relationships rather than on elemental abundances alone. The isotopic approach allows one to best discern the effects of atmospheric contamination, which can be quite large, and it circumvents some of the ambiguities that result from using the concentrations measured in rocks to interpret the inter-elemental fractionations that occur during igneous processes. The observational data upon which the review is based are presented in diagrams that form an integral part of the chapter. The key points provided by the figures are described in the review, but a detailed description of what they readily show for all the individual localities is not always given, in order to maintain a focus on the general characterization of mantle source reservoirs. 3

4

Vapor-melt partitioning. Noble gas solubility in basaltic melt decreases with increasing atomic mass, and is directly related to the atomic radius of the gas (Jambon et al. 1986; Lux 1987; Broadhurst et al. 1992; Shibata et al. 1998). At 1400°C the experimentally determined values for mid-ocean ridge basalt are 56, 25, 5.9, 3.0 and 1.7×10-3 std cm3/g-bar for He, Ne, Ar, Kr and Xe, respectively (Jambon et al. 1986). Degassing of basaltic melts will therefore lead to significant fractionation in the relative abundance of the noble gases, with the exsolved (volatile) phase enriched in the heavier noble gases, such as Ar and Xe, compared to the lighter species, such as He and Ne. The residual melt will show the opposite effect, with a preferential depletion in the heavier species compared to the lighter ones. The low solubilities lead to elevated concentrations

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in the vesicle phase of basalts. Mid-ocean ridge basalt glasses generally exhibit equilibrium vesicle/melt partitioning for helium (Kurz and Jenkins 1981) and for argon (Marty and Ozima 1986). Noble gas solubility is only weakly dependent on temperature (Lux 1987) but it depends on melt composition. The compositional control is well described by the ‘ionic porosity’ (Carroll and Stolper 1993; Carroll and Draper 1994) or by the ratio of nonbridging oxygens to silicon (Shibata et al. 1998). The ionic porosity is the difference between the unit cell volume of a material and the calculated volume of the anions and cations, and provides an integrated measure of the interstitial sites available within the melt structure. The logarithm of noble gas solubility is linearly correlated with ionic porosity. The solubility also shows an increasing sensitivity to small changes in ionic porosity as the size of the gas atom increases. This observation led Carroll and Stolper (1993) to suggest that, as melt structure becomes more tightly packed, the availability of interstitial sites capable of accommodating the larger atoms decreases dramatically. Noble gas solubility also depends on H2O and CO2 content of silicate melts (Paonita et al. 2000; Nuccio and Paonita 2000). Paonita et al. (2000) applied a novel method to study these effects, by adding a known amount of He-bearing glass to their experimental runs. The solubility of He was determined over a range of mixing proportions of H2O and CO2 in a rhyolite and in a trachybasalt. The He solubility is strongly influenced by H2O content, showing about a factor of 3 increase with the addition of 3 wt % H2O, apparently because the addition of water increases the availability of sites in the melt that accommodate noble gases. Solubility also increases exponentially with atomic size due to the addition of H2O (Nuccio and Paonita 2000), so while Xe is less soluble than He in anhydrous melts, Xe and He solubilities are nearly the same when several percent H2O is dissolved. The effect of CO2 is more uncertain but it appears to be the opposite of H2O, showing a decrease in the solubility of He by a factor of ~1.5 with addition of 0.05 wt % CO2 (Nuccio and Paonita 2000). The major volatile composition of the melt therefore affects the relative degassing behavior of the noble gases. During extended degassing, a CO2–rich anhydrous magma will retain its dissolved He more efficiently than a H2O–rich magma (Paonita et al. 2000). Variations in the amount of H2O and CO2 that were initially present in variably degassed melts may be partly responsible for the observed range of CO2/3He in oceanic basalts (see the Helium subsection, Relation to major volatiles). The effect of pressure on noble gas solubility in silicate melts is currently an active area of investigation. Chamorro-Perez et al. (1998) reported an order of magnitude decrease for Ar solubility in an olivine melt near 5 GPa, corresponding to mantle depths near 150 km. They inferred that melt densification makes it impossible to accommodate Ar in interstitial sites near this pressure. This is a surprising result, because it would imply that Ar is moderately compatible at depth, and partial melting would not be an effective means of mantle degassing. However, more recent data indicate that the Ar clinopyroxene/silicate melt partition coefficient is relatively constant at ~4×10-4 for pressures up to at least 8 GPa (Chamorro et al. 2002), so there is no structural change in the melt over that pressure range. Recent work on Ar and Xe solubility in synthetic melts also does not show a decrease in solubility even at pressures of 11 GPa. Instead, it appears that Ar solubility increases to about 6 GPa, above which it reaches a threshold concentration of 0.8 wt % (Schmidt and Keppler 2002). An important consideration in applying the experimental results for solubility and partitioning should be the very low abundance of noble gases in the mantle. Crystal-melt partitioning. Compared to noble gas solubility in melts, much less is known about their solubility in minerals. Several different approaches can be used to determine a crystal/melt partition coefficient D (where D = weight concentration in

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crystal/weight concentration in melt) but each one has its limitations. The presently available data are contradictory and the D values range by orders of magnitude, from 1. A thoughtful summary of the issues involved in these estimates is presented by Carroll and Draper (1994). A few experimental studies of noble gas mineral/melt partitioning have been carried out (Hiyagon and Ozima 1982, 1986; Broadhurst et al. 1990, 1992). These studies show a very wide range in partition coefficients for some minerals, and their applicability to mantle melting is difficult to evaluate. Hiyagon and Ozima (1982, 1986) powdered their run products to physically separate the glass from the crystals, so there are questions of atmospheric adsorption, the extent to which all adhering glass could be removed from crystals, and the possible presence of fluid inclusions. They obtained olivine/melt values of DHe ≤ 0.07, DAr = 0.05-0.15 and DXe ≤ 0.3. Broadhurst et al. (1990, 1992) used separate containers for minerals and melts and were extremely careful to use inclusion-free starting materials, but they also obtained a wide range of D values. They observed a weak increase in D from He through Xe for each of the minerals studied (forsterite, diopside, anorthite and spinel) and suggested that this was related to the increasing polarizability with atomic number. They also suggested that the wide range in partition coefficients is related to the number of lattice defects, because variation in the density of interstitial sites in the minerals on the scale needed to explain the range seemed unreasonable. All of these experimental studies obtained D values indicating that the noble gases, especially the heavier species, are more compatible than is usually assumed. Given the experimental difficulties, these D values should be considered as upper limits. A second approach for obtaining crystal/melt partition coefficients is to use naturally occurring glass-mineral pairs. This approach has been used to study partitioning between olivine and basalt melt (Marty and Lussiez 1993; Kurz 1993; Valbracht et al. 1994). In two studies that used similar approaches to analyze olivine-rich basalts from the MidAtlantic Ridge, Marty and Lussiez obtained a value of DHe≤0.008 and DAr≤0.003, while Kurz (1993) obtained DHe ≤0.0058. These investigators analyzed olivine and glass from the samples both by crushing in vacuum, to release He trapped in inclusions and bubbles, and by fusion, to release the trapped plus the dissolved components. Marty and Lussiez (1993) showed that the olivine and glass have the same 3He/4He ratio, and appear to be in chemical equilibrium based on their Fe/Mg ratio, although this was questioned by Hiyagon (1994a). The He and Ar released from this glass is dominated by the vesicle gas fraction (Marty and Lussiez 1993), and the He and Ar released from the olivine by crushing appears to be dominated by gas in shrinkage bubbles associated with trapped melt inclusions (Marty and Lussiez 1994). The latter observation would support the notion that the olivine crystals grew before the magma was vapor-saturated. If this was not true then the investigated partitioning involves three phases and is much more complicated. The presence of melt and fluid inclusions in crystals and the possibility of vesicle loss from a magma following crystallization are limitations to using natural samples, and led Marty and Lussiez (1993) and Kurz (1993) to present their D values as minima. In situ laser analysis is becoming increasingly important in addressing crystal/melt partitioning of noble gases. In a preliminary study using a UV laser ablation microprobe, Brooker et al. (1998) obtained a range of Ar partition coefficients from 0.013-0.14 for olivine and 0.0016-0.589 for clinopyroxene. The low values are probably more realistic and less affected by possible adsorption effects, early partial melting of the crystals and the presence of fluid inclusions. Although the partitioning behavior of the noble gases between minerals and melts is currently poorly understood and further work is needed, the least ambiguous results

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indicate that the noble gases have D values below 1. Given the low abundances of the noble gases in the mantle, the experimental data give some indication that mineral defects may play an important role. Both the experiments and the naturally occurring mineral/melt pair studies indicate that He behaves as a highly incompatible element. U and Th are also known to be highly incompatible, having bulk D values on the order of 10-3 or less, so the behavior of He relative to U and Th still needs to be established. Given that the lowest D value determined for He is likely to be a maximum, the large differences in 3He/4He between the mantle sources for MORBs and OIBs can be taken to reflect differences in their degassing history. Mantle structure and noble gases The chemical structure of the Earth’s mantle is directly related to the style of mantle convection, and the debate over whole mantle vs. layered mantle convection has gone on for decades. There are several scales of mantle convection indicated by seismic tomography, gravity and geochemistry (Anderson 1998b). The largest scale is controlled by the pattern of cooling associated with subducting plates, while the smaller scales (4001000 km) are probably controlled by the depth of phase transitions and the thickness of the upper mantle low viscosity region. Seismic evidence now clearly shows that some subducting slabs penetrate below the 660 km seismic discontinuity (van der Hilst et al. 1997), so this depth can no longer be viewed as a strong barrier to mass transport between the upper and lower mantle. Some investigators take this as sufficient evidence for whole mantle convection. In this case, the isotopic differences between depleted MORBs and enriched OIBs might be explained by a mantle that contains large scale blobs of chemically enriched or primitive material (e.g., Manga 1996; Becker et al. 1999). The origin of such blobs is unclear, and detecting them (if they exist) is currently beyond the resolution capabilities of seismology. Alternatively, the mantle may have some form of layered structure. If so, then buoyant upwellings produced at thermal or chemical boundary layers (mantle plumes) will be an important mechanism by which deep material is brought close to the Earth’s surface, where it partially melts to form ocean island basalt magmas. Mantle plumes have been implicated in the origin of many ocean islands since the discovery of plate tectonics (Morgan 1971). Bathymetric tracks of volcanoes such as the Hawaiian-Emperor chain indicate that these plume sources move laterally much more slowly than the plates, so the depth of origin must lie below the convecting upper mantle, although the exact depth is currently not well constrained. Such plumes most likely originate from boundary layers, perhaps as deep as the core-mantle boundary, and they are expected to entrain small amounts of material from the underlying reservoir. In actuality there is a slow relative motion between the hotspots on the Earth’s surface that is explained by advection of the plumes by large-scale mid-mantle flow. This mid-mantle flow is generally toward ridges and opposite in direction to the flow field in the upper mantle as indicated by plate motion directions (Steinberger and O’Connell 1998). In some cases, most notably East Africa, the upwelling also appears to be much broader than expected for a narrow plume conduit (Ritsema et al. 1999). Obviously much less is known about the deep mantle than the upper mantle. The style and vigor of mantle convection are described by the Rayleigh number (Ra). Ra is estimated to be 107 for the uppermost mantle; the estimate for the deep mantle is more uncertain but it may be several orders of magnitude smaller, mostly due to its higher viscosity (Tackley 1998). Consequently, the mixing time for the upper mantle is relatively rapid compared to the deep mantle, and compositional heterogeneities have a greater likelihood of surviving deeper in the Earth’s interior (Gurnis and Davies 1986). Nevertheless, state-of-the-art coupled convection-degassing models that incorporate a high viscosity lower mantle, a phase transition at the base of the upper mantle, and temperature- and pressure–dependent rheology, currently fail to produce an isolated, high

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He/4He region in the Earth (van Keken and Ballentine 1998, 1999).

Recent seismic tomography seems to indicate that some type of boundary may be present at a depth near 1700 km in the lower mantle (van der Hilst and Kárason 1999; Kellogg et al. 1999). Below this depth the mantle may be denser than the overlying mantle due to differences in bulk composition. However, the deeper material could contain more of its original inventory of heat producing elements, so it is hot and only slightly more dense than mantle at the same depth lying on the adiabat (the pressuretemperature path for material that expands or contracts without gaining or losing heat). This configuration is dynamically stable but it may lead to the development of significant topography on the surface of this layer. The resulting small density differences (~0.5%) would be sufficient to inhibit mixing, but the layer’s surface could respond dramatically to down-going slabs and rising thermal plumes (Kellogg et al. 1999). This model is consistent with laboratory studies of thermochemical convection in a chemically heterogeneous fluid that has a stratified density and viscosity structure (Davaille 1999). These lab experiments show that hot domes oscillate vertically through the fluid while thin plumes rise from their upper surface. There are other ‘layered’ models. For example, the ‘perisphere’ model places all OIB sources in a thin, shallow enriched layer beneath the lithosphere (Anderson 1995). It accounts for the presence of enriched basalts in continental rifts, but it is difficult to reconcile with the voluminous volcanism at some ocean islands. Many of the proposed models do not readily satisfy the noble gas observations that support the preservation of relict primitive mantle. Some contrasting ideas on mantle convection and its structure and evolution that consider the noble gas constraints are given in papers by Kellogg (1992), Davies and Richards (1992), Albarède (1998), Coltice and Ricard (1999) and Tackley (2000). Two books on mantle dynamics are also now available (Davies 1999; Schubert et al. 2001). Intriguingly, there is no compelling evidence for the survival of primitive mantle based on refractory element ratios or the isotopes of Sr, Nd and Pb in ocean island basalts (Hofmann 1997). Hofmann et al. (1986) showed that certain elemental ratios, such as Ce/Pb and Nb/U, had uniform and similar values in both MORBs and OIBs. These Ce/Pb and Nb/U ratios are distinct from values for either the primitive mantle or the continental crust, indicating that the chemical signature of crust extraction over geologic time has mostly been homogenized throughout the mantle. Therefore, the observed Sr, Nd and Pb isotopic differences between MORBs and OIBs must be the result of more recent processes. Hofmann and White (1980) convincingly established that the extreme isotopic compositions of OIBs are probably controlled by plate-tectonic recycling, in which OIBs are produced by the heating and melting of subducted slabs. Zindler and Hart (1986b) and Weaver (1991) also demonstrated that the OIBs with the most radiogenic Pb isotopes were probably dominated by recycled oceanic crust, while other enriched OIB endmembers probably contain a small percentage of recycled sediments (terrigenous vs. pelagic) mixed into their mantle source region. In the layered mantle model, OIB magmas from intraplate hotspots with high He/4He come from deep mantle source regions that are convectively isolated from the upper mantle source of MORBs. These deep mantle regions contain some proportion of relatively undifferentiated, primitive mantle. OIBs that do not have elevated 3He/4He may originate from shallower source regions. The upwelling, deep OIB mantle source also supplies material to the MORB mantle. The steady-state model for the upper mantle is a current paradigm in noble gas geochemistry (e.g., Allègre et al. 1986/1987; O’Nions 1987; Kellogg and Wasserburg 1990; O’Nions and Tolstikhin 1994, 1996; Porcelli and Wasserburg 1995a,b). This steady-state is reached through a balance of deep mantle input from below, slab subduction from above, and radiogenic production. Further discussion 3

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of the steady-state model is provided by Porcelli and Ballentine (2002, this volume). HELIUM Significance Helium isotope measurements in ocean ridge and island basalts provide some of the most basic geochemical information on mantle source reservoirs. More helium isotope analyses have been performed for oceanic volcanic rocks than for other noble gas species, and helium isotopes have played a leading role in the study of mantle heterogeneity. Helium isotope analyses are readily performed by modern mass spectrometers because there is a general absence of atmospheric contamination in samples due to the low concentration of helium in air (5.24 parts per million by volume at standard temperature and pressure). There are 2 naturally occurring isotopes of helium. 3He is much less abundant than 4He; for example, the atmospheric 3He/4He ratio (RA) is 1.39×10-6 (Mamyrin et al. 1970; Clarke et al. 1976). Nearly all of the terrestrial 4He has been produced as α-particles from the radioactive decay of 238U, 235U and 232Th over geological time, while nearly all of the 3He is primordial. Because helium undergoes gravitational escape from the thermosphere and has an atmospheric residence time of 1 to 10 million years, it is not recycled by plate tectonics to the Earth’s interior, and this makes the 3He/4He ratio unique among isotopic tracers of mantle sources involved in volcanism (Lupton 1983). By far the most important terrestrial source of 3He is degassing from the Earth’s interior. Excess 3He in volcanic rocks was first reported by Krylov et al. (1974) and Lupton and Craig (1975). The presence of this 3He in mantle-derived materials has profound implications; it means that the Earth is still outgassing volatiles that were trapped at the time of its accretion more than 4500 Ma ago. Other sources of 3He to the atmosphere include the auroral precipitation of solar wind, direct accretion from cosmic rays, and the flux of cosmic dust and meteorites (Lupton 1983). Small, but measurable amounts of 3He are produced in rocks at the Earth’s surface by high energy cosmic rays, predominantly from spallation of O, Si, Mg and Fe atoms, providing a means for determining surface exposure ages and erosion rates (Kurz 1986). Very small amounts of 3He are also produced during radioactive decay of U and Th as a result of neutron interactions with Li, by the reaction 6Li(n,α)3Η→3Ηe (Morrison and Pine 1955). The neutrons are produced by α-particle interactions on target elements such as Mg, Si and O in the host rock. This results in a low 3He/4He production ratio (typically 10-5) in mantle-derived lavas and fluids, to low values (~10-8) in continental regions due to increased amounts of radiogenic 4He. The general pattern in oceanic basalts is one in which MORBs show a relatively small range in 3He/4He (8.75±2.14 RA; Table 2), while OIBs are much more variable, and extend to values that are higher than the MORB mean by more than a factor of 4. Table 2. Helium isotopes in MORB glasses. Location n

Median Mean

Standard Skewness deviation

Atlantic 236 8.08

9.58

2.94

1.10

Pacific

245 8.14

8.13

0.98

0.04

Indian

177 8.24

8.49

1.62

1.73

All

658 8.11

8.75

2.14

1.92

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Radiogenic production The amount of 4He produced in a closed system, at secular equilibrium and over time t is given by 4

He* = 238U{8 (eλ238t - 1) + (7/137.88) (eλ235t - 1) + 6 κ (eλ232t - 1)}

(1)

where 238U is the amount present, λ238, λ235 and λ232 are the decay constants for 238U, 235U and 232Th (0.155125, 0.98485 and 0.049475 Gy-1, respectively), 1/137.88 is the present ratio of 235U/238U and κ is 232Th/238U. The range of κ values in the Earth’s mantle is about a factor of 2, from κ of ~2 to ~4 (Allègre et al. 1986; O’Nions and McKenzie 1993). By convention, the helium isotope ratio is often expressed as 3He/4He, i.e., as the non-radiogenic to radiogenic isotope, which causes some inconvenience when formulating changes due to radioactive ingrowth. The 3He/4He ratio in igneous samples, R, is usually expressed relative to the atmospheric ratio, i.e., as R/RA. Marine air provides a convenient and useful standard given its isotopic homogeneity for 3He/4He. Mid-ocean ridge basalts Global variability. Helium isotope variations along mid-ocean ridges have been extensively studied, and the variability along ridges is important to understand in the context of convective mixing and melt generation in the upper mantle. The difference in 3 He/4He ratio commonly observed between MORBs and OIBs is accepted by most investigators as evidence for two distinct mantle source regions. This viewpoint has been challenged by Anderson (2000a,b; 2001) on the basis of a statistical comparison of MORB and OIB data sets. The choice of basalt samples included in these data sets directly impacts the accuracy of any conclusions about the mantle, so it is important to discuss this choice in some detail and to make it wisely. Anderson (2001) chose to include lavas from Iceland as mid-ocean ridge basalts, on the grounds that the MidAtlantic Ridge passes through its center. He also included near-axis seamounts and backarc basin basalts. The inclusion of Iceland certainly affects the mean and variance in any such comparison, because the 3He/4He ratios near its center are higher than those observed anywhere along the global ocean ridge system. The inclusion of seamounts and back-arc basin basalts affects the variance much more than the mean, due to the presence of lavas that experienced shallow-level degassing and significant lowering of their 3 He/4He ratio by interaction with seawater or altered wallrock (e.g., Graham et al. 1988; Staudacher and Allègre 1989; Hilton et al. 1993). Anderson did not carry out a similar data compilation of the OIB data set, but used a mean value for 23 ocean islands summarized by Allègre et al. (1995) based on 276 individual OIB analyses. The OIB data included lavas from very different stages of evolution in the volcanic history of the islands, and the sources of the original data were not given in detail. Anderson (2001) compared the mean value for these 23 islands with the mean for ~500 individual MORB glasses, so the two data sets were also not treated in the same way. In this section I have made a first-order attempt to objectively estimate the mean 3He/4He, its variance and skewness, for mid-ocean ridges from the Atlantic, Pacific and Indian Oceans, and for all ridges collectively. I arrive at different values for the mean and standard deviation compared to Anderson’s work (Table 2). Mid-ocean ridge basalts display a relatively narrow range of 3He/4He, while oceanisland basalts are more variable and often extend to higher values. The full range in MORBs is between 1 and 18 RA. This range excludes subaerial lavas from Iceland, but not submarine lavas from the ridges to its south or north (Reykjanes or Kolbeinsey). In making this comparison, I also have not included seamounts or back-arc basin basalts, because their clear tectonic association with ridge mantle is questionable. I have included only one analysis from any individual sampling locality (i.e., one sample per dredge or

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rock core), as a very coarse attempt to avoid sampling bias from the few areas that have been studied in considerably more detail. I have also excluded a handful of glass samples with 3He/4He < 2.5 RA, which, given their differentiated and highly degassed nature ([He] 0.705 (White 1985). It is important to continually bear in mind that crustal material should have extremely low 3He/4He (~0.1 RA) if it is degassed and more than 10-100 million years old. Yet the lowest 3He/4He ratios observed at ocean island localities suggested to contain such recycled components are about 3-5 RA, or 300-500 times larger than the expected value. If recycled material is present in an OIB mantle source region, then the fraction of He from such a component must be extremely low for any of the island systems studied so far (e.g., Farley 1995). The islands of Tristan da Cunha and Gough in the South Atlantic are notable because they have Pb, Nd and Sr isotope characteristics similar to hypothetical values for the bulk silicate Earth. Kurz et al. (1982a) showed that 3He/4He was 5-6 RA at these islands, lower than MORB values (Fig. 10a), indicating that a primitive mantle reservoir is not involved. On this basis, Kurz et al. (1982a) suggested that recycled crust was present in the Tristan and Gough mantle source regions. In the Azores, 3He/4He ranges between 3.5 RA and 11.3 RA (Kurz et al. 1990; Moreira et al. 1999), and there is a good covariation of 3He/4He with Pb isotopes at the

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b

Figure 10. (a) He-Sr, (b) He-Nd, (c,d) He-Pb isotopic relations for mid-ocean ridge and ocean island basalts (after Graham et al. 1998 and references therein). The fields shown encompass paired analyses of the same samples in all cases, except that where boxes are shown they encompass the range of values at those localities. Selardalur, in northwest Iceland, is currently the locality having the highest measured 3He/4He ratio in a lava for which Sr, Nd and Pb isotope data are also available (Hilton et al. 1999).

Noble Gas Isotope Geochemistry of MORBs and OIBs

275 c

d

Figure 10, continued. (c,d) He-Pb isotopic relations for mid-ocean ridge and ocean island basalts. The 208Pb*/206Pb* is the isotopic ratio corrected for primordial Pb, and is related to the time-integrated Th/U ratio, designated κPb (Allègre et al. 1986).

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scale of the archipelago, indicating significant He isotope heterogeneity within the mantle source region (Moreira et al. 1999). The lowest 3He/4He is associated with elevated 207 Pb/204Pb in basalts from Sao Miguel (which are among the highest observed at any ocean island), while the highest 3He/4He occurs in lavas from Terceira that have relatively radiogenic 206Pb/204Pb and typical OIB 207Pb/204Pb. Moreira et al. (1999) interpret these variations with a 3-component mixing model involving lower mantle (having high 3He/4He) that was entrained into a mantle plume containing recycled crust (with radiogenic 206Pb/204Pb). The low 3He/4He signature at Sao Miguel is attributed to shallow level mixing between the plume and km-size rafts of continental lithosphere that were delaminated into the shallow mantle during Jurassic rifting and opening of the North Atlantic. The Cape Verde Islands show a pattern of 3He/4He and 206Pb/204Pb variation that is similar to that of the Azores, although with slightly less radiogenic Pb (Christensen et al. 2001). 3He/4He ratios in the Cape Verdes are both higher and lower than are typically found in MORBs. Values range between 3.2 and 13.8 RA in primitive lavas from the islands of Santo Antão and Fogo, and there is no systematic variation with age. These variations appear to be broadly consistent with the 3-component model of Moreira et al. (1999) for the Azores, but in detail the lithospheric component involved in the Cape Verdes must be different (Christensen et al. 2001). At Heard Island the 3He/4He ratios are bimodal (Hilton et al. 1995). The Laurens Peninsula volcanic series has 3He/4He between 16.2 and 18.3 RA, while the Big Ben series lavas have 3He/4He between 5 and 8.4 RA. Heard Island shows a range in 87Sr/86Sr and 206Pb/204Pb that is among the largest seen at a single ocean island (Barling et al. 1994), in some cases with extremely radiogenic Sr (87Sr/86Sr > 0.706). The Laurens Peninsula lavas are relatively homogeneous in Sr-Nd-Pb isotopes, while the Big Ben series is heterogeneous. Strongly curvilinear Pb-Nd and Pb-Sr isotope arrays, and linear Pb-Pb arrays, indicate that the genesis of Heard lavas involves binary mixing (Barling and Goldstein 1990). The bimodal distribution of 3He/4He does not follow this simple relationship, however, and Hilton et al. (1995) suggested that the low 3He/4He ratios were derived by contamination from radiogenic He in the Kergeuelen lithosphere, which may partly have a continental origin (Barling et al. 1994). Based on the He isotope results for Heard Island, Hilton et al. (1995) called into question whether the low 3He/4He seen at other ocean islands, such as Tristan da Cunha, Gough and St Helena could have an origin from recycled material, because those islands are characterized by relatively few He isotope analyses, and some of the phenocrysts have low He contents which could make them more susceptible to contamination by radiogenic He. Relatively low 3He/4He ratios also typify islands with the most radiogenic 206Pb/204Pb ratios (the ‘HIMU’ component, or high μ, where μ = 238U/204Pb) such as St. Helena and the Cameroon Line in the South Atlantic, and the Cook-Austral islands in the South Pacific (Graham et al. 1992a; Hanyu and Kaneoka 1997; Barfod et al. 1999). The HIMU component has been variably interpreted to originate from recycled ocean crust (Hofmann and White 1982; Zindler et al. 1982; Chauvel et al. 1992; Hanan and Graham 1996; Hanyu and Kaneoka 1997), carbonatite metasomatism (Tatsumoto 1984), intra-mantle differentiation by silicate melts (Halliday et al. 1990; Barfod et al. 1999), recycled and metasomatized continental lithosphere (McKenzie and O’Nions 1995), and recycled oceanic lithosphere (Moreira and Kurz 2001). The significance of low 3He/4He values at these ocean island localities is further complicated, because low ratios might also be produced by interaction with the oceanic lithosphere through which OIB magmas ascend, or by prolonged storage of degassed magma at crustal levels (Zindler and Hart 1986a; Graham et al. 1988). Thorough studies of “low3 He/4He” islands are needed to resolve many of these issues.

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At the present time, the myriad of mantle components inferred from the Sr-Nd-Pb isotopes cannot be adequately placed into a reservoir framework within the mantle. Many of the components appear to be ubiquitous. For example, the enriched type of mantle present at some ocean islands also appears to comprise a worldwide pollutant of the MORB mantle (Hanan and Graham 1996). Furthermore, while it is obvious that the MORB mantle is well characterized for 3He/4He, it is currently not possible to identify a single value of 3He/4He (or even a narrow range of values) that characterize each of the OIB isotopic end-members. One possible exception to this is the relatively uniform and low 3He/4He ratios (2-sigma above the air ratio. Darker symbols for the southern MAR data set crudely form a vector between the MORB line and solar compositions; these are lavas from the Shona and Discovery hotspot-influenced sections of the ridge, where 3He/4He ratios range up to 15 RA (Moreira et al. 1995; Sarda et al. 2000); see text for further discussion. Dashed lines depict mixing between air and mantle Ne for Iceland, Loihi, Réunion, Kerguelen and MORB (MAR popping rock), drawn through the highest 20Ne/22Ne or the most precisely determined compositions for each locality. The extrapolated 21Ne/22Ne ratio, corresponding to solar 20Ne/22Ne at each locality, is 0.035 for Iceland, 0.039 for Loihi, 0.043 for Réunion, 0.053 for Kerguelen and 0.075 for MORB.

A good estimate of the 21Ne/22Ne ratio of the mantle source can be determined from the realization that Ne in a volcanic rock is a binary mixture between air and mantle (e.g., Moreira et al. 1995). This estimate is similar to that determined by a ternary deconvolution using 20Ne/22Ne and 21Ne/22Ne ratios for air, primordial (here assumed to be solar for illustration) and nucleogenic end-members (e.g., Honda et al. 1993a), because the nucleogenic 22Ne production is minor compared to 21Ne. The extrapolated 21 Ne/22Ne corresponding to a solar 20Ne/22Ne ratio, 21Ne/22NeE, can be determined for a line that passes through the air and any data point on the Ne three-isotope diagram, as illustrated in Figure 12. Its numerical value is given by 21

Ne/22NeE = (21Ne/22NeM - 21Ne/22NeA)/ƒ22 + 21Ne/22NeA

(3)

where E, S, A and M refer to extrapolated, solar, air and measured values, respectively, and where ƒ22 is the proportion of mantle-derived Ne in a sample, i.e.,

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20Ne/22NeM - 20Ne/22NeA 20Ne/22NeS - 20Ne/22NeA .

(Note: 20Ne/22NeS = 13.8, 21Ne/22NeS = 0.0328, 20Ne/22NeA = 9.8 and 21Ne/22NeA = 0.029).

Figure 12. Schematic illustration for determining 21Ne/22NeE, the extrapolated 21Ne/22Ne ratio of the mantle. S, A, M and E correspond to solar, air, measured and extrapolated values.

Mid-ocean ridge basalts In a diagram of 20Ne/22Ne vs. 21Ne/22Ne, ocean ridge basalts form a well defined array that passes through the composition of air and extends to values of 20Ne/22Ne ≥ 12.5 and 21Ne/22Ne ≥ 0.07 (Fig. 11). This array is usually interpreted as a mixing line between upper mantle Ne and air Ne. How and when this contamination by air occurs is conjectural. Some investigators propose that it takes place during magma ascent and eruption (e.g., Farley and Poreda 1993), or that it was introduced by ancient subduction (e.g., Sarda et al. 2000); others argue that it is largely an artifact introduced during sample preparation (e.g., Ballentine and Barfod 2000). The variability of repeat analyses by stepwise heating makes it evident that some allowance should be made for the effects of air contamination in every analysis, and the measured Ne isotope ratio should always be taken as a minimum estimate for the mantle source composition. Global variability. In general, the Ne isotope measurements have relatively large and correlated analytical uncertainties. The scatter about the MORB Ne isotope array therefore makes it difficult to determine whether the upper mantle is characterized by a single Ne isotope composition. It seems likely that the whole mantle is characterized by a uniform value of 20Ne/22Ne, but this is an unproven assumption. Because there are no known reactions that produce 20Ne to a significant extent, variable 20Ne/22Ne in the mantle seems implausible unless nucleogenic production of 22Ne is significant and spatially variable. Two MORB suites, from the southern East Pacific Rise (Niedermann et al. 1997) and the southern Mid-Atlantic Ridge (Moreira et al. 1995), each show a decrease in 20Ne/22Ne with increasing 21Ne/22Ne, with a vector that points approximately to the solar Ne isotope composition (especially for the southern Mid-Atlantic Ridge

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samples with higher 3He/4He ratios near the Shona hotspot; see Fig. 11). Based on this evidence, Niedermann et al. (1997) speculated that the mantle could vary spatially in the production of 22Ne (and consequently in its 20Ne/22Ne ratio), due to a heterogeneous distribution of fluorine and 22Ne production via the reaction 19F(α,n)22Na(β+)22Ne. However, the recent estimates by Yatsevich and Honda (1997) indicate that 22Ne production in the mantle is 0.704) cannot be explained by mixing of either recycled crust or MORB mantle with a high 3He/4He-source such as that for Loihi or Iceland. Hanyu et al. (2001) suggest that the Réunion plume source is therefore distinct from the Iceland and Loihi sources, and that the mantle contains more than one relatively undegassed reservoir, each having elevated but different 3He/4He ratios. This is an intriguing hypothesis that warrants further consideration, especially in light of developing models for deep mantle stratification that involve large domains having a complex topography that responds to convection and pressure from subducting slabs (e.g., Kellogg et al. 1999). Another possibility to bear in mind, however, is that some decoupling of He and Ne isotopes may occur if the mantle He/Ne ratio is fractionated during melting processes. Solar hypothesis. The steep OIB correlations on the three-Ne isotope diagram, especially for basalts from Iceland and Loihi Seamount, trend toward Ne isotope compositions resembling either solar wind or the solar component commonly found in meteorites (Ne-B). It is not possible at the present time to distinguish unambiguously between these two possibilities, and the implications for early Earth evolution are quite different. The highest values of 20Ne/22Ne yet measured are ~12.5 for terrestrial lavas from several localities, including Loihi Seamount, Iceland and the popping rocks. This led Trieloff et al. (2000) to suggest that the Earth’s initial Ne is best described by Ne-B, a component produced in meteorites by irradiation from the solar wind. Following this reasoning, Trieloff et al. (2000) argued that the Earth acquired its Ne while the young sun was still very active, during which time small planetesimals were thoroughly irradiated

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prior to their assembly to form the Earth. In contrast, Ballentine et al. (2001) argued that, because all basalt analyses can be expected to contain some air neon, a solar Ne isotopic composition (20Ne/22Ne of ~13.8) best represents the Ne trapped within the Earth. This conclusion is based on the correlation between Ne and Ar isotopes seen in step crushing analyses of the MAR popping rock (Moreira et al. 1998; see Fig. 14, below) and its extrapolation to the very high 40Ar/36Ar measured by laser extraction from individual bubbles in the same lavas (Burnard et al. 1997), for which precise Ne isotope measurements were not possible. Following this reasoning, the Earth would have solar Ne isotope composition, and gravitational capture of a dense primitive atmosphere within the solar nebula during Earth accretion is a logical model (Mizuno et al. 1980). ARGON Significance Argon isotopes can provide powerful constraints on the formation of the atmosphere, and on geodynamics and isotopic evolution of the mantle. Unlike He which is lost from the atmosphere, Ar has accumulated over Earth history. The primordial 40Ar/36Ar ratio is also extremely low, so the amount of 40Ar initially present in the Earth can be reasonably taken as zero. Transport of 40Ar to the atmosphere involves volcanic degassing, hydrothermal circulation through the crust, and erosion of continental crust that releases radiogenic Ar generated by decay of 40K. The half-life of 40K (1.25 Gy) is relatively short compared to the age of Earth, making the abundance of Ar in the atmosphere along with the 40Ar/36Ar of the atmosphere and upper mantle useful parameters for unraveling the history of degassing and plate tectonic recycling (e.g., Turekian 1959; Hamano and Ozima 1978; Fisher 1978; Sleep 1979; O’Nions et al. 1979; Hart et al. 1979; Sarda et al. 1985; Allègre et al. 1986/1987; Turner 1989; Albarède 1998; Coltice et al. 2000; Porcelli and Ballentine 2002). As one example, in the limiting case where all the 40Ar generated in the crust has been released to the atmosphere, 30% of the Ar in the atmosphere would be from the crust and 70% from the mantle (Turcotte and Schubert 1988). Under these circumstances, only one-quarter of the whole mantle would be outgassed. As another example, under assumptions of a layered mantle (in which the 660 km seismic discontinuity represents the boundary) with known K content, about one-half of the terrestrial 40Ar inventory is still retained in the deep Earth (Allègre et al. 1996). These conclusions depend directly on the K concentration of the Earth; a lower K/U ratio than commonly assumed leads to the conclusion of a more degassed deep Earth (Davies 1999). Radiogenic production In a closed system, the amount of radiogenic 40Ar produced over time t is given by 40

Ar* = a·b·K (eλ40t – 1)

(4)

where a is the natural abundance of 40K (0.0117%), b is the branching ratio of 40K decay to 40Ar by electron capture (0.1048), K is the amount of potassium present, and λ40 is the total decay constant (0.554 Gy-1). Atmospheric contamination Obtaining meaningful Ar isotope analyses on basalts is analytically challenging, because contamination by atmospheric components is notorious. Potential sources of this contamination include seawater, altered wall rock, or air itself. Basalt glasses may show complicated variations within an individual glassy rim, with lower 40Ar/36Ar in both the outermost glassy part as well as in the more crystalline interior (Kumagai and Kaneoka 1998). The atmospheric contaminant in MORBs appears to be significantly less in

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vesicles, because all measurable 36Ar in the popping rock is accounted for by the amount released during the analysis of vesicle-free glass (Burnard et al. 1997), and there is no correlation between 3He and 36Ar from basalt crushing analyses (Ballentine and Barfod 2000). This suggests that for MORBs the contamination may not be occurring by crustal assimilation, but rather by adsorption or entrapment following eruption. In the case of many ocean islands, phenocrysts are separated from subaerially erupted lavas for noble gas analysis. Such samples usually show a large range in measured 40Ar/36Ar due to variable proportions of atmospheric and magmatic components (Farley and Craig 1994). The air component is not readily removed by physical or chemical treatment. It is also present to a large extent in fluid inclusions, along with the magmatic gas, as revealed by laser fusion analyses (Farley and Craig 1994; Burnard et al. 1994a). This indicates that the contamination may also occur by assimilation of crust, or by direct addition of seawater or air to a magma (Farley and Craig 1994). Argon isotope ratios in OIBs that resemble air are most likely due to contamination, and should not be taken as evidence for a deep mantle reservoir having near-atmospheric Ar isotope composition (Patterson et al. 1990). In fact, where OIBs have been analyzed that have high 3He/4He, they can range to moderately high 40Ar/36Ar ratios, up to ~8000 (Farley and Craig 1994; Burnard et al. 1994a,b; Trieloff et al. 2000). Ballentine and Barfod (2000) reviewed the effects of atmospheric contamination in MORBs and OIBs in detail. Intriguingly, they showed that the modern air contaminants released during crushing are preferentially located in sites that are more easily ruptured than those containing magmatic gas. This suggests that much of the contamination occurs through the annealing of microfractures, following air entrapment that occurs either during sample recovery from seafloor depths or during sample preparation in the laboratory. This contaminant gas should, in principle, be separable from magmatic gas by stepwise crushing or heating, although a diagnostic test of when it is completely removed remains to be fully developed. Measured 40Ar/36Ar ratios should always be considered minimum values in young basalts. It is commonly assumed that all 36Ar present in a MORB or OIB sample could be derived from air contamination. This is a reasonable assumption provided that the 40 Ar/36Ar is >3000, as seems likely for mantle source regions. This assumption allows a minimum estimate for the amount of magmatic 40Ar present in a sample. This amount of radiogenic Ar derived from the mantle, denoted 40Ar*, is given by 40

Ar* = (40Ar/36ArM -40Ar/36ArA) 36ArM

(5)

where M designates the measured ratio and A the air ratio (40Ar/36ArA = 295.5). There is no straightforward way to estimate a sample’s complement of primordial 36Ar derived from the mantle. For example, if one assumes a mantle 40Ar/36Ar ratio (such as 40 Ar*/36ArP = 40,000, where p refers to primordial Ar), any correlation involving 36ArP (such as 3He/36ArP) will be inherently the same as one involving 40Ar*, since both are directly correlated by the assumed 40Ar*/36ArP ratio. Mid-ocean ridge basalts Global variability. The mantle 40Ar/36Ar ratio is variable due to time-integrated variations in K/36Ar. MORBs show 40Ar/36Ar ratios that range up to 40,000. The very highest values have been measured using a laser to rupture individual vesicles in the MAR popping rock (Burnard et al. 1997). Values up to 28,000 have been measured for these same samples by incrementally heating or crushing bulk samples of the glass in vacuum (Staudacher et al. 1989; Moreira et al. 1998). Very high ratios, with relatively large uncertainties, have also been measured in glassy basalts from the Mid-Atlantic Ridge (40Ar/36Ar = 42,400±9700; Marty and Humbert 1997) and from 10°S on the

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southern East Pacific Rise (40Ar/36Ar>35,000, limited by the absence of detectable 36Ar; Fisher 1994). Ratios above 25,000 have also been measured in MORB glasses from the MAR near the Kane transform (Sarda et al. 1985), from the South Atlantic (Sarda et al. 2000), and from the N Chile Ridge in the eastern Pacific (Niedermann and Bach 1998). It is tempting to suggest that the whole upper mantle might be characterized by a 40Ar/36Ar value near 30,000 based on this similarity. However, individual analyses of the popping rock show the highest measured 40Ar/36Ar (>20,000) during the last two steps of crushing which release the lowest amounts of 36Ar (2σ above the air ratio. Data sources are given in the caption to Figure 11.

There is no systematic behavior between 40Ar/36Ar and 3He/4He for the upper mantle based on the MORB analyses. 40Ar/36Ar ratios for MORBs are quite variable, ranging by 2 orders of magnitude, while 3He/4He varies by about a factor of 2 (Fig. 13). For the case of the popping rock, the highest 40Ar/36Ar ratios are also associated with the highest measured 20Ne/22Ne (Fig. 14), consistent with the lowest degree of atmospheric contamination in those analyses. Moreira et al. (1998) inferred that the maximum 40 Ar/36Ar of the MORB mantle is 44,000, based on extrapolation of the popping rock 40 Ar/36Ar-20Ne/22Ne trend to solar 20Ne/22Ne = 13.8 (Fig. 14). High 40Ar/36Ar ratios are also found in other samples with lower 20Ne/22Ne, such as from the N Chile Ridge and southern MAR. Much of the variation observed in Figure 14 must be due to air contamination. Different degrees of air contamination would only produce a single trajectory however, so the scatter in the data indicates a significant variation in the Ne/Ar ratio for either, or both, of the mixing end-members (i.e., the atmosphere-derived contaminant and the magmatic gas). For example, to account for the data scatter, any binary mixing between magmatic gas and unfractionated air would require that the

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22

Ne/36Ar ratio of the magmatic gas ranges from an order of magnitude larger to an order of magnitude smaller than the air ratio. From a simultaneous consideration of Ne-Ar and Ne-Xe isotope systematics, Harrison et al. (2002) showed that contamination usually involves a fractionated atmospheric component having 22Ne/36Ar significantly less than air, such as might be expected if water-rich or adsorbed components play a role. This seriously complicates a full interpretation of even the best available heavy noble gas analyses. Measured 40Ar/36Ar ratios should be considered minimum estimates for their mantle source.

Figure 14. 40Ar/36Ar vs. 20Ne/22Ne. Data sources are given in the caption to Figure 11. Plotted points include individual analysis steps (crushing or heating), but only when the measured Ne and Ar isotope ratios were >2σ above the air ratio.

Regional variability. Given the strong potential effect of contamination by atmospheric components, it is difficult to quantitatively assess the spatial variability in 40 Ar/36Ar along ridges. Sarda et al. (1999a) showed that the maximum 40Ar/36Ar ratio along the Mid-Atlantic Ridge, measured by step heating of basalt glasses, varies systematically with indicators of mantle heterogeneity such as Pb isotopes, with one endmember defined by the Azores hotspot. The correlations hold for the North and South Atlantic considered as a whole. Sarda et al. (1999a) reasoned that this ruled out shallow level contamination processes, and favored mixing that involved depleted upper mantle and recycled materials having low 40Ar/36Ar and radiogenic Pb. This would require that the recycled material has a K/36Ar ratio less than upper mantle values, i.e., that atmospheric noble gases are recycled to the mantle by subduction. Burnard (1999b) challenged this explanation however, and argued that shallow level contamination could explain the regional correlation between Ar and Pb isotopes. Radiogenic Pb signatures are typically found along axial topographic highs, where ridges are often influenced by nearby island hotspots. Low 40Ar/36Ar is also found along shallower ridge sections where degassing at shallower levels can lead to an increased susceptibility to contamination, and this can produce a correlation between 40Ar/36Ar and axial depth (Burnard 1999b;

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Burnard et al. 2002). Sarda et al. (1999b) suggested that such a process should produce erratic results rather than a systematic trend. Whether the regional variation in Ar isotopes along ocean ridges carries information about variations in the upper mantle is still debated. Chlorine contents in suites of MORB glasses studied for Ar isotopes could potentially resolve this debate (e.g., Michael and Cornell 1998). Relation to other volatiles. The N2/36Ar ratio varies by several orders of magnitude in MORB glasses, and is correlated with 40Ar/36Ar (Marty 1995). This means that the nitrogen abundance correlates with the amount of 40Ar, due to two-component mixing between a mantle-derived end-member having high 40Ar/36Ar and an atmospheric endmember having 40Ar/36Ar = 295.5. The trend of increasing N2/36Ar with 40Ar/36Ar in MORB glasses can be extrapolated to an upper mantle 40Ar/36Ar > 30,000, and gives N2/36Ar >2.2×106 (Marty 1995), which is two orders of magnitude higher than the value in the modern atmosphere. This could mean that small amounts of nitrogen have been recycled to the volatile-depleted upper mantle. Marty and Humbert (1997) and Marty and Zimmerman (1999) subsequently showed that the high 40Ar/36Ar ratios of the upper mantle are accompanied by a depleted 15N signature relative to air, with δ15N = -3 to –5 %0. This indicates that only a very limited amount of nitrogen recycling could have occurred, and that atmospheric nitrogen is not derived solely by outgassing of the upper mantle. Instead, it seems possible that the early atmosphere underwent volatile fractionation of N relative to 36Ar. The high degree of correlation between N2 and 40Ar abundances in MORB glasses would then indicate some parallel degassing of N2 and 40Ar over geologic time (Marty 1995). A detailed model of mantle degassing, early atmospheric dissipation and fractionation, and recycling fluxes using combined He, Ne, Ar and N isotope constraints from oceanic basalts has been presented by Tolstikhin and Marty (1998). Solar hypothesis. Ratios of noble gas isotopes that have only a primordial component, such as 38Ar/36Ar, are key in any attempt to identify potential solar components within the Earth. Anomalously low 38Ar/36Ar was first recognized by Niedermann et al. (1997) in their measurements of East Pacific Rise basalts. They did not place much significance on those results given that only a few of them were outside the 2-sigma uncertainty of the air ratio (air 38Ar/36Ar = 0.188). Low ratios, down to 38Ar/36Ar = 0.185, were also measured for the deeply erupted glassy basalts at Loihi Seamount (Valbracht et al. 1997). The low 38Ar/36Ar ratios at Loihi were correlated with nonatmospheric and elevated 20Ne/22Ne, providing preliminary evidence for a solar Ar component in the mantle. (The mean value of 38Ar/36Ar is 0.175 in solar wind). Subsequently, Pepin (1998) took the results from both the EPR and Loihi studies as promising evidence for a solar-like Ar component in the mantle, and discussed some of the implications for planetary accretion and atmospheric origin. The low 38Ar/36Ar ratios in the EPR samples are not correlated with high 40Ar/36Ar, however, as would be expected from binary mixing between air and mantle components (Kunz 1999). Indeed, some of those low 38Ar/36Ar ratios were measured simultaneously with atmospheric 40 Ar/36Ar ratios. It should also be recognized that very small analytical artifacts in the original studies might be the cause for the low 38Ar/36Ar ratios. Mass fractionation effects during gas extraction or analysis must be completely removed, and this is especially difficult to verify in the presence of relatively large amounts of 40Ar. At this point in time it seems possible that a solar Ar component is present within the mantle, but the analytical evidence is weak and it remains unproven. Ocean island basalts The highest 40Ar/36Ar ratios at ocean islands are found in xenoliths from Samoa, where values range up to 21,500 (Farley et al. 1994). The origin of the high 40Ar/36Ar in

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these xenoliths is unclear, however, because the fluid inclusions show strong metasomatic overprints that occurred in the upper mantle (Burnard et al. 1998). The highest 40Ar/36Ar ratios obtained on ocean island basalt glasses are from Loihi Seamount, where they range up to 8300 (Trieloff et al. 2000). In the most detailed study of Ar isotopes at a single ocean island, Farley and Craig (1994) obtained a maximum 40Ar/36Ar of 8000 for olivine phenocrysts from Juan Fernandez, a high 3He/4He hotspot. Bulk crushing and laser fusion of individual crystals were carried out for forty splits of olivine from a single lava, and a variety of chemical and physical pre-treatments were attempted. The relation between 40Ar/36Ar and 3He/4He or 20Ne/22Ne in ocean island basalts and xenoliths is generally poor (Figs. 13 and 14). The apparent difference in 40Ar/36Ar– 3 He/4He systematics between MORBs and OIBs was suggested by Kaneoka (1983, 1985) to indicate a layered mantle structure, in which OIB sources have higher 3He/4He and 40 Ar/36Ar ratios that are nearly the same as the atmosphere. More recent data showing 40 Ar/36Ar up to ~8000 in OIBs now makes a model involving OIB source reservoirs of near atmospheric composition untenable. High 40Ar/36Ar values, particularly in the Loihi glasses (Trieloff et al. 2000), are associated with 20Ne/22Ne ratios comparable to the highest values obtained in terrestrial basalts, evidence that these Loihi analyses are not very strongly affected by air contamination. The mantle source for OIBs from localities such as Iceland, Hawaii and Réunion clearly has a 40Ar/36Ar ratio higher than the atmosphere (Trieloff et al. 2000; Burnard et al. 1994b). It is not possible to assume that all air contamination has been eliminated even in these first rate analyses. A ratio of 40 Ar/36Ar ≥ 8000 is arguably the best estimate for the mantle source of these OIBs (Fig. 14), although it is not currently possible to rule out the presence of a deep OIB source reservoir, having somewhat lower 40Ar/36Ar, that generates OIB magma that subsequently undergoes slight contamination by volatiles derived from the shallower MORB mantle. KRYPTON Krypton isotopes in oceanic basalts have not been very diagnostic of mantle processes. The Kr isotopic composition of oceanic basalts is typically the same as modern air. Minor production of 83Kr, 84Kr and 86Kr occurs from the spontaneous fission of 238U, but it has not yet been detected in young oceanic basalts, due to the small fission yields (109 years). Some of the observed He-Ne isotope variation suggests that mixing, either between mantle reservoirs or between different magmas, may have occurred relatively recently. For example, the trend may be produced by interaction of OIB source material, derived from the deep mantle and having solar-like 21Ne/22Ne and low 4He/3He, with MORB source material, derived from the upper mantle and having 4He/3He = 84,600 (3He/4He = 8.5 RA) and 21Ne/22Ne = 0.075 (Moreira and Allègre 1995; Moreira et al. 2001). If r is defined as the concentration ratio [3He/22Ne]MORB/[3He/22Ne]OIB, where the subscripts designate the end-member components on Figure 18, a mixing line which best describes the data from Iceland would have r between 1 and 10. Compared to the OIB end-member, the MORB endmember must be enriched in He relative to Ne in this instance. This could indicate that any such mixing beneath Iceland is between a MORB magma that was previously degassed (leading to a higher He/Ne ratio as predicted from solubility considerations), and an undegassed, plume-like magma (Moreira et al. 2001). However, the observed scatter in Figure 18 for oceanic basalts also indicates a significant variation in the 3 He/22Ne ratio from one location to another, and this might be due to differences in the type of mixing. For example, the Loihi data would be better described by r values 40 RA, supporting the existence of mantle plumes from deep in the Earth. Best model cases are Iceland and Hawaii. Temporal and spatial variations at ocean island hotspots may be weakly associated with variations in plume flux and associated melting processes, but there is no systematic global behavior between plume flux and 3He/4He signal. Some ocean island hotspots, such as Réunion, may have relatively constant and intermediate 3He/4He ratios indicating a duration of 107-108 years. This suggests the possibility of deep mantle reservoirs having variable 3He/4He.

3. The mantle is characterized by elevated ratios of 20Ne/22Ne and 21Ne/22Ne compared to the atmosphere. Mid-ocean ridge basalts and ocean island basalts show distinct 20 Ne/22Ne-21Ne/22Ne trends that result from differences in the dilution, by primordial Ne, of the nucleogenic 21Ne that is produced in their respective mantle sources. The

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MORB trend passes through the composition of air and extends to values 20 Ne/22Ne≥12.5 and 21Ne/22Ne≥0.07. OIB trends are much steeper; in the case Iceland they pass through the composition of air and extend to values 20 Ne/22Ne≥12.5 and 21Ne/22Ne≤0.035. It seems likely that the whole mantle characterized by a uniform value of 20Ne/22Ne, but this has not been proven.

of of of is

4. MORB and OIB samples show a range of measured 40Ar/36Ar that is mostly due to variable proportions of atmospheric and magmatic components. The highest 40 Ar/36Ar ratio measured in a MORB is ~40,000. The correlation between 40Ar/36Ar and 20Ne/22Ne for the MAR popping rock, when extrapolated to a solar 20Ne/22Ne = 13.8, indicates a possible upper limit of 44,000 for 40Ar/36Ar in the upper mantle. High 3He/4He ocean islands range to moderately high 40Ar/36Ar ratios, up to ~8000. This may be the minimum value appropriate for the mantle source of OIBs. 5.

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Xe/130Xe anomalies are present in MORBs from all the major ocean basins. Small anomalies also occur in 131,132,134,136Xe/130Xe ratios, and these are correlated with the 129 Xe/130Xe anomalies. The MORB excesses in 129Xe and 131,132,134,136Xe are derived from extinct radioactivity of 129I, and fission of 238U and extinct 244Pu. The largest MORB 129Xe/130Xe and 136Xe/130Xe ratios, and the most precisely determined, are for the popping rock, which range up to 7.73 and 2.57, respectively. The maximum 129 Xe/130Xe of the MORB mantle may be near 8.2, based on extrapolation of the popping rock Xe-Ne isotope trend to solar 20Ne/22Ne = 13.8 (Fig. 16). Ocean island basalts and xenoliths show smaller Xe isotope anomalies than MORBs. These small excesses in radiogenic and fissiogenic Xe appear to be correlated, but given the analytical uncertainties this is not firmly established. It is also possible that the OIBs have been contaminated by MORB-type Xe during magma ascent.

6. Coupled He-Ne isotope systematics reveal that the highest values of 3He/4He at ocean island localities are accompanied by the lowest values of 21Ne/22NeE, the 21 Ne/22Ne ratio corrected for air contamination and extrapolated to solar 20Ne/22Ne. The general 3He/4He-21Ne/22NeE behavior (Fig. 18) suggests that either mixing between OIB and MORB components has occurred, or that the Earth’s mantle is characterized by heterogeneity in 3He/22Ne ratio that is about a factor of 20. This heterogeneity may derive from fractionation processes associated with melting and degassing, or it may be primordial heterogeneity preserved since planetary accretion. The mean time-integrated 3He/22Ne ratio, inferred from coupled He-Ne isotope systematics, is similar for MORBs and OIBs (~9 and ~6, respectively) and appears to be higher than the ratio in the solar wind (~4) (Fig. 17). 7. Relatively recent fractionations dominate the measured ratios of 4He/40Ar*, 3 He/22NeS, 4He/21Ne* and 4He/136Xe* in MORBs and OIBs. These fractionations are controlled by magma degassing and by partial melting effects. Notable exceptions are deeply erupted basalts from Loihi Seamount and the suite of basalt glasses from Iceland studied by Trieloff et al. (2000), and the Mid-Atlantic Ridge popping rock studied by Moreira et al. (1998). These samples contain noble gases with radiogenic abundance ratios resembling values expected for the mantle. Abundance ratios of the primordial noble gases (3He/36Ar, 22Ne/36Ar and 130Xe/36Ar) in the mantle sources of these basalt suites are significantly fractionated from solar ratios. The 3He/22Ne ratio is an exception, and is within a factor of three of the solar ratio. SUMMARY Variations in the isotope compositions of noble gases are related to processes controlling the distribution of K, U and Th, the major heat producing nuclides in the Earth. Mid-ocean ridge basalts have 3He/4He ratios between 3 and 15 RA, while ocean

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island basalts have ratios between 3.5 and 43 RA. The ubiquitous presence of ‘excess’ 3 He in mantle-derived rocks establishes that primordial volatiles are still escaping from the Earth’s interior. In addition, 3He/4He ratios above 15 RA are present only at ocean island localities, indicating a lower time-integrated (U+Th)/3He in their mantle source regions. The whole mantle is also characterized by elevated ratios of 20Ne/22Ne and 21 Ne/22Ne compared to the atmosphere. The 20Ne/22Ne ratio appears to be relatively uniform throughout the mantle, resembling either solar wind or the solar component trapped in meteorites and the atmosphere of Jupiter. The 21Ne/22Ne ratio is variable in the mantle due to nucleogenic production of 21Ne in reservoirs having different (U+Th)/22Ne ratios. This nucleogenic contribution is relatively larger in the MORB mantle source compared to some OIB sources (most notably at Iceland, Hawaii and Réunion), leading to higher 21Ne/22Ne in MORBs. The mantle 40Ar/36Ar ratio is variable due to variations in K/36Ar, with MORBs displaying the highest 40Ar/36Ar ratios, up to ~40,000. The upper mantle source of MORBs is therefore characterized by higher radiogenic/primordial noble gas isotope ratios due to its degassed nature and consequent higher parent/daughter ratios (e.g., (U+Th)/3He and K/36Ar). The relatively undegassed nature of the mantle source for some ocean islands is also clearly recognized from 3He/4He, 20Ne/22Ne21 Ne/22Ne and 40Ar/36Ar in oceanic basalts. The contrast between MORB and OIB noble gas isotope compositions is the most fundamental geochemical evidence for some mode of stratification within the Earth’s mantle. Ocean island basalts from localities such as Hawaii, Iceland and Réunion have a deeper mantle source than mid-ocean ridge basalts. The near solar-like Ne isotope compositions, the high 40Ar/36Ar ratios, and the nonatmospheric 129Xe/130Xe and 136Xe/130Xe in oceanic basalts also provide basic clues on the origin and evolution of the atmosphere. The atmosphere appears to have ‘formed’ within the first 50-70 Ma of Earth’s history, probably coincident with the period of terrestrial core formation (Halliday and Lee 1999). The time-integrated ratio of the primordial nuclides 3He and 22Ne in the mantle may be inferred from coupled He-Ne isotope systematics. This time-integrated ratio, , is independent of measured concentrations but rests on assumptions for the initial isotope composition of the solid Earth, usually taken to be solar. MORBs and OIBs show considerable overlap using this approach, and on average have ratios between 6 and 9, seemingly higher than the solar ratio of 4. Other methods for comparing relative abundance ratios in basalts depend on measured noble gas concentrations, and are subject to more uncertainty from elemental fractionations that occur both in nature and during sample preparation or analysis. Key ratios estimated in this way include 3 He/22Nesolar, 4He/21Ne*, 4He/40Ar* and 4He/136Xe* (where the * designates a radiogenic or nucleogenic component after correction for air contamination). Abundance ratios computed in this way may be compared to expected mantle source values, assuming radioactive equilibrium and known values of K/U and Th/U in the mantle source. Differences from the expected mantle values provide information about the style of magma degassing (e.g., bulk vs. fractional) and about crystal/magma partitioning during partial melting. Much of our knowledge about the noble gas geochemistry of the mantle relies on the analyses of a limited number of samples from three key areas; the popping rock from 14°N on the Mid-Atlantic Ridge, submarine basalt glasses from Loihi Seamount, and subglacial basalt glasses from Iceland. It is amazing in retrospect how much information about the Earth’s mantle is contained in the noble gas isotope compositions of these few samples. We should anticipate that much more remains to be learned, as analytical techniques improve and as other key sample localities are discovered.

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Staudacher T, Sarda P, Richardson SH, Allègre CJ, Sagna I, Dmitriev LV (1989) Noble gases in basalt glasses from a Mid-Atlantic Ridge topographic high at 14°N: Geodynamic consequences. Earth Planet Sci Lett 96:119-133 Steinberger B, O’Connell RJ (1998) Advection of plumes in mantle flow: implications for hotspot motion, mantle viscosity and plume distribution. Geophys J Intl 132:412-434 Stolper EM, Holloway JR (1988) Experimental determination of the solubility of carbon dioxide in molten basalt at low pressure. Earth Planet Sci Lett 87:397-408 Stuart FM, Ellam RM, J. HP, Fitton JG, Bell BR (2000) Constraints on mantle plumes from the helium isotope composition of basalts from the British Tertiary Igneous Province. Earth Planet Sci Lett 177:273-285 Sturm ME, Klein E, M., Graham DW, Karsten J (1999) Age constraints on crustal recycling to the mantle beneath the southern Chile Ridge: He-Pb-Sr-Nd isotope systematics. J Geophys Res 104:5097-5114 Sun S-S, Jahn BM (1975) Lead and strontium isotopes in post-glacial basalts from Iceland. Nature 255:527 Sun S-S, Tatsumoto M, Schilling J-G (1975) Mantle plume mixing along the Reykjanes Ridge axis: Lead isotopic evidence. Science 190:143-147 Swindle TD, Caffee MW, Hohenberg CM, Taylor SR (1984) I-Pu-Xe dating and the relative ages of the Earth and Moon. In Hartmann WK, Phillips RJ, Taylor GJ (eds) Origin of the Moon. Lunar and Planetary Institute, Houston, p 331-357 Tackley PJ (1998) Three-dimensional simulations of mantle convection with a thermo-chemical basal boundary layer: D” ? In Gurnis M, Wysession ME, Knittle E, Buffett BA (eds) The Core-Mantle Boundary Region, 28. American Geophysical Union, Washington, DC, p 231-253 Tatsumoto M, Unruh DM, Stille P, Fujimaki H (1984) Pb, Sr and Nd isotopes in oceanic island basalts. In Proc 27th Intl Geol Congr 2, Geochemistry and Cosmochemistry, p 485-501 Taylor RN, Thirlwall MF, Murton BJ, Hilton DR, Gee MAM (1997) Isotopic constraints on the influence of the Icelandic plume. Earth Planet Sci Lett 148:E1-E8 Thompson L (1980) 129Xe on the outgassing of the atmosphere. J Geophys Res 85:4374-4378 Tolstikhin IN, Marty B (1998) The evolution of terrestrial volatiles: a view from helium, neon, argon and nitrogen isotope modeling. Chem Geol 147:27-52 Trieloff M, Kunz J, Clague DA, Harrison D, Allègre CJ (2000) The nature of pristine noble gases in mantle plumes. Science 288:1036-1038 Trull TW (1994) Influx and age constraints on the recycled cosmic dust explanation for high 3He/4He ratios at hotspot volcanos. In Matsuda J (ed) Noble Gas Geochemistry and Cosmochemistry. Terra Scientific, Tokyo, p 77-88 Trull TW, Nadeau S, Pineau F, Polvé M, Javoy M (1993) C-He systematics in hotspot xenoliths: implications for mantle carbon contents and carbon recycling. Earth Planet Sci Lett 118:43-64 Turcotte DL, Schubert G (1988) Tectonic implications of radiogenic noble gases in planetary atmospheres. Icarus 74:36-46 Turekian KK (1959) The terrestrial economy of helium and argon. Geochim Cosmochim Acta 17:37f Turner G (1989) The outgassing history of the Earth’s atmosphere. J Geol Soc 146:147-154 Valbracht PJ, Staudacher TJ, Malahoff A, Allègre CJ (1997) Noble gas systematics of deep rift zone glasses from Loihi Seamount, Hawaii. Earth Planet Sci Lett 150:399-411 Valbracht PJ, Honda M, Staudigel H, McDougall I, Trost AP (1994) Noble gas partitioning in natural samples: results from coexisting glass and olivine phenocrysts in four Hawaiian submarine basalts. In Matsuda J (ed) Noble Gas Geochemistry and Cosmochemistry. Terra Scientific, Tokyo, p 373-381 Valbracht PJ, Honda M, Matsumoto T, Mattielli N, McDougall I, Ragettli R, Weis D (1996) Helium, neon and argon isotope systematics in Kerguelen ultramafic xenoliths: implications for mantle source signatures. Earth Planet Sci Lett 138:29-38 van der Hilst R, Kárason H (1999) Compositional heterogeneity on the bottom 1000 kilometers of Earth’s mantle: toward a hybrid convection model. Science 283:1885-1888 van der Hilst RD, Widiyantoro S, Engdahl ER (1997) Evidence for deep mantle circulation from global tomography. Nature 386:578-584 van Keken PE, Ballentine CJ (1998) Whole-mantle versus layered mantle convection and the role of a high-viscosity lower mantle in terrestrial volatile evolution. Earth Planet Sci Lett 156:19-32 van Keken PE, Ballentine CJ (1999) Dynamical models of mantle volatile evolution and the role of phase transitions and temperature-dependent rheology. J Geophys Res 104:7137-7151 van Keken PE, Ballentine CJ, Porcelli D (2001) A dynamical investigation of the heat and helium imbalance. Earth Planet Sci Lett 188:421-434 Vance D, Stone JOH, O’Nions RK (1989) He, Sr, and Nd isotopes in xenoliths from Hawaii and other oceanic islands. Earth Planet Sci Lett 96:147-160 Verma SP, Schilling J-G (1982) Galapagos hot spot-spreading center system 2. 87Sr/86Sr and large ion lithophile element variations (85°W-101°W). J Geophys Res 87:10838-10856

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Walker RJ, Morgan JW, Horan MF (1995) Osmium-187 enrichment in some plumes: Evidence for coremantle interaction ? Science 269:819-822 Weaver BL (1991) Trace element evidence for the origin of ocean-island basalts. Geology 19:123-126 Wetherill GW (1954) Variations in the isotopic abundances of neon and argon extracted from radioactive minerals. Phys Rev 96:679-683 Wetherill GW (1975) Radiometric chronology of the early solar system. Ann Rev Nucl Sci 25:283-328 White WM (1985) Sources of oceanic basalts: radiogenic isotopic evidence. Geology 13:115-118 White WM, Schilling J-G (1978) The nature and origin of geochemical variation in Mid-Atlantic Ridge basalts from the central North Atlantic. Geochim Cosmochim Acta 42:1501-1516 White WM, McBirney AR, Duncan AR (1993) Petrology and geochemistry of the Galápagos islands: portrait of a pathological mantle plume. J Geophys Res 98:19533-19563 White WM, Schilling J-G, Hart SR (1976) Strontium isotope geochemistry of the central North Atlantic: evidence for the Azores mantle plume. Nature 263:659Yatsevich I, Honda M (1997) Production of nucleogenic neon in the Earth from natural radioactive decay. J Geophys Res 102:10291-10298 Zindler A, Hart SR (1986a) Helium: problematic primordial signals. Earth Planet Sci Lett 79:1-8 Zindler A, Hart SR (1986b) Chemical geodynamics. Ann Rev Earth Planet Sci 14:493-571 Zindler A, Jagoutz E, Goldstein SL (1982) Nd, Sr and Pb isotope systematics in a three-component mantle: a new perspective. Nature 298:519-523 Zindler A, Hart SR, Frey FA, Jakobsson SP (1979) Nd and Sr isotope ratios and rare earth element abundances in Reykjanes Peninsula basalts: Evidence for mantle heterogeneity beneath Iceland. Earth Planet Sci Lett 45:249-262

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9 Noble Gases and Volatile Recycling at Subduction Zones David R. Hilton Geosciences Research Division Scripps Institution of Oceanography La Jolla, California 92093 [email protected]

Tobias P. Fischer Department of Earth and Planetary Sciences Northrop Hall, University of New Mexico Albuquerque, New Mexico 87131

Bernard Marty Centre de Recherches Petrographiques et Geochimiques (CRPG) 15 Rue Notre-Dame des Pauvres, B.P. 20 54501 Vandoeuvre les Nancy Cedex, France

INTRODUCTION Volatiles are lost from the Earth's mantle to the atmosphere, hydrosphere and crust through a combination of subaerial and submarine volcanic and magmatic activity. These volatiles can be primordial in origin, trapped in the mantle since planetary accretion, produced in situ, or they may be recycled—re-injected into the mantle via material originally at the surface through the subduction process. Quantifying the absolute and relative contributions of these various volatile sources bears fundamental information on a number of issues in the Earth Sciences ranging from the evolution of the atmosphere and hydrosphere to the nature and scale of chemical heterogeneity in the Earth’s mantle. Noble gases have a pivotal role to play in addressing the volatile mass balance between the Earth’s interior and exterior reservoirs. The primordial isotope 3He provides an unambiguous measure of the juvenile volatile flux from the mantle (Craig et al. 1975). As such, it provides a means to calibrate other volatiles of geological and geochemical interest. A prime example is the CO2 flux at mid-ocean ridges (MOR): by combining estimates of the 3He flux at MOR with measurements of the CO2/3He ratio in oceanic basalts, Marty and Jambon (1987) derived an estimate of the CO2 flux from the (upper) mantle. The approach of using ratios (involving noble gas isotopes) has also been extended to island arcs. Marty et al. (1989) found significantly higher CO2/3He ratios in arc-related geothermal fluids than observed at mid-ocean ridges, consistent with addition of slabderived CO2 to the mantle wedge. Sano and Williams (1996) scaled the CO2 flux to 3He, showing that the output of CO2 at subduction zones was comparable in magnitude to that at spreading ridges. Therefore, for CO2 at least, subduction zones also represent a major conduit for the loss of volatiles from the solid Earth. However, the process of subduction is also the principal mechanism by which materials (including volatiles) are returned to the mantle. For this reason, constraining volatile fluxes and inventories at subduction zones is crucial to understanding volatile budgets on the Earth and the nature of recycling between the mantle and the atmosphere, hydrosphere and crust. In this contribution, we focus on the noble gas systematics of subduction zones. First, we review the various methodologies of sampling noble gases (and other volatiles) using fluids and rocks in both the subaerial and submarine environments. The aim is to 1529-6466/02/0047-0009$05.00

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give the reader background details on how the noble gas database has been accumulated and some of the issues of concern regarding data integrity. We continue by cataloguing the noble gas systematics of both arcs and back-arcs (the output regions)—pointing out similarities and differences in the worldwide record. We highlight some of the major controversies in the interpretation of the database. Finally, we concentrate on the issue of volatile mass balance at subduction zones, emphasising the role of the noble gases (especially 3He) in constraining both the output fluxes of various volatile species and their provenance—both at individual arcs as well as on a global scale. We discuss implications of the volatile mass balance (or imbalance in some cases) for the recycling history of terrestrial volatiles between the Earth’s internal and external reservoirs. SAMPLING FOR NOBLE GASES Noble gases are found in all types of fluids and rock samples. For example, fumarolic gas discharges, bubbling hot springs, groundwaters, and natural gases are prime sampling media for noble gases, as are submarine glasses and various minerals which can be crushed or melted in vacuo to release their trapped volatiles. In this section, we review the principal means of sampling noble gases at subduction zones. Volcanic and geothermal fluids Volcanic and geothermal gas discharges are widely exploited for sampling magmatic volatiles. Various sampling techniques and strategies have been developed depending upon the circumstances of the gas discharge—particularly the temperature. In all cases, precautions are undertaken to avoid, or at least minimize, atmospheric contamination. High temperature fumaroles on passively degassing subaerial volcanoes provide the opportunity to sample volatiles released directly from a magma body. Lower temperature gases associated with bubbling hot springs, usually located on the flanks of volcanoes, allow for the sampling of noble gases released via hydrothermal systems. To facilitate the transfer of high temperature (>400°C) gases into sampling containers, a silica glass tube (~1-inch diameter) is normally inserted into a gas vent. In the case of vents with discharge temperatures ranging from the boiling point of water to ~400°C, a titanium tube (~1-inch diameter) is used. Gases bubbling into hot springs are sampled by placing a plastic funnel under water—if possible, at the bottom of the spring. The silica/titanium tube or the funnel is connected via a glass connector and silicone or Tygon® tubing to a sampling container, which is used for sample storage and transfer to the analytical facility. The following types of sampling containers are in common usage: • Lead-glass flask (~50 cm3 volume) with either one or two vacuum stopcocks. Leadborate glass is used because of its low diffusivity for helium compared to Pyrex (helium diffuses 5 orders of magnitude more slowly through lead-borate glass than through Pyrex at 25°C; Norton 1957). Bottles with two stopcocks (one on each end) allow for the flushing of the bottle with sample prior to collection. The inlet section of a single stopcock flask (previously evacuated to UHV) is usually configured in a “Y” shape to allow flushing of air from the connecting tubes prior to opening the valve and admitting the gas to the flask. • The second type is the “Giggenbach” bottle used by most gas geochemists to obtain the total chemistry of gas discharges. The sampling bottle is an approximately 200 cm3 evacuated glass flask (Pyrex or lead glass) equipped with a Teflon stopcock. The flask contains 50-80 cm3 of 4-6 N NaOH solution (Giggenbach and Goguel 1989). The solution absorbs the reactive gases (CO2, SO2, H2S, HCl, HF) allowing build up of a large partial pressure of the remaining (non-reactive) gases (N2, H2, O2, CO, hydrocarbons and noble gases) in the headspace volume. An upside-down position of

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the bottle during storage, with the solution covering the Teflon stop-cock, helps prevent possible air contamination of the sample during periods of storage. Gas splits for noble gas analyses are taken from the headspace of the sample bottle. The third sampling container consists of a copper tube (~10 cm3) that is crimped using a cold-welder sealing device or pinched off using refrigeration clamps after collection of the sample (see Weiss 1968; Kennedy et al. 1985). The clamps are designed to seal the tube by cold welding and are left in place until sample extraction. Copper is virtually impervious to helium, facilitating secure sample storage; however, interaction with high temperature, acid gases results in the formation of a mixture of copper sulphide, copper sulphate and copper chloride, which can impede the formation of a leak-tight cold seal (F. Goff written comm., 2001).

During sample collection, it is critical to avoid air contamination. Usually this is not a problem at high flow-rate fumaroles as the sampling tube, silicone rubber tubing and sampling bottles are rapidly flushed with the discharging gas. At lower flow-rate fumaroles and bubbling hot springs, more care must be taken to effectively flush the sampling system. This is best achieved by leading a tube from the out-flowing side of the sample bottle into water, allowing the gas to flow through the sample system and to bubble into water. Alternatively, a vacuum hand-pump may be used to facilitate flushing. Depending on the vigor of the gas discharges, flushing of the sampling system can take between ~5 and 40 minutes. Silicone tubes should be kept as short as practical in order to minimize the volume that needs to be flushed with gas prior to sample collection. Gases from bubbling hot springs can also be sampled using the water-displacement method as described by Craig (1953), Mazor and Wasserburg (1965) and Kennedy et al. (1988). A funnel is submerged in the spring and the entire sampling apparatus is purged with spring fluids using a hand-pump. The funnel is then inverted over the upwelling gas bubbles and the gas is allowed to displace all the water in the apparatus—with the exception of the copper-tube sampler, which is connected to the flow-through sampling line via a “Y” connection. After a steady flow of gas is attained through the sampling line, the liquid from the copper-tube sampler is displaced. This technique maximizes the volume ratio of the displacing gas to that of the displaced liquid and minimizes the effect of back solution of gas into cooler fluid as it is displaced. While filling the copper-tube with gas, the outside of the tube may be cooled with water (or snow/ice) in order to lower the vapor pressure of the water. Holding the tube vertically during sampling allows for the condensed phase to flow out and a steady flow of condensed liquid and non-condensable gas is maintained. This technique, therefore, facilitates collection of non-condensable gas at ambient pressure and cooler than ambient temperatures (see Kennedy et al. 1988). Hot spring waters and groundwaters Noble gases dissolved in groundwaters and hot spring waters can be sampled in a fashion similar to that of gas discharges from fumaroles. The most common sampling device is a copper tube, either cold welded or crimped shut using refrigeration clamps. A Tygon® tube is connected to one end of the copper tube and inserted as deeply as possible into the mouth of the spring. A second Tygon® tube is attached to the other end of the copper tube with the discharging water flowing through the tube assembly. After flushing the system for a few minutes and tapping the copper tube lightly, e.g., with a wrench or screwdriver to release air bubbles from the walls of the copper tube, the tube is crimped shut. To avoid air contamination, the out-flowing end of the tube should be crimped first followed by the up-stream end. Water samples can also be collected using evacuated flasks or “Giggenbach” bottles (without the caustic solution). As the spring water fills the evacuated bottles, dissolved

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gases will ex-solve and create a head-space in the bottle. Volatiles of interest, including the noble gases, can be withdrawn from the head-space volume for analysis. Water samples from deep hot spring pools can be sampled at depth by use of a sampling device attached to an extension (aluminum) pole (Kennedy et al. 1988). This technique minimizes air contamination due to entrainment of circulating air-saturated waters from the surface of the pool. A copper tube is fastened to the aluminum pole with a Tygon® tube attached to the bottom end of the copper tube. A second Tygon® tube is attached to the top end of the copper tube and to the pole, and left open to the atmosphere. A spring-loaded Nylon clamp is attached to the pole below the copper tube and triggered by a stainless steel cable that allows for the release of the clamp and the pinching of the Tygon® tube. A second Nylon clamp is attached to the pole above the copper tube. For sample collection, the sampling device is kept vertical and lowered into the pool. At depth, the spring-loaded clamps are released, pinching the Tygon® tubes. The device is brought back to the surface, maintaining the hydrostatic head of the sampling depth. At the surface, the copper tube is pinched shut using a cold welding device or refrigeration clamps. Groundwater pumped from wells is normally sampled using copper tubes that are connected to the well by Tygon® tubing. Caution must be exercised to sufficiently flush the sample device prior to crimping the copper tube. A “Y” connection may be used to release excess water pressure during and after pinching of the copper tube. Geothermal wells Gases from geothermal wells are collected in the same fashion as fumarole discharges. At geothermal wells, however, it is necessary to use a steam separator in order to efficiently sample the non-condensable gases, which partition into the steam phase. The collected sample then consists of steam plus gas. Natural gases Natural gases are generally CH4-rich and at high pressure, making the sampling devices described above impractical. Natural gases are collected into 500 cm3 or 1000 cm3 “Whitey” stainless steel gas cylinders sealed at both ends with high-pressure valves. The pre-evacuated cylinders are either attached directly to the well-head or to the vaporside of a gas-liquid separator. Once attached, the cylinders are opened, flushed with sample gas and sealed. In the laboratory, the cylinders are attached to a vacuum line and aliquots of the sample are transferred into copper tubes (Hiyagon and Kennedy 1992). The partitioning of noble gases from oil or water into the gas phase is a function of the solubility of the noble gases in the liquid phase, which, among other factors, depends on the atomic mass of the gas. Therefore, fractionation may occur if samples are collected from a separator. According to Ballentine et al. (1996), between 82% and 96% of He, Ne, and Ar, is transferred into the gas phase at the separator. This confirms that noble gases strongly partition into the gas phase, resulting in minimal isotopic fractionation during the steam separation process. Gas fields with high gas/(water+oil) ratios further reduce noble gas (elemental) fractionation during collection of the samples at the separator (Torgersen and Kennedy 1999). Mafic phenocrysts and xenoliths Mafic minerals contained in volcanic rocks (phenocrysts) or in xenoliths are widely exploited in noble gas studies as they frequently contain fluid and/or melt inclusions that trap noble gases. Olivines and pyroxenes are the most commonly utilized minerals. Whole-rock samples are first crushed to 0.5 to 2 mm or larger, depending on the size of the crystals. The olivines and pyroxenes are then separated using a Frantz-Isodynamic magnet separator, followed by hand-picking under a binocular microscope to remove any

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adhering matrix. Mineral separates are then cleaned in distilled water, methanol or acetone. After cleaning and drying, the separates can be processed for noble gases. The two extraction techniques generally used to decrepitate inclusions in the samples are vacuum crushing and vacuum melting. Both methods, however, integrate volatiles from all types and sizes of inclusions in the sample. The relatively recent application of lasers to noble gas studies of mafic crystals (e.g., Burnard et al. 1994) offers the exciting possibility that individual inclusions, or trails of inclusions, can be targeted to reveal spatial and/or temporal variations preserved within single crystals. There are two general categories of crushing devices whereby crushing takes place either on-line or off-line (see Hilton et al. 1999 for a description of the two types of crusher and a comparison of He isotope results produced by both). In the case of on-line devices, vacuum is maintained between the crusher and the mass spectrometer inlet line during the crushing or pulverizing of the sample. This is not the case for off-line crushers. Indeed, many samples processed for He-isotopes prior to the early-mid 1990s were processed by off-line “ball mills”: in these cases, careful monitoring of possible air leakage (e.g., by neon analysis) was deemed essential to ensure the integrity of the helium isotope results (e.g., Hilton et al. 1992). A further potential problem with off-line ball-mills relates to their exceptional efficiency in crushing samples. Prolonged crushing times of mafic crystals can lead to the release of He components extraneous to the magmatic system under scrutiny, e.g., cosmogenic and/or radiogenic He sited in the mineral matrix (Hilton et al. 1993a; Scarsi 2000). Nowadays, most investigators use online crushing devices,with many adopting the added precaution of minimizing crushing times (Hilton et al. 1993a, 1999). Submarine glasses Submarine glasses from a variety of tectonic environments (including arcs and backarcs) are a widely exploited medium for obtaining noble gas isotopic compositions and abundances. Rapid quenching of lavas as they are extruded onto the seafloor traps volatile phases in the glassy rinds. The normal procedure (e.g., Hilton et al. 1993b) involves selecting pristine, glass chips and shards by hand-picking under a binocular microscope. Only glass pieces with no visible signs of alteration are selected for analysis. The glasses are then cleaned ultrasonically with acetone and/or distilled water to remove any adhering matrix and/or alteration phases. Sizes of glass chips generally range from 1to 5-mm diameter. As in the case of mafic minerals, samples can be processed by vacuum melting and/or crushing. Problematic issues of noble gas analysis Given the wide range of sampling media available for noble gas studies and the various means of extracting trapped gases for analysis, it is instructive to scrutinize results to test the integrity of both a particular sampling medium as well as the methodology or experimental approach. Mafic minerals (olivine and clinopyroxene) are widely used to gain insight into the helium isotope systematics of the magmatic source region. However, there are processes that may compromise the integrity of 3He/4He results obtained using mafic crystals. These principally relate to addition of extraneous (radiogenic) helium through interaction with pre-existing crust or wallrock. The following observations are pertinent in helping to recognize these effects: 1. A correlation between measured 3He/4He ratio and helium content for co-genetic suites of samples may reveal late-stage addition of crustal (radiogenic) helium to helium-poor samples. Hilton et al. (1995) noted that He-poor crystals in some ocean-island phenocrysts had low 3He/4He ratios consistent with addition of

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radiogenic He. Due to the higher diffusivity of helium in pyroxenes (Trull and Kurz 1993), this observation is more likely observed in pyroxenes than in olivines. 2. A correlation between 3He/4He ratios and an index of magmatic differentiation. For example, Gasparon et al. (1994) found lower 3He/4He ratios in more highly evolved phenocrysts (phenocrysts with a lower Mg #) from the Sunda arc, Indonesia. 3. The occurrence of isotopic disequilibrium between coexisting phenocryst pairs. Marty et al. (1994) reported higher 3He/4He ratios in olivine phenocrysts compared to pyroxene or hornblende for Mt. Etna volcano, an observation that is consistent with addition of late-stage, radiogenic or atmospheric helium. The same effect is seen in arc-related phenocrysts from the Lesser Antilles (Van Soest et al. 2002). Submarine glasses can also exhibit noble gas variations depending on the method chosen for gas extraction. Crushing releases volatiles sited in vesicles whereas melting integrates vesicle gas and gas dissolved in the glass phase. Depending upon the degree of vesicularity, noble gases can fractionate from each other as a function of their relative solubilities in silicate melt (Jambon et al. 1986; Carroll and Webster 1994). A direct corollary of partitioning of noble gases between vesicle and melt is that the melt phase can become extremely depleted in noble gases. As a consequence, the isotopic composition of this residual gas (in the magma/glass) may become susceptible to modification by addition of volatiles resulting from radiogenic and nucleogenic reactions (e.g., magma aging; Zindler and Hart 1986). Alternatively, extraneous volatiles could be addedtovolatile-poor magma immediately prior to eruption thereby modifying a magmatic noble gas signature by one characteristic of crustal and/or seawater/atmospheric contamination (see Hilton et al. 1993a). When comparing different sampling media directly, large differences in 3He/4He ratios have been observed. For example, at Cerro Negro volcano, Nicaragua, olivine phenocrysts erupted in 1992 have 3He/4He ratios of 3.5±0.5 RA (Fischer et al. 1999a) whereas fumarole gases collected days after the eruption have ratios of 6.8 RA (Sano and Marty 1995). An unreasonably high residence time of approximately 1 Ma is needed to lower 3He/4He ratios in the phenocrysts from 6.8 to 3.5 RA assuming a magma U-content of 0.3 ppm and a He-content of 1.5×10-9 cm3 STP/g olivine (and a He partition coefficient of 0.008; Marty and Lussiez 1993). This time period is much longer than the age of the Cerro Negro volcanic system (Fischer et al. 1999a). For this active volcanic system, therefore, the noble gas ratios of fumarolic discharges may be more representative of the present-day mantle source, with the 1992 activity erupting olivine phenocrysts having trapped helium from an earlier magmatic episode. NOBLE GAS SYSTEMATICS OF ARC-RELATED VOLCANISM The study of noble gases in arcs began with the seminal work of Mamyrin et al. (1969) who reported 3He/4He ratios greater than air for geothermal fluids from the Kurile Islands. This early observation of primordial (mantle-derived) 3He in arc volcanoes has been confirmed by numerous subsequent studies (e.g., Craig et al. 1978; Sano and Wakita 1985; Poreda and Craig 1989) which have repeatedly emphasized the role of the mantle wedge in dominating the helium budget in the great majority of cases (the Banda arc of Indonesia being a notable exception; see below). It was also realized through noble gas studies that other contributors to the arc volatile inventory (the subducting slab and/or arc crust) could be traced through either the relative abundance of a major volatile phase to a noble gas or by the isotopic composition of a particular noble gas. An example of the former would be the high N2/Ar ratios of island arcs tracing addition of slab-derived sedimentary N2 to the source region (Matsuo et al. 1978; Kita et al. 1993): in the latter

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case, it has been argued that low 3He/4He ratios in the Andes indicate (upper) crustal additions to the volatile budget (Hilton et al. 1993a). In this section, we review noble gas systematics of arc-related volcanism worldwide. Helium isotope studies dominate because most arc products are erupted subaerially, and air contamination is a relatively minor (correctable) problem for helium: this is not the case for Ne-Ar-Kr-Xe isotope systematics. Consequently, this section is weighted towards reporting observations of helium isotope variations in arc-related minerals and fluids. However, we summarize also the available database for neon, argon and xenon isotopes (todateKr shows only air-like isotopic compositions). Finally, we consider the limited database of the relative abundances of the noble gases in arc-related products. Helium isotope systematics of arc-related volcanism At present, there are close to 1000 reported 3He/4He ratios from a variety of sampling media associated with arc-related volcanism. Table 1 summarizes the helium isotope database representing 26 individual arc segments worldwide (see Fig. 1 for locations). In this compilation, we focus our attention on the active arc only as detailed across-arc studies (e.g., Sano and Wakita 1985) have long established the link between the presence of mantle-derived helium and the locus of magmatic activity. Furthermore, we include in Table 1 all analyses irrespective of sampling medium—the only exception is duplicate analyses of the same sample (in which case the highest value has been selected). We have made no attempt to choose ‘representative’ values for a particular volcanic system: all analyses irrespective of distance from volcano summit (geothermal fluids) or age of lava flow (phenocrysts) are included. However, it should be noted that there is evidence of a geographic control on geothermal fluid 3He/4He ratios located around some individual volcanic centers: the highest values are found close to the eruptive vents and lower, more radiogenic values occur away from the volcanic centers (Sano et al. 1984; Williams et al. 1987; Marty et al. 1989; Hilton et al. 1993a; Van Soest et al. 1998). To facilitate discussion, we tabulate the range of 3He/4He ratios (maximum and minimum values) together with the average value (with 1 standard deviation from the mean). With reference to Table 1, the following points are emphasized: 1. The highest arc-related 3He/4He ratios fall within the range normally associated with depleted (N-type) MORB mantle (i.e., 8±1 RA; see Farley and Neroda 1998 and Graham 2002, this volume). The highest values (8.8 to 8.9 RA) are reported for 4 arc segments: the southern Lesser Antilles (Pedroni et al. 1999), the Colombian Andes (Sano et al. 1997) and the Sunda arc system in both Sumatra (Gasparon et al. 1994) and Bali (Hilton and Craig 1989). 2. In nine of the arc segments, the highest 3He/4He ratio fails to reach the lower limit of the MORB range, i.e., the highest value is MORB (i.e., higher than 8±1 RA) are observed at 3 of the back-arc basins: the Manus Basin (15.2 RA), the Lau Basin (22.1 RA) and the North Fiji Basin (10.3 RA). Such high values are normally considered diagnostic of a deep-seated mantle plume component (Craig and Lupton 1976), and the respective authors in the three cases above have adhered to this interpretation. For example, Poreda (1985) and Poreda and Craig (1992) suggested that channeling of plume material from the nearby Samoa hotspot could explain the 3He/4He ratios at Rochambeau Bank towards the northern end of the Lau Basin (see Turner and Hawkesworth (1998) for further discussion on mantle flow between Samoa and the Lau Basin). Similarly, Macpherson et al. (1998) explained high 3He/4He ratios in the Manus Basin as reflecting a mantle plume originating at the underlying core-mantle boundary. While noting the presence of a “3He-rich, hotspot-like component” in the North Fiji Basin, Ishibashi et al. (1994) did not comment on its origin. In all three cases, however, there was no attempt to relate the high 3He/4He ratios to shallow processes associated with the back-arc, so it appears safe to assume that the high values observed in these regions bear little or no information on the subduction process. Whereas the majority of the Manus Basin samples (particularly from the Manus Spreading Centre and the Extensional Transform Zone) are characterized by 3He/4He ratios greater than those of MORB, this is not the case for lavas from the Lau Basin or the North Fiji Basin. Indeed, a significant number of samples from both basins have

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Table 4. Summary of helium isotope variations in back arc-related volcanics and hydrothermal fluids. Location (segment) Lau Basin Rochambeau Bank Peggy Ridge Northern Basin Central Spreading Center East Spreading Center Valu Fa Ridge Total (Lau Basin)

R/RA (max)

R/RA (min)

R/RA*(mean) (± 1 S.D.)

N$

References1

22.1 10.9 8.1 11.2 8.9 7.8 22.1

11.0 1.2 5.9 7.5

16.3 ± 5.3 8.3 ± 3.1 7.2 ± 0.7 8.6 ± 0.8

1.2 1.2

5.7 ± 2.6 8.5 ± 3.8

5 8 9 14 1 10 47

[1-2] [1-2] [2-4] [2-4] [4] {3, 5] [1-5]

Mariana Trough North Central South Total (Mariana Trough)

8.6 8.5 8.8 8.8

5.7 0.85 7.9 0.85

7.8 ± 6.8 ± 8.3 ± 7.8 ±

0.8 2.8 0.3 1.4

15 8 15 38

[6-7] [1, 7]] [8] [1, 6-8]

North Fiji Basin 160oN segment Triple Junction Area 15 oN segment N-S segment 174 o E segment Total (North Fiji Basin)

8.8 10.0 9.0 10.3 8.8 10.3

8.6 7.9

8.7 ± 0.1 9.1 ± 0.7

3.8 8.1 3.8

8.4 ± 1.5 8.4 ± 0.9 8.7 ± 1.1

3 16 1 13 6 36

[9] [9-10] [9] [9-11] [9] [9-11]

Manus Basin Seamount Extensional Transform Zone Manus Spreading Center Southern Rifts East Manus Rift Total (Manus Basin)

15.2 13.5 12.7 8.8 6.4 15.1

0.92 0.67 3.4 0.61 0.61

10.1 ± 4.3 10.1 ± 4.2 3.7 ± 2.4 8.5 ± 4.7

1 7 17 2 8 35

[12] [12] [8, 12] [12] [8, 12-13] [8, 12-13]

22.1

0.61

8.4 ± 3.2

156

Summary Footnotes to Table 4 * $ 1.

[1-13]

Includes only air- or ASW-corrected 3He/4He data except where no He/Ne ratios are given. N = number of individual results incorporated into mean. In the case of duplicate analysis, the higher number is selected. Helium isotope data sources: 1= Poreda (1985), 2 = Poreda and Craig (1992), 3 = Hilton et al. (1993b), 4 = Honda et al. (1993b), 5 = Bach and Niedermann (1998), 6 = Ikeda et al. (1998), 7 = Sano et al. (1998), 8 = Macpherson et al. (2000), 9 = Nishio et al. (1998), 10 = Ishibashi et al. (1994), 11 = Marty and Zimmermann (1999), 12 = Macpherson et al. (1998), 13 = Marty et al. (2001).

336

Hilton, Fischer & Marty

MORB-like 3He/4He values (see references in Table 4). Specifically, lavas with 3He/4He values of8±1 RA are erupted in the central Lau Basin as well as along all spreading centers of the North Fiji Basin. Therefore, in addition to a contribution from a deep mantle plume, the other principal source of helium in active back-arc basins is the same as that supplying mid-ocean ridges. This conclusion reinforces that derived from the majority of island arc studies—namely, that the mantle is the predominant source of helium in subduction zone environments with little or no contribution from the subducted slab. There are, however, a number of samples in each of the four back-arc basins that have 3He/4He values significantly lower than the MORB range. Oceanic glasses with 3 He/4He ratios ≤1 RA (implying that ≥80% of the helium is radiogenic in origin; Andrews 1985) are observed along the Peggy Ridge and Valu Fa Ridge (Lau Basin), the central Mariana Trough and throughout the Manus Basin (Table 4). Assuming that, in all cases, the glasses used in these studies were young enough to rule out post-eruptive in-growth of 4He, then any radiogenic helium must have been inherited prior to eruption of the lavas onto the sea floor. Three explanations have been advanced to account for the occurrence of these low 3He/4He values: 1. The mantle wedge is depleted in helium so that it cannot mask/dilute radiogenic helium derived from the subducted slab (subducted sediments or underlying oceanic basement). 2. Magmas are contaminated by radiogenic helium trapped in crust through which the magmas are erupted. Contamination is accentuated by pre-eruptive degassing of magmas which lowers absolute concentrations thereby making magmas more susceptible to contamination. 3. The mantle wedge is enriched in U- and Th-series isotopes derived from the subducted slab so that the wedge itself is characterized by 3He/4He ratios < MORB. The first explanation was suggested by Poreda (1985) to explain a small number of low 3He/4He ratios (1000. There appears consensus on the origin of the high ratios—namely, the mantle wedge, even though the absolute ratios fall far short of MORB-type values (up to 40,000; Burnard et al. 1997). This relatively narrow range in 40Ar/36Ar ratios for BABB (296 to 5000) reflects the ubiquitous presence of air-derived argon in all samples, which acts to lower mantle 40Ar/36Ar values. Debate on the origin of the low 40Ar/36Ar ratios in back-arc basins (subduction-related or upper crustal contamination) mirrors that for arc localities. As all samples in Table 5 are oceanic glasses, and therefore erupted in a submarine setting, the most pressing concern is seawater contamination. Honda et al. (1993b) noted a correlation between absolute abundances of krypton and xenon (which are relatively enriched in seawater) with total water content, and suggested interaction with seawater as one possible explanation for the atmospheric-like heavy noble gas systematics. They also pointed to a correlation between Kr and Xe and Ba/Nb ratio (a key slab fluid tracer), opening up the possibility that seawater interaction may have occurred in the down-going slab and not necessarily close to the site of eruption. In this way, the heavy noble gases (and by implication the low 40Ar/36Ar ratios) would be indicative of volatiles recycled through the subduction zone. Bach and Niedermann (1998) have also taken up this point as they observed the same type of correlations between absolute abundances of noble gases and both water contents and Ba/Nb ratios for samples from the Valu Fa Ridge (Lau Basin). The alternative explanation, that low 40Ar/36Ar ratios in back-arc basins trace

Noble Gases and Volatile Recycling at Subduction Zones

339

shallow-level contamination of arc-rifted crust, was advanced by Hilton et al. (1993b) who argued that such crust has the necessary attributes of radiogenic He, low 4He/40Ar* values (where 40Ar* is radiogenic argon) and low 40Ar/36Ar ratios (~ 300) resulting from ageing and modification by hydrothermal alteration. Interaction of previously-degassed primary magmas with such crust would result in the superimposition of a slab-like signature (high water, high Ba/Nb) onto crustal noble gases characteristics, and result in geochemical features such as seen along the Valu Fa Ridge. As in the case of argon isotope variations in arc-related volcanism, considerable controversy still surrounds assigning provenance to potential end-member contributors. Table 5. Summary of 40Ar/36Ar isotope variations in back arc-related volcanics. Location (segment)

Lau Basin Northern Basin Central Spreading Center East Spreading Center Valu Fa Ridge Total (Lau Basin) Mariana Trough North Central South Total (Mariana Trough) North Fiji Basin N-S segment Total (North Fiji Basin) Manus Basin Seamount Extensional Transform Zone Manus Spreading Center Southern Rifts East Manus Rift Total (Manus Basin) Summary

40

Ar Ar (max) 36

40

Ar Ar (min) 36

40

Ar Ar (mean) (± 1 S.D.)

36

N$

Refs1

5300

444

2758 ± 2337

7

[1-2]

4900

1206

2429 ± 1140

10

[1-2]

488

293

337 ± 66

11

[1, 3]

5300

293

1647 ± 1686

29

[1-3]

4480

295

1647 ± 1686

12

[4-5]

4400

338

2

[5]]

2734

296

613 ± 636

15

[6]

4480

295

927 ± 1139

29

[4-6]

3339

532

1512 ± 1583

3

[7]

3339

532

1512 ± 1583

3

[7]

1

[8]

2680

296

1203 ± 1046

6

[8]

1764

295

846 ± 563

9

[8]

1

[8]

7

[8-9]

461

1

1310

274

[2]

301

278

293 ± 8

2680

274

770 ± 701

24

[8-9]

5300

274

1148 ± 1312

85

[1-9]

Includes only air- or ASW-corrected 3He/4He data except where no He/Ne ratios are given. N = number of individual results incorporated into mean. In the case of duplicate analysis, the higher number is selected. Helium isotope data sources: 1 = Hilton et al. (1993b), 2 = Honda et al. (1993b), 3 = Bach and Niedermann (1998), 4 = Ikeda et al. (1998), 5 = Sano et al. (1998), 6 = Hilton et al. (unpublished), 7 = Marty and Zimmermann (1999), 8 = Shaw et al. (unpubl), 9 = Marty et al. (2001).

* $

340

Hilton, Fischer & Marty

Finally, we note that deviations from atmospheric-like krypton isotope ratios have not been found for any back-arc basin: however, there are two reports of anomalous xenon isotope variations for the Mariana Trough. Ikeda et al. (1998) found coupled 134 Xe/130Xe and 129Xe/130Xe deviations from air—similar to those found in MORB, whereas Sano et al. (1998a) reported only 129Xe excesses (relative to 132Xe). These signatures reinforce the idea that the mantle wedge is the principal source of volatiles in the Mariana Trough particularly where circumstances limit the amount of atmospherederived contributions. At all other localities, only atmospheric-like xenon isotope ratios have been found. RECYCLING OF VOLATILES AT SUBDUCTION ZONES: A MASS BALANCE APPROACH In the previous two sections, we considered the noble gas systematics of arc and back-arc regions. We now turn our attention to the exploitation of the noble gases in understanding volatile mass balance at subduction zones. Our focus is on the ability of noble gases to both quantify mantle-degassing rates and identify volatiles of differing provenance. Specifically, we discuss the utility of noble gases in determining the volatile output via subduction zones from various sources—both the intrinsic output of noble gases themselves as well as the other (major) volatile phases that comprise the terrestrial atmosphere. In addition, we adopt a new approach to calculate volatile outputs for individual arc segments worldwide. This allows us to assess the volatile mass balance for a number of convergent margins worldwide. In this way, we can determine more realistically how the range in volatile input parameters affect volatile output. Summing the input versus output parameters globally, we consider the long-term recycling efficiency of both the major volatiles and the noble gases, and discuss implications for Earth evolution. The global volatile output at arc volcanoes: The 3He approach There have been numerous attempts at estimating volatile fluxes associated with arcrelated volcanism. In this respect, the noble gas isotope that has received most attention is 3 He as its primordial origin makes it an unambiguous tracer of mantle-derived volatiles. Therefore, if the arc flux of 3He can be established, it would lead to the derivation of other volatile fluxes by simple measurement of the ratio xi/3He, where xi is any volatile species discharging from volcanoes (CO2, SO2, H2S, HCl, N2, etc.). Two distinct approaches have been taken to estimate the 3He flux from arc volcanoes: 1. Using the relatively well constrained figure of mid-ocean ridge degassing flux (~1000 mol3He/yr; Craig et al. 1975), together with the assumption that the magma production rate of arcs is 20% that of MOR (Crisp 1984), an arc 3He flux of ~200±40 mol/yr is derived (Torgersen 1989). This method further assumes that the 3He content of magma in the mantle wedge is the same as that beneath spreading ridges. 2. Using total 3He fluxes calculated for subaerial volcanism and estimating the fraction contributed by arc volcanism. For example, Allard (1992) derived an estimate for the total flux of 3He into the atmosphere by subaerial volcanism to be 240-310 moles/yr (based upon integrating the CO2 flux from 23 individual volcanoes worldwide and coupling this flux with measurements of the CO2/3He ratios). Of the total subaerial 3He flux, he suggested that approximately 70 mol/yr was arc-related. Adopting a similar approach, Marty and LeCloarec (1992) used polonium-210 (210Po) s the flux indicator (along with the 210Po/3He ratio) to estimate a total subaerial volcanic 3He flux of 150 mol/yr—of which over half (>75 mol/yr) was due to arc volcanism.

Noble Gases and Volatile Recycling at Subduction Zones

341

It should be noted that neither approach attempts a direct measurement of the arc 3He flux. Whereas scaling to estimated magma production rates is used in the first instance, the second methodology relies on knowledge of an absolute flux of some chemical species from volcanoes together with a measurement of the ratio of that species to 3He. The most widely used species to derive absolute chemical fluxes from subaerial volcanoes is SO2 using the correlation spectrometer technique (COSPEC) (Stoiber et al. 1983). Carbon dioxide (see Brantley and Koepenick 1995) as well as 210Po (Marty and LeCloarec 1992) have also been used to calibrate absolute fluxes of other species albeit to a much lesser extent. In Table 6, we compile various estimates of the volatile flux from arcs for a number of noble gas isotopes (3He, 4He and 36Ar) and major volatile phases (N2, CO2 and H2O). The reader is referred to the literature given in the footnote for details of the methodology adopted to derive each estimate and the inherent assumptions involved. For comparison, we also compile estimates of the total subaerial volcanic fluxes. Note that all flux estimates in Table 6 are made on a global basis, i.e., the fluxes are assumed to represent the integrated output from all arcs worldwide. Although such fluxes are essential in addressing large-scale geochemical questions (e.g., mass transfer of volatiles through the upper mantle; Porcelli and Wasserburg 1995), they are of limited use in assessing the state of mass balance (input via the subducting slab versus output via the arc and backarc) at individual arcs worldwide, and the (localized) effect of variations in the type and amount of subducted sediment on volatile output. For these questions, an estimate of volatile fluxes at individual arc segments is required. In the next section, we adopt an integrated flux approach (using SO2) to produce volatile flux estimates but limit ourselves to distinct arc segments. A test of the success of this approach will be whether the sum of the individual arc fluxes equals that derived on a global basis (Table 6). Volatile output at individual arcs—SO2 and the power law distribution In this section, we provide estimates of the 3He (and other volatile) fluxes from individual arc segments worldwide. To achieve this objective, we use time-averaged SO2 flux measurements from 43 passively degassing arc volcanoes (Andres and Kasgnoc 1998) coupled with our own compilation of approximately 700 volcanic and hydrothermal gas compositions. It must be borne in mind that the accuracy of individual COSPEC measurements may vary from 10-40% (Stoiber et al. 1983) and this error will propagate through to the final estimates. The time-averaged volcanic SO2 emission rates were compiled from approximately 20,000 individual measurements dating from the early 1970s to 1997 (Andres and Kasgnoc 1998). Based on this extensive data set, Andres and Kasgnoc (1998) found that SO2 emissions during eruptions contributed only 1% to the total volcanic SO2 flux. This conclusion supports earlier work of Berresheim and Jaeschke (1983) who showed that 90% of the SO2 is emitted during non-eruptive events. Although high SO2 flux rates undoubtedly occur during eruptive periods, most eruptions are short-lived and extrapolation of the flux rates to longer time intervals are probably questionable. Therefore, in our treatment of SO2 fluxes in this section, we feel justified in using fluxes which were derived for passively degassing volcanoes only. Furthermore, we recognize that it is impossible to measure the emissions from every single volcano; therefore, this necessitates adoption of an extrapolation procedure to account for the (small) SO2 flux that is not measured directly. We have followed the methodology of Brantley and Koepenick (1995) who found that if the distribution of volcanic emissions follows a power law, thenthere is an empirical relationship between the cumulative number of volcanoes (N) having or exceeding a given emission rate (f) and the emission rate itself. The relationship takes the form:

342

Hilton, Fischer & Marty N = a * (f )-c

(1)

where N is the number of volcanoes having an emission rate ≥ f and a and c are constants (the value of a is determined by the choice of units of f). If the constant c < 1, the total volcanic flux (ftot) can be calculated using the following approximation: ftot = f1 + f2 +….… fN [ (c/{1-c})*(N+1)*(N/{N+1})1/c]

(2)

Table 6. Compilation of volatile fluxes (mol/yr) of arc-related and global subaerial volcanism. Species

Arc Flux

Reference

3

200±40 70±25 >75

1 2 3

92 2.0±0.13 (x 107) 1.23 x 107 1.9 x 108 4.7 x 105 3.1 x 1012 0.3±0.2 (x1012) 0.5±0.4 (x1012) 0.7 x 1012 1.5 x 1012 2.5 x 1012

5 1 5 6 6 7 4 10 2 8 9

1.6 x 1012 6.4 x 108 2.0 x 1010 8.0 x 1012

14 15 16 17

4

He

He

40

Ar Ar CO2 36

N2 H2O # 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17

Reference Torgersen (1989) Allard (1992) Marty and Le Cloarec (1992) Marty et al.(1989) This work (Table 9) This work (Table 9) Sano and Williams (1996) Varekamp et al. (1992) Marty B. and Tolstikhin I. N. (1998) Williams et al. (1992) Gerlach (1991) Brantley S. L. and Koepenick K. W. (1995) LeGuern F. (1982) This work (Table 9) Sano et al. (2001) This work (Table 9) This work

Global Subaerial Flux

Reference

275±35 240-310 150 3-150

1 2 3 4

3.75±1.35 (x 108)

1

0.77±0.58 (x 1012) 1.8 x 1012 1.5 x 1012 3.3 x 1012 2.5±0.5 (x1012) 1.1 x 1012 5.5 x 1012

10 11 2 3 12 13 9

2.8 x 109

15

55-550 x 1012

13

Method Scaling to MOR 3He flux 3 He/CO2 ratios plus CO2 flux 210 Po/SO2 – CO2/SO2 and CO2/3He ratios CO2/3He and CO2/SO2 ratios plus SO2 flux Summation of individual arcs (arc = 5.4RA) Summation of individual arcs (arc 40Ar/36Ar = 401) C/3He plus 3He flux C/3He plus mass flux CO2/3He ratios plus magma emplacement C/S plus S flux Flux measurement Flux measurement Flux measurement Summation of individual arcs N/3He plus 3He flux Summation of individual arcs H2O/CO2 = 50 and CO2 flux from ref. 14

Noble Gases and Volatile Recycling at Subduction Zones

343

where f1 is the largest flux, f2 is the second largest flux, and fN refers to the Nth largest flux. Brantley and Koepenick (1995) found that the global emission of SO2 is governed by a power law distribution (c = 0.8) so that an estimate of the total SO2 flux can be made using a relatively small number of volcanoes. They cautioned, however, that only larger volcanoes (f ≥ 320 Mmol/yr) followed a power law distribution. In spite of this caveat, we have attempted to use the power law distribution to estimate the SO2 fluxes for individual arcs worldwide. Andres and Kasgnoc (1998) have tabulated the SO2 flux for 49 volcanoes representing 11 distinct arc systems, and we have used these estimates to derive the total fluxes for each of these arcs. We adopt two different approaches: 1. We have assumed that the global value of c (= 0.8; Brantley and Koepenick 1995) is applicable to each individual arc segment. Therefore, we can calculate ftot from Equation (2) above and the tabulated fluxes for volcanoes from each arc. 2. We calculate a different c-value for use with each individual arc segment. In the case of Central America, for example, we plot log (N) versus log (fSO2) for 11 volcanoes, and assuming a linear correlation (power law distribution), derive c from the gradient (Fig. 4). This value of c is used in Equation (2) to derive the total flux for that arc.

Figure 4. Log-log plot of cumulative frequency of number of Central American volcanoes (N) having SO2 flux equal or greater than f (following methodology of Brantley and Koepenick 1995). Data from Andres and Kasgnoc (1998).

In Table 7, we show the results for 11 arc segments using these two approaches. Summing the individual arc fluxes, we obtain SO2 flux estimates (in mol/yr) of 2.80×1011 (c = 0.8) and 2.62×1011 (various c-values). Both values compare well with previous estimates using variants of the same (power law) approach—2-3×1011 mol/yr (Brantley and Koepenick 1995); 2.92×1011 mol/yr (Stoiber et al. 1987) and 2.4×1011 mol/yr (Berresheim and Jaeschke 1983). Although the agreement is encouraging, one point of concern is that there appears a large discrepancy between summing the individual fluxes at some arcs and the total (extrapolated) flux. For example, the sum of the measured fluxes for the Andes lies between 43% and 60% of the total extrapolated flux (depending upon the chosen c-value). For comparison, Andes and Koasgnoc (1998) found that—on a global basis—the ‘additional’ extrapolated flux amounted to 19% of the total i.e., the sum of the measured fluxes was 81% of the total flux. It is noteworthy that calculating c-values for individual arcs seems to decrease the discrepancy between summed and extrapolated fluxes. Additionally, it seems that the large extrapolation appears to affect mostly arcs with small numbers of volcanic SO2 fluxes (e.g., Andes, Antilles, New Zealand, Philippines; see Table 7). Therefore, while accepting the flux extrapolations give a useful first-order estimate

344

Hilton, Fischer & Marty Table 7. Summary of SO2 fluxes (×109 mol/yr)— measured (meas) and extrapolated (extrap) using power-law distribution. # volcanoes SO2 flux1

∑ meas

Andes Alaska-Aleutians Antilles Central America Indonesia Italy Japan Kamchatka-Kuriles New Zealand Papua New Guinea

fluxes

Extrap flux2

% meas extrap

C-coefficient3 (r2-value)

Extrap flux4

% meas extrap

3 3 1 11 4 3 8 3 2 5

28.2 0.41 1.31 20.3 1.64 27.2 20.0 1.64 4.67 35.8

65.9 0.56 4.41 25.1 2.76 29.8 29.6 2.51 15.3 85.4

42.8 73.3 29.7 80.9 59.2 91.4 67.6 65.5 30.5 42.0

-0.70 (0.77) -0.34 (0.84) -0.46 (0.72) -0.55 (0.80) -0.23 (0.92) -0.60 (0.98) -0.41 (0.92) >1 -0.84 (0.92)

47.3 (0.34) 21.3 1.88 27.0 23.3 1.67 100.9

59.6 95.5 86.9 85.9 98.2 35.5

Philippines

2

5.12

18.2

28.1

-

-

-

GRAND TOTAL

45

279.5

52.2

261.85

55.8

Arc Segment

1. 2. 3. 4. 5.

146

SO2 flux measurements from Andres and Kasgnoc (1998) Extrapolated flux using methods of Brantley & Koepenick (1995) assuming constant c = global avg value = 0.8. Recalculated c-constant using only data from specific arc (r2 = correlation coefficient on linear regression). Recalculated extrapolated flux using C-coefficient in previous column. Where missing (or in parenthesis), values from column 4 are used in summation.

of the SO2 flux for individual arcs, we caution that not all individual arc flux estimates can be treated with the same degree of confidence. Clearly, the situation will improve as more SO2 flux measurements become available. It is interesting that the highest SO2 emissions are from volcanic arcs in the Papua New Guinea -SW Pacific region—8.5 to 10.1 (×1010) mol/yr. However, we note again that the sum of individual flux is significantly less (35-42%) than the extrapolated values, and that the flux extrapolation is weighted by one large value—from Bagana in the Solomon Islands. At the other extreme, the lowest emissions are from the AlaskaAleutians arc with a SO2 flux of only 5.6×108 mol/yr, or less than 1% of the largest value. In Table 8 we report a compilation of volcanic gas chemistries (as molar ratios) using over 700 individual analyses of arc-related volcanic and hydrothermal fluids. In an attempt to obtain a more representative picture of the chemical variability, we present the database as median ratios which gives less weight to extreme values. In this compilation, St is the total S, i.e., SO2 + H2S, and N2,(ex) is the amount of “excess” N2 after correction for N2 contributed from meteoric sources in the volcanic edifice. The “excess” N2 was calculated using: (3) [N2]ex = [N2] measured – (40 × [Ar] measured) where square brackets designate concentration, and the numeric value 40 implies that the measured argon is derived solely from air-saturated water in equilibrium with the atmosphere, i.e., N2/Ar = 40 (see Fischer et al. 1998). With reference to Table 8, there are two outliers in the CO2/St column—the Philippines arc (104.5) and New Zealand (28.7)—which are both significantly higher than typical ‘arc-like’ values of ~5 (Giggenbach 1996). We suggest that the available database in these two cases may be biased to mainly low-temperature hydrothermal fumaroles and geothermal fluids that

721

GRAND TOTAL

28.7 104.5

5.7 6.1 5.0 2.7 4.3 8.5 6.5 1.7

CO2/St

53.0 115.4

CO2/ N2 61.1 33.8 64.6 34.1 60.9 95.9 18.9 14.1 102.3 140.0

CO2/ N2 (ex) 73.2 72.8 195.4 95.9 62.3 24.1 23.8 1.05 0.90

CO2/He (×105) 2.92 2.00 1.26 0.81 1.47 1.45 0.76 0.95 4.03 13.6

CO2/Ar (×103) 10.7 4.61 n.a. 8.96 7.28 72.0 4.39 2.82 [60-62] [12, 63-65]

[1-4] [5-7] [8-10] [11-18] [19-25] [12, 26-31] [15, 32-46] [47-59]

Refs.

1 = Fischer et al. (1997), 2 = Lewicki et al. (2000), 3 = Sturchio et al. (1993), 4 = Giggenbach et al. (1986), 5= Symonds et al. (1990), 6 = Sheppard et al. (1992), 7 = Motyka et al. (1993), 8 = Chiodini et al. (1996), 9 = Hammouya et al. (1998), 10 = Brombach et al. (2000), 11 = Menyailov et al. (1986a), 12 = Giggenbach (1992), 13 = Taran (1992), 14 = Rowe et al. (1992), 15 = Goff and McMurtry, (2000), 16 = Giggenbach and Corrales (1992), 17 = Janik et al. (1992), 18 = Taran et al. (1998), 19 = Allard (1983), 20 = Giggenbach et al. (2001), 21 = LeGuern (1982), 22 = Poorter et al. (1993), 23 = Fischer unpubl., 24 = Delmelle et al., (2000), 25 = Sriwana et al. (2000), 26 = Chiodini et al. (1995), 27 = Cioni and D'Amore (1984), 28 = Goff unpublished, 29 = Rogie et al. (2000), 30 = Minissale et al. (1997), 31 = Chiodini (1994), 32 = Mizutani and Sugiura (1982), 33 = Symonds et al. (1994), 34 = Giggenbach et al. (1986), 35 = Nemoto (1957), 36 = Mitzutani (1962), 37 = Mitzutani and Sugiura (1982), 38 = Symonds et al. (1996), 39 = Mitzutani (1966), 40 = Giggenbach and Matsuo (1992), 41 = Mizutani et al. (1986), 42 = Matsuo et al. (1974), 43 = Shinohara et al. (1993), 44 = Ohba et al. (1994), 45 = Ohba et al. (2000), 46 = Kiyosu and Kurahashi (1984), 47 = Fischer et al. (1998), 48 = Taran et al. (1991), 49 = Menyailov and Nikintina, (1980), 50 = Kirsanova et al. (1983), 51 = Taran et al. (1987), 52 = Taran (1985), 53 = Taran and Korbalev (1995), 54 = Taran et al. (1992), 55 = Menyailov et al. (1988), 56 = Menyailov et al. (1986b), 57 = Taran (1992), 58 = Taran et al. (1997), 59 = Taran et al. (1995), 60 = Giggenbach (1995), 61 = Christensen (2000), 62 = Giggenbach et al. (1993), 63 = Reyes et al. (1993), 64 = Delmelle et al. (1998), 65 = Giggenbach and Poreda (1993).

Footnotes for Table 8 Gas compositional data sources:

165

84 9

214 48

Andes Alaska-Aleutians Lesser Antilles Central America Indonesia (Sunda) Italy Japan KamchatkaKuriles New Zealand Philippines

70 15 44 94 42 70 83 37

# Volcanic centers 5 6 2 10 10 13 9 16

# Analyses

Arc Segment

Table 8. Compilation of gas chemistries (median molar ratios) for 10 arc segments worldwide.

Noble Gases and Volatile Recycling at Subduction Zones 345

would act to preferentially remove sulfur species (Giggenbach 1996). Therefore, in our calculations of volatile fluxes (below), we assume a value of 5 for the CO2/St ratio of both these arcs.

346

Hilton, Fischer & Marty Table 9. Compilation of gas fluxes (mol/yr) for 11 arc segments worldwide.

Arc Segment

SO21 (× 109)

CO2 (× 109)

N2 total (× 108)

N2 (ex)2 (×108)

He (×104)

3

Andes Alaska-Aleutians Lesser Antilles Central America Indonesia (Sunda) Italy Japan Kamchatka-Kuriles New Zealand** Papua New Guinea/ SW Pacific* Philippines** TOTAL (11 arcs)

47.3 0.56 4.41 21.3 1.88 27.0 23.3 1.67 15.3 100.9

268.8 3.43 21.8 57.5 8.13 230.0 150.8 2.77 76.6 435.8

44.0 1.01 3.38 16.9 1.33 24.0 79.6 1.97 14.4 71.5

36.7 0.47 2.94 0.85 62.6 1.16 7.48 45.4

91.9 1.71 17.4 71.3 5.53 158.8 198.5 2.90 73.2 296.4

6.91 0.13 1.31 5.36 0.42 11.9 14.9 0.22 5.50 22.3

25.1 0.74 6.42 1.12 3.20 34.3 0.98 19.0 59.9

18.2 261.8

91.3 1346.9

7.9 266.0

6.52 164.1

101.5 1019.1

7.63 76.6

6.73 157.5

TOTAL (Global)&

315.7

1624.2

320.8

197.9

1228.9

92.4

189.9

He3

Ar (×106)

1. From Table 7 (column 7) 2. Excess N2 calculated using equation 3 (see text). 3. Calculated using average arc 3He/4He ratio = 5.37 RA (Table 1) * Flux calculations assumes Indonesia chemistry (from Table 8) ** Assumes CO2/SO2 molar ratio = 5 & Assumes additional 20.6% to all fluxes (see text)

In Table 9 we couple our estimates of the SO2 flux at various arcs (Table 7) with our compilation of the gas chemistry (Table 8) to produce an estimate of the fluxes of a number of volatile species (CO2, N2, N2,ex, 4He, 3He and Ar) for 11 arcs worldwide. The largest emitter of volatiles (in an absolute sense) is the arc systems of Papua New Guinea and the South-west Pacific although it must be cautioned that these fluxes are derived using data from the Indonesia arcs—no gas chemistry is yet available for these regions. The northern Pacific arcs (Alaska-Aleutians and Kamchatka-Kuriles) are relatively poor emitters of volatiles, and this presumably reflects the virtual absence of carbonate sediment in this region (see Table 10). The flux estimates in Table 9 are based on measured SO2 emissions of 45 individual arc volcanoes from 11 regions. However, a significant number of arcs have only little or no SO2 flux data—these include South Sandwich, Mexico, Cascadia, Izu-Bonin, Ryuku, Marianas, Makran, Andanan, East Sunda, Vanuatu and Tonga. This paucity of data is somewhat surprising as there are a total of 106 historically active volcanoes in these regions (Simkin and Siebert 1994). If we make a rather crude analogy to Central America, where approximately one-third of the 32 historically active volcanoes emit measurable SO2, we would expect ~35 of these 106 volcanoes to contribute to the global SO2 flux. These 35 volcanoes would be expected to contribute a total of ~0.54×1011 mol/yr if the median flux of the measured volcanoes (1.54×109 mol/yr) is an appropriate figure for their individual SO2 fluxes. This figure can be added to the sum of SO2 fluxes for the eleven measured arcs (2.62×1011 mol/yr) to yield an estimate of 3.2×1011 mol/yr for the total flux of SO2 from all arc volcanoes worldwide.

1.98E+15

4.64E+15

sediment subducted (g/y) Ocenic crust subducted(g/yr)

6.24E+03 9.64E+04 1.96E+02 1.42E+03 1.21E-03 8.72E-02

4.79E+03

2.26E+05

1.51E+02

3.32E+03

9.28E-04

2.04E-01

2.53E-05

CO2 sed

CO2org

CO2 crust

N2 sed

N2crust

Hesed

Hecrust

Arsed

Arcrust

36

36

2.78E-04

3.29E-05 5.25E-04

2.31E-04

1.65E-01

8.48E-03

2.68E+03

1.38E+03

1.82E+05

4.38E+04

1.06E+05

1.29E+05

3.64E+05

3.28E+05

3.75E+15

1.93E+14

1800

Vanuatu 103

1.90E-04

6.89E-05

5.96E-02

2.53E-03

9.68E+02

4.10E+02

6.59E+04

1.30E+04

5.98E+04

3.83E+04

1.08E+05

1.77E+05

1.36E+15

5.74E+13

1000

E.Sunda 67

3.81E-04

8.02E-05

1.20E-01

2.94E-03

1.95E+03

4.78E+02

1.33E+05

1.52E+04

9.68E+03

4.46E+04

1.26E+05

1.81E+05

2.72E+15

6.69E+13

2010

Java 67

1.42E-04

1.64E-04

4.45E-02

6.00E-03

7.23E+02

9.74E+02

4.92E+04

3.10E+04

3.35E+03

9.09E+04

2.58E+05

4.13E+05

1.01E+15

1.36E+14

1000

Sumatra 50

1.27E-04

3.87E-04

4.01E-02

1.42E-02

6.50E+02

2.30E+03

4.43E+04

7.32E+04

3.01E+03

2.15E+05

6.08E+05

1.24E+06

9.10E+14

3.22E+14

1500

Andaman 30

9.42E-05

3.43E-04

2.96E-02

1.26E-02

4.80E+02

2.04E+03

3.27E+04

6.49E+04

2.63E+05

1.90E+05

5.39E+05

6.69E+05

6.73E+14

2.86E+14

950

Makran 35

3.95E-04

2.64E-05

1.24E-01

9.67E-04

2.02E+03

1.57E+02

1.37E+05

4.99E+03

1.61E+04

1.47E+04

4.15E+04

2.99E+04

2.82E+15

2.20E+13

1550

Philip 90

2.29E-04

2.18E-05

7.21E-02

7.98E-04

1.17E+03

1.30E+02

7.97E+04

4.12E+03

0.00E+00

1.21E+04

3.43E+04

4.17E+04

1.64E+15

1.81E+13

1350

Ryuku 60

6.80E-05

2.22E-05

2.14E-02

8.13E-04

3.47E+02

1.32E+02

2.36E+04

4.20E+03

0.00E+00

1.23E+04

3.49E+04

5.14E+04

4.86E+14

1.85E+13

800

Nankai 30

1.88E-04

7.07E-05

5.92E-02

2.59E-03

9.61E+02

4.21E+02

6.54E+04

1.34E+04

2.69E+04

3.93E+04

1.11E+05

1.71E+05

1.35E+15

5.89E+13

1400

Marianas 47.5

1.49E-04

7.02E-05

4.67E-02

2.58E-03

7.59E+02

4.18E+02

5.17E+04

1.33E+04

3.85E+04

3.90E+04

1.11E+05

1.22E+05

1.06E+15

5.85E+13

1050

Izu-Bon 50

2.38E-04

5.78E-05

7.48E-02

2.12E-03

1.21E+03

3.44E+02

8.26E+04

1.09E+04

0.00E+00

3.21E+04

9.10E+04

1.64E+05

1.70E+15

4.82E+13

800

Japan 105

Subduction Rate: from Plank and Langmuir (1998) and references therein. Trench Length: from Plank and Langmuir (1998) and references therein. Sediment subducted: calculated using subduction rate, trench length, sediment thickness and density for each arc (Plank and Langmuir, 1998). Oceanic crust subducted: calculated using subduction rate, trench length (Plank and Langmuir, 1998) and assuming a thickness of the oceanic crust of 7 km and a density of 2.89 g/cm3.

Footnote for Table 10

6.50E-04

1.83E+04 0.00E+00

1.41E+04

0.00E+00

H2O Available

5.70E+04

H2O Crust

5.18E+04

3.68E+04

3.98E+04

FLUXES (Mmol/yr) H2O sed

1400 2.74E+13

1350

2.11E+13

Trench Length (km)

Kerm 70

Tonga 170

Trench Subd rate (mm/yr)

Table 10. Volatile input at subduction zones.

Noble Gases and Volatile Recycling at Subduction Zones 347

1.97E+04 0.00E+00 6.71E+03 4.87E+04 2.11E+02 7.15E+02

5.68E+04

0.00E+00

1.93E+04

1.46E+05

6.08E+02

2.15E+03

H2O Available

CO2 sed

CO2org

CO2 crust

N2 sed

N2crust

1.19E+05

Arcrust

36

3.34E-04

8.11E-05

1.05E-01

2.97E-03

1.70E+03

4.83E+02

1.16E+05

1.54E+04

0.00E+00

4.51E+04

1.28E+05

1.76E+05

2.38E+15

6.76E+13

1900

62

Aleut

1.59E-04

9.44E-05

4.98E-02

3.46E-03

8.09E+02

5.62E+02

5.51E+04

1.79E+04

0.00E+00

5.25E+04

1.49E+05

1.86E+05

1.13E+15

7.87E+13

800

70

Alaska

1.29E-04

1.54E-04

4.05E-02

5.65E-03

6.57E+02

9.17E+02

4.48E+04

2.92E+04

0.00E+00

8.55E+04

2.42E+05

4.58E+05

9.20E+14

1.28E+14

1300

35

Cascadia

2.14E-04

2.10E-05

6.71E-02

7.70E-04

1.09E+03

1.25E+02

7.42E+04

3.98E+03

0.00E+00

1.17E+04

3.31E+04

3.99E+04

1.53E+15

1.75E+13

1450

52

Mexico

3.16E-04

9.24E-05

9.94E-02

3.39E-03

1.61E+03

5.50E+02

1.10E+05

1.75E+04

2.38E+05

5.13E+04

1.45E+05

1.01E+05

2.26E+15

7.70E+13

1450

77

Centam

2.08E-04

3.90E-05

6.54E-02

1.43E-03

1.06E+03

2.32E+02

7.23E+04

7.39E+03

9.62E+04

2.17E+04

6.14E+04

7.31E+04

1.49E+15

3.25E+13

1050

70

Colomb

4.25E-04

3.07E-05

1.34E-01

1.13E-03

2.17E+03

1.83E+02

1.48E+05

5.82E+03

4.41E+04

1.71E+04

4.84E+04

4.06E+04

3.03E+15

2.56E+13

1500

100

Peru

4.53E-05

5.72E-06

1.42E-02

2.10E-04

2.31E+02

3.41E+01

1.57E+04

1.08E+03

0.00E+00

3.18E+03

9.01E+03

9.44E+03

3.24E+14

4.77E+12

800

20

SSand

2.72E-05

4.51E-06

8.55E-03

1.65E-04

1.39E+02

2.68E+01

9.45E+03

8.53E+02

3.18E+02

2.50E+03

7.09E+03

1.66E+04

1.94E+14

3.76E+12

400

24

N.Ant

2.72E-05

3.79E-05

8.55E-03

1.39E-03

1.39E+02

2.26E+02

9.45E+03

7.18E+03

4.26E+02

2.11E+04

5.97E+04

1.14E+05

1.94E+14

3.16E+13

400

24

S. Ant

6.10E+03

2.30E+03

1.92E+06

8.43E+04

3.11E+10

1.37E+10

2.12E+12

4.35E+11

9.06E+11

1.28E+12

3.62E+12

5.30E+12

TOTAL (mol/a)

Fluxes: calculated using amount of sediments subducted or amount of oceanic crust subducted and concentrations of volatile species in the sediments or oceanic crust listed below. H2O sed: water in the sediments, excluding pore waters. Calculated using the amount of water stored in the sediments for each individual arc (Plank and Langmuir, 1998) H2O crust: amount of water stored in the oceanic crust - 3.4 wt% (Schmidt and Poli, 1998). H2O available: amount of water available (1.2 wt%) in the zone of arc magma generation (Schmidt and Poli, 1998) CO2 sed: amount of (carbonate) CO2 in sediments of each individual arc ( Plank and Langmuir, 1998) CO2 org: amount of reduced organic CO2 stored in sediments. An average concentrations of 1 wt% is used (Bebout, 1995). CO2 crust: amount of CO2 stored in the oceanic crust (carbonate CO2). A value of 0.214 wt% is used (Alt and Teagle, 1999). N2 sed: amount of N2 in oceanic sediments - 0.01 wt% (Bebout, 1995). N2 crust: amount of N2 in oceanic crust - 0.001 wt% (Bebout, 1995). He sed: amount of 4He in oceanic sediments. An average value of 1.0 x10-6 cm3 STP/g is used (Staudacher and Allègre, 1988). He crust: amount of 4He in oceanic crust. An average value of 1.0 x10-6 cm3 STP/g is used (Staudacher and Allègre, 1988). 36 Ar sed: amount of 36Ar in oceanic sediments. An average value of 2.7 x10-8 cm3 STP/g is used (Staudacher and Allègre, 1988). 36 Ar crust: amount of 36Ar in oceanic crust. An average value of 3.2 x10-9 cm3 STP/g is used (Staudacher and Allègre, 1988).

1.40E-04

3.54E-05

1.02E-04

Arsed

36

4.21E-04

1.30E-03 4.41E-02

3.75E-03

1.32E-01

Hesed

Hecrust

5.57E+04

2.91E+05

1.61E+05

1.00E+15

2.95E+13

H2O Crust

3.00E+15

550

90

Kamchat

FLUXES (Mmol/yr) H2O sed

8.51E+13

Ocenic crust subducted(g/yr)

1650

90

Kurile

sediment subducted (g/y)

Trench Length (km)

Subd rate (mm/yr)

Trench

Table 10, continued. Volatile input at subduction zones.

348 Hilton, Fisher & Marty

Noble Gases and Volatile Recycling at Subduction Zones

349

To estimate the global flux of the volatiles other than SO2 we need to make the assumption that the additional SO2 added (~20%) scales proportionally to the other volatile species. In effect, the volatile ratios (e.g., CO2/St etc) of the unmeasured volcanoes must equal the average values of the eleven arcs where data is available. Under this assumption, global estimates of various volatile fluxes are given in the last row of Table 9. With respect to 4He (Table 7), we note that the total global arc flux is 1.2×107 mol/yr, which gives a total 3He flux of approximately 92 mol/yr—for a mean arc 3He/4He ratio of 5.4 RA (from Table 1). This estimate of the 3He flux from arcs falls within a factor of 2 of that by Torgersen (1989) but agrees within error with two other estimates— Allard (1992) estimated 70 mol/yr and Marty and LeCloarec (1992) 75 mol/yr (Table 6). This consistency bodes well for our approach of assuming a power law distribution to volatile fluxes at individual arcs, and for our admittedly crude method of estimating fluxes from non-represented arcs. Similarly, our estimate for the total arc flux of CO2 (1.6×1012 mol/yr) shows remarkable agreement with other published values (Table 6). As we discuss in the next section, our estimate of the global arc N2 flux is higher than that of Sano et al. (2001) due the availability of a significantly more extensive (and presumably more representative) database of gas chemistries. Using helium to resolve volatile provenance Volatile flux estimates derived in the previous section, both for arcs individually as well as arc-related volcanism globally, make no distinction as to the source or provenance of the volatiles. However, in order to assess the chemical mass balance between output at arcs and input associated with the subducting slab, the total arc output flux must be resolved into its component structures. In this way, the fraction of the total output that is derived from the subducted slab can be quantified and compared with estimates of the input parameter. As we show in this section, helium has proven remarkably sensitive in discerning volatile provenance. We use CO2 and N2 to illustrate the case. CO2 provenance at arcs. The CO2/3He ratio is significantly higher in arc-related terrains compared to mid-ocean ridge (MOR) spreading centers. MOR spreading centers have CO2/3He ratios ~2×109 (Marty and Jambon 1987; Marty and Tolstikhin 1998) whereas island arcs have CO2/3He values ≥ 1010 (Marty et al. 1989; Varekamp et al. 1992, Sano and Marty 1995, Sano and Williams 1996; Van Soest et al. 1998). Such high values have been used to argue for addition of slab carbon to the source region of arc volcanism. However, in addition to the slab (both the sedimentary veneer and underlying oceanic basement), there are other potential contributors to the total carbon output—the mantle wedge and/or the arc crust through which magmas traverse en route to the surface. Distinguishing between these various sources is possible by considering carbon and helium together (both isotopic variations and relative abundances). In the first instance, a number of workers (Varekamp et al. 1992; Sano and Marty 1995; Sano and Williams 1996) approximated the C-output at arcs using a three endmember model consisting of MORB mantle (M), and slab-derived marine carbonate/limestone (L) and (organic) sedimentary components (S). Sano and Marty (1995) used the following mass balance equations: (13C/12C)o= fM(13C/12C)M + fL(13C/12C)L + fS(13C/12C)S

(4)

1/(12C/3He)o = fM /(12C/3He) M + fL /(12C/3He) L + fS/(12C/3He) S

(5)

fM + fS + fL = 1

(6)

where O = observed and f is the fraction contributed by L, S and M to the total carbon output. It should be noted that application of these equations first involves conversion of

350

Hilton, Fischer & Marty

carbon isotope ratios (in the δ-notation) to absolute 13C/12C values (see Sano and Marty 1995). Then it is possible to determine the relative proportions of M-, L- and S-derived carbon in individual samples of arc-related geothermal fluids. Appropriate end-member compositions must be selected, and both Sano and Marty (1995) and Sano and Williams (1996) suggest δ13C values of -6.5‰, 0‰, and -30‰ (relative to PDB) with corresponding CO2/3He ratios of 1.5×109; 1×1013 and 1×1013 for M, L, and S respectively. The fractions derived for M-, L- and S-derived carbon are particularly sensitive to the choice of the mantle CO2/3He ratio as well as the sedimentary δ13C value. Based on the analysis of arc-related geothermal samples from 30 volcanic centers worldwide and utilizing high, medium and low temperature fumaroles, Sano and Williams (1996) estimated that between 10 and 15% of the arc-wide global CO2 flux is derived from the mantle wedge—the remaining 85-90% coming from decarbonation reactions involving subducted marine limestone, slab carbonate and pelagic sediment. Subducted marine limestone and slab carbonate supply the bulk of the non-mantle carbon—approximately 70-80% of the total carbon—the remaining ~10-15% is contributed from subducted organic (sedimentary) carbon. With this approach, therefore, it is possible to attempt more realistic volatile mass balances at arcs. It is noteworthy, however, that most studies ignore the arc crust as a potential source of carbon. Although this omission may not be significant in intra-oceanic settings, this is unlikely to be the case at all localities (see discussion regarding the Lesser Antilles; see next section). N2 provenance at arcs. Using an approach analogous to that for carbon, Sano and co-workers (Sano et al. 1998b, 2001) have recently directed attention at understanding the nitrogen cycle at subduction zones. Again, the problem is to identify and quantify the various contributory sources to the volcanic output: however, a major concern in this case is atmospheric nitrogen. There are three major sources of nitrogen at subduction zones: the mantle (M), atmosphere (A) and subducted sediments (S), and each has a diagnostic δ15N value and N2/36Ar ratio. Therefore, observed (O) variations in these two parameters for individual samples can be resolved into their component structures using the following equations (Sano et al. 1998b): (δ15N)o= fM(δ15N)M + fA(δ15N)A + fS(δ15N)S

(7)

1/(N2/36Ar)o = fM/(N2/36Ar) M + fA/(N2/36Ar)A + fS/(N2/36Ar)S

(8)

fM + fS + fA = 1

(9)

where fM is the fractional contribution of mantle-derived nitrogen, etc. Note that the noble gas isotope 36Ar is used in this case since degassing is not expected to fractionate the N2/36Ar ratio due to the similar solubilities of nitrogen and argon in basaltic magma— this is not the case for helium and nitrogen, and degassing corrections may be necessary if the N2/3He ratio is used (Sano et al. 2001). In the above scheme, end-member compositions are generally well constrained: the mantle and sedimentary end members both have N2/36Ar ratios of 6×106 (air is 1.8 × 104) but their δ15N values are distinct. The upper mantle has a δ15N value of -5±2‰ (Marty and Humbert 1997, Sano et al. 1998b) whereas sedimentary nitrogen is assumed to be +7±4‰ (Bebout 1995, Peters et al. 1978) (air has δ15N = 0 ‰). The wide difference in δ15N between the potential end-members makes this approach a sensitive tracer of N2 provenance. Gas discharges from island arc volcanoes and associated hydrothermal systems have N2/ Ar ratios that reach a maximum of 9.7×104, with δ15N values up to +4.6‰ (Sano et al. 2001). This would indicate that a significant proportion (up to 70%) of the N2 could be derived from a subducted sedimentary or crustal source. The situation is reversed in the 36

Noble Gases and Volatile Recycling at Subduction Zones

351

case of BABB glasses, which have significantly lower δ15N values (-2.7 to +1.9 ‰): this implies that up to 70% of the nitrogen could be mantle-derived. A first-order conclusion from this observation is that N2 is efficiently recycled from the subducting slab to the atmosphere and hydrosphere through arc and back-arc volcanism, with the flux probably weighted towards the arc flux. After correction for the atmospheric contribution, Sano et al. (1998b) estimate the total flux of N2 through subduction zones (arc and back-arc regions) worldwide at 6×108 mol/yr (by normalizing to the arc 3He flux). Interestingly, this value is almost 40-times lower than our estimate of Nex output (2×1010 mol/yr; Table 9) based on integrating fluxes from different arcs. This is a consequence of Sano et al. (2001) adopting a N2/3He ratio of 5.6×106 for arc volcanics based on a relatively small database of 11 geothermal samples. Our database has >700 geothermal and hydrothermal fluid analyses with a median N2/3He ratio of 2.6×108 —approximately a factor of 40 higher. Problematic issues regarding estimates of volatile output at arcs There are two major issues of concern with the approach of using CO2/3He and N2/36Ar (or N2/3He) ratios in combination with δ13C and δ15N values to constrain the sources of volatiles at arcs. The first issue is the selection of end-member isotopic and relative elemental abundances—this factor has a profound effect on the deduced provenance of the volatile of interest. The second is the assumption that various elemental (and isotopic) ratios observed in the volcanic products are representative of the magma source. Both have the potential to compromise the accuracy of the output flux estimates. In the case of CO2, the methodology of Sano and co-workers assumes that subducted marine carbonate and sedimentary organic matter can be distinguished as potential input parameters based solely on their perceived C-isotopic compositions prior to subduction (0‰ versus -30‰ respectively). However, this approach ignores the anticipated evolution of organic-derived CO2 to higher δ13C as a function of diagenetic and/or catagenetic changes experienced during subduction (Ohmoto 1986). In the southern Lesser Antilles, for example, Van Soest et al. (1998) calculated that >50% of the total carbon would be assigned to an organic, sedimentary origin if an end-member S-value of -10‰ were chosen, as opposed to 1 where S = solubility), the CO2/3He ratio in the melt phase will evolve to lower values as a function of fractionation style (Rayleigh or Batch) and extent of degassing. Measured CO2/3He ratios in fluids during periods of tholeiitic volcanism were low (~ 5×108). In contrast, CO2/3He ratios ~1010 were recorded in other active periods, which is consistent with degassing of alkalic magmas (in this

352

Hilton, Fischer & Marty

case, SHe/SCO2 < 1). In order to obtain a meaningful estimate of the carbon budget in arcs, therefore, the initial (pre-degassing) CO2/3He ratio is of prime importance, and it is often assumed that the measured CO2/3He ratio equates to the initial magmatic value—as shown above for Loihi Seamount, this may not necessarily be the case. The same issues of degassing-induced changes to elemental ratios apply also to the N2-He-Ar systematics used to resolve the provenance of nitrogen in arcs (Sano et al. 2001). A related concern is that of isotopic fractionation of carbon (or nitrogen) during subduction and/or magma degassing. Sano and Marty (1995) have concluded that arc-related high-temperature fluids are likely to preserve the δ13C values of the (magmatic) source based on comparisons of δ13C values in fluids and phenocrysts. Furthermore, they cite evidence of overlapping δ13C values between high- and medium-to-low-enthalpy hydrothermal fluids, leading to the general conclusion that any fractionation induced by degassing and/or interactions within the hydrothermal system must be minimal. On the other hand, Snyder et al. (2001) have argued that geothermal fluids in Central America have experienced 1 to 2‰ shifts in δ13C resulting from removal of bicarbonate during slab dewatering and/or by precipitation of calcite in the hydrothermal system. It should be noted, however, that even if observed values of δ13C in arc-fluids are fractionated, the magnitude of the isotopic shifts proposed by Snyder et al. (2001) will make an minor difference only to calculations involving the source of carbon. To date, there is no evidence of nitrogen isotopic fractionation during magmatic degassing (Marty and Humbert 1997). Volatile output at the back-arc and fore-arc regions In addition to supplying volatiles that are lost via arc-related volcanism, the subducting slab may also contribute volatiles to both the back-arc and fore-arc regions. To complete a realistic mass balance for subduction zones, therefore, it is essential to quantify volatile fluxes at the back-arc and fore-arc. As we discuss below, both fluxes are severely underconstrained at present. Volatile data, encompassing both major volatile phases (CO2, N2) and noble gases (He and Ar) are available for a number of back-arc basins including the North Fiji Basin, Mariana Trough and Manus Basin (Ishibashi et al. 1994, Nishio et al. 1998, Sano et al. 1998a, Marty and Zimmermann 1999, Sano et al. 2001). A wide range in CO2/3He ratios has been reported—from 2.5×108 (North Fiji Basin) to 2.2×1011 (Manus Basin). There are considerable differences in the treatment of this database. On the one hand, Nishio et al. (1998) reporting data for the North Fiji Basin argue that the value of the mantle CO2/3He ratio in this region is 2×108: in this way, they calculate that samples with higher CO2/3He ratios (up to 9.2×109) contain up to 90% slab-derived carbon. On the other hand, Sano et al. (1998a) recognize that measured CO2/3He ratios in the Mariana Trough could be fractionated during degassing. After taking account of changes in the ratio due to degassing, the measured CO2/3He ratios correct to a value close to that of MORB (~2×109) implying that the carbon is wholly of mantle derivation. The different interpretations have profound effects on the estimated output flux from back-arc basins. If we scale the mid-ocean ridge 3 He flux (1000 mol/yr for ~ 6×104 km of ridges) to the length of back-arc ridges (~2.4×104 km; Uyeda and Kanamori 1979) and use the extreme CO2/3He ratio of the North Fiji Basin (9.2×109) as representative of a back-arc basin source region (with a slab contribution), then the CO2 flux from back-arc basins (globally) is 3.7×1012 mol/yr—of which up to 90% (3.3×1012 mol/yr) could be derived from the slab. This estimate would make the back-arc CO2 flux comparable with that for MOR (2.2±0.9×1012 mol/yr; Marty and Tolstikhin 1998), and significantly higher than most estimates for arc volcanism (Table 6). Alternatively, if we accept the arguments of Sano et al. (1998a), the output of slab-derived CO2 via back-arc volcanism is essentially zero. The same uncertainty in back-arc flux rates applies also to N2. Limited data on

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BABB glasses from the Mariana Trough, the Manus and North Fiji Basin show N2/3He ratios ranging from 1.16×106 to 1.37×107 and δ15N values varying between -2.7 and +1.9‰ (Sano et al. 2001). Therefore, depending upon the chosen end-member composition, the sediment contribution to the back-arc N2 flux can vary between zero and close to 40 %. In the latter case, a significant fraction (one-third) of the available slab-derived nitrogen is recycled to the surface via back-arc volcanism as opposed to two-thirds via the arc (Sano et al. 2001). Although volatile flux estimates from back-arc regions vary greatly, the situation is even more poorly constrained for fore-arc regions. We are unaware of any flux estimates for either major volatiles or noble gases. This is in spite of abundant circumstantial evidence for fluid venting in fore-arc regions e.g., serpentinite diapirism in the Marianas (Fryer et al. 1985) and mud volcanism at various accretionary prisms worldwide (Brown 1990). Indeed, there are reports of 3He/4He ratios > crustal production rates (~0.05 RA; Andrews 1985) in fore-arc regions such as the Nankai Trough (1.9 RA; Kastner et al. 1993), the Cook Inlet of Alaska (0.8 RA; Poreda et al. 1988) and Kavachi volcano in the Solomon Islands (6.9 RA; Trull et al. 1990)—this would indicate a direct output of mantle-derived volatiles. However, there are no estimates of the scale of this phenomenon. The alternative approach to estimating fore-arc losses of volatiles is somewhat circular and based upon mass balance. For example, the estimated carbon load of incoming material at trenches worldwide (sedimentary carbonate, sedimentary organic carbon and oceanic meta-basalts) is 1.2, 0.8, and 3.4 (×1012 mol/yr) respectively (Bebout 1995, Alt and Teagle 1999, Plank and Langmuir 1998). This total (5.4×1012 mol/yr) far exceeds that of the carbon released via arc magmatism (~3×1012 mol/yr; Table 6), implying that a substantial amount of carbon is either released at the fore-arc or by-passes the zone of magma generation and is recycled into the mantle. Whereas there is considerable leeway in all these estimates, the uncertainty is compounded by debate on whether the zone of magma generation acts as a barrier to further subduction of volatiles (Staudacher and Allègre 1988) or if various C-bearing phases are stable at the likely P-T conditions below the arc so that deep(er) subduction can occur (cf. Kerrick and Connolly 2001). In this latter scenario, if the input of carbon to the mantle via subduction matches output via MOR (i.e., steady-state upper mantle and no losses to the back-arc) then approximately 2×1012 mol/yr (the MOR output; Marty and Tolstikhin 1998) by-passes the zone of magma generation leaving a residual 0.4×1012 mol/yr potentially available for loss via the fore-arc. There is considerable latitude in these figures (and assumptions inherent in the approach): for these reasons, therefore, we caution that constraining volatile losses at the fore-arc by this mass balance approach be treated with skepticism. The volatile input via the trench In this section, we provide estimates of the volatile flux input—of both noble gases (He, and Ar) and major volatiles (CO2, N2 and H2O)—via the trench. We consider both the subducted sediment load as well as the underlying crustal basement (of thickness 7 km). Note that the volatile input parameter can be calculated for an extensive listing of subduction zones: however, the output parameter via arcs (Table 9) is well constrained by flux data for a limited number of arcs only. For these arcs, at least, we have a direct comparison between input and output for specific volatile species. However, by extrapolating the (limited) output data to arcs worldwide, we can address the question of volatile recycling between the mantle and crust, hydrosphere and atmosphere from a global perspective. In Table 10, we compute the volatile fluxes into the subduction zone for individual arcs using (a) available noble gas, H2O, CO2 and N2 concentration data for oceanic

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sediments and oceanic (crustal) basement, and (b) estimates of sediment and crust fluxes at each arc segment. In this compilation, it is important to note that the fluxes refer to volatiles potentially entering the subduction zone at the site of the trench—the effects of sediment off-scraping and/or underplating (Von Huene and Scholl 1991) on the volatile fluxes are not considered. The following points are relevant in the computation of the volatile flux data (Table 10): 1. There is limited noble gas data available for oceanic crust. However, Staudacher and Allègre (1988) report He and Ar data for 2 fine-grained basalt samples from the Indian Ocean. We note that although we adopt these data, the samples are of Jurassic age (~108 Ma) so their noble gas contents may not necessarily be appropriate for all types of subducting oceanic crust. Similarly, there is limited noble gas data available for sedimentary material, and we again use concentrations from Staudacher and Allègre (1988) obtained for sediments from the South West Indian Ridge. With the exception of helium, noble gas contents of oceanic sediments are generally higher than oceanic basalts. 2. There are 3 separate values for the water content of the subducting slab. The sediment H2O contents are given on an arc-to-arc basis by Plank and Langmuir (1998) and represent bound (not porewater) contents of various drill-core lithologies proximal to the arc in question. A value of 3.4 wt % is used for water bound in oceanic crust assuming a 50:50 mix of basalt and gabbro (Schmidt and Poli 1998). Finally, a value of 1.2 wt % H2O and termed ‘available water’, is considered ‘available’ for the generation of arc mamas following water loss from dehydration and diagenetic processes associated with the early stages of subduction (Schmidt and Poli 1998). 3. In an analogous fashion to water, the CO2 concentrations in sediments come from an arc-by-arc compilation of carbonate contents of various drill core lithologies (Plank and Langmuir 1998). The amount of reduced (organic) carbon in sediments is taken as 1% based on average contents of metamorphosed sediments of the Catalina Schist (Bebout 1995). A value of 0.214 wt % is used for the amount of CO2 stored in altered oceanic crust (Alt and Teagle 1999). 4. The N2 concentrations are average values of oceanic sediments (0.01 wt %) and crust (0.001 wt %) (from Bebout 1995). In their compilation of the characteristics of oceanic sediments, Plank and Langmuir (1998) has shown that both the amount and chemical composition of sediments being subducted are highly variable: we emphasize that the same observation is true also for the major volatile components, H2O and CO2. However, due to the paucity of data, it is currently not possible to evaluate if there is any regional variability in the noble gas contents of oceanic sediments and their crystalline basements. The input fluxes of various volatile species (Table 10) can now be compared to various output fluxes through arc volcanism (Table 9) to assess the extent of volatile mass balance. At this stage, we ignore possible volatile losses at the back-arc given the large uncertainty in actual values. Also, we note that in the case of subducted carbon, it is important to distinguish between reduced sedimentary carbon, sedimentary carbonate and carbonate of the altered oceanic crust.

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Table 11. Fractional contributions of carbonate- (L), Sedimentary/organic- (S) and MORB- (M) derived CO2 to total arc CO2 discharges*. L

S

M

0.56 Nd 0.71 0.59 0.74 Nd 0.67 0.51 0.85 Nd Nd

0.33 Nd 0.17 0.31 0.13 Nd 0.13 0.27 0.10 Nd Nd

0.09 Nd 0.03 0.06 0.12 Nd 0.18 0.22 0.06 Nd Nd

Arc Segment Andes Alaska-Aleutians Lesser Antilles Central America Indonesia Italy Japan Kamchatka-Kuriles New Zealand Papua New Guinea Philippines

Nd = no data available (for Alaska-Aleautians, Italy, PNG and the Philippines). * Endmember compositions for M, L and S used in the computations are δ13C (-6.5‰, 0‰ and -20‰) with corresponding CO2/3He ratios of 1.5 × 109; 1 × 1013; 1 × 1013, respectively.

Volatile mass balance at subduction zones Before comparing the output to input fluxes for the various volatile species, it is necessary first to resolve the carbon output via the arc into its components parts reflecting carbon provenance from the mantle wedge, and slab-derived carbonates and organic carbon. In Table 11, we report % contributions to the CO2 output using the mass balance equations of Sano and Marty (1995) but adopting a δ13C value for the sedimentary end member (S) of -20‰. Bebout (1995) shows convincing evidence of a progressive increase in δ13C of carbonaceous material with increasing metamorphic grade in the Catalina Schist, and Van Soest et al. (1998) discuss the effect of adopting a higher δ13C value for the S-end member on the calculation of CO2 provenance. Resolution of the CO2 output into M-, L- and S- components is shown in Figure 5 for arcs where sufficient data is available. In each case, we report the median values for M, L and S. The limestonederived component dominates at all localities. In Tables 12 (major volatiles) and 13 (He and Ar) we compute ratios of output to input fluxes for the various individual arcs. The output fluxes are given in Table 9—modified in the case of carbon for provenance (L, S, and M; see Table 11). The input fluxes are given in Table 10. It is noteworthy that output/input ratios vary significantly among individual arcs. We point out the following features of interest in Tables 12 and 13. 1. The sedimentary output of CO2 (S) significantly exceeds the input of organic sedimentary carbon for 4 of the 9 arcs where data are available (Andes, Japan, New Zealand and the Philippines). For these localities, an additional source of organic carbon is required: the uppermost arc crust seems the only likely possibility. At the other arcs, the input of organic carbon exceeds the output via the arc, and suggests that either sediment accretion and/or off-scraping of subducting sediments could prohibit the transport of organic carbon to the zones of magma generation. For Central America, the ratio is very close to unity suggesting that the output from the arc is nearly balanced by the input of organic sedimentary carbon.

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Figure 5. Ternary plot of CO2 provenance in arc-related volcanism (data from Table 11). L, S, and M represent limestone, sediment and mantle wedge respectively.

Table 12. Ratio of output (via arc) to input (via trench) for major volatile species (CO2, N2 and H2O) for selected arc segments. CO21 Arc Segment

S OUT/ Org IN

8

L OUT/ Car4 IN

L OUT/

(Car+ crst5) IN

N2 (excess)2

H2O3

OUT/

OUT/

(Sed + crst) IN

IN total6

OUT/

IN avl.7

6.71

1.07

0.42

1.00

60.1

346.4

Alaska-Aleutians9

0.028

∞$

0.010

0.013

0.27

1.76

Lesser Antilles10

0.46

20.8

0.79

-

5.52

46.2

1.02

0.14

0.097

0.14

11.7

56.0

0.023

0.462

0.031

0.021

0.42

3.0

Japan

1.80

∞

1.22

4.03

29.6

234.9

Kamchatka-Kuriles12

0.029

∞

0.0073

0.031

0.22

1.81

New Zealand13

1.29

∞

0.68

0.46

35.2

209.3

2.38

4.20

0.44

0.30

63.9

310.5

Andes

Central America Indonesia (Sunda)

11

Italy

Papua New Guinea/ SW Pacific Philippines 1. 2. 3. 4. 5. 6. 7. 8. 9. 10. 11. 12. 13. $

CO2 output resolve into L-, S-, and M-components (from Table 11). N2 excess = non-atmospheric N2 (assumes all argon is air-saturated water derived – equation 3 in text). H2O output assumes arc H2O/CO2 ratio = 50 (Symonds et al., 1994). Car = sedimentary carbonate input (CO2 sed - from Table 10) Crst = oceanic crust carbonate input (CO2 crust – from Table 10). Total water = sediment- and basement-hosted water only (excludes pore water). Water available at zone of magma generation (1.2 wt% - Schmidt and Poli, 1998). Input from Columbia and Peru arcs (Table 10). Includes input from Aleutians and Alaska arc (Table 10). Includes input from both northern and southern Lesser Antilles (Table 10). Includes input from Java and Sumatra (Table 10). Includes input from Kuriles and Kamchatka arcs (Table 10). Input taken for Kermadec arc (Table 10). Infinity (zero carbonate input).

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Table 13. Ratio of output (via arc) to input (via trench) for noble gas isotopes (3He, 4He, 36Ar and 40Ar) for selected arc segments. 3 1

Arc Segment

Andes Alaska-Aleutians Lesser Antilles Central America Indonesia (Sunda) Italy Japan Kamchatka-Kuriles New Zealand Papua New Guinea/ SW Pacific Philippines 1. 2. 3.

He OUT/ IN 488.8 11.5 1002 744.9 34.6

4

He OUT/ IN2 4.55 0.11 9.32 6.94 0.32

36

Ar OUT/ IN 88.9 2.75

40

Ar3 OUT/ IN 103.3 3.20

39.1 3.63

45.5 4.22

2767 17.4 888.8

25.8 0.16 8.28

288.5 3.49 152.0

335.3 4.06 176.7

879

8.19

39.7

46.2

See Table 12 for arcs used in input computation. Input 3He/4He ratio = 0.05 RA (radiogenic helium) Output 40Ar/36Ar ratio = 401 (Table 4); input 40Ar/36Ar ratio = 345 (average of sediment and oceanic crust – from Staudacher and Allègre, 1988).

2. The output of carbonate-derived CO2 (L) cannot be balanced with the amount of sedimentary carbonate CO2 being subducted except for the Central American and Sunda (Indonesia) volcanic arcs. This would suggest that, in both these localities, sediment-derived carbonate may be efficiently transported to the zones of magma generation. When carbonate CO2 from the subducting altered oceanic crust is taken into account, the output can be supplied solely from the slab: indeed, at all localities, except Japan, only a fraction of the input CO2 is necessary to supply the output. These figures reinforce the notion that subduction zones act as conduits for the transfer of carbon into the mantle (Kerrick and Connolly 2001). 3. In general, the amount of excess (non-atmospheric) N2 being emitted from arc volcanoes is less than the amount of N2 being subducted (with the exception of the Andes and Japan). This would imply that a significant amount of subducted N2 may not reach the zones of arc magma generation and/or it is retained and lost to the (deeper) mantle. The observation of large increases in the C/N ratio of Catalina Schists as a function of increasing metamorphic grade (Bebout 1995)—interpreted as enhanced devolatilization (loss) of N2—would seem to favor the former possibility. 4. The flux of H2O from each arc segment greatly exceeds the amount of water available at the zone of arc magma generation (1.2 wt %; Schmidt and Poli 1998). Even when the total amount of water potentially subducted is taken into account (IN total; Table 12), for the most part (excepting Alaska-Aleutians, Indonesia and KamchatkaKuriles), output via the arc exceeds input via the trench. Water from the mantle wedge and/or the arc crust must contribute to the output in these cases. Alternatively, the ‘excess’ water may be meteoric in origin although it must first become incorporated by magma (and contribute to the high ‘magmatic’ H2O/CO2 ratio of 50). 5. The 3He output flux exceeds the input parameter in all cases. This is true even in

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cases (Alaska-Aleutians, Indonesia and Kamchatka-Kuriles) where the potential input of 4He exceeds the output via the arc. The observation that at each locality the 3 He/4He ratio of the output flux approaches the MORB value indicates that little, if any, of the slab-derived helium reaches the zones of magma generation. 6. A similar pattern is observed for 36Ar; output via the arc greatly exceeds potential input via the trench. The output volatile flux appears dominated by atmospherederived 36Ar (thereby validating its use to calculate ‘excess’ N2). Likewise, output of 40 Ar far exceeds potential input. In the case that all the 36Ar is atmospheric, the excess 40Ar (to raise the average arc 40Ar/36Ar value to 401; Table 4) can be either crustal (Kennedy et al. 1991) or mantle (Patterson et al. 1994) in origin. Global implications of volatile recycling at subduction zones In this final section, we extend our considerations of subduction zone processing of volatiles (input versus output fluxes; see previous section) to address the issue of longterm volatile exchange between the surface of Earth and the mantle. We follow the approach of Marty and Dauphas (2002) by relating the ratio of the volatile flux through arc volcanism and the amount of volatiles carried by subducting plates (Farc/Fsub) to the Mean Degassing Duration (MDD) of a particular volatile species. The MDD (in years) is the ratio of the total surface inventory of a volatile element (ocean + crust + atmosphere) to the present-day mantle flux, approximated by the MOR flux (assuming that plume degassing is minor compared to ridges). As discussed in general terms in the previous section, a Farc/Fsub ratio close to unity implies little or no recycling of subducted volatiles into the mantle beyond the zone of magma generation, a ratio of 1 implies that additional contributions to the volatile flux (other than from subducting sediments and oceanic crust) are required. A MDD value less than the age of Earth implies rapid recycling of volatiles between the mantle and surface reservoirs. A MDD value greater than 4.55 Ga suggests either a decreasing degassing rate with time or the occurrence of a volatile component at the Earth's surface not derived from the mantle, or both. Global averages of Farc/Fsub ratios together with MDD values are given in Table 14 and plotted against each other in Figure 6.

Figure 6. Volatile flux ratio (Farc/Fsub) vs. mean degassing duration (MDD) for four major (S, C, H2O, N2) and two noble-gas (He, Ar) volatile species following approach of Marty and Dauphas (2002).

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Table 14. Average global volatile fluxes out of volcanic arcs (Farc) and mid-ocean ridges (FMORB), and into subduction zone (Fsub) together with total surface inventories (atmosphere+oceans+crust). VOLATILE

Farc

Fsub

mol/yr

mol/yr

H 2O

8.12E+13

8.92E+12

SO2

3.16E+11

CO2 (L+S)

Farc/Fsub

Surface Invent. (mol)

FMORB

Surface/FMORB

mol/yr

yr

9.10

5.80E+23

2.90E+13

2.0E+10

(4.11E+12)

0.077

5.33E+20

(2.67E+12)

2.0E+08

1.45E+12

3.46E+12

0.42

7.00E+21

2.20E+12

3.2E+09

N2 (S + Crst)

1.98E+10

4.48E+10

0.44

2.74E+20

1.00E+09

1.8E+11

3

HE

9.24E+01

1.40E-01

660

9.14E+13

1.00E+03

9.1E+10

4

He

1.23E+07

2.00E+06

6.2

3.61E+18

8.93E+07

4.0E+10

40

Ar

1.90E+08

2.90E+06

65.5

1.64E+18

4.00E+07

4.1E+10

36

Ar

4.74E+05

8.40E+03

56.4

5.55E+15

1.00E+03

5.55E+12

Footnote to Table 14 Farc:

Fsub:

total global flux of volatile species from subduction zones - calculated by summing individual arc volatile fluxes and adding additional 20.6% (see Table 9). The water flux was calculated using the CO2 flux (Table 9) and assuming a H2O/CO2 ratio of 50. The CO2 (L + S) – limestone- + sediment-derived CO2 is the total arc CO2 flux (Table 9) multiplied by 0.89 (the average L+S value for 7 arcs worldwide – Table 11). 36Ar flux is calculated using the total Ar flux (Table 9) and an average 40Ar/36Ar ratio of 401 (Table 4). total global subduction of volatile species. Calculated by summing flux values of individual arcs (Table 10).

Surface Invent.: global inventory of volatile species currently at the Earth's surface. H2O: from Marty and Dauphas (2002) SO2: from Marty and Dauphas (2002) CO2: from Zhang and Zindler (1993) N2: from Pepin (1991) 3 He: from Ozima and Podosek (1983) – assumes 20Ne inventory of atmosphere and planetary 3 He/20Ne ratio of 0.0315. Present atmosphere abundance is lower because of escape to space. 4 He: from Ozima and Podosek (1983) – radiogenic 4He released into atmosphere taken as 2.2 times atmosphere 40Ar abundance. Present abundance is much lower because of escape to space. 36 Ar &40Ar: from Ozima and Podosek (1983) FMORB: flux of volatiles from the mantle, approximated by MORB flux. H2O: from Jambon and Zimmermann (1990) SO2: from Alt et al. (1985) CO2 :from Marty and Tolstikhin (1998) N2: using 3He flux (1000 mol/yr) from Craig et al. (1975) and MORB N2/3He ratio of 1 x 106 (Sano et al., 2001). 3 He: from Craig et al. (1975) 36 Ar: calculated from the MORB 3He flux and MORB 36Ar/3He ratio = 1 (Tolstikhin and Marty 1998).

In Figure 6, we observe a generally positive correlation between Farc/Fsub and MDD indicative of decreasing efficiency in atmosphere-mantle exchange in the order SO2, CO2, H2O, 3He ≈ 4He ≈ 40Ar, N2 (ex), 36Ar. Sulfur has a Farc/Fsub ratio 1000 K temperature difference develops between the two portions of the mantle, and so overturns will cause large variations in upper mantle temperatures. However, the available geological record suggests that there has been a relatively uniform mantle cooling rate of ~50-57 K/Ga (Abbott et al. 1993; Galer and Mezger 1998). Further, numerical models simulating 3He/4He variations in a mantle that has undergone this transition do not reproduce the observed 3He/4He distributions seen today (van Keken and Ballentine 1999). The lower boundary layer reservoir

The boundary layer at the core-mantle boundary (CMB) has been explored as a reservoir for high 3He/4He ratio He, in the context of whole mantle convection. It has been suggested that subducted oceanic crust could accumulate there and form a distinct chemical boundary layer, accounting for the properties of the D′′ layer (Christensen and

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Hofmann 1994). Altered ocean crust is strongly depleted in 3He and may be enriched in U, and so is likely to have relatively radiogenic He. However, a complementary harzburgitic lithosphere that is depleted in both U and He also develops during crust formation. Coltice and Ricard (1999) examined a model in which separate reservoirs of altered ocean crust and depleted harzburgitic lithosphere are maintained below the bulk convecting mantle. In this model (Fig. 9C), there are mass fluxes of altered ocean crust to a D′′ reservoir, and of subducted depleted harzburgitic lithosphere to a residual depleted mantle (RDM) located immediately above D′′. In order to account for the depletion of the bulk mantle, extraction of the continental crust and degassing to the atmosphere are also considered. Transfer of trace elements occurs by bulk mass fluxes that have fixed chemical fractionation factors relative to the source reservoir, so that chemical fluxes are determined by the mass fluxes. The U-Pb, Rb-Sr, and U,Th-He systems are all tracked, and the isotopic composition of all the reservoirs are fixed by observations, with D′′ material being identified with the source of HIMU hotspots and RDM with that of Loihi. The model was not extended to the other noble gases. Since the model was fully and clearly developed, it is possible to examine its implications in some detail here. 1)

2) 3) 4)

5)

Model reservoirs and assumptions. The bulk mantle has a constant volume (a free parameter; depends on D′′ and RDM volumes). The starting U concentration is that of the bulk silicate Earth and that of He is calculated for a closed system with a Loihi 3He/4He ratio. The main He outflow is to the atmosphere by degassing a constant volume with time (i.e., first order degassing). U is largely transferred to the continental crust at a set first-order removal rate. Smaller fluxes of He and U to D′′ and RDM (via the altered ocean crust and depleted oceanic lithosphere, respectively) also occur. Small inputs occur from subduction of continental crust and from D′′ and RDM. Present isotopic compositions are those seen in MORB. The model continental crust grows linearly over 4.5 Ga to the present volume. It only serves to deplete the bulk mantle of U. Return to the bulk mantle by subduction occurs at a rate set by those of continental growth and mantle outflow. The atmosphere grows progressively and serves only to deplete the bulk mantle of He. Since He is lost to space, the atmospheric He abundance is not a constraint. The D′′ layer grows linearly, with the present volume taken as a free parameter. There is a constant mass inflow from subduction, with a U concentration related to the bulk mantle U concentration by a large enrichment factor operating during formation of ocean crust. There is a mass flux out (dependent upon the growth rate of the layer). Present isotopic compositions are those seen in HIMU basalts. The residual depleted mantle (RDM) has a constant volume, which is a free parameter. Mass outflow is specified from plume buoyancy fluxes. The mass inflow contains He and U, with concentrations related to those of the bulk mantle by specified small enrichment factors operating during formation of oceanic lithosphere. The present 3He/4He ratio is the highest seen at Loihi, while the starting value is solar.

Model results. The primary conclusion of the model is that Loihi and HIMU characteristics (for Sr, Pb, and He) can be generated from subducted components sequestered in two deep reservoirs. A nonlinear inversion was used to maximize the fit to the observations of continent concentrations and reservoir isotope ratios within the bounds of assigned uncertainties in the fluxes and fractionation coefficients. This was done using the U-Pb and Rb-Sr systems as well. In the best fit results, the D′′ layer is 250-km thick (5% of the total mantle) and the RDM is 500-km thick.

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The main issue here is whether this model provides a plausible mechanism for maintaining a reservoir with high 3He/4He. The characteristics of He in the RDM Loihi source are determined by two components. The He initially incorporated into this RDM reservoir had a solar 3He/4He ratio and a concentration equal to that for a closed system reservoir that would presently have high 3He/4He OIB ratios. Therefore, without other fluxes, this would simply be a ‘primitive’ reservoir. The subducted depleted lithosphere has a He concentration fixed at 1% that of the bulk mantle. It may have contained a small concentration of He when it is formed prior to extensive degassing of the bulk mantle, and this will have had high 3He/4He ratios. Later it essentially only dilutes the concentration of He in the RDM. Helium is lost according to a 1.8 Ga half-life (for the best-fit results), and so the initial He has been reduced by 0.18. The U concentration decreased from 20 ppb to 3.5 ppb, and so was reduced by a similar factor. Therefore, the initial U/He ratio was not changed significantly and the reservoir evolved isotopically like a closed system, although concentrations were reduced. Changes in the U/He ratio of the bulk mantle have little effect, because the transfer from the bulk mantle to the RDM is small. Model shortcomings. The difficulty with the model as an explanation for mantle He variations lies in the nature of the RDM reservoir. Subducted depleted material is too depleted to supply OIB alone, so that a fixed reservoir is necessary to sequester high initial He concentrations. However, there is no simple geophysical explanation for initially maintaining a fixed volume reservoir and so for intimately associating subducted and primitive material. The issue then becomes the same as that for preserving a relatively undegassed reservoir in the Earth: how to convectively isolate such a region. Whether or not this region is diluted by depleted material is then secondary. Therefore, while such depleted material may explain other characteristics such as Sr and Pb isotopes (see Coltice and Ricard 1999), the issue of how to maintain a He reservoir with a high 3 He/4He ratio remains. Also, the model does not address the evolution of other noble gases. While the 40Ar budget might be accommodated, additional reservoirs would need to be added to store solar Ne and radiogenic Xe separate from the atmosphere, and the evolution of the bulk mantle then would require modification. However, these changes may be possible. The upper boundary layer reservoir

As an alternative to assigning high OIB 3He/4He ratios to an isolated lower mantle enriched in He, Anderson (1998a) argued that these ratios represent He incorporated within lithosphere that is highly depleted in U and so maintains the isotopic composition at the time of trapping. This may include harzburgitic residues that have been invaded by CO2 and accompanying noble gases. Such a reservoir would be protected from destruction by convective mixing, and would then only require a single underlying mantle He composition. While this may plausibly occur, there are difficulties with accounting for most OIB signatures in this way. For example, in the case of Loihi, the underlying crust is 80-100-Ma old. If the He was derived from the MORB-source mantle in the past, then the upper mantle had a 3He/4He ratio of 37 RA at 100 Ma ago. Any reasonable extrapolation of upper mantle 3He/4He ratios to the past would require Earth values far in excess of the initial, solar nebula value at 4.5 Ga. It should be noted that lithospheric xenoliths generally do not show high 3He/4He ratios, and while this may reflect imprinting by fluids not long before eruption (see Dunai and Porcelli 2002), a preexisting higher 3He/4He ratio would only be conjectural. Overall, a compelling case for long-term lithospheric storage of He with high 3He/4He ratios has not been made. Deeper or ‘abyssal’ layering

Kellogg et al. (1999) have developed and numerically tested a model in which

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mantle below ~1700 km has a composition, and so density, that is sufficiently different from that of the shallower mantle to largely avoid being entrained and homogenized in the overlying convecting mantle. This model has generated a great deal of interest because of its ability to preserve a region in the mantle behind which the radioelements and primitive noble gases can be preserved, while accommodating many geophysical observations. For example, it is argued that the depth of the compositional change varies, allowing slab penetration to the core mantle boundary in some locations while providing a barrier into the lower mantle elsewhere. Supporting tomographic evidence for a significant number of slabs being disrupted at 1700 km is given by van der Hilst and Kàrason (1999). More recently, this supporting evidence has been questioned because of the loss of tomographic resolution in this portion of the mantle (Kàrason et al. 2001). Although Kellogg et al. (1999) suggest that this layer will be hard to detect seismically because of its neutral thermal buoyancy, irregular shape and small density contrast, Vidale et al. (2001) argue that seismic scattering nevertheless should be resolvable, and yet is not observed. Also, if the overlying mantle has the composition of the MORB source, then the abyssal layer must contain a large proportion of the heat-producing elements, and must efficiently remove heat from the core. Therefore, a minimum 12.621.4 TW of mantle heat must cross from the layer into the convecting mantle, neglecting any component of secular cooling (see Heat and helium section). It is not yet clear what this effect would have on the thermal stability of the layer or temperature contrast with the overlying mantle. Overall, if such a feature were found to be viable, it must be incorporated into a comprehensive noble gas model. This could include some of the features described above of other layered mantle models. Heterogeneities preserved within the convecting mantle

Early conceptual models for preserving geochemical heterogeneities in a regime of whole mantle convection discussed the possibility of convecting material passing through and sampling compositionally different regions. Two end member models have been postulated, one in which ‘blobs’ or ‘plums’ of enriched material are passively entrained in the convecting mantle to provide OIB–source material (Davies 1984). The other has been called ‘penetrative convection’ (Silver et al. 1988), in which downgoing cold material drops into a compositionally different lower mantle layer. The slabs, on heating at the core mantle boundary, regain positive buoyancy and on return to the surface entrain a small portion of the deeper reservoir. In this way, lower mantle material is provided to either OIB-source or is mixed into the MORB-source mantle. Regarding the first case, numerical models have shown that more viscous ‘blobs’ can be preserved in a convective regime if they are at least 10 to 100 times more viscous than the surrounding mantle (Manga 1996). In these simulations, the high viscosity blobs tend to aggregate, leading to the formation of large-scale heterogeneities from smaller ones. Becker et al. (1999) have investigated the dynamical, rheological and thermal consequences of such blobs containing high radioelement and noble gas concentrations to account for the bulk Earth U budget and the high 3He/4He ratios seen in OIB. The higher heat production and resulting thermal buoyancy within these blobs must be offset by a combination of increased density (~1%) and small size (30 RA for the last ~2 Ga (Seta et al. 2001). If high ratios have not been present since that time, then the domains with high 3 He/4He ratios must have been preserved since then. 3) The source must have started out with sufficiently high 3He to be able to supply OIB after depletion. For example, 10% melting of a mantle source, assuming a He silicate/melt partition coefficient of 8 × 10-3 (Marty and Lussiez 1993), will leave only 8% of the He in the residue. Thus, melting of highly depleted MORB-source mantle will leave a component that will not readily impart a distinctive isotopic signature to OIB. This is a particular problem when it is considered that such high 3 He/4He components must dominate the He signature in OIB source regions where there is clear evidence for recycled oceanic crust that likely contains high 4He concen-trations, such as in Hawaii (Hauri et al. 1994). Therefore, the initial mantle source must be more gas-rich. 4) There must be a mechanism for the preferential involvement of this previously melted material at OIB. 5) The constraints provided by Xe isotopes indicate that the noble gases in the mantle cannot have evolved in a single depleting mantle reservoir (see Xe isotopes and a nonresidual upper mantle section). Therefore, other reservoirs must be invoked to

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complement these domains. At present, it appears that this mechanism might explain some of the He isotope variations seen in mantle-derived materials, although it is not clear if it can explain the highest 3He/4He ratios. However, a comprehensive model is necessary that explores the total volume of the high 3He/4He domains and the rates at which they are generated and destroyed, the isotopic evolution of the surrounding mantle, and the evolution of other rare gases, to fully evaluate this idea. Nonetheless, it is clear that this mechanism cannot be simply appended to a geophysical or lithophile element model for whole mantle convection. Storage of noble gases in the core

The core has often been cited as a possible source of He with high 3He/4He ratios, in response to the difficulties of storing this He in the mantle. Any He in the core would likely have been trapped during core formation and unaccompanied by U and Th, and so would now have a primordial 3He/4He ratio. It has been argued that there is insufficient 3 He in the core to supply the mantle, because measured silicate-metal noble gas partition coefficients are very low (Matsuda et al. 1993). However, calculation of the abundances of noble gases in the core requires not only the relevant partition coefficients during core formation, but also what quantity of noble gases were available during partitioning. These parameters depend upon whether the mantle was solid or molten as the core segregated. Whether the core now can supply the noble gases seen in the mantle must also consider how reintroduction into the mantle occurs. These issues are considered below, following Porcelli and Halliday (2001). Jephcoat (1998) observed that Xe forms a high-density solid at deep mantle pressures. Sinking of these solids could transport Xe, and possibly some fraction of other noble gases incorporated into this structure, to the deep mantle or core. However, it has not been demonstrated either that such a scarce element (~20 fg/g) is sufficiently insoluble in other high pressure phases to form a separate phase, or that it can form separate grains in the mantle sufficiently large to sink. This mechanism for transport to the core therefore cannot be evaluated further. Partitioning of noble gases. Matsuda et al. (1993) reported molten metal/ silicate melt partition coefficients of D(He)Fe/LS ~ 0.04 at 5 kbar and D(He)Fe/LS ~0.01 at 20 kbar (both at 1600°C). Differences between noble gases were not found, although because analytical errors were large, the possibility that noble gases are fractionated during partitioning cannot be discounted. Much lower values for all the noble gases of ~10-3 were reported for 60 kbar (at 1600°C) and of ~3 × 10-4 for 100 kbar (≈ 300 km depth). Comparable values were obtained using a lherzolite silicate melt for pressures between 5 and 60 kbar (Sudo et al. 1994). There are no data available for the deeper mantle, although it has been shown experimentally that Xe does not alloy with iron at coremantle boundary pressures (Caldwell et al. 1997). A partition coefficient for He in the mantle of D(He)Fe/LS = 1 × 10-4 is taken here for further discussion.

There are no data available for the partitioning of noble gases between solid silicates and liquid metal. However, there are some measurements for olivine and silicate melts at surface pressures. Marty and Lussiez (1993) found solid silicate/ silicate melt partitioning values of D(He)SS/LS ≤ 8 × 10-3 for coexisting olivines and glasses. Noting that the molten metal/solid silicate partition coefficient is D(He)Fe/SS = D(He)Fe/LS / D(He)SS/LS; then if D(He)Fe/LS = 1 × 10-4 and D(He)SS/LS = 8×10-3, then the partitioning between liquid metal and solid silicate is D(He)Fe/SS = 1 × 10-2. Availability of noble gases during core segregation. The light noble gases in the mantle appear to be solar, and the most likely mechanisms for capture of such gases are

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by solar irradiation of accreting material and gravitational capture and dissolution of nebula gases (see Pepin and Porcelli 2002). If the mantle were solid during core segregation, then the noble gases present could only have been derived from irradiated solid material. While the irradiation models require further development, it may be possible to obtain a sufficient quantity to account for the atmospheric inventory (Sasaki 1991; Podosek et al. 2000). Porcelli and Halliday (2001) used a reference value of 6×1010 atoms 3He/g for this process. If the mantle were molten at the time of core formation, noble gases provided by accreting solid materials would have been lost to the atmosphere. The model for providing deep mantle noble gases by dissolution of a gravitationally captured atmosphere then must be invoked to provide noble gases within the Earth (Mizuno et al. 1980; Harper and Jacobsen 1996; Porcelli et al. 2001). In this scenario, the Earth reaches much of its present mass prior to dissipation of the solar nebula, and so gravitationally captures a dense atmosphere. Due to accretional energy, temperatures sufficiently high to melt a large portion of the Earth are reached. Noble gases of solar composition are dissolved from the atmosphere and convected into the deep mantle. This mechanism may be able to provide ~7×1012 atoms 3He/g in the mantle (Mizuno et al. 1980; Woolum et al. 2000; Porcelli and Halliday 2001) or possibly more. Concentrations in the core. If He is partitioned from solid silicates to liquid metal percolating through the deep mantle, initial concentrations must have been provided by irradiated accreting material. The core represents ~32% of the bulk mass of the Earth; if this was removed from the mantle by a single batch process, then He

CFe = HeiCSS / [0.32+(1-0.32)/D(He)Fe/LS]

(14) Hei

10

3

If the concentration of the solid mantle was initially CM0 = 6 × 10 atoms He/g and D(He)Fe/SS = 1 × 10-2, then the resulting core concentration is HeCFe = 1 × 109 atoms 3 He/g. If He was provided by equilibration between liquid metal and a magma ocean at lower pressures, the initial mantle concentration was provided by a hot dense atmosphere. In this case, the relevant partition coefficient of D(He)Fe/LS is lower than that for equilibration with solid silicate, but there may have been greater concentrations of noble gases available from the overlying atmosphere. Then, He

CFe = D(He)Fe/LS HeC LS -4

HeO

(15) 12

If D(He)Fe/LS = 10 and CM0 = 7 × 10 3 He/g, similar to the value calculated above.

3

atoms He/g, then

He

8

CFe = 7 × 10 atoms

Transport back into the mantle. Various mechanisms can be considered for the transport of noble gases back into the mantle. 1) Expulsion of noble gases from the outer core as the solubility limits are exceeded due to core crystallization. However, the saturation limits of both He and Ne, which would be present in the core at a ratio comparable to the solar ratio of 4He/20Ne = 820 (Pepin et al. 1999), are unlikely to be reached together. Therefore, He and Ne that would be provided from the core this way are likely to be highly fractionated. The evidence that mantle He/Ne ratios are not dramatically fractionated from the solar ratio appears to discount this as a viable mechanism. 2) Diffusion across the core-mantle boundary. However, there are no appropriate experimental diffusion coefficient data available to calculate the flux of He across the mantle boundary layer. Large fractionations between the noble gases also may be expected, although there is no data available to quantify this. Therefore, this possibility cannot be evaluated.

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3) Partitioning into overlying partial melts. The ultra-low velocity zone at the coremantle boundary may reflect the presence of mantle melt (see Garnero 2000; Ohtani and Maeda 2001), and partitioning from the core back into the overlying mantle may occur if conditions are favorable. However, a flux cannot be easily calculated without constraints on partition coefficients and the volume and residence time of melts at the core-mantle boundary. 4) Bulk transfer of core material into the mantle (e.g., Macpherson et al. 1998). This provides the most straightforward method for transporting He. Mass transfer may occur into the mantle by exsolution of oxides from the core (Walker 2000), although the noble gas flux cannot be calculated without knowledge of the relevant partitioning behavior during exsolution. For simple entrainment of material from the ÝC = He F HeCFe . Using the OIB 3He flux of HeF = 5 × 1024 core, the mass flux is M 3 He/yr and a core He concentration of HeCFe = (0.7-1) × 109 atoms 3He/g, then M& C = (5-7) × 1015 g/yr is obtained. This is equivalent to (2-3) × 1025g over 4.5 Ga, or ~0.6-0.8% of the upper mantle. Is the core a plausible reservoir? These calculations indicate that, using presently available constraints, the core remains a plausible source of the 3He found in OIB. Measured high 3He/4He ratios would then be the result of mixtures of He from the core that has a solar nebula 3He/4He ratio with radiogenic 4He from the mantle. However, the implications for other geochemical parameters also must be considered. Other elements that would be affected by bulk transfer of core material into the mantle include the Pt-group elements and volatiles that may be relatively abundant in the core, such as H and C. While some limits on the amount of transferred core material are provided by these (Porcelli and Halliday 2001), it should be noted that the amount needed to sustain the 3He flux is still very uncertain, and it is possible that upward revisions in the amount of noble gases within the early Earth, or in the metal/silicate partition coefficients, could reduce the required flux of core material by an order of magnitude or more.

Note that storage of He in the core remains only one component of a noble gas model that can describe the range of noble gas observations. The core has only been evaluated as a possible storage of 3He. The incorporation in the core of other noble gases, and their relative fractionations, cannot be clearly evaluated without more data. Also, the distribution of radiogenic nuclides such as 40Ar, 129Xe, and 136Xe that are produced within the mantle must be explained with a model that fully describes the mantle reservoirs. While these issues may be tractable, a comprehensive model that incorporates a core reservoir remains to be formulated. It should be emphasized that the core does not completely explain the distribution of He isotopes, because the issue of the 4He-heat imbalance is not addressed at all by this model. Subduction of meteoritic He

Anderson (1993) has proposed that subduction of interplanetary dust material in seafloor sediments is the source of mantle 3He. Such material has high concentrations of implanted solar wind He due to its small grain size and so high relative surface area. However, the amount of material being deposited in pelagic sediments is presently about 104 times too low to account for the flux of 3He released at mid-ocean ridges. While the flux of 3He from OIB may be up to 102 times lower, only a small fraction of the subducted sediments can be assumed to be supplying OIB. Therefore, it was suggested that the 3He supply to the Earth was much higher in the past. However, it has been argued that this supply has been insufficient throughout Earth history (Trull 1994), although this may still be open to debate. It should be noted that any higher past fluxes must have been composed of material that is sufficiently small to have avoided heating and degassing during atmospheric entry and surface impact (see Flynn 2001). Further, laboratory

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measurements indicate that He is lost from dust particles by diffusion at low temperatures (Hiyagon 1994) and so will be lost from downgoing sediments during subduction. Allègre et al. (1993) pointed out that the He/Ne ratio of meteoritic dust is much lower than that of the mantle, and so while the Ne in the mantle could be supplied by this mechanism, the 3He cannot. However, the flux of extraterrestrial material to the Earth is also too low to supply the necessary Ne (Stuart 1994). Thus, in the absence of either direct evidence or plausible arguments for sufficiently high past fluxes of material of the required composition, this mechanism is not considered to be a possible dominant source of mantle noble gases (Farley and Neroda 1998). CONCLUSIONS

A fundamental issue that all mantle models address is how to maintain separate reservoirs to account for observed mantle noble gas isotopic variations. None have been able to accommodate all noble gas constraints as well as geophysical observations. However, considerable progress has been made in understanding the processes that may be involved in the distribution, transport, and evolution of mantle noble gases. Overall assessment of the models Models with a boundary to convection at 670 km. 1. The limited interaction mantle models that involve extraction of the atmosphere from the upper mantle do not appear to be viable due to the lack of a simple relationship between atmospheric and MORB Ne and Xe. The models also are opposed to geophysical arguments for whole mantle convection. 2. The steady state upper mantle is compatible with much of the presently available noble gas constraints. However, the present formulation, with a mantle boundary at a depth of 670 km, stands contrary to geophysical evidence for substantial mass fluxes across this depth. The mechanism for interactions between mantle reservoirs is promising, however, and might be adapted to other configurations, although not necessarily without time-dependent distributions. 3. The idea of a break up of previously layered mantle reconciles the geochemical characteristics that can be explained by long-term evolution in a layered mantle (according to one of the layered mantle models) with geophysical observations regarding recent mantle behavior. However, numerical modeling of the thermal state of the mantle does not support this. Models with noble gas reservoirs in other mantle layers. 4. The lower boundary layer model avoids some of the geophysical problems of larger convective layers, but must resolve how such a layer is initially stabilized with gasrich material. Also, it must be combined with other mechanisms to account for the mantle evolution of all the noble gases. However, the model illustrates how a deep boundary may be important. 5. The upper mantle boundary layer model is not supported by any compelling observations, and its consideration emphasizes the need to store He with high 3 He/4He ratios, as well as Xe isotopes, in the deep mantle. 6. A deeper (>670 km) mantle layer provides a potential basis for a comprehensive noble gas model. However, there are geophysical difficulties with the thermal stability of such a layer and the lack of tomographic evidence for such layering.

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Models with heterogeneities within a whole mantle convection regime. 7. Dispersed domains of mantle heterogeneities, comprising a large proportion of the mantle but preferentially sampled at OIB, could provide a basis for a comprehensive noble gas model. However, this also needs further geophysical investigation into their long-term survival, particularly if they contain high concentrations of heatproducing elements. The mechanisms discussed for interaction between reservoirs in layered mantle models could be applied here. 8. The production of domains that are depleted but have high He/U ratios may also explain some He isotopic variations, although these domains do not appear to have enough 3He to impart the necessary isotopic signatures to OIB. This model also requires further development to include the other rare gases. Models with extra-mantle noble gas reservoirs. 9. The core remains a possible source of isotopically distinctive He. However, further work is required, especially on noble gas partition coefficients into core material, both to support this and to formulate the evolution of other noble gases within the mantle.

Overall, the main obstacle to a comprehensive description of mantle noble gas evolution is in finding a configuration for distinct mantle domains that is compatible with geophysical observations. In this regard, a major difficulty is the heat-helium imbalance. Models involving mantle stratification can account for this, but have been discounted on other geophysical grounds. Overall, no other adequate explanation for this imbalance has been proposed. More widespread identification of this as a major unanswered question for the distribution of U and the transport of both He and heat may provoke more geophysical investigation into possible solutions. Another problem requiring resolution is the nature of the high 3He/4He OIB source region. Most models equate this with undepleted, undegassed mantle, although some models invoke depletion mechanisms. However, none of these have matched the end member components seen in OIB lithophile isotope correlations. It remains to be demonstrated that a primitive component is present and so can dominate the He and Ne isotope signatures in OIB. Some persistent misconceptions

As understanding of terrestrial noble gas geochemistry has evolved, various erroneous conclusions have persisted both within the field and in related fields. 1) The very radiogenic Ar and Xe isotope ratios of the upper mantle demand early degassing of the mantle. This is a model-dependent conclusion based on the assumption that upper mantle noble gases are residual from atmosphere degassing. However, Xe isotope systematics precludes such a relationship (Ozima et al. 1985). Since other models, while not necessarily correct, can account for the observed Xe isotope variations, it is clear that the isotopic evidence can be interpreted in various ways. Nonetheless, early transfer of volatiles to the atmosphere probably did occur and was caused by impact degassing. 2) Subduction of heavy noble gases must be very limited. This has been based on models of the isolation of the upper mantle or arguments about preservation of nonatmospheric 129Xe/130Xe ratios in the mantle. In fact, upper mantle nonradiogenic Xe isotopes could be dominated by subducted Xe and admixed with very radiogenic Xe, and some models explicitly incorporate subducted Xe fluxes. Until more is conclusively known about Ar and Xe isotopic variations in the mantle, subduction must be considered a potentially important process.

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3) The processes that have dominantly formed the atmosphere continue at present. Although primordial noble gases continue to degas, their isotopic compositions do not match those of the atmosphere and limit their contribution to a small fraction of the present atmospheric inventory. Volatile species continue to be added to the atmosphere, but the dominant inputs occurred earlier. The formation of the atmosphere cannot be easily extrapolated from the present, but must be understood by quantitative modeling of past processes. Important parameters that are still unknown

There are, of course, many questions regarding terrestrial noble gases that remain to be explored. Some of the issues that are critical to making advances in global models of noble gas behavior are: 1) Subduction. In order to quantify the subduction flux of noble gases, including Ne, further constraints are required on the incorporation of noble gases in subducting materials, and their behavior during melting. Are there phases that can carry these elements to greater depths? Are noble gases added to the mantle wedge in metasomatizing fluids or melts? 2) Partitioning into the core. If He with high 3He/4He ratios can be stored in the core, one of the reasons for requiring a separate reservoir in the mantle is removed. More partition coefficient data at different pressures is required, along with further understanding of the conditions of core formation. Progress in geophysics, mineral physics, and petrology is necessary to determine the mechanisms for reintroduction of noble gases to the mantle. 3) Silicate partition coefficients. While noble gases are clearly highly incompatible, quantitative partitioning data is required to consider the evolution of melt source regions, particularly the ratios of noble gases to parent K, U and Th. 4) Heavy noble gas characteristics in OIB. Variations in Ar and Xe isotopic compositions are not yet clearly documented. Constraining the compositions associated with high 3He/4He ratios is clearly critical for mantle models. 5) Mantle Ar and Xe nonradiogenic isotope compositions. It is still an open question whether there are solar or fractionated heavy noble gases in the mantle. 6) The origin of noble gases. The initial sources and trapping mechanisms of noble gases affected the initial character and distribution of noble gases, and are still poorly understood. These provide the starting conditions for mantle evolution models. 7) Heat and He. A major hurdle in reconciling mantle convection patterns with noble gas distributions is uncovering the mechanism responsible for separating heat and 4 He. 8) The connections between the distributions and evolution of noble gases and major volatiles have not been fully explored. This requires determining how initial distributions may have been related. Subduction rates also likely play a large role in major volatiles. 9) Mantle convection. Geophysical debates regarding the feasibility of maintaining chemical distinct domains or layers in the mantle need resolution. 10) Noble gas concentrations in different mantle reservoirs. Determining the noble gas concentrations of different mantle source regions has not been possible through direct measurement of volcanic materials. Understanding and correcting for magmatic degassing and secondary gas loss processes remains a challenge. While many aspects of mantle noble gas geochemistry and atmospheric evolution

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remain to be explored, further progress will also be made as interpretations respond to advances in related fields such as geophysics, early solar system evolution, and trace element geochemistry. Conversely, noble gases will continue to provide strong constraints on planetary evolution as more encompassing theories are assembled. ACKNOWLEDGMENTS

Reviews by Don Anderson, Geoff Davies, Darryl Harrison, David Hilton, Manuel Moriera, Mario Trieloff, and Rainer Wieler, and a particularly thorough review by Eleanor Dixon, have greatly improved this paper. This work was supported by the ETH and the Swiss National Science Foundation. REFERENCES Abbott DA, Burgess L, Longhi J, Smith WHF (1993) An empirical thermal history of the Earth's upper mantle. J Geophys Res 99:13835-13850 Ahrens TJ (1993) Impact erosion of terrestrial planetary atmospheres. Ann Rev Earth Planet Sci 21: 525-555 Ahrens TJ, O'Keefe JD, Lange MA (1989). Formation of atmospheres during accretion of the terrestrial planets. In Origin and evolution of planetary and satellite atmospheres. In Atreya SK, Pollack JB, Matthews MS (eds) University of Arizona Press, Tucson, p 328-385 Albarède (1998) Time-dependent models of U-Th-He and K-Ar evolution and the layering of mantle convection. Chem Geol 145:413-429 Allard P (1992) Global emissions of helium by subaerial volcanism. Geophys Res Lett 19:1479-1481 Allègre CJ (1997) Limitation on the mass exchange between the upper and lower mantle: the evolving convection regime of the Earth. Earth Planet Sci Lett 150:1-6 Allègre CJ, Staudacher T, Sarda P, Kurz M (1983) Constraints on evolution of Earth's mantle from rare gas systematics. Nature 303:762-766 Allègre CJ, Staudacher T, Sarda P (1986) Rare gas systematics: formation of the atmosphere, evolution and structure of the Earth's mantle. Earth Planet Sci Lett 87:127-150 Allègre CJ, Sarda P, Staudacher T (1993) Speculations about the cosmic origin of He and Ne in the interior of the Earth. Earth Planet Sci Lett 117:229-233 Allègre CJ, Moreira M, Staudacher T (1995a) 4He/3He dispersion and mantle convection. Geophys Res Lett 22:2325-2328 Allègre CJ, Poirier J-P, Humler E, Hofmann AW (1995b) The chemical composition of the Earth. Earth Planet Sci Lett 134:515-526 Allègre CJ, Hofmann AW, O'Nions RK (1996) The argon constraints on mantle structures. Geophys Res Lett 23:3555-3557 Althaus T, Niedermann S, Erzinger J (2000) Noble gas studies of fluids and gas exhalations in the East Carpathians, Romania. Chem Erde (Geochem) 60:189-207 Anders E, Grevesse N (1989) Abundances of the elements: meteoritic and solar. Geochim Cosmochim Acta 53:197-214 Anderson DL (1993) Helium-3 from the mantle—primordial signal or cosmic dust? Science 261:170-176 Anderson DL (1998a) The helium paradoxes. Proc Natl Acad Sci USA 95:4822-4827 Anderson DL (1998b) A model to explain the various paradoxes associated with mantle noble gas geochemistry. Proc Natl Acad Sci USA 95:9087-9092 Anderson DL (2000a) The statistics of helium isotopes along the global spreading ridge system and the central limit system. Geophys Res Lett 27:2401-2404 Anderson DL (2000b) The statistics and distribution of helium in the mantle. Intl Geol Rev 42:289-311 Anderson DL (2001) A statistical test of the two reservoir model for helium isotopes. Earth Planet Sci Lett 193:77-82 Armstrong RL (1968) A model for the evolution of strontium and lead isotopes in a dynamic Earth. Rev Geophys 6:175-199 Armstrong RL (1991) The persistent myth of crustal growth. Austral J Earth Sci 38:613-630 Azbel YA, Tolstikhin IN (1993) Accretion and early degassing of the earth—Constraints from Pu-U-I-Xe isotopic systematics. Meteoritics 28:609-621. Bach W, Niedermann S (1998) Atmospheric noble gases in volcanic glasses from the southern Lau Basin: origin from the subducting slab? Earth Planet Sci Lett 160:297-309 Ballentine CJ, Barfod DN (2000) The origin of air-like noble gases in MORB and OIB. Earth Planet Sci Lett 180:39-48

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Production, Release and Transport of Noble Gases in the Continental Crust Chris J. Ballentine Department of Earth Sciences The University of Manchester Manchester, M13 9PL, United Kingdom [email protected]

Pete G. Burnard Division of Geological and Planetary Sciences MS 100-23, California Institute of Technology Pasadena, California 91125 INTRODUCTION Noble gases within the crust originate from three main sources: the atmosphere, introduced into the crust dissolved in groundwater; the mantle, in regions of magmatic activity; and those produced in the crust by the result of radioactive decay processes. The continental crust contains approximately 40% of the terrestrial radioelements (Rudnick and Fountain 1995) that produce noble gases and, after the mantle and the atmosphere, forms the third major terrestrial noble gas reservoir (neglecting the core). In addition to these sources, contributions from interplanetary dust particles (IDP), cosmic ray interaction with the crustal surface and anthropogenic noble gases can in some cases be a significant source of noble gases in crustal materials. The use of noble gases to understand the role of fluids in different geological settings relies on their low natural abundance and chemical inertness. The low abundance of noble gases in crustal systems and their distinct isotopic character means that contributions from these different sources can often be resolved and quantified. With this, information is gained about the source of associated fluids, the environment from which they originated the physical manner in which they have been transported to the sampling site and the different phases that may have interacted within the crustal fluid system. This is only possible, however, with a detailed understanding of the processes that control the concentration and isotopic composition of the noble gases in different crustal environments. The first part of this chapter deals with the three different mechanisms of noble gas production within the crust—radiogenic, nucleogenic, fissiogenic (Fig. 1). We show how production ratios are affected not only by the source region radioelement concentration, but in the case of nucleogenic reactions, also by the spatial distribution and concentration of the target elements. For completeness we consider cosmogenic noble gas production rates and Interplanetary Dust Particle (IDP) accumulation. We consider the total crustal budget and quantify how much of these differently sourced noble gases have, or can, contribute to the crustal fluid system. We then discuss how release from the minerals in which they are produced occurs by recoil, diffusion, fracturing and mineral alteration. The character of the mineral in which they are produced, the release process, the thermal regime and the ability of the surrounding fluid regime to transport the released noble gases, all play a role in fractionating the crustal noble gas elemental pattern and in determining their respective flux from the deeper crust into shallow systems. We detail how and where magmatic noble gases are introduced into the crust and discuss the relationship of 4He and 3He/4He with heat flow.

1529-6466/02/0047-0012$05.00

DOI:10.2138/rmg.2002.47.12

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Figure 1. Schematic diagram of 238U decay to 206Pb, showing fission and α decay. Fission results in the direct formation of heavy Kr and Xe isotopes and neutrons. Most α particles stabilise to become 4 He, although a small portion react directly with local (0-40 μm) light nuclei in (α,n) reactions to create nucleogenic noble gases such as 21Ne in the reaction 18O(α,n)21Ne. Less than 20% of the subsurface neutron flux is produced directly by fission (e.g., Yatsevich et al. 1997). The remainder is dominated by production from the (α,n) reactions. The neutrons in turn also react with nuclei on a 0.1-1 m length scale. Typical reactions include the (n,α) route producing, for example, 3He in the reaction 6Li(n,α)3H (β-) → 3He. The nuclei available for reaction in the 0-40 μm range of an αemitter determine the production rate of (α,n)-derived noble gases. The greater length-scale of the neutron penetration distance means that noble gases produced by the (n,α) route are not sensitive to mineral-scale elemental heterogeneity.

RADIOGENIC, NUCLEOGENIC, AND FISSIOGENIC NOBLE GASES The subsurface neutron flux and reaction probability Neutron interaction with atomic nuclei is directly linked to the production of several noble gas isotope species in the crust. Before considering the production of crustal noble gases it is essential to first understand the factors controlling the subsurface neutron flux. There are three main types of reaction that produce neutrons in the crust: Cosmic ray interactions; spontaneous fission; and alpha particle interaction with light nuclei. Cosmic ray interaction is only important within the top few meters of the crust (Niedermann 2002, this volume) and we neglect this source of neutrons here. Neutrons are produced by (α,n) reactions within 0-40 μm of 235,238U and together with a small contribution from the spontaneous fission of 238U to give: n = (α,n)U + (α,n)Th + (sf,n)U

232

Th, (1)

where n is the total number of neutrons and (α,n)U, (α,n)Th, (sf,n)U are the neutrons produced by 235,238U, 232Th (α,n) reactions, and spontaneous fission of 238U respectively. Spontaneous fission neutron contributions from 232Th and 235U are negligible (Andrews and Kay 1982; Morrison and Pine 1955). The spontaneous fission neutron yield of 238U is 2.2±0.3 neutrons per fission (Morrison and Pine 1955), and the decay constant for spontaneous 238U fission is 8.6 × 10-17 yr-1 (Eikenberg et al. 1993). This compares with the 232Th neutron yield of 2.5 neutrons per fission with a fission decay constant estimated

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to be 8 MeV (elements with Z > 16) have a Coulomb barrier above the natural upper limit and are not therefore a significant source of subsurface neutrons. For an m1(α,n)m2 reaction Q = [(m1 + mα) – (m2 + mn)] c2

(5)

where m1 , mα , m2 and mn are the masses of m1, the α-particle, m2 and neutron respectively and c the speed of light. The reaction threshold for an endothermic reaction, Eth, is given by: Eth = –[(m1 + mα) / m1] Q

(6)

and is the minimum kinetic energy an α-particle must have in order for the reaction to be energetically favorable. Even though the Coulomb barrier is reached, 16O, 28Si, which constitute some 75% of the Earth’s crust by weight, as well as 40Ca (also excluded by the Coulomb barrier) and 24Mg do not therefore significantly participate in natural (α,n) reactions. As demonstrated above, the reaction cross section is dependant on the energy of the reacting α-particle. To calculate the neutron production of a compound, it is necessary to know the neutron yield and the mass stopping ability of each element, the latter of which is also dependant on the α energy. Experimentally derived neutron yields exist for the range of naturally occurring α-particle energies (e.g., Feige et al. 1968; Jacobs and Liskien 1983; West and Sherwood 1982) and can be fitted to a polynomial to allow calculation of the yield at any alpha energy. From the neutron yield of the constituent elements, Yi, the neutron yield of a particular compound, Yc, can be calculated from Yc = Σ Wi Sim YI m

(7)

where Wi and Si are the mass fraction and mass stopping energy for alpha particles of element i. With the data supplied by Feige (1968), and the assumption that neutron

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contributions from O and Ca can be neglected an empirical U and Th derived neutron flux for an homogenous media can be calculated following: (α,n)U = 0.01 [U] {13.8[Na] + 5.4[Mg] + 5.0[Al] + 1.31[Si] + 2.0[C]} g-1 yr-1 -1

(8) -1

(α,n)Th = 0.01 [Th] {6.0[Na] + 2.45[Mg] + 2.55[Al] + 0.56[Si] + 0.83[C]} g yr (9) where [U] and [Th] are the respective concentration of U and Th in ppm whereas the concentration terms for Na, Mg, Al, Si and C are their percentage concentrations in the rock. The total subsurface neutron density, N, can be derived by combining Equations (1), (3), (8) and (9) to give N(neutrons g-1 yr-1) = 0.01[U] {13.8[Na] + 5.4[Mg] + 5.0[Al] + 1.31[Si] + 2.0[C]} + 0.01[Th] {6.0[Na] + 2.45[Mg] + 2.55[Al] + 0.56[Si] + 0.83[C]}+ 0.4788 [U] (10) This is similar to the equation presented by Andrews (1985) but differs in neglecting Ca (see earlier discussion) instead of C as a significant neutron source.

Figure 2. Martel et al. (1990) show the energy dependence of neutron production from (α,n) reactions for selected mineral and rock compositions using the data of Gorshkov et al. (1962). Neutron fluxes calculated assuming element homogeneity, would significantly overestimate the flux if the radioelements are concentrated in accessory phases such as uraninite. [Used by permission of Elsevier, from Martel et al (1990), Chemical Geology, Vol. 88, Fig. 5, p. 215.]

Using the light element neutron yield from Feige (1968) and the stopping energy distribution from Ziegler (1977), Martel et al. (1990) have calculated the α−energydependent neutron yield for selected mineral and rock compositions (Fig. 2). The results show the dependence of U and Th siting on the neutron flux produced in a rock, also showing that neutron production probability falls to zero at energies of 1% of the total capture probability for at least one rock type have not been tabulated. Rock composition used is from Parker (1967). *Average crust values from Rudnick and Fountain (1995); http://earthref.org/GERM/

The calculation of noble gas production rates involving fast neutron interactions is more complex, and requires a detailed calculation of the energy spectrum of the produced neutrons, and the rate and probability with which they thermalize. For example, the reactions 24Mg(n,α)21Ne and 25Mg(n,α)22Ne both have reaction energy thresholds in the range 2-4 MeV. Rison (1980) in calculating the Ne isotope yields from these reactions, estimated the neutron spectrum from the element (α,n) reactions as a δ function at the energy (Eα + Q), where Eα is the energy of the impinging alpha particle and Q is the reaction energy release. The number of neutrons was estimated from experimentally determined yields of (α,n) reactions. Yatsevich and Honda (1997) note that this approach does not take into account α-particle energy loss before interaction and ignored reaction channels with the excited states of residual nuclei. Consequently the calculated Ne isotope yields are significantly exaggerated. Yatsevich and Honda (1997) rectify this by estimating the energy spectrum of neutrons from the thick-target angle-integrated spectra of neutrons emitted in (α,n) reactions on light elements measured by Jacobs and Liskien

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487

(1983). This data is for the energy interval 4.0-5.5 MeV, and has been extended by fitting the data to a constant form function and incremental extrapolation to 9.0MeV to enable interpolation over the different natural α-energy intervals. Combining this approach with a neutron transport model incorporating both elastic and inelastic scattering plus residual nuclei scattering results in a neutron spectrum with a significantly lower average energy than the earlier work of Rison (1980) (Fig. 4). This in turn results in 21Ne and 22Ne yields by the 24Mg and 25Mg reactions calculated by Yatsevich and Honda (1997) being approximately five times lower than the earlier values of Rison (1980).

Figure 4. Yatsevich and Honda (1997) calculate the energy spectrum of neutrons produced from U and Th (α,n) reactions in an homogenous mantle over 4.5 Ga. They calculate a lower energy spread than the simplified calculation of Rison (1980) because of the inclusion of the effect of elastic and, to a lesser extent, inelastic scattering of the neutrons. Calculated yields for fast neutron reactions, such as with 24,25Mg that have a 2-4 MeV threshold, are therefore much lower with the revised neutron energy spectrum. [Used by permission of the American Geophysical Union, from Yatsevich and Honda (1997), Journal of Geophysical Research, Vol. 104, Fig. 2, p. 10294.]

Helium The present day 4He production in the crust is dominated by the α-decay of the U and 232Th decay chains, and is therefore directly proportional to the concentration of these radioelements in the crust. Although some α-particles are produced by a variety of crustal nuclear reactions, these sources of 4He production are many orders of magnitude smaller and can be neglected; similarly α-particles consumed by (α,n) reactions within the crust are a very small proportion of the α-particles produced. The 4 He produced from each radio-isotope, R, expressed as a function of present day concentration in the rock is given by 235,238

4

He atoms g-1 yr-1 = Xr [R] (NA/Ar) × 10-6 (eλt - 1) × yieldr

(12)

where Xr = fractional natural abundance of isotope R, NA = Avogadro’s number (6.023 × 1023), Ar = molar mass of R (g), λr = decay constant of R (yr-1), yield = number of α particles emitted in the complete decay chain, [R] = the concentration of R in ppm and t = age (yr). For 238U, 235U and 232Th, λ238 = 1.55 × 10-10, λ235 = 9.85 × 10-10, λ232 = 4.95 × 10-11 (Steiger and Jäger 1977), with an α yield in each decay chain of 8, 7 and 6 and a natural abundance of X238 = 0.9928, X235 = 0.0072 and X232 = 1.000 respectively, the number of atoms of 4He produced in 1 gram of rock per year becomes 4 3

He atoms g-1 yr-1 = (3.115 × 106 + 1.272 × 105) [U] + 7.710 × 105 [Th]

(13) 6

He production within the crust is dominated by thermal neutron capture by Li in the reaction 6Li(n,α)3H (β-)3He. Other reactions are detailed in Mamyrin and Tolstikhin (1984, p. 101), but for average crustal compositions, these have a 3He yield at least four

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orders of magnitude lower than the 6Li route and are neglected here. The rate of 3He production therefore, is directly proportional to the subsurface thermal neutron density (Eqn. 10), the number of neutrons that reach thermal energy level (Pth = 0.8) and the relative capture cross section of Li (Eqn. 11). Taking the composition of a homogenous average upper crust (Table 2) as a working example, we obtain: 3

He = 0.8 × {0.01[2.8] (13.8[2.89] + 5.4[1.33] + 5[8.08] +1.31[30.9] + 2[0.324]) + 0.01[10.7] (6.0[2.89] + 2.45[1.33] + 2.55[8.08] + 0.56[30.9] + 0.83[0.324]) + 0.4788 [2.8]} × 2.05 × 10-4 / 9.79 × 10-3 = 1.87 × 10-1 atoms g-1 yr-1

Table 2. Composition of average crust.* Weight fraction element Element Li B C O Na Mg Al Si Cl K Ca Ti Cr Mn Fe Co Ni Sm Gd U Th

Atomic Mass 6.94 10.81 12 16 22.99 24.31 26.92 28.09 35.45 39.1 40.078 47.87 52 54.94 55.845 58.93 58.69 150.36 157.25 238.03 232.04

Lower Crust 6.00E-06 3.71E-06 5.88E-04 4.69E-01 1.93E-02 4.28E-02 8.78E-02 2.44E-01 5.28E-05 4.98E-03 6.72E-02 4.79E-03 2.15E-04 7.74E-04 5.88E-02 3.80E-05 8.80E-05 2.80E-06 3.10E-06 2.00E-07 1.20E-06

Middle Crust 7.00E-06 3.22E-06 7.18E-03 4.87E-01 2.37E-02 2.05E-02 8.19E-02 2.83E-01 5.28E-05 1.67E-02 3.64E-02 4.20E-03 8.30E-05 7.74E-04 4.48E-02 2.50E-05 3.30E-05 4.40E-06 4.00E-06 1.60E-06 6.10E-06

Upper Crust 2.00E-05 9.56E-06 3.24E-03 4.75E-01 2.89E-02 1.33E-02 8.04E-02 3.09E-01 2.32E-04 2.82E-02 3.00E-02 3.00E-03 3.50E-05 6.20E-04 3.15E-02 1.00E-05 2.00E-05 4.50E-06 3.80E-06 2.80E-06 1.07E-05

*Rudnick and Fountain (1995); Parker (1967) Wederpohl 1995; Gao et al. 1997; http://earthref.org/GERM/

Table 3. Neutron production, Li capture probability, and 3He/4He in average crust. Lower Crust Pth 0.8 U(α,n) neutrons/g/yr 2.51E-01 Th(α,n) neutrons/g/yr 6.98E-01 9.58E-02 Uf neutrons/g/yr F6Li 6.99E-03 3

He atoms/g/yr He atoms/g/yr

4

3

He/4He

Middle Crust 0.8 1.95E+00 3.42E+00 7.66E-01 8.05E-03

Upper Crust 0.8 3.58E+00 6.25E+00 1.34E+00 2.09E-02

5.84E-03 1.57E+06

3.95E-02 9.89E+06

1.87E-01 1.73E+07

3.71E-09

3.99E-09

1.08E-08

3

(14)

4

The He and He production rate for average lower, middle and upper crust compositions (Table 2) are shown in Table 3. It should be noted that the 3He/4He ratio is independent of [U] for constant U/Th. Perhaps the most interesting observation from these calculations is the low value of 3He/4He, even for the Li enriched upper crust. When we consider that the effect of radioelement heterogeneity (the preferential siting of U and Th in accessory phase minerals) would be to significantly reduce the neutron flux producing the 3He (Figs. 2 and 3), these 3He/4He values must be considered to be upper limits for radiogenic production in average crust. Comparison with measured 3 He/4He in whole rock samples, mineral separates and associated fluids is non-trivial. This is due to several factors: (i) calibration and measurement of 3He/4He ratios 2.8 Ga) by orders of magnitude. In samples where mantle and cosmogenic 3He contributions to sedimentary rocks are discounted, variations from predicted ratios by a factor of three are ascribed to fractional release, with one exception, anhydrite, apparently requiring a fractionation factor of 20 (Tolstikhin et al. 1996). Table 4. Selected measured and calculated 3He/4He in granites.

Sample ID

Location

WR Orthoclase Biotite Plagioclase Quartz Amphibole Zircon 1-WR 2-WR 4-WR 43-WR 27-WR ss1-WR ss1-fldspar ss1-mica ss1-quartz rh11g-wr rh12d-wr1 rh12dwr2 urananite1 Water WR Water

Rapakiwi Rapakiwi Rapakiwi Rapakiwi Rapakiwi Rapakiwi Rapakiwi Ukraine Ukraine Ukraine Tuva Caucasus Carnmenellis, UK Carnmenellis, UK Carnmenellis, UK Carnmenellis, UK Carnmenellis, UK Carnmenellis, UK Carnmenellis, UK Carnmenellis, UK Carnmenellis, UK Stripa, Sweden Stripa, Sweden

Concentration 3He/4He (ppm) Measured Li U Th 36 6 620 10 12 37

6.7 2.5 7.2 5.9 5.1 3.5 990 4.8 9.5 11 1.6 4.5 11.8

35 2.7 4.8 5.3 3.0 12.7 150 44 42 85.4 10 14 15.1

386 16.4 300 10.2

8.4 6.6

11 44.1

33

37 38 26 10 17 353

1.60E-08 4.00E-07 1.20E-07 5.00E-08 1.50E-07 4.50E-08 2.00E-08 1.60E-08 1.90E-08 3.00E-09 8.00E-09 6.00E-08 2.10E-09 2.38E-09 4.62E-09 8.40E-10 2.10E-09 1.54E-09 4.90E-09 9.80E-10 3.50E-08

3

He/4He Calc*

3

He/4He Calc** Ref

1.60E-08

1.60E-08 1.70E-08 1.40E-08 7.00E-09 9.00E-09 8.00E-08

2.00E-08

6.40E-08 6.70E-08 6.70E-08

1.80E-08 1.80E-08 1.80E-08

5.94E-09 5.70E-09

* Assuming element homogeneity WR = Whole rock ** Assuming radioelements concentrated in uraninite grains with φ>100μm

1 1 1 1 1 1 1 2 2 2 2 2 3 3 3 3 3 3 3 3 3,4 5,6 5

1) Gerling et al. (1971) (in Mamyrin and Tolstikhin 1984) 2) Tolstikhin and Drubetskoy (1977) (in Mamyrin and Tolstikhin 1984) 3) Martel et al. (1990) 5) Andrews et al. (1989a) 4) Hilton et al. (1985) 6) Andrews et al. (1989b)

Granite systems provide an example where mantle influences appear to be minimal, and cosmogenic effects can be safely ruled out. This allows the effect of heterogeneity and release effects to be assessed. In Table 4 we have compiled selected 3He/4He data from different granite systems that also have theoretical values for comparison. With the exception of sample 27-WR, all measured whole rock values are either within error of calculated 3He/4He values or significantly lower. Two important observations can be made: i) more than one granite whole rock sample matches both measured and calculated 3 He/4He; and ii) the Stripa granite, which is the only system where the calculated neutron flux is confirmed with a measured neutron flux (Andrews et al. 1986), contains groundwater with 3He/4He indistinguishable from predicted production (Andrews et al. 1989a;b). Although a small data set, an important inference is that many systems display no resolvable fractionation from theoretical values.

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Where fractionation is observed, the ‘nugget’ effect must be first ruled out. This is when a whole rock sample contains an accessory mineral concentration in excess of the average rock. Accessory phases, such as uraninite, contain high concentrations of He with low 3He/4He (Table 4) and will result in a 3He/4He lower than the average bulk rock. In the case of the Carnmenellis granite (Martel et al. 1990) this cause can be ruled out by multiple whole rock analyses giving the same 3He/4He value. If the measured uraninite size of ~50-μm radius is representative of the radioelement distribution, this rock would be expected to produce 3He/4He = 1 × 10-8. Nevertheless, measured whole rock and mineral separate analyses are a factor of three too low to be accounted for by radioelement heterogeneity reduced flux alone (Table 4). An additional factor of three to five is required and is probably release related (Martel et al. 1990). A complementary 3 He-enriched fluid would be expected to be observed. This is indeed seen in the associated groundwater (Table 4). It would appear that the order of magnitude difference between the whole rock and groundwater 3He/4He in the Carnmenellis granite can be ascribed to a combination of radioelement heterogeneity and preferential release of 3He into the surrounding fluid system. If however, the average uraninite grain size is an order of magnitude smaller, all of the difference between whole rock and groundwater 3He/4He has to be ascribed to preferential 3He release into the surrounding groundwater. In summary, the effect of radioelement heterogeneity on 3He/4He in most systems is probably small. There is some evidence from the Carnmenellis system that fractionation of 3He/4He due to preferential 3He release could be as high as a factor of 10, but taking into account the evidence for radioelement element heterogeneity this is more reasonably, at most, a factor of 3 (e.g., Tolstikhin et al. 1996). In large crustal fluid systems reasonably sampling ‘average’ crust, 3He/4He ratios in excess of 1-3 × 10-8 (i.e., more than three times the upper crust value) are due to a resolvable 3He excess from sources external to the crust (e.g., Marty et al 1993). Neon The production of Ne isotopes in the crust is entirely due to nucleogenic routes. Recognized by (Wetherill 1954), the only significant production routes are 17,18 O(α,n)20,21Ne, 19F(α,n)22Na(β+)22Ne, 24,25Mg(n,α)21,22Ne, 23Na(n,α)20Ne, 19F(α,p)22Ne (Yatsevich and Honda 1997). Their rate of production is therefore related to radioelement and target-element concentrations as well as the distribution of the target element with respect to any radioelement heterogeneity. The first investigations of Ne production rates in the crust include the work by Sharif-Zade et al. (1972), Shukolyukov et al. (1973) and Verkhovskiy and Shukolyukov (1976a,b). In addition to (α,n) production, Rison (1980) also investigated (n,α) production rates. This work however, used oxygen (α,n) yields from Feige (1968) which results in a 20,21Ne yield twice as high as more recent measurements of the (α,n) yields by West and Sherwood (1982). More recently it has also been shown by Yatsevich and Honda (1997) that the neutron yields from the (n,α) routes calculated by Rison (1980) are also too high by a factor of five because of an overestimation of the neutron energy spectrum (Fig. 4). Using the revised (α,n) yields from West and Sherwood (1982), Hünemohr (1989) and Yatsevich and Honda (1997) have calculated the Ne (α,n) yields for average crust and mantle materials. Using only the reaction cross sections Leya and Wieler (1999) have also estimated Ne production in average crust. These production rates are summarized in Table 5. Yatsevich and Honda (1997) show that with the updated Mg (n,α) yields, production by this route accounts for 0.6Ra). Data without full He, Ne and Ar abundance and isotopic information are not included. No correlation is observed between 4He/40Ar and 4He/21Ne in this filtered data set. The data distribution approximates a Gaussian distribution with a mean observed crustal 4He/21Ne production rate of 1.71±0.09 × 107. This compares with the theoretical estimates for average crust of between 2.02-2.64 × 107 (Table 5). At present it is unclear why the crustal record preserves a 21Ne production rate 15-35% greater than that the theoretical calculations would predict. Argon 40

Ar production in the crust is dominated by the decay of 40K, and is therefore directly proportional to the K concentration. 40K has a branched decay mode, producing 40 Ca by beta decay and 40Ar by electron capture, with decay constants of λβ =

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4.962 ×10-10 yr-1 and λe = 0.581 × 10-10 yr-1 respectively (Steiger and Jäger 1977). The total decay constant of 40K is therefore given by: λΚ = λβ + λe = 5.543 × 10-10 yr-1

(18)

(It should be noted that this value has recently been re-determined to give λΚ = (5.463±0.054) × 10-10 yr-1, although usage of this newer value is not yet widespread (Begemann et al. 2001; Kelley 2002)). The fraction of 40K that decays to 40Ar is therefore λe/λK and the 40Ar production rate can be expressed as a function of K (in ppm), [K], by: 40

Ar atoms g-1 = XK [K] × 10-6 (NA/AK) ( λe / λΚ ) (eλt-1)

(19)

where XK = fractional natural abundance of 40K = 1.167 × 10-4 , NA = Avogadro’s number (6.023 × 1023), AK = molar mass of K = 39.964g, and t = age (yr). The present day production can therefore be expressed as: 40

Ar atoms g-1 yr-1 = 102.2 [K]

(20)

Combining Equation (20) with Equation (13), the term for He production in the crust, the 4 He/40Ar ratio in the crust is defined by the (U+Th)/K ratio, where: 4

He/40Ar = {(3.115 × 106 + 1.272 × 105) [U] + 7.710 × 105 [Th]} / 102.2 [K] (21)

Using the crustal compositions from Table 2, Equation (21) predicts present day He/40Ar production ratios in the lower, middle and upper crust of 3.09, 5.79 and 6.0 respectively, with a production weighted average (Table 2) of 5.7. In principle measured crustal 4He/40Ar ratios can provide a test of these K/U ratios (Dymond and Hogan 1973), an issue that has recently provided significant controversy (e.g., Albarède 1998). In practice 4He/40Ar ratios are sensitive to fractionation during thermal release from their respective mineral sites (Ballentine et al. 1994; Mamyrin and Tolstikhin 1984), as well as subsequent transport and phase related fractionation processes (Ballentine et al. 1991). We have adopted in the Ne section above the assumption that 4He/40Ar ratios within the range 4 to 6 are relatively unaffected by these process to filter out fractionated samples, but with the implicit assumption that the K/U ratio of the crust is well defined. We discuss later the information about release, transport and phase fractionation available by considering deviations from predicted crustal values. Also see Ballentine et al. (2002). 4

36

Ar production in the crust is small compared to the ambient background of atmosphere-derived 36Ar introduced into the crust dissolved in groundwater and is usually neglected. Although a small amount of muon-induced 36Ar occurs close to the surface, the principle route of production is the β-decay of 36Cl (Fontes et al. 1991; Hünemohr 1989). 36Cl has a half-life of 3.01 × 105 yr and decays to 36Ar with a branching ratio, R, of 0.95. 36Cl is only produced in the crust by the thermal neutron reaction 35 Cl(n,γ)36Cl (Bentley et al. 1986; Fontes et al. 1991). The fraction of thermal neutrons captured by any one element, Fi, is given by Equation (11). It should be noted that the reaction cross section given for Cl in Table 1 is the combined probability for 35Cl and 37 Cl with thermal neutron capture cross sections and relative abundances of 43 and 0.43 barns and 75.77% and 24.23% respectively. These values give F35Cl = 0.02206 for average upper crust. For a system in steady state the number of 36Cl atoms present is given by: 36

Cl atoms g-1 = Pth N F35Cl / λ36

(22)

where Pth is the probability of a neutron reaching thermal energy, N is the neutron density given by Equation (10) and λ36 the decay constant of 36Cl. Once this equilibrium has been attained the production of 36Cl is equal to its rate of decay to 36Ar, giving the following 36 Ar production rate:

Production of Noble Gases in the Continental Crust 36

497

Ar atoms g-1 yr-1 = Pth N R F35Cl

(23)

For average upper crust (Table 2), N = 10.6 neutrons g-1 yr-1 (Table 3), giving a 36Ar production rate of 0.19 atoms g-1 yr-1. This compares with 40Ar production in the upper crust of 2.93 × 106 atoms g-1 yr-1, to give a crustal 40Ar/36Ar production ratio of 1.54 × 107. Fontes et al (1991) discuss how 36Ar rates can be a significant factor in specific Cl-U-Thrich environments and further discuss cosmogenic rates of 36Ar production. 38 Ar excess relative to 36Ar observed in early studies of U- and Th-rich minerals led workers to investigate the 35Cl(α,p)38Ar and 41K(n,α)38Ar production routes of 38Ar (Fleming 1953; Wetherill 1954). Hünemohr (1989) and Eikenberg et al. (1993) in addition investigated the production via 37Cl(n,γ)38Cl(β-)38Ar. 41K(n,α)38Ar is energetically un-favorable and can be discounted. Although energetically favorable, the thermal neutron interaction with 37Cl has a reaction cross-section of only 0.43 barns and a natural abundance of 24.23%. The 38Ar production rate from thermal neutrons is given by: 38 Ar atoms g-1 yr-1 = Pth N F37Cl (24) For average upper crust (Table 2) F37Cl = 6.7 × 10-5 to give a 38Ar production rate of 6 × 10-4 atoms g-1 yr-1. No modeling of the 35Cl(α,p)38Ar production rates have been made, although the ratio of the nuclear cross sections for 35Cl(α,p)38Ar and 19F(α,n)22Ne is predicted from Woosley et al. (1975) to be ~0.2 for α-energies between 4 and 8 MeV (Eikenberg et al. 1993). Both Hünemohr (1989) and Eikenberg et al. (1993) show that 38 Ar/22Ne excesses are correlated with 35Cl/19F ratios in (U+Th)-rich minerals with a gradient consistent with this estimate. The relationship between relative F and Cl αreaction cross sections enables an empirical equation for 38Ar production rates to be derived from Equation (17): 38

Ar =

3.98 ×10 -24 19 [Cl ] ⎛ 0.340 0.139 [Th]⎞ × 0.3465× 0.2423× 0.2 × × [U] + ⎟ ⎜ 37 16 ⎝ 0.479 0.479 3 ⎠ 0.02

= {[Cl](0.76[U] + 0.104[Th]} × 10-25

(25) 38

[Cl], [U] and [Th] are concentrations in ppm and the production rate of Ar in cm3 STP g-1 yr-1. For average upper crust (Table 2) this gives a production rate of 0.002 atoms g-1 yr-1. Inclusion of 37Cl(n,γ) derived 38Ar results in a total 0.0026 atoms g-1 yr-1, to give a 38Ar/36Ar crustal production ratio of ~0.014. This value is calculated assuming an homogenous distribution of elements in the upper crust. Unlike 36Ar, the 38Ar production is dominated by α-particle reactions and the production rate of 38Ar will be sensitive to element heterogeneity and the siting of Cl relative to U+Th in the crust. This is directly analogous to 22Ne production rates and is illustrated in Cl and (U+Th)-rich minerals by observed 38Ar/36Ar production ratios in excess of 14.7 (Eikenberg et al. 1993), some 103 times higher than ratios predicted for an elementally homogenous crust. Although some natural gases show 38Ar/36Ar ratios in excess of air ratios (Ballentine 1991), pointing to average crustal production ratios in excess of calculated average values, no systematic assessment of the 38Ar production rate in the crust exists. Estimates of nucleogenic contributions to the atmosphere based on average crustal production rates will underestimate the nucleogenic 38Ar contribution, while the effect of element heterogeneity on 38Ar/36Ar ratios in the mantle has yet to be assessed. Small amounts of the unstable gases 37Ar and 39Ar are produced in the crust through the thermal neutron reactions 39K(n,p)39Ar and 40Ca(n,α)37Ar (Lehmann et al. 1993; Pearson Jr. et al. 1991). 37Ar and 39Ar have half-lives of 34.95±0.08 days (Renne and Norman 2001) and 269 years respectively. Applications of these tracers are reviewed by Kipfer et al. (2002, this volume).

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Krypton and xenon Fission of 238U provides the dominant mechanism for the production of 83,84,86Kr and 129,131,132,134,136 Xe in the crust today, while production of the shielded isotopes 80,82Kr and 124,126,128,130 Xe can be neglected. Contributions from the spontaneous fission of 232Th and thermal or fast neutron induced fission of 235U, 238U and 232Th, must also be taken into consideration. Although there has been considerable past interest in the determination of the various fission spectra and yields, recent investigation of the U-Xe-Kr dating tool pioneered by Shukolyukov et al. (1974), has resulted in a re-determination of many of these values (Eikenberg et al. 1993; Ragettli et al. 1994; Wieler and Eikenberg 1999) (Table 7). Fission products of 244Pu and the decay products of 129I, both now extinct, contribute important Kr and Xe isotopic components to both the terrestrial mantle and atmosphere (Porcelli and Ballentine 2002). Kr and Xe derived from these extinct radionuclei are not produced in the crust hence do not contribute to the crustal system except where carried in as components of magmatic or atmosphere-derived fluids. It is convenient to consider the production rate of 136Xe by the various routes, and scale the yield of the other Kr and Xe fissiogenic isotopes to this product (Table 7). 136Xe production by spontaneous fission in the crust is directly proportional to the 238U concentration. 238U has a branched decay mode, producing 234Th by α-decay with λα = 1.55 × 10-10 yr-1 (Steiger and Jäger 1977), and spontaneous fission producing amongst other isotopes 136Xesf with λsf ~ 9 × 10-17 yr-1. Because λsf 1, reducing the solubility of i. Although these effects are commonly called ‘salting in’ and ‘salting out’ of the non electrolyte, from Equation (10) it can be seen that these are no more than empirically derived changes in the activity of the dissolved gas. Measurements of noble gas solubilty in water and NaCl solutions have focused on low temperatures (0-40°C) and salinities up to that of seawater (Weiss 1970, 1971; Clever 1979a,b). This data is used in noble gas paleotemperature investigations and is

Tracing Fluid Origin, Transport and Interaction in the Crust

545

Table 2: Noble gas Setchenow coefficients, ki(T) Species

G1

G2

G3

He -10.081 15.1068 4.8127 Ne -11.9556 18.4062 5.5464 Ar -10.6951 16.7513 4.9551 Kr - 9.9787 15.7619 4.6181 Xe T2 -14.5524 22.5255 6.7513 Constants from Smith and Kennedy (1982) to fit ki(T) = G1 + G2/(0.01T) + G3ln(0.01T), where T is temperature in Kelvin and salinity is in units of mol-1(see Eqn. 8)

Figure 4. Setchenow coefficients calculated from Table 2 as a function of T. Only He reaches a minimum in the experimental range, limiting the ability of this data set to be extrapolated to higher temperatures.

tabulated in Kipfer et al. (2002, this volume). To expand the dataset available, specifically to investigate the range of NaCl concentrations and temperatures found in natural brines, Smith and Kennedy (1983) measured the solubilities of the noble gases in 0 to 5.2 molar NaCl solutions between 0 and 65°C. The coefficients and equation used to fit the experimental data are shown in Table 2. The variation of the Setchenow coefficients ki, with temperature T is shown in Figure 4. Over the experimental temperature range, only kHe reaches a minimum. Although the data for the heavier noble gases are fit to the same functional form, it is not possible to assess the error when extrapolating to temperatures higher than 65°C for any gas except He. Although the work of Smith and Kennedy only investigates the effect of NaCl on noble gas solubility, they note that the contribution by individual ions should be additive and in dilute brines it should be possible to estimate the salt effect of multi-electrolyte solutions. While no data exists for Mg++ and Ca++ ions, data for KI solutions show that kAr is independent of the electrolyte species (Ben-Naim and Egel-Thal 1965), suggesting that an NaCl ‘equivalent’ concentration provides a reasonable value from which to calculate the Setchenow coefficient. This relationship has been used in multi-ion mixtures such as seawater and for more concentrated solutions such as the Dead Sea brines (Weiss 1970; Weiss and Price 1989). The effect of non-electrolytes such as other dissolved gases in solution is more difficult to assess due to the lack of empirically derived data. Under near surface conditions, most non-electrolytes are relatively insoluble and will have negligible effect. However, at greater pressures significant amounts of major gas species may be in solution. For example at 190atm and 70°C (hydrostatic pressure and temperature at 2km depth), about 3cm3(STP) of CH4 saturate 1 cm3 of pure water to give a 0.13M CH4 solution (Price 1981). The assumption must be made therefore, that interactions between

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gaseous non-electrolytes must be small and do not significantly affect the activity coefficient of the noble gases. However, it has been shown that, for example, the CH4 saturation concentration in water does significantly decrease when small quantities of CO2 are present (Price 1981). While the data is not available to assess the magnitude of effect these different major species may have on noble gas solubility, this illustrates the potential shortcoming in assuming that no significant non-electrolyte interactions occur in gas-saturated solutions. Noble gas solubility in water and oil Water. The solubility of noble gases in water has received considerable attention from physical chemists investigating molecular models of solution in liquid water. Eley (1939) first considered the process of noble gas solution as consisting of a two step mechanism involving the creation of a cavity in the fluid. Ben-Naim and Egel-Thal (1965) described the thermodynamic behavior of aqueous solutions of noble gases in terms of a two-structure model and discuss the origin of the ‘stabilized structure of water’ by the noble gases, and the ‘degree of crystallinity’ of the water caused by the addition of electrolytes and non electrolytes to the solution. Comprehensive reviews are given by Wilhelm et al. (1977) and Ben-Naim (1980) with further discussion on the solvation structure of water by Stillinger (1980) who models the clumping tendency of strain free polyhedra cages formed on the solution of non polar gases. A detailed understanding of the noble gas behavior in water has resulted in a statistical thermodynamic model for the solubility of noble gases at varying temperatures based on the distributions of molecular populations among different energy levels (Braibant et al. 1994). Early laboratory determinations of noble gas solubility were neither comprehensive nor over large temperature ranges. Benson and Krause (1976) produced the first complete data set for noble gas solubilities in pure water for the temperature range 0-50°C, but as only helium reaches a minimum in this range no extrapolation from this data is possible to higher temperatures. Potter and Clyne (1978) increased the data set by investigating solubilities up to the critical point of water. However, this work was subject to some error, as shown by the subsequent work of Crovetto et al. (1982) and confirmed by Smith (1985) both of whom have fitted their solubility data to curves with a third order power series between 298K and the critical temperature of water. The fit from Crovetto et al. (1982) has been taken here for Ne, Ar, Kr and Xe while the solubility of He relative to Ar has been taken from Smith (1985) to calculate the Henry’s constants for high temperature aqueous systems (Table 3, Fig. 5). Table 3. Henry's constants for noble gases in water Species

He Ne Ar Kr Xe

Ao

-0.00953 -7.259 -9.52 -6.292 -3.902

A1

0.107722 6.95 8.83 5.612 2.439

A2

A3

0.001969 -1.3826 -1.8959 -0.8881 0.3863

-0.043825 0.0538 0.0698 -0.0458 -0.221

Coefficients for Ne, Ar, Kr and Xe from Crovetto et al. (1981) to fit the equation: ln(Ki) = Ao + A1/(0.001T) + A2/(0.001T)2 + A3/(0.001T)3 where Ki is Henry’s constant in GPa. Coefficients for He are from Smith (1985) to fit the equation: ln(FHe) = Ao + A1/(0.001T) + A2/(0.001T)2 + A3/(0.001T)3 where FHe = (XHe/XAr)liquid/(XHe/XAr)gas. X is the mol fraction. T is temperature in Kelvin. Valid temperature range is from 273K to the critical point of water. 1GPa = 9870 atm. For water, Ki(atm) = 55.6 Kim(atm Kg/mol), where Kim is Henry’s coefficient expressed in terms of molality.

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Figure 5. Henry’s coefficients for noble gases in water, calculated from Table 3 following the molality convection, plotted as a function of temperature. The valid temperature range of this data set is 273K to the critical point of water.

Although it is usual to express the Henry’s constant in units of pressure (Eqn. 1), to enable comparison between oil (in which the mole fraction is difficult to calculate) and water systems, we use here the molality convention where: Kmi = ΦiPi/Ci

(11)

Φi is the gas fugacity coefficient, Pi the partial pressure of i in atm, and Ci the number of moles of i in 1000 g of the liquid phase (water). Oil. Solubility studies of the noble gases in crude oil have mostly been limited to empirical approximations as a function of oil density and temperature (e.g., Zanker 1977; ASTM 1985). The most comprehensive study to date remains that of Kharaka and Specht (1988), who have taken two crude oils of different density and experimentally determined the solubility of He, Ne, Ar, Kr and Xe over the temperature range 278 to 373 K. The respective solubilities are fitted to a linear equation in the form Log(Kmi) = A + BT, where A and B are the experimentally determined coefficients, T the temperature in °C and Kmi is the Henry’s constant following the molality convention (Eqn. 11). The coefficients are given in Table 4 and plotted as a function of temperature in Figure 6. Table 4: Henry's constants for noble gases in oil

Species He Ne Ar Kr Xe

Heavy Oil (API=25) A B 3.25 -0.0054 3.322 -0.0063 2.121 -0.0003 1.607 0.0019 1.096 0.0035

Light Oil (API=34) A B 3.008 -0.0037 2.912 -0.0032 2.03 0.001 1.537 0.0014 0.848 0.0052

The solubility constant of the noble gases in oil is dependent on oil density. Coefficients taken from Kharaka and Specht (1987) for two oils fit the equation Log(Kim) = A + BT where the solubility constant Kim follows the molality convention and is in units of atm Kg mol-1. T is temperature in oC. The valid temperature range for the determination is 25-100 oC.

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Figure 6. Henry’s coefficients for noble gases in ‘light’ oil (API gravity = 34o), calculated from Table 4 following the molality convention, plotted as a function of temperature. Noble gas solubility constants in more dense oil are higher (Table 4). The valid temperature range of the experimental determination is 298-373 K, although with no apparent minima, extrapolation to higher temperatures is probably reasonable.

The molality convention is followed because crude oil consists of various mixtures of different molecules that make the determination of the gas mole fraction almost impossible to determine. Unlike the studies of noble gas solubilities in water, there is no comprehensive study that has investigated the noble gas solubility in oil as a function of oil density and temperature. While a linear relationship between solubility based on the two oils measured can be assumed (e.g., Ballentine et al. 1996) and to a first order is supported by the empirical approximations (Zanker 1977; ASTM 1985), this remains a limiting factor in the application of noble gas solubility studies involving an oil phase. PHASE EQUILIBRIUM AND FRACTIONATION OF NOBLE GASES Liquid-gas phase partitioning of noble gases Noble gas partitioning and solubility fractionation between equilibrated subsurface phases was initially studied by Goryunov and Kozlov (1940) and further studied (Zartman et al. 1961; Bosch and Mazor 1988; Zaikowski and Spangler 1990; Ballentine et al. 1991; Ballentine et al. 1996; Hiyagon and Kennedy 1992; Pinti and Marty 1995; Torgersen and Kennedy 1999; Battani et al. 2000). Recent reviews are by O'Nions and Ballentine (1993), Ballentine and O'Nions (1994) and Pinti and Marty (2000). Under equilibrium conditions, the distribution of noble gases between gas and liquid phases is given by Henry’s Law (Eqn. 1). Following Goryunov and Kozlov (1940) and Zartman et al. (1961) and assuming ideal behavior in the gas phase, another form of this equation is derived when the concentration i in the gas phase Cgi is related to the concentration in the liquid phase phase Cli by Cgi = Kid Cli

(12) d

Henry’s constant in this form, Ki , is dimensionless. For a fixed volume of gas and liquid, the number of moles in the gas phase [i]g is related to the total number of moles present [i]T and the liquid to gas volume ratio, Vl/Vg, by [i]g = [i]T (Vl/Vg Kid + 1) –1 Taking the limits as Vl/Vg → 0, then [i]g → [i]T, and as Vl/Vg → , then [i]g → 0.

(13)

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549

It can be seen that Henry’s constants can be dimensionless (Kid, Eqn. 12) expressed in units of pressure, (Ki, Eqn. 2) or take the molaltity format, Kim (Tables 3 and 4). A similar form of Equation (13) for Ki or Kim can be simply derived taking due account of units. For example we consider Henry’s constant, Ki, expressed in units of atm and water as the liquid phase. In this case the mole fraction, xi , can be related to the concentrations in terms of the water density, ρH2O (g/cm3) and the water volume, VH2O (cm3) where xi = 18 ([i]T – [i]g]) (ρH2O VH2O ) –1

(14)

The partial pressure of i, pi, in the gas volume Vg can be expressed in terms of the concentration of i in the gas phase at temperature T (K) assuming that at STP (1 atm, 273 K), 1 mole of gas occupies 22400 cm3 (ideal gas behavior) by pi = [i]g (22400 T) (273 Vg) –1

(15)

Substitution of Equations (14) and (15) into Equation (2) and rearranging gives ⎛ 22400 T ρ H2O VH2O ⎞ [i]g = [i]T ⎜ +1⎟ γi ⎝ 18 × 273 φi Ki Vg ⎠

−1

(16)

Equation (13), but with Henry’s constant in units of molality, Kim (Kg atm/mol), can be similarly derived to give ⎛ ⎞ 22400 T ρ l Vl [i]g = [i]T ⎜ +1⎟ γi m ⎝ 1000 × 273 φi Ki Vg ⎠

−1

(17)

where ρl is the density of liquid l (g/cm3) at the system pressure and temperature, T. This form enables the partitioning of species i between gas and liquid phases to be calculated for any water/gas equilibrium, taking due account of any non-ideal behavior in species i in either the gas or liquid phases. In the simplest case it is possible for example to calculate the volume of gas with which groundwater has equilibrated from one noble gas concentration determination in the water phase, and an estimate of the conditions under which equilibration took place. Similarly, the volume of water with which a gas phase has equilibrated can be quantified from the determination of the concentration of one noble gas in the gas phase. This is discussed in more detail in the following sections. Liquid-liquid phase partitioning of noble gases The partition coefficient Di between two phases for any species i is defined as the equilibrium concentration of i in one phase relative to the other. For two separate liquid phases (we consider here oil and water) the relationship between Di and the Henry’s constants is given as Di = Cioil/CiH2O = Kim(H2O)/Kim(oil)

(18)

where Cioil, CiH2O, Kim(H2O) and Kim(oil) are the number of moles of i in 1000 g of oil, 1000 g of water and the Henry’s constants (atm Kg/mol) of i in water and oil respectively. Substituting Cioil = [i]oil/1000xVoilρoil, CiH2O = [i]H2O/1000 × VH2OρH2O, and [i]T = [i]oil+[i]H2O into Equation (18) (where ρoil, ρH2O, [i]T, [i]oil and [i]H2O are the density of oil and water (g/cm3), total number of moles of i, the number of moles of i in the oil phase and the number of moles of i in the gas phase respectively), and rearranging gives ⎛V ρ ⎞ Km [i]oil = [i ]T ⎜ H2O H20 mi (oil) + 1⎟ ⎝ Voil ρoil K i (H2O) ⎠

−1

(19)

This equation form enables the partitioning of species i between any two liquid phases to be calculated as a function of liquid density and relative volumes. Similar to the gas-

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liquid system, in the simplest case, determination of the concentration of one groundwater-derived noble gas concentration in either the water or the oil phase enables quantification of a system oil/water volume ratio. This is discussed in detail in the following sections. Relative fractionation Although the absolute concentration and distribution of noble gases between different phases is often useful if the original noble gas concentration in one fluid phase is known (e.g., atmosphere-derived noble gases in groundwater), in many systems the concentration may not be so well determined, but a reasonable estimate of the initial relative concentrations of noble gases in a fluid may be available (e.g., crustal-radiogenic or diluted air/groundwater-derived noble gases). In this case it is often convenient to investigate the relative change in noble gas ratios, or fractionation, from the predicted ratio. By convention fractionation is usually assessed relative to 36Ar. Gas-liquid. From Equation (13), the relative fractionation between, for example, species i and Ar in the gas phase, ([i]/[Ar])g, can be related to the original ratio in the system ([i]/[Ar])T, after Bosch and Mazor (1988), by ⎛ Vg 1 ⎞ ⎜ + d ⎟ ⎛ [i] ⎞ ⎛ [i] ⎞ ⎝ Vl KAr ⎠ ⎜ ⎟ =⎜ ⎟ ⎝ [Ar ]⎠ g ⎝ [Ar ]⎠ T ⎛ Vg 1 ⎞ + d⎟ ⎜ ⎝ Vl K i ⎠

(20)

Taking the limits as Vg/Vl → ∞ then ([i]/[Ar])g → ([i]/[Ar])T, and as Vg/Vl → 0 then ([i]/[Ar])g → ([i]/[Ar])T(Kdi/KdAr). Therefore as Vg/Vl becomes small, the noble gases are fractionated proportionally to their relative solubilities in the liquid phase, or more precisely including the effects of non-ideality on the solubility: ⎛ [i] ⎞ ⎜ Ar ⎟ ⎝ [ ]⎠ g

γi d K Φi i As Vg/Vl → 0, → γ = Fgas ⎛ [i] ⎞ Ar KdAr ⎜ Ar ⎟ Φ Ar ⎝ [ ]⎠ T

(21)

where Fgas is [i]/[Ar] ratio in the gas phase normalized to the original system ratio to give a fractionation factor. An Fgas of 1 indicates that no fractionation from the original system value has occurred. Liquid-liquid. In a similar fashion the relative change in the [i]/[Ar] ratio in a liquidliquid system can be assessed. For example, the high solubility of noble gases in oil relative to water can potentially result in significant and distinct fractionation of the noble gases where equilibrium has occurred between these two phases. Following Bosch and Mazor (1988), the [i]/[Ar] ratio in the oil phase, ([i]/[Ar])oil, is related to the original [i]/[Ar] ratio, ([i]/[Ar])T, the oil/water volume ratio Voil/VH2O and the solubility of the noble gas species in each phase, where ⎛ [i] ⎞ ⎜ ⎟ ⎝ [Ar ]⎠ oil

⎛ V K dAr(oil) ⎞ oil + d ⎜⎜ ⎟⎟ ⎛ [i] ⎞ ⎝ VH2O KAr(H2O) ⎠ =⎜ ⎟ ⎝ [Ar ]⎠ T ⎛ Voil K di(oil) ⎞ + d ⎟⎟ ⎜⎜ ⎝ VH2O K i(H2O) ⎠

Maximum fractionation occurs in the oil phase when

(22)

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as Voil/VH20 → 0,

⎛ K dAr (oil) ⎞ ⎛ [i] ⎞ ⎜⎜ d ⎟⎟ ⎜ Ar ⎟ ⎝ [ ]⎠ oil ⎝ K Ar (H2O) ⎠ → = Foil ⎛ [i] ⎞ ⎛ K di (oil) ⎞ ⎜ Ar ⎟ ⎟⎟ ⎜⎜ d ⎝ [ ]⎠ T ⎝ K i (H2O) ⎠

551

(23)

The magnitude of the fractionation seen in either phase is therefore proportional to the ratio between the noble gas relative solubilities in the two liquid phases. Equilibrium fractionation of the Ne/Ar ratio: An example. From Equations (20) and (21) and the solubility data presented in Tables 2 and 3, the fractionation of the noble gases between gas and liquid phases can be calculated for a range of Vg/Vl ratios and at temperatures and salinities appropriate to those in an active sedimentary basin. Figure 7a shows the maximum Ne/Ar fractionation in water and gas phases for pure water and a 5M NaCl brine. For pure water, maximum fractionation in the gas phase is 3.4 at 290 K and decreases with increasing temperature, salinity and Vg/Vl. For example, at 290 K and in equilibrium with a 5M NaCl brine, Ne/Ar fractionation in the gas phase has a maximum value of 2.5. The effect of phase equilibrium with an oil phase and the resulting magnitude of fractionation that can occur is illustrated in Figures 7b and 7c. The relative solubilities of the noble gases in an oil phase have a greater range than in water. This range increases with increasing oil density. The effect of this increased difference in solubility between the noble gases is an increase in the magnitude of fractionation that can occur in an oil/gas/water system. For example a gas phase in equilibrium with ‘Heavy’ crude oil (API = 25) at 330 K as Vg/Vl aproaches zero will have a Ne/Ar fractionation factor of 7.1. This decreases to a maximum of 4.3 for a light crude oil (API = 34) at the same temperature (Fig. 7b). Fractionation in any gas phase associated with oil decreases with increasing temperature, Vg/Vl and decreasing oil density. Equilibration between water and an oil phase causes maximum fractionation in the oil phase as the salinity of the water phase approaches saturation and with increasing oil density when Voil/Vwater approaches zero (Fig. 7c). Similarly, maximum fractionation in the water fractionation in the water phase as Voil/Vwater approaches infinity, and as the salinity of the water phase approaches saturation. Unlike liquid/gas phase fractionation, which increases with decreasing temperature, water/oil fractionation reaches a maximum at moderately low temperatures. This occurs, for example, in a pure water/ ‘light’ oil (API = 34) system at 310 K, with a maximum Ne/Ar fractionation of 0.51 and 1.96 in the oil and water phases respectively. This can be compared with a pure water/ ‘heavy’ oil (API = 25) system where at 286 K a maximum Ne/Ar fractionation of 0.27 and 3.69 is obtained in the oil and water phases respectively. Rayleigh fractionation A simple dynamic model. The maximum magnitude of noble gas fractionation that can occur when two phases have been equilibrated is summarized for Ne/Ar in Figure 7. Although the ‘phase equilibrium’ model demonstrates the effect of the physical conditions in a system on the limits of noble gas fractionation, the phase equilibrium model represents only one end-member of the processes that may be occurring in a dynamic subsurface fluid environment. To convey some sense of the relevance of the phase equilibrium model in a dynamic system it is useful to consider the extent to which noble gases partition and fractionate between phases when a gas bubble passes through a column of liquid (Ballentine 1991; Fig. 8).

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Figure 7. (A) The dependence of Ne/Ar fractionation between water and a gas phase is shown as a function of temperature, salinity and the gas/water volume ratio (Eqns. 20, 21) modified from Ballentine et al. (1991). The maximum equilibrium fractionation in the gas phase occurs when the water phase salinity and temperature is low and as Vg/Vw approaches 0. A graduated scale between pure water and 5 M NaCl brine is shown to illustrate the effect of changing salinity. Graduated scales are also shown to illustrate the effect of changing the Vg/Vw ratio for both the pure water and a 5M NaCl brine. Fractionation of the Ne/Ar ratio in the water phase is the inverse of that in the gas phase, with maximum fractionation occuring at low temperature and salinity as Vg/Vw approaches infinity. (B) The dependence of Ne/Ar fractionation in an oil/gas phase system is shown as a function of temperature (faint line is ‘light’ oil, API = 34; dark line ‘heavy’ oil, API = 25). Maximum fractionation in the gas phase occurs as the Vg/Voil ratio approaches zero. Maximum fractionation occurs in an oil/gas system at low temperature and as the oil density increases.

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Figure 7 caption, continued. Fractionation in the oil phase is the inverse of that occuring in the gas phase, with maximum fractionation occuring as the Vg/Voil ratio approaches infinity. (C) The dependence of Ne/Ar fractionation in an oil/water phase system (‘light’ oil, API = 34) is shown as a function of temperature. Solid lines represent the limit of experminetal data, dashed lines an extrapolation. A fractionation maximum occurs at low temperature, 310 K for ‘light’ oil, and high salinity. In the water phase maximum fractionation occurs at high salinity, high oil density and as the Vwater/Voil ratio approaches zero. Maximum fractionation in the oil phase occurs at high salinity, high oil density and as the Vwater/Voil ratio approaches infinity.

Figure 8. (A) A water column is divided into fifty equal unit cells and it is assumed there is no liquid or dissolved gas between cells. Each cell originally has the noble gas content of airequilibrated water and all calculated Ne/Ar ratios are normalized to this value to obtain a fractionation factor F. The column temperature is taken to be 325 K, which for pure water gives KNe = 133245 atm and KAr = 55389 atm. A gas bubble of constant volume is passed sequentially through the column, equilibrium assumed to occur in each water cell and the Ne and Ar partitioned into the respective gas and water phases (Eqn. 16). The evolution of the Ne/Ar ratio in the gas bubble (bold) and each water phase increment (Faint) is shown for different gas/water volume ratios, Vg/Vl. The gas bubble Ne/Ar ratio approaches the maximum fractionation value predicted for a gas/water phase equilibrium where as Vg/Vl → 0, F → KNe/KAr. The cell Vg/Vl ratio only determines the rate at which this limit is approached. (B) The same water column with a fixed cell Vg/Vl ratio of 0.01. n subsequent bubbles are passed through the column and the He/Ne distribution between phases calculated at each stage. The gas bubble Ne/Ar ratio evolution for n = 1, 10, 20 and 30 is shown in bold, together with the residual Ne/Ar in the water column cells (faint lines). All gas bubbles approach the limit imposed by the phase equilibrium model. The water phase is fractioned in the opposite sense and is fractionated in proportion to the magnitude of gas loss following the Rayleigh fractionation law (Eqn. 24).

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As a starting point, the liquid can be taken to be water that has equilibrated with air to obtain its noble gas content. Furthermore, it is assumed that the liquid is saturated with respect to the dominant gas species forming the bubble/gas phase. The column is divided into cells and it is assumed that there is no transport of dissolved gases or fluid between the cells. When a gas bubble, initially with no noble gas content, is introduced into the first cell the distribution of both Ne and Ar can be calculated from Equation (16) assuming complete equilibration between the gas and fluid in that cell only. The volume of the bubble is assumed to be constant and, now with a noble gas content, is moved to the next cell. Equilibrium is again assumed, and the resulting distribution of Ne and Ar between the gas and liquid phases calculated. In this manner the Ne and Ar concentrations and Ne/Ar ratio can be calculated for the gas phase and each water cell as the bubble is sequentially passed through the unit cells of the liquid column. The Ne/Ar ratios in both gas bubble and modified water cell are normalized to the air-equilibrated water ratio originally in each water cell to obtain a fractionation factor, F. The effect of varying Vg/Vl ratios in each increment is illustrated in Figure 8a. No gas bubble exceeds the fractionation limit predicted by the equilibrium solubility model where as Vg/Vl → 0, Fgas → KNe/KAr (Eqn. 21). When this value has been reached in the gas phase, the Ne and Ar concentrations in the bubble are in equilibrium with the noble gas content of the unmodified water phase. Neither gas nor water phase will be modified as this bubble passes through additional water ‘cells’ in the column. The only effect of the varying Vg/Vl ratios is the rate at which equilibration between the bubble and water cells is reached. After the first bubble has modified the Ne/Ar content of the water column, another bubble with the same volume and with no initial noble gas content can be passed into the column. The Ne and Ar concentration and Ne/Ar ratio in both the water cell and gas bubble can again be calculated for each cell as the bubble progresses up through the water column. In this manner for any number, n, of subsequent bubbles the noble gas distribution and evolution of each water cell and bubble can be calculated. In Figure 8b, the Ne/Ar fractionation of the water and gas phase is shown for a column with a cell Vg/Vl = 0.001 for n=1, 10, 20 and 30 bubbles. For n = 1, the plotted fractionation in both phases is the same as that shown in Figure 8a. Subsequent gas bubbles initially inherit a lower and lower Ne/Ar ratio as n increases because they equilibrate with the water cells fractionated by the previous bubbles. Nevertheless, the gas bubbles all approach the same F(Ne/Ar) limit predicted by the solubility equilibrium model as they progress through the column. The water column, after each bubble has passed through, becomes increasingly more depleted in noble gas content and retains a more and more fractionated Ne/Ar ratio as the dissolved noble gases preferentially partition into the gas phases. The noble gas ratio in the water phase exceeds the fractionation predicted by the equilibrium model, and is related to the fraction of gas remaining by the Rayleigh fractionation law ⎛ [i] ⎞ ⎛ ⎞ = [i] P (α −1) ⎝ [Ar ]⎠ water ⎝ [Ar ]⎠ o

(24)

P is the fraction of Ar remaining in the liquid (water) phase, ([i]/[Ar])o, the original liquid phase i/Ar ratio and α is the fractionation coefficient given for a gas/liquid system where:

α = (Kiliquid/KArliquid)

(25)

Similarly, the Rayleigh fractionation coefficient used to determine the magnitude of fractionation in a water phase that has interacted with an oil phase (instead of gas) is given by

α = (Kiwater KAroil)/(Kioil KArwater)

(26)

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The K variables are the solubilities of the noble gas i and Ar in oil and water. In summary: • The phase equilibrium model limits the maximum noble gas fractionation that will occur in a gas phase migrating though groundwater. • By direct analogy, noble gas fractionation in an oil phase migrating through groundwater will be similarly limited by the phase equilibrium model. As the phase equilibrium value is approached in either the gas or the oil phase, quantitative information from the magnitude of fractionation about the volume of water that has equilibrated with the non-water phase will be lost (although minimum volumes can be inferred). The path length and time required for either the gas or oil phase to achieve the phase equilibrium limit will depend on the relative availability of the groundwater for equilibration, which in turn is controlled by factors such as porosity, tortuosity and interconnectivity of the rock matrix, as well as the residence time of both the gas or oil and groundwater phases. It is also important to consider scale. For example, the ‘bubble’, can be considered to be as large as a gas or oil field, and its movement only relative to the water. In this case an oil or gas field that equilibrates with an active groundwater system cannot be distinguished from groundwater equilibration during oil or gas phase migration from source rock to trap. In contrast to a migrating gas or oil phase, the residual groundwater phase will be fractionated following Rayleigh fractionation (Eqn. 24; Fig. 9). • Although extensive fractionation in the residual groundwater phase can occur in the opposite sense to that of the migrating phase, the absolute concentration of the fractionated noble gases in the water phase is very much reduced. • Strongly fractionated noble gas ratios can be transferred to the gas or oil phase from previously ‘stripped’ groundwater, but any such oil or gas can only contain low water-derived noble gas concentrations. Re-solution and effervescence A mechanism proposed to increase the magnitude of fractionation in a gas phase over that predicted by the phase equilibrium model limit is that of a multi-stage process of re-solution and effervescence (Zartman et al. 1961). In the simplest case a gas phase, containing a significant concentration of fractionated noble gases in equilibrium with airequilibrated groundwater, could be re-dissolved by a change in physical conditions such as an increase in pressure, a decrease in temperature and salinity, or mixing with unsaturated (with respect to the major gas phase) water. This would create a local increase in both the groundwater noble gas concentration and the magnitude of noble gas fractionation in solution. Subsequent formation of a gas phase in equilibrium with the modified groundwater would show a fractionation relative to air-equilibrated water in excess of that predicted by the single stage equilibrium solubility model. We can consider the volumes and concentrations required for this process through a worked example. In Figure 10, water and gas volumes have been chosen to produce significant fractionation in the final gas phase by a two-stage process of re-solution and effervescence, with a Vg/Vl ratio of 0.001 at each stage. The water is taken to be seawater, which has equilibrated with air at 20°C, now at conditions typical of 700-m depth: a temperature of 310K, hydrostatic pressure (68atm) and with a high salinity (5M NaCl) (KNe = 4.72×105 atm and KAr = 2.33×105 atm). The water is equilibrated with a gas phase and produces a 20Ne/36Ar fractionation value of 1.77 in the gas phase, which can be compared to the phase equilibrium limit of KNe/KAr=2.03 as Vg/Vl → 0. It is assumed that this volume of gas is then re-dissolved in a small volume of unaltered air-equilibrated seawater. The 20Ne/36Ar ratio of the fluid is now 1.72 times greater than the unaltered

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water and, when equilibrated with a gas phase under the same conditions as the first stage, produces a 20Ne/36Ar fractionation in the gas phase of F = 3.00. This is ~50% greater than the maximum fractionation for a single stage process. Zartman et al. (1961) suggest that this process may be repeated several times to produce highly fractionated noble gas patterns.

Figure 9. Fractionation of the Ne/Ar and Xe/Ar ratios in groundwater is shown by process and interacting phase. Conditions are taken to be 330 K and 2 M NaCl salinity. Water-Gas batch equilibration is shown as a function of Gas/Water volume ratio (Vg/Vl), as well as fraction of Ar remaining in the water phase, P-Ar. Batch noble gas distribution between water and gas phases is calculated following Equation (17). Rayleigh fractionation (faint line) of the water phase by gas is calculated following Equations (24) and (25). The graduated scale shows the fraction of Ar remaining in the groundwater phase. Water-Oil batch equilibration is also calculated for a light crude oil (API = 41, density = 0.82g/cm3, Kharaka and Specht 1988) following Equation (19). The change in Ne/Ar and Xe/Ar ratios is shown as a function of Oil/Water volume ratio (Voil/Vwater), as well as fraction of Ar remaining in the water phase, P-Ar. Rayleigh fractionation (faint line) of the water phase by gas is calculated following Equations (24) and (26). The graduated scale shows the fraction of Ar remaining in the groundwater phase. Tie lines between batch and Rayleigh fractionation resulting in the same P-Ar are also shown for gas-water and oil-water systems. On the same figure we have taken fresh water equilibrated with air at 1atm pressure at 293 K as our reference composition. The figure shows the effect on the fractionation values of variable recharge temperature and air addition. The effect of either a gas phase or an oil phase passing through groundwater will be distict and in principle enables quantification of the respective oil/water or oil/gas ratios.

The first point for consideration in this two-stage model is the re-solution of the gas phase. We use CH4 as an example. 65 m3 CH4 (STP) occupies 1 m3 at a depth of 700 m (68 atm and 310 K). From Price (1981), 14.6 m3 CH4 (STP) saturates 10 m3 of pure water under these. ~40 m3 of pure water, unsaturated with respect to CH4, are therefore required to re-dissolve this gas (more if this were a brine). If this volume is air-equilibrated water, the additional (and unfractionated) noble gas content would have the effect of lowering the final 20Ne/36Ar fractionation value to F = 1.57, less than the maximum fractionation value predicted for a single stage equilibrium. Notwithstanding more complex processes, such as the addition of a phase with no initial noble gas content, it would seem that the

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process of re-solution of the major gas phase provides a major limiting factor when advocating a re-solution and effervescence model. If the fluid in equilibrium with the final gas phase is air-equilibrated water, the concentration of fractionated noble gases in solution must be significantly higher than

Figure 10. Re-solution and effervescence—a 3-stage worked example. Initial: A volume of seawater is equilibrated with air at 20°C and 1 atm, and then buried. The temperature and salinity and temperature is increased and the salinity modified Henry’s constants, KiMki, calculated. Stage 1: The water is equilibrated with a gas phase and the noble gas distribution between the phase calculated (Eqn. 16). Stage 2: The gas phase is re-dissolved in a volume of seawater that has an unmodified noble gas composition. Stage 3: The volume of water containing the re-dissolved noble gases is equilibrated with a gas phase containing no noble gases and the noble gas distribution between the phases re-calculated. The fractionation achieved in the gas phase is approximately 50% higher than that predicted by the phase fractionation model limit, but requires huge quantities of water. The physical process of re-solution of gases is also problematic (see text). When invoking a multi-stage process of re-solution and effervescence to account for highly fractionated noble gas ratios, the geological implications of mass balance and mechanism of re-solution must be very carefully considered.

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typical meteoric groundwaters. This would also require a proportionally high concentration of noble gases in the gas phase (although this could be masked by subsequent dilution by air-noble gas free gas addition). This is demonstrated in the worked example (Fig. 10), where the final gas phase has an order of magnitude greater concentration of 20Ne and 36Ar than the gas in equilibrium with unaltered air-equilibrated seawater. • Subsurface fluid phases with high noble gas concentrations as well as noble gas fractionation in excess of single step phase fractionation limits may provide field evidence for multi-stage processes of re-solution and effervescence. • The volume of the highly fractionated gas relative to the volume of original groundwater is however, very small. Orders-of-magnitude less fractionated gas is produced relative to the volume of liquid than in a single-stage water/gas equilibrium. • Both the physical re-solution of the major gas phase and also the large mass of the fluid phase required appear to preclude resolution and effervescence as a significant mechanism to fractionation noble gases beyond the soluble equilibrium limit without very careful consideration of the geological context. Multiple subsurface fluid phases The solubility equilibrium model limits the noble gas concentration and relative fractionation in a fluid ‘receiving’ noble gases originally associated with another fluid phase. The residual concentration and fractionation of the noble gases in the ‘donating’ fluid is controlled by Rayleigh fractionation or batch equilibrium depending on whether the system is open or closed to loss of the ‘receiving’ fluid phase (Fig. 9). As soon as a third or more phases are involved, the system become more complex, but is nevertheless still controlled by Rayleigh and batch equilibration limits. For the purpose of discussion we consider a system in which the original fluid phase is groundwater containing airderived noble gases at concentrations fixed during recharge. If the water noble gas content is altered by equilibration with an oil phase, subsequent equilibration of a gas phase with either the oil or water will result in a fractionation value in the gas phase that will reflect the salinity, temperature, oil density and gas/water/oil volume ratios and type of equilibration—either open or closed system The limits imposed by closed system interaction between water, gas and oil phases on the range of the gas phase 20Ne/36Ar ratios, originally derived from the air-equilibrated water, can be assessed (Bosch and Mazor 1988). For example at 310 K, KNe and KAr are 8471 atm kg/mol and 4155 atm kg/mol respectively in a 5M NaCl brine, and KoilNe and KoilAr are 622 atm Kg/mol and 117 atm Kg/mol. Maximum positive Ne/Ar fractionation in a gas phase occurs when the oil phase equilibrates with a small volume of water, transferring the noble gas content of the water into the oil phase with minimal fractionation. Subsequent equilibration of the oil with a small volume of gas will produce a fractionation value of F(Ne/Ar)gas = 622/117 = 5.3 (Eqn. 21), compared with the maximum fractionation of F(Ne/Ar)gas = 8471/4155 = 2.03 predicted for a water/gas system under the same conditions (Fig. 7a). Maximum gas phase Ne/Ar fractionation, in the opposite direction, occurs when the oil phase equilibrates with a large volume of water to produce F(Ne/Ar)oil = (117/4155)/(622/8471) = 0.31 (Eqn. 23). The addition of one extra phase, crude oil, to a gas/water system more than doubles the range of fractionation that can occur in any associated gas phase, from between F(Ne/Ar)gas = 1.0 to 2.03 to between F(Ne/Ar)gas = 0.31 to 5.3. In the case of an open system, fractionation in the residual phase will be even more extreme, and reflected in much lower concentrations (Battani et al. 2000). While it is possible to envisage a myriad of different interactions between water, gas and oil phases, depending on the order of interaction, open or closed system behavior and the relative fluid volumes, it is not a

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useful exercise at this point to consider all of the possibilities, as other factors may also play a role in constraining the physical model development. For example, stable isotope information from the hydrocarbon gases may rule out the involvement or association with an oil phase (Schoell 1983), and systems must be considered on a case-by-case basis. Diffusion or kinetic fractionation In a gas phase. Gaseous diffusion processes can generate both elemental and isotopic fractionation in natural gases. Marty (1984) reviews the processes that can affect noble gases after Present (1958) and distinguishes among: a) free-molecule diffusion; b) mutual diffusion; and c) thermal diffusion. (a) Free-molecule diffusion takes place when a gas is traveling through a conduit in which gas-wall collisions are more frequent than gas-gas collisions. For this to occur the conduit diameter must be smaller than the mean free path of the gas atoms. Because the mean speed of particles is proportional to m-1/2 , where m is the mass of the gas atom, in the case of a binary mixture the lightest component will be enriched at the outlet of a conduit. It is shown that the fractionation coefficient, α, between two elements of masses m1 and m2 is approximated by:

α = [(m2/m1)1/2]

(27)

For a system depleted by a free-molecule diffusive process the Rayleigh fractionation law can again be applied. Taking m1 as the mass of gas i and m2 as the mass of Ar: ⎛ [i] ⎞ ⎛ ⎞ = [i] P ( α−1) ⎝ [Ar ]⎠ gas ⎝ [Ar ]⎠ o

(28)

where P is the fraction of Ar remaining in the gas reservoir, ([i]/[Ar])o, the original gas phase i/Ar ratio and α is the fractionation coefficient in Equation (27). It should be noted that for natural gases the dimension of the conduit needs to be very small (diameter < 10-8 m). Because the mean free path of the gas atoms is proportional to temperature and inversely proportional to pressure, this will decrease with depth. Fractionation through free molecular diffusion therefore, will only be significant in special circumstances and when the pressure of the system is low (approaching atmospheric pressure). (b) Mutual diffusion describes the diffusion of two or more gas species when the dominant interactions are gas-gas collisions. When a gas i with mass m1 diffuses through a gas with an average molecular mass mg, the diffusion coefficient of i is proportional to [m1 × mg) /(m1 + mg)]-1/2 = (m1*)-1/2

(29)

where m1* is the ‘reduced mass.’ The mean velocity of gas i is proportional to its diffusion coefficient. Therefore, for a second gas j with a mass of m2 the relative velocity between i and j provides the fractionation coefficient α that can be used in Equation (28), where

α = (m2*/m1*)1/2 = (m2/m1)1/2 × [(m1 + mg)/(m2 + mg)]1/2

(30)

These equations are only strictly relevant in the subsurface to a single phase gas system undergoing diffusive loss and remains to be applied in any noble gas study. (c) Thermal diffusion occurs when a gas mixture is in a non-equilibrium state because of a tendency for lighter molecules to be concentrated at the high temperature

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boundary. It is shown that the fractionation factor, f, in a steady state is: f = (T’/T)β

(31)

where T’ and T are the highest and lowest temperature and β is the element or isotope pair specific thermal diffusion parameter. Isotope enrichment can only be obtained by preferential extraction of either the cold or hot gas (e.g., Clusius and Dickel 1938). Marty (1984) considers various scenarios in the subsurface where isotopic fractionation may occur in light of the results of Nagao et al (1979, 1981). From both theoretical considerations (see above) and observed isotopic anomalies, this author concluded that the most likely cause of rare gas isotopic fractionation in natural gases is mutual diffusion taking place between atmospheric gases in soil pores and volcanic/geothermal gases ascending through the upper level of the crust and sediments. In water. The two principle controlling factors for gas diffusion processes in liquids are the gas mass and its activation energy for diffusion. The diffusion activation energy in turn is controlled by the extent of interaction of the gas molecule or atom with the liquid phase. For noble gases, because they are monatomic and have a stable electron shell, there is little interaction with water and the rate of diffusion is almost entirely controlled by their respective masses. This is in contrast with species such as CO2 and CH4 where interaction occurs with water molecules through induced dipole-dipole moments. Because this is in addition to mass, these species diffuse significantly more slowly in water than noble gases of similar mass (Table 5, Fig. 11). Gas

Table 5. Diffusion coefficients* in water (Jähne et al. 1987). Medium Ea A 1σ error 1σ error

He Ne Ar ** Kr Xe Rn H2 CH4 CO2

water seawater water water water water water water seawater water water

Kj/Mol

%

10-5 cm2/s

%

11.70 12.02 14.84 17.30 20.20 21.61 23.26 16.06 14.93 18.36 19.51

5 5 8 10 3 5 11 3 9 4 2

818 886 1608 3141 6393 9007 15877 3338 1981 3047 5019

2.1 1.8 3.5 5 1.6 3.5 11 1.6 4.3 2.7 1.3

* To fit equation D = A exp(-Ea/RT) where D is the diffusion coeficient, T the temperature in Kelvin, and R the gas constant ** Extrapolated, see text

The most complete study of noble gas diffusion rates in water remains the experimental determination by Jähne et al. (1987). In this work the diffusion coefficient in water was determined for systems between 0 and 35°C and the results expressed in terms of the diffusion constant, A (cm2/s), and diffusion activation energy, Ea (Kj/Mol) (Table 5), to provide a temperature dependent expression for the determination of the gas diffusion coefficient, D (cm2/s), at variable temperature following D = A e-Ea/RT

(32)

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561

R and T are the gas constant and temperature in Kelvin respectively. For any one temperature it can be shown that the diffusion coefficient for the noble gases are well correlated with the square roots of their masses (Fig. 11). Although Ar has not been experimentally determined in this study, this clear relationship enables the values of A and Ea for Ar to be readily interpolated from the other noble gas values. The interpolated values for Ar are also shown in Table 5. The correlation of diffusion coefficient with the square root of their masses also allows the relative mass fractionation of noble gases to be calculated using Equations (27) and (28).

Figure 11. Diffusion coefficients for noble and selected active gases measured in water at 25°C, shown as a function of mass (after Jähne et al 1987).

Jähne et al. (1987) also investigated isotopic specific diffusion coefficients for 3He and the change in δC13(CO2) during diffusive gas loss from water. The increase in He diffusivity for the 3He compared to 4He was in agreement with the ratio of the square-root of their masses. This result provides further supporting evidence that the diffusion coefficients for individual isotopic noble gas species can reasonably be determined as a function of mass from Table 5 for variable temperatures. This is in contrast with the results for the study of δC13(CO2), which showed a fractionation factor far lower than the value predicted from the square root of the reduced mass. This discrepancy indicates that in the case of active gases the difference is not just an effect of mass but of the isotope specific interaction energy with the water molecules. The effect of salinity on gas diffusion rates is not quantitatively determined except for He and H2 in seawater (Table 5), showing a reduction in diffusion rate with an increase in salinity. In a water-filled porous medium. As soon as diffusion in a water-filled porous medium is considered, the effects of porosity, tortuosity, permeability as well as any interaction with the porous medium must also be considered. This enables us to define the first order ‘apparent’ diffusion coefficient, Da (cm2/s), where: Da = D/(R×Rf)

(33)

Rf is the retardation factor caused by the porous medium geometry and R the retardation factor due to physical or chemical interaction between the gas and porous medium. For the case of noble gases these latter interactions are usually insignificant and R = 1. Rf, determined for different rock varies over orders of magnitude but has not been correlated with physical rock properties to enable an assessment of all rock types. For example, Rf values of between 1.5 and 2.5 have been estimated for He and Ar in deep-sea sediments

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with a porosity of 70-80% (Ohsumi and Horibe 1984). Krooss and Leythaeuser (1988) have measured diffusion coefficients for light hydrocarbons in sedimentary rocks and calculated Rf values of between 20-50 for argillaceous sediments. Bourke et al. (1989) using both iodine and tritiated water diffusion in London clays calculates a similar value (Rf ~ 30). This is in contrast with Neretnieks (1982) who reports Rf values ranging from 100 to 1000 in compacted bentonite for hydrogen and methane diffusion. Rebour et al. (1997) review the literature describing gas diffusion in a porous medium as a ‘double’ porosity process. In this model, gas diffusion is affected by the increase in water viscosity when in the close vicinity of clay minerals. This produces an environment in which the gas diffusion rate is expected to be variable in the porous network depending on the local tortuosity and grain-size distribution. In modeling this type of system, diffusion is considered to occur along a direct pathway. These ‘fast’ routes interconnect ‘slow’ regions, into and out of which gas also diffuses. Experimental work by the same authors (Rebour et al. 1997) determines Rf = 200 for a clayey marl from Paris basin Callovo-Oxfordian sediments that have a porosity and permeability of 23% and 10-22 m-2, respectively. The isotopic fractionation in hydrocarbon gases by diffusive processes is a topic of current research (e.g., Zhang and Krooss 2001; Prinzhofer and Pernaton 1997), but is masked in natural systems by the variability of source values. As yet no work linking noble gases with stable isotope fractionation has been undertaken. • Despite the uncertainty regarding absolute rates of diffusion for noble gases in a water-filled medium, the relative rates remain a direct function of mass. In principle, for example, the extent of diffusive gas loss for any reservoir can be determined by the magnitude of fractionation of known noble gas elemental ratios using Equation (24) and the appropriate mass fractionation coefficient (Eqn. 27). RESOLVING DIFFERENT NOBLE GAS COMPONENTS IN CRUSTAL FLUIDS Terrestrial noble gases are dominated by three reservoirs: the atmosphere, crust and mantle. The isotopic compositions of noble gases produced by radioactive decay in the crust are distinct from noble gases derived from the mantle, which in turn are distinct from those in the atmosphere (Fig. 1). • Notably, the isotopes of 20Ne, 36Ar, 82Kr and 130Xe are not produced in significant quantities by radioactive processes in the crust (Ballentine and Burnard 2002, this volume) and, in the absence of a magmatic contribution, are almost entirely dominated by atmosphere-derived sources. Two-component mixing. The atmosphere however, does contain a significant amount of noble gases, such as 21Ne, 40Ar and 136Xe that are also derived from crustal/radiogenic sources. To distinguish which of these species are derived from the immediate crustal system from those that are derived from atmosphere sources, the isotopic ratios can be compared with the atmospheric ratio to identify the crustal ‘excess.’ In a two-component crust/air mixture [21Ne]crust = [21Ne]tot x [1-(21Ne/20Ne)air/(21Ne/20Ne)s] 40

40

40

36

40

36

Xe)air/(

136

[ Ar]crust = [ Ar]tot x [1-( Ar/ Ar)air/( Ar/ Ar)s] 136

[

136

Xe]crust = [

Xe]tot x [1-(

136

Xe/

130

Xe/

130

(34) (35)

Xe)s]

(36)

The subscripts crust and tot refer to the crustal and total concentrations while the subscripts air and sample refer to the isotopic composition of the atmosphere (Ozima and Podosek 1983) and sample respectively.

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U and Th decay in both the mantle and crust to produce 4He. However, the Earth’s mantle has also preserved a significant quantity of ‘primitive’ 3He during accretion (e.g., Porcelli and Ballentine 2002, this volume). Because 3He is not produced in significant quantities by radioactive decay processes, mantle-derived He has a far higher 3He/4He ratio than crustal sources and even small magmatic additions to crustal fluid systems are readily resolvable (Poreda et al. 1986; Oxburgh et al. 1986). In contrast to the heavier noble gases, He, because of thermal escape from the atmosphere, has only a low abundance in the atmosphere. It is nevertheless necessary to correct any measured He isotopic composition for air-derived contributions by using the observed air-derived 20Ne concentration following (Craig et al. 1978) where (3 He/ 4 He)c =

( 3 He/ 4 He)s ×( 4 He/ 20 Ne)s /( 4 He/ 20 Ne)air −(3 He/ 4 He) air ( 4 He/20 Ne) s /(4 He/ 20 Ne) air − 1

(37)

Subscripts c, s and air refer to the corrected, measured and air-derived ratios, respectively. The (4He/20Ne)air elemental ratio, unlike isotopic ratios, is subject to elemental fractionation. In applications where the air correction is large and/or critical the (4He/20Ne)air ratio has to be determined with care. For example, in a groundwater that has not undergone phase fractionation this value can be determined from the recharge temperature and estimate of air in excess of recharge equilibrium or ‘excess air’ (e.g., Kipfer et al. 2002, this volume). In less critical applications where the air-derived component reasonably has a groundwater origin the measured (4He/20Ne)air = 0.288 in 10°C air-equilibrated water is often used (e.g., Craig et al. 1978). In many old groundwaters and hydrocarbon fluids, (4He/20Ne)s is large and the correction is negligible. In this case (3He/4He)c ≈ (3He/4He)s. Once corrected for atmosphere-derived He, the 3He/4He ratio represents the sum of only two components; the crust and the mantle. The contribution of crustal 4He is then given by [4 He]crust =

[ 4 He] tot × [( 3 He/ 4 He) mantle −( 3He/ 4He) c ] [(3 He/ 4He) mantle −( 3 He/ 4 He)crust ]

(38)

Subscripts mantle, crust and c refer to the mantle, crust and air-corrected values. Although (3He/4He)crust is well defined (~1×10-8, Ballentine and Burnard 2002, this volume), the choice of (3He/4He)mantle has to be made with care: 3He/4He for local subcontinental lithospheric mantle recorded in mantle xenoliths ranges between 8.54×106 to 6.53×10-6 (Dunai and Baur 1995; Dunai and Porcelli 2002, this volume), input from the convecting mantle typical of that supplying mid ocean ridges would have higher values at 3He/4He = 1.12×10-5 (e.g., Graham 2002, this volume), while mantle regions beneath the crust influenced by a high 3He/4He plume, such as Yellowstone USA (Kennedy et al. 1985), may have yet higher 3He/4He values. Quoted errors need to reflect the degree of end-member uncertainty. Three-component mixing. When a significant magmatic component is present in a crustal fluid, in addition to the ubiquitous air-derived noble gases, there will also be a significant contribution from purely crustal radiogenic sources. In principle a similar approach to the resolution of the three-component He mixture can be taken (Eqns. 37, 38), by first correcting for atmosphere-derived contributions by reference to an unambiguously air-derived isotope. As above, the accuracy of this correction is entirely dependent on how well the elemental ratio of the air-derived pair is known. In many systems of interest elemental fractionation may have occurred and this approach for a single sample is then no longer appropriate. In the case of Ar, negligible 36Ar contributions to either crust or mantle components enable the atmosphere 40Ar to be corrected following Equation (35), although in this case

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the 40Ar excess is the sum of crustal and mantle contributions. Negligible or unresolvable differences among mantle, crust and air ratios of 38Ar/36Ar make it impossible for this isotope pair to be effectively used in resolving the mantle and crustal 40Ar components. In the case of Ne, the 21Ne/22Ne and 20Ne/22Ne ratios of all three components are significantly different, and given three isotopes and three components the contribution from each source to each isotope can be calculated (e.g., Ballentine and O'Nions 1992) where [20]air = 20 Total

(39) 21 x - 21 + 22 22 mntl 22 rad 20 meas 21 - 21 + 22 x 22 mntl 22 rad 20 air

21 22 21 22

x 20 rad 22 x 20 rad 22

- 21 mntl 22 - 21 mntl 22

x 20 mntl 22 x 20 mntl 22

rad rad

+ 21 20 + 21 20

x 20 - 20 meas 22 rad 22 mntl x 20 - 20 air 22 rad 22 mntl

[21]mntl =

21 Total

(40)

20 - 20 + 22 x 22 air 22 rad 21 meas 20 - 20 + 22 x 22 air 22 rad 21 mntl

21 22 21 22

x 20 22 x 20 air 22 air

- 20 22 - 20 rad 22 rad

x 21 22 x 21 air 22 air

rad rad

+ 20 21 + 20 21

x 21 - 21 22 rad 22 air x 21 - 21 mntl 22 rad 22 air meas

[22]rad = 22 Total

(41) 20 x 21 - 20 x 21 + 21 x 22 air 22 mntl 22 mntl 22 air 22 meas 20 x 21 - 20 x 21 + 21 x 22 air 22 mntl 22 mntl 22 air 22 rad

20 22 20 22

- 20 22 - 20 mntl 22 mntl

air air

+ 20 22 + 20 22

x 21 - 21 22 air 22 mntl 21 - 21 x rad 22 air 22 mntl meas

The parentheses subscripts crust, air, mntl and meas refer to the isotopic composition of the crust, air mantle and sample respectively. The square bracket subscripts crust, mntl, air, and Total, refer to the concentration of the crust, mantle and air component relative to the total isotopic contribution, respectively. Xe is another noble gas isotopic system where the crust, mantle and air end-members are significantly different. Substitution of the appropriate Xe isotopic end-member compositions into Equations (39)-(41) enables the end-member contributions to be derived for the three component Xe isotopic system. Element ratio mixing lines. Although it is not possible to resolve mantle and crustal Ar contributions in a single sample, this is possible with multiple samples from environments in which the elemental ratios from the respective end-member sources are constant and unaffected by subsequent fractionation. This is achieved by extrapolation, of an element-ratio/isotope-ratio mixing line to the isotope-defined end-members. For example, a plot of 3He/4He vs 40Ar*/4He (where there are negligible air contributions to He and 40Ar* is the 40Ar corrected for air-derived 40Ar, Eqn. 35), represents an isotope ratio and elemental pair whose component parts have only two sources – the crust and the mantle. For a system in which the mantle and crustal components have constant 40Ar/4He, a mixing line will be defined. Extrapolation to mantle and crustal 3He/4He end-member isotope compositions enables resolution of the respective 40Ar/4He component ratios (e.g., Stuart et al. 1995) (Fig. 12). • This technique is applicable to all noble gas systems that allow reduction to two component element-ratio/isotope-ratio mixing lines. 40

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Figure 12. 40Ar*/4He vs 3He/4He measured in Dae Hwa (S. Korea) W-Mo deposit fluid inclusions after Stuart et al. (1995). Extrapolation of the mixing line defined by the samples from different mineralization zones to known end-member 3He/4He isotopic compositions enables end-member 4He/40Ar ratios to be determined. In this example the mantle 40Ar/4He = 0.69±0.06 and is typical of unfractionated samples from the mantle (e.g., Graham 2002, this volume), in contrast the crustal 40Ar/4He = 0.007. The latter value is far higher than crustal production ratio of ~0.2 and typical of a fluid derived from shallow cool regions of the crust (e.g., Ballentine and Burnard 2002, this volume).

Isotope ratio mixing lines. In cases where the elemental ratios may have been variably altered, by for example phase fractionation, the element-ratio mixing lines discussed above will not be preserved. Isotopic ratios are unaffected by this form of fractionation and isotope-ratio only mixing lines can be constructed. Ballentine (1997) for example, use this approach to identify the mantle He/Ne ratio of magmatic fluids in natural gases (Fig. 13). In this study a data inversion was used to identify the best fit mixing line and crustal 21Ne/22Ne production ratio for natural gases from gas fields around the world. The Ne isotopic composition, corrected for the atmospheric contribution using Equations (39)-(41), then represents a two component mix of mantle and crustal-derived Ne. Similarly, air contributions to the 3He/4He are negligible and this ratio represents a mix of mantle and crustal-derived He. The air-corrected Ne, (21Ne/22Ne)c, plotted against 3He/4He falls on a single mixing line for almost all samples and defines the mixing constant, r, where:

r = (4He/21Ne)crust/(4He/21Ne)mntl

(42)

(4He/21Ne)crust and (4He/21Ne)mntl are the 4He/21Ne ratio of the crust and mantle components before mixing respectively. Further examples of isotope mixing relationships and three-dimensional approaches when resolving three components from the isotope systematics are detailed in the section Description and analysis of multi-component noble gas mixtures in ore fluids. NOBLE GASES IN HYDROCARBON GAS AND OIL RESERVOIRS

In the context of playing a role in hydrocarbon exploration or field development noble gas studies are in their infancy. In principle noble gas fractionation in groundwater can provide a sensitive and quantitative tool with which to identify both natural gas

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3

He/4He (R/Ra)

Figure 13. Plot of (21Ne/22Ne)c vs. measured 3He/4He (R/Ra) for natural gases containing a significant mantle-derived component (Ballentine 1997). All data falls within 2σ error of a remarkably well defined mixing line between crustal and mantle derived endmember values. The hyperbolic mixing constant, D, determined from the mixing line is defined by D = (He/Ne)crust/(He/Ne)mantle. As (He/Ne)crust is reasonably well defined (e.g., Ballentine and Burnard 2002, this volume) the mixing line enables the magmatic He/Ne ratio in these crustal systems to be calculated. This is quite distinct from the mixing line involving an unfractionated mantle source bounded by a mantle He/Ne ratio with ‘Solar’ or ‘Popping rock’ values (hatched region). The difference has been ascribed to only partial degassing and solubility related fractionation of the noble gases in the magmatic source supplying volatiles to the crustal system. This appears to be a common feature in many extensional systems (see also Fig. 18).

(Zaikowski and Spangler 1990) and oil migration pathways—in the absence of either the gas or oil phase (Fig. 9). In practice, the acquisition of groundwater samples that are uncontaminated by drill fluids from boreholes is difficult for commercial reasons. The extension of this type of study to groundwater trapped in fluid inclusions is at the limit of current experimental techniques and remains unexplored. Most studies to date have focused on the information available from noble gases that are readily determined in natural gas and oil samples from actively producing fields. Of particular interest is the quantification of the amount of water that has equilibrated with an oil or a gas phase. This provides an understanding of the role that groundwater has played in both the secondary migration of oil and gas (where primary migration is expulsion from the source rock, and secondary migration is transport to the reservoir) and diagenetic processes associated with groundwater movement during hydrocarbon migration or emplacement. An equally important use of noble gases has been in identifying the origin of non-hydrocarbon gases such as CO2 and N2 by relating these species to mantle-derived and crustal-radiogenic gases, respectively.

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Identifying and quantifying groundwater/gas/oil interaction The origin of atmosphere-derived noble gases in the subsurface. The introduction of atmosphere-derived noble gases into the subsurface is dominated by noble gases dissolved in groundwater. In the case of meteoric waters this occurs at recharge. The physical processes controlling the concentrations of noble gases in the meteoric groundwater phase are well constrained and include temperature, altitude, and salinity at recharge in addition to a small portion of air in excess of recharge equilibrium values (Kipfer et al. 2002, this volume). Figure 9 shows the typical range of Ne/Ar and Xe/Ar ratios in fresh water normalized to air-equilibrated water at 10°C.

Water associated with the sedimentary burial process (‘formation’ water) or density driven sinking plumes from highly saline lakes have both also equilibrated with the atmosphere, contain dissolved air noble gases, and contribute to the subsurface inventory (Zaikowski et al. 1987). These latter sources remain poorly constrained. Podosek et al. (1980, 1981) have shown that atmosphere-derived Kr and Xe can be preferentially trapped in shales. Recently, Torgersen and Kennedy (1999) have found correlated Kr and Xe enrichments in oil-associated natural gases that are consistent with this trapped sedimentary origin of atmosphere-derived Kr and Xe. The range, type and condition of rock sequences that preferentially trap Xe and Kr, and the conditions of their release, remain poorly constrained (Table 6). Phase fractionation. In a simple two-phase system, the recipient phase equilibrating with the groundwater will be the sampled oil or gas. The magnitude of fractionation in the oil or gas from the original groundwater values will therefore be controlled by the equilibrium solubility law (Eqns. 20, and 22) and reflect the subsurface conditions of temperature, water salinity and gas/water or oil/water volume ratio. The most uncertainty Table 6. Fractionated atmosphere-derived Ne, Kr and Xe measured in carbon-rich crustal rocks (after Torgersen and Kennedy, 1999). Reference

F(22Ne)

F(84Kr)

F(132Xe)

[36Ar] ×10-8 cm3 STP/g

Bogard et al. 1965

4.6-1010

0.7-1145

1600-45000

0.15-45

Frick and Chang 1977

7.6-15.6

220-261

2955-4345

n.a.

Podosek et al 1980

0.1-11.3

0.8-20.5

4-1300

0.1-1.0

in making this calculation is attached to the estimation of the elemental ratios before fractionation occurred. For example in the case of a meteoric water-gas system, even though a reasonable estimate of recharge temperature may sometimes be made, the amount of ‘excess air’ can be highly variable (Fig. 9). In principle, given four observables (Ne, Ar, Kr, Xe) three unknowns can be resolved using a data inversion technique similar to that developed to investigate paleotemperature calculations from noble gas concentrations in groundwater (Ballentine and Hall 1999; Kipfer et al 2002, this volume). If reasonable estimates of the subsurface conditions can be made the variables may include for example, temperature of groundwater recharge (derived from the gas phase albeit with large errors!) excess air and the gas/water volume ratio. One of the principle advantages of such an approach is the rigorous propagation of errors through to the derived values, including errors due to measurement as well as errors associated with an incomplete model formulation. Development of this technique is the focus of current research (Ballentine et al. 1999). Without a full data inversion workers to

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date have typically relied on elemental ratio pairs and an assumed range for their original value. In this respect the non atmosphere-derived noble gases can play a role in testing the robustness of a calculation because the magnitude of fractionation of non atmospherederived noble gases in the groundwater will be subject to the same fractionation process and show coherent fractionation with the atmosphere-derived gases (Fig. 14). The underlying assumptions in using this robustness check are two-fold: (1) the radiogenic noble gas composition has not itself been subject to fractionation (e.g., Ballentine and Burnard 2002, this volume); and (2) the radiogenic noble gases were indeed in the fluid system prior to the fractionating process.

Figure 14. Plot showing radiogenic 21Ne*/40Ar* vs 20Ne/36Ar and 4He/21Ne* vs. 20Ne/36Ar in the Hajduszoboszlo gas field, Hungary, after Ballentine et al. (1991). Left hand side figure: The crustal derived 21Ne*/40Ar* fractionates coherently with the groundwater-derived 20Ne/36Ar ratio (Line B) showing that these two differently sourced noble gas pairs were mixed before the fractionating processes. Right hand side figure: In contrast the crustal-derived 4He/21Ne* shows little fractionation from the predicted source ratio (Lines A). This is explained by a solubility based fractionation process, because although Ne and Ar have very different solubilities in groundwater the solubilities of He and Ne in water under subsurface conditions are similar. This will result in negligible fractionation of the He/Ne ratio.

Identifying the phase of transport. The hydrogeologic system can play an important role in the transport of gas from source to trap and identification of this process can influence exploration strategies (Toth 1980; Toth and Corbett 1986). Identification of whether or not groundwater has played a significant role in natural gas transport is possible by considering the concentration of noble gases derived from the groundwater that are now in the natural gas. For example, if we consider a groundwater that has not undergone phase separation or equilibration, its concentration of atmosphere-derived noble gases, as discussed above, fall in a relatively tight range and remain constant. The saturation limit for a major gas species (we consider here CH4) is a function of pressure and temperature (depth) as well as water salinity. The point at which CH4 saturation is reached will result in CH4 gas phase formation. The noble gases will partition between the gas and water phase. Because the volume of the gas phase at this stage is small relative to the water the relative abundance of the noble gases will be fractionated following Equation (20). The solubility of Ar however is similar to that of CH4 under typical subsurface conditions. The gas phase will therefore preserve an unfractionated CH4/36Ar ratio of the groundwater at the point (depth) at which saturation occurred (Ballentine et al. 1991; Fig. 15).

Gas reservoirs that preserve the ‘saturation’ CH4/36Ar ratio consistent with their

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formation depth may have been transported to the trapping site dissolved in the groundwater. Case studies

We review here selected case studies to illustrate some of the main applications of noble gases to natural gas and oilbearing systems. Groundwater and natural gas transport, the Pannonian basin, Hungary. The Pannonian basin forms the Great Hungarian Plain and is an extensional basin formed in the Middle Miocene that has developed into a series of deep basins separated by shallower basement blocks. Oil and gas fields are found through-out the basin, with most to-date being Figure 15. CH4/36Ar plotted as a function of depth for gas found on the basin margins or fields in the Vienna basin, Austria (Ballentine 1991), above the uplifted basement Pannonian basin, Hungary (Ballentine et al. 1991) and the blocks. Several noble gas Po Basin, Italy (Elliot et al. 1993). These values are compared with the ‘saturation’ CH4/36Ar value calculated studies have focused on the for seawater containing 7.5×10-7 cm3 (STP) 36Ar, a basin aquifer systems (Oxburgh salinity of 0.23M NaCl equivalent, a temperature gradient et al. 1986; Martel et al. 1989; of 0.03 K/km and at hydrostatic pressure (Solid line). The Stute and Deak 1989; Stute et dashed line is the saturation value if the salinity increases al. 1992). Noble gas studies of to 3 M NaCl equivalent. The gases in the Pannonian and Po basin studies lie on the saturation line for their depth, the natural gas reservoirs are by are closely linked to the groundwater system, and may Ballentine and O'Nions (1992), have exsolved from solution. These contrast with the Sherwood Lollar et al. (1994, Vienna basin gases that have had far less contact with the 1997). We focus here on the groundwater system. Hajduszoboszlo gas field study by Ballentine et al. (1991). This field is a stacked gas field producing at various intervals between 700- to 1300-m depth and occupies a portion of the sub-basin high to the north west of the Derecske sub-basin. 3He/4He, 20Ne/22Ne, 21Ne/22Ne and 40Ar/36Ar were determined as well as He, Ne and Ar abundance. The 3He/4He ratios of between 0.18 to 0.46Ra show that between 2 to 5% of the He is mantle derived. 21Ne/22Ne and 40Ar/36Ar ratios range between 0.299-0.46 and 340-1680. These are all in excess of atmosphere derived values (0.0290 and 295.5) due to the addition of crustal radiogenic 40Ar and 21Ne (40Ar* and 21Ne*) contributing 3-15% and 32-82% of the 21Ne and 40Ar respectively. The greatest contribution of crustal-derived gases is clearly correlated with the deepest samples.

Coherent fractionation is observed in a plot of 21Ne*/40Ar* (crustal component) against 20Ne/36Ar (atmosphere-derived) pointing to a fractionation process operating on the system after mixing of these differently sourced species (Fig. 14). Plotting 4He/21Ne* (crustal component) against 20Ne/36Ar no coherent fractionation is seen (Fig. 14). This is the classic pattern for a solubility equilibrium fractionation process. Negligible fractionation between 4He/21Ne* is observed because their solubilities in water under reservoir conditions are very similar. There is no resolvable mass fractionation, and diffusion as a significant transport process in this system can therefore be neglected.

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Making the assumption that the natural gas originally contained no 36Ar, the 36Ar now in the natural gas requires a minimum volume of groundwater to have interacted with the gas phase. This is a minimum estimate because this assumes quantitative degassing of the groundwater phase, yet the observed 20Ne/36Ar fractionation discussed above occurs during partial degassing. The volume of water estimated from typical groundwater 36Ar concentrations would occupy 1000 km3 of rock at 15% porosity, some 670 times larger than the current reservoir volume. • The mass balance clearly indicates an interaction between the gas now in the field and the regional groundwater system. • CH4/36Ar ratios between 1.1 to 3.5×106 in the natural gas are indistinguishable from the ‘saturation’ CH4/36Ar ratio for the depth of production and therefore consistent with gas transportation to the trapping site dissolved in the regional groundwater system (Ballentine et al. 1991; Fig. 15). Hydrocarbon migration and water-oil interaction in a quiescent basin: The Paris Basin, France. The Paris Basin provides a unique opportunity to study the characteristics of hydrocarbons that originated from a common source-rock lithology and migrated into different sedimentary layers, where they subsequently interacted with fluids of contrasting compositions. As described in the section Noble gases in ancient groundwaters and crustal degassing, the Paris Basin is a post-Variscan intra-cratonic basin in which a multi-layered aquifer system has developed. Of interest in this context are the Upper Triassic (Keuper) Chaunoy fluvial sandstones and the Middle Jurassic oolitic limestone reservoirs. The Middle Jurassic aquifer is separated from the Triassic aquifer by 400-700 m of low-permeability Lower Jurassic (Lias) mudrocks and shales, which are the source-rocks of oils in the Paris Basin. The Middle Jurassic contains lowenthalpy geothermal waters (from 50 to 80°C) and oil accumulations. The Triassic aquifer contains waters at temperatures up to 120°C associated with oil accumulations. Oil primary migration took place from the Liassic shales upward to the Jurassic limestone and laterally to the Triassic sandstone, during the Paleocene-Oligocene time (Espitalié et al. 1988). Vertical faults, affecting the Mesozoic cover of the Paris Basin and reactivated by tectonic post-Alpine stresses, have played an important role in oil secondary migration. These vertical faults constitute the preferential pathways for oil flow through the Lias. An important hydrodynamic component flowing in the Jurassic and the Triassic aquifers and contemporary with oil migration seems to have affected both the distribution of oil pools in the Paris Basin (Poulet et Espitalié 1987) and the loss of ~90% of the hydrocarbon (Espitalié et al. 1988). Cross-formational fluid flow in the Paris Basin is also apparent in the common source of salinity (halite deposited in the eastern part of the Triassic aquifer) for both Jurassic and Triassic groundwaters (Worden & Matray 1995).

Both groundwater and oil accumulations have been studied for noble gases (Marty et al. 1993; Pinti and Marty 1995, 1998; Pinti et al. 1997). The main noble gas feature in the Paris Basin fluids is the presence of a resolvable mantle-derived noble gas component, which is weaker than the one measured in the younger and tectonically active Pannonian Basin (Ballentine et al. 1991). An important in situ contribution of radiogenic noble gases is masking progressively the traces of a large-scale fluid flow, which affected the Paris Basin probably in early Tertiary time. It is likely that this episodic fluid flow introduced deep-seated mantle-derived and radiogenic noble gases into the basin, and possibly triggered the hydrocarbon primary and secondary migration within the basin (Pinti and Marty 1998). The relationship between the helium 3He/4He ratios and the total amount of 4 He in the basement, Trias and Middle Jurassic groundwaters, suggests that there are at least three sources of helium occurring in the Paris Basin. The first source is fluids circulating in the southern crystalline basement and characterized by high 3He/4He ratios (0.12-0.14 Ra) due to addition of mantle-derived 3He. The second source is Triassic

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groundwaters located at the center of the basin, which is characterized by 3He/4He ratios intermediate between the basement and the Middle Jurassic fluids (3He/4He = 0.08 Ra). The third source is water located east of the Middle Jurassic aquifer, with low 3He/4He isotopic ratios (3He/4He = 0.02 Ra) resulting from the production of helium in the local reservoir’s rocks (Bathonian-Callovian limestones). The transport of radiogenic helium, argon and neon (and associated mantle-derived helium) from the Trias to the Middle Jurassic aquifer is apparent in the distribution of the radiogenic 4He/40Ar* and 21Ne*/40Ar* isotope ratios among Triassic and Middle Jurassic oil-field brines (see section Noble gases in ancient groundwaters and crustal degassing and Fig. 28, below). The 4He/40Ar* and 21Ne*/40Ar* isotope ratios clearly show a correlation and indicate mixing between the Trias and the Middle Jurassic groundwaters. The variation of the radiogenic noble gas isotope ratios can be attributed to the initial ratio of the parent elements 238,235U, 232Th and 40K in minerals and rocks, which varies for different lithologies, or to preferential diffusion of 4He and 21Ne* relative to 40Ar* from the mineral to the fluid phase. This in turn depends on the thermal and tectonic regime of the basin. The Trias groundwaters show 4He/40Ar* ratios of 4-7 and 21Ne*/40Ar* ratios of 2.5-4×10-7. These ratios could correspond to a source having a K/U ratio of about 35,000 and which releases He, Ne and Ar in water close to their production ratio. This source could be the Triassic sandstones, the crystalline basement, or both. The second source of radiogenic noble gases has high 4He/40Ar* ratios of 40 and 21Ne*/40Ar* ratios of 65×10-7 and could correspond to the carbonate, which is characterized by very low K/U ratios. In the Middle Jurassic oils, the elemental fractionation of atmosphere-derived noble gases was found to be consistent with oil/water phase equilibrium partitioning (Pinti and Marty, 1995). In the Triassic oils, the noble gas fractionation trends indicate a more complicated history, notably involving degassing of hydrocarbons previously equilibrated with groundwaters. Pinti and Marty (1995) have interpreted this degassing episode as the result of processes of gas stripping due to oil washing (Lafargue and Barker, 1988). Calculations indicated that both Middle Jurassic and Triassic oils have seen much larger quantities of waters with oil/water ratios possibly ranging between 0.2 and 0.01, whereas the present-day oil/water ratios in the Middle Jurassic and Triassic oil fields average ~ 1. Assuming a mean groundwater residence time of few Ma in the center of the Paris Basin (Marty et al., 1993), where most of the oil accumulations reside, and an integrated mean oil/water ratio in oil reservoirs lower by one order of magnitude than those presently observed, then the residence time of oils in their reservoirs should also be an order of magnitude higher than those of flowing waters and could be of the order of ~20-40 Ma. Such a figure is in qualitative agreement with current estimates for the timing of oil migration in the Basin (Paleocene-Oligocene, Poulet et Espitalié 1987; Espitalié et al. 1988). Groundwater and diagenesis, the Magnus oil field, North Sea. The Magnus oilfield is located in the East Shetland Basin, northern North Sea. The field consists of a single oil phase with no associated gas cap and contained an estimated in-place oil reserve of 2.65×108 m3 oil (STP). Hydrocarbon accumulation occurs in Middle Jurassic sandstones located on the dipping flank of a tilted Jurassic fault block. Petrographic and isotopic evidence from diagenetic minerals show that minerals in the crest of the reservoir grew in pore water containing significantly more meteoric water than those down dip, which are dominated by seawater (Emery et al. 1993; Macaulay et al. 1992). Cementation of the Magnus sandstone appears to have occurred concurrently with reservoir filling at ~72-62 Ma (Emery et al. 1993). Models addressing the role of groundwater in effecting regionally observed cementation of oil-bearing systems appeal to either local dissolution

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and reprecipitation or require the regional flow of groundwater (e.g., Bethke et al. 1988; Gluyas and Coleman 1992; Aplin et al. 1993). Because the Magnus system filling occurred at the same time as the cementation it is reasonable to assume that the noble gases in the oil phase preserve a record of the groundwater volumes during the quartz precipitation. Depressurization during production results in both gas and oil phases being present at the surface. A pilot study determined that He, Ne and Ar are almost quantitatively partitioned into the gas phase under separator conditions. Given the flow rate of both oil and gas, the analyses of the gas phase alone enables an accurate reconstruction of the subsurface and single oil phase noble gas composition. Gas samples across the Magnus field were taken and the He, Ne and Ar isotopic ratios and abundances in the oil were determined (Ballentine et al. 1996). Both the He and the Ne isotope systematics require a contribution from a mantle source. If the mantle end-member is modeled using midocean ridge values (Graham 2002, this volume), 2.3 to 4.5% of the 4He and 4.3 to 6.2% of the 21Ne in the Magnus oil is mantle-derived. The remainder of the 4He and 9.0 to 12.0% of the 21Ne is crustal-radiogenic, and the remaining 21Ne is atmosphere-derived. The quantity of radiogenic noble gas associated with the Magnus oil/groundwater system can only be accounted for by production predominantly from outside the volume of the Magnus Sandstone aquifer/reservoir drainage area and the associated Kimmeridge Clay source rock formation and together with the mantle-derived noble gases, provides strong evidence for cross formational communication with deeper regions of the crust. This is not the case for the groundwater-derived noble gases. The 20Ne and 36Ar have been input into the oil phase by interaction with an airequilibrated groundwater. Because the Magnus oil field has no gas cap, this is a simple two-phase system. Similarly, because the groundwater, with the exception of a small amount of meteoric water incursion at the crest of the system, is dominated by seawater there is no variable excess air component to consider. In principle, knowing the original noble gas concentration in the seawater and the temperature and salinity of water on equilibration with the Magnus oil, together with either the 20Ne or 36Ar concentration and their respective solubilities in the oil enable Equation (19) to be used to determine the system oil/water volume ratio. In practice the noble gas solubility database for different oils is limited, and KmAr(oil) and KmNe(oil) are not known for the Magnus oil. Nevertheless noble gas solubility data is available for two oils of different density (Kharaka and Specht 1988; Table 4). If it is assumed that at any one temperature and over a small density range the relative change in solubility of both Ne and Ar is proportional, an equation can be developed that links the KmAr(oil) to KmNe(oil) in the Magnus oil (Ballentine et al. 1996). This then leaves two unknowns, the linked solubility term and the oil/water volume ratio. With two separate equations for 20Ne and 36Ar derived from the general Equation (19), these can be solved. • Noble gas partitioning between a seawater-derived groundwater and the oil phase at the average Magnus Sandstone aquifer temperature requires a subsurface seawater/oil volume ratio of 110(±40) to account for both the 20Ne and 36Ar concentrations in the central and southern Magnus samples. • The volume of groundwater that has equilibrated with the Magnus oil is indistinguishable from the static volume of water estimated to be in the down-dip Magnus aquifer/reservoir drainage volume. This suggests that the Magnus oil has obtained complete equilibrium with the groundwater in the reservoir drainage volume, probably during secondary migration, and further suggests that the concurrent cementation of the Magnus sandstone aquifer has occurred with little or no large-scale movement of groundwater through the aquifer system. Tracing the CO2 source in the west Texas Permian basin, USA. CO2 in natural

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gases can originate from a number of sources including methanogenesis and oil field biodegradation, kerogen decarboxilation, hydrocarbon oxidation, decarbonation of marine carbonates and degassing of magmatic bodies. The δ13C(CO2) signature can be used to distinguish between these different sources, with the exception of magmatic and carbonate-derived CO2, which have overlapping δ13C(CO2) (Jenden et al. 1993; Sherwood Lollar et al. 1997). Although methanogenesis or oil field biodegradation can sometimes result in gas fields with up to 40% CO2 by volume, kerogen decarboxilation and hydrocarbon oxidation rarely result in gas containing more than a few percent CO2. This is in distinct contrast to decarbonation/magmatic sources, which can result in gas fields containing up to 100% by volume CO2. Noble gases can be used to distinguish between decarbonation/magmatic sources (Fig. 16; Sherwood Lollar et al. 1997; Ballentine et al. 2001).

Figure 16. CO2/3He vs. fraction of CO2. The main figure shows the range of CO2/3He values found in pure magmatic samples. CO2/3He values above this range, irrespective of CO2 content, can only be attributed to a CO2 source containing no 3He and provides an unambiguous identification of crustal-sourced CO2. Values within this range or below contain a magmatic CO2 component but have been subject to possible CO2 loss, dilution (e.g., addition of CH4 or N2), and/or crustal CO2 addition (after Sherwood Lollar et al. 1997). Inset shows the values found in CO2 rich natural gases in the JMBB field, west Texas Permian basin, which vary within the magmatic range (after Ballentine et al. 2001). Vectors A, B and C show the effect of crustal CO2 addition, CO2 loss through reaction or precipitation and dilution respectively. Near constant δ13C(CO2) rules out either loss or addition of CO2 and requires the range to be due to magmatic source variation (Fig. 18).

The west Texas Permian basin was formed as a result of the late Palaeozoic collision of South America with North America that also resulted in the Marathon-Ouachita orogenic belt as well as widespread interior continental deformation. The Val Verde basin is a foreland sub-basin of the west Texas Permian basin, and lies between the Central basin platform and the Marathon thrust belt (Fig. 17). Natural gas reservoirs show a systematic regional increase in CO2 content towards the Marathon thrust belt,

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varying from an average of about 3% in the basin center to as high as 97% on the foredeep margin of the thrust belt. The main producing formation in the JM-Brown Bassett (JMBB) field is brecciated Ordovician Ellenberger dolomite. This field reflects the regional spatial trend in natural gas CO2 content, with samples increasing from 20% to 55% CO2 towards the Marathon thrust belt. The remaining gas is dominated by CH4.

Figure 17. Showing coherent spatial variation in %CO2 and CO2/3He in CO2-rich natural gases in the Val Verde basin, part of the west Texas Permian basin (after Ballentine et al. 1991). Arrows show the direction of the regional increase in CO2 content and CO2/3He ratio towards the Marathon thrust belt. Inset shows the location of the Val Verde basin relative to the major Permian uplift and basinal features. Basins: 1, Delaware; 2, Midland; 3, Palo-Duro; 4, Anadarko; 5, Arkoma; 6, Ft Worth; 7, Kerr. Uplifts: A, Sierra Diablo; B, Central basin; C, Ozona; D, Concho arch; E, Llano; F, Devils River.

Samples from across the field were analyzed for their C, He, Ne and Ar isotopes as well as the abundance of He, Ne, Ar and major gas species. 3He/4He varies between 0.24 and 0.54 Ra. If it is assumed that the magmatic and crustal components have 8.0 and 0.02 Ra composition respectively, the measured ratios correspond to between 3.2 and 6.8% of the helium being derived from the mantle. In reality, sub-continental mantle 3He/4He ratios are believed to be slightly lower than the mantle supplying mid-ocean ridges (Dunai and Baur 1995) and the percentage mantle contributions to the JMBB are

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therefore a lower limit. Similarly from the 21Ne/22Ne and 40Ar/36Ar ratios the crustal components can be resolved (4He*, 21Ne* and 40Ar*) showing 4He*/40Ar* and 4 He*/21Ne* ratios indistinguishable from average crustal production (Ballentine and Burnard 2002, this volume). This result indicates that no fractionation process has operated on these species either during release from their source, on transport, or during their residence time in the trapping structure (eliminating for example, significant diffusive loss since trapping). It was not possible to investigate the groundwater-derived noble gases due to significant amounts of air contamination on sampling (sampling procedures used were for stable isotopes, not noble gases).

Figure 18. Evolution of δ13C(CO2) and CO2/3He calculated for the gas phase of a degassing magma body (after Ballentine et al. 1991). Two models are shown: (i) The composition of the gas evolving from the magma by a Rayleigh fractionation process; and (ii) The composition of this gas in an accumulating reservoir. The model undegassed magma is taken to have δ13C(CO2) = -4.7‰ and CO2/3He = 2×109, and is within the range estimated for the mantle source. The tick marks are the percentage loss of CO2 from the magma body. δ13C(CO2) fractionation between a CO2 gas phase and magma is taken to be 2‰ (Mattey 1991), and the relative solubility of He/CO2 = 5 (Bottinga 1991). Both CO2/3He and δ13C(CO2) of the JMBB field are consistent with partial degassing of the source magma body (Ballentine et al. 2001). In this context, samples with the highest CO2/3He are from the earliest stages of outgassing.

The 3He/4He correlates directly with percent CO2, showing clear two-component mixing between the hydrocarbon gas containing crustal-derived He and a CO2 component with an elevated 3He/4He. CO2/3He for all samples are within the magmatic range but vary systematically with percent CO2 (Fig 16). Various models were investigated to account for the CO2/3He variation, including crustal CO2 addition, precipitation or a combination of the latter combined with CH4 addition/dilution. None of these models were able to satisfy both the very small variation in δ13C(CO2) and the mixing vectors shown in Figure 16. A magma-degassing model was constructed that accounted for both δ13C(CO2) and CO2/3He (Fig. 18). In the context of a magma-degassing model (Figs. 13 and 18), samples with the highest CO2/3He are from the earliest stages of outgassing and are located closest to the Marathon thrust belt. A simple filling model in which reservoirs closest to the magma source are filled and then diluted by subsequent outgassing (lower CO2/3He) predicts that the highest CO2/3He ratios are furthest from the degassing magma,

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ruling out the Marathon thrust belt as the source of magmatic CO2 in the Val Verde basin. Tertiary volcanism is associated with the Basin and Range province to the west of the Val Verde basin, and some 100 km away from the JMBB study. Although this is a potential source of magmatic volatiles, this source is not consistent with the inferred direction of filling (Fig. 17). The increase in CO2 content as the MTB is approached can be accounted for if CO2 emplacement pre-dates CH4 generation in the hydrocarbon `kitchens' to the north of the gas fields. Assuming simple filling, the traps closest to these `kitchens' would have the highest CH4 content. CO2 charging, therefore, pre-dates the onset of hydrocarbon generation in the basin, which occurred about 280 Myr ago. Maximum uplift of the Central basin and Ozona platforms (Fig. 17) occurred between 310 and 280 Myr ago in response to the MTB loading. Associated deep volatile release would provide the appropriate timing, mechanism and required spatial consistency to be the source of the magmatic CO2 preserved in the Val Verde basin. • The Rayleigh fractionation model proposed to account for regional CO2/3He variation provides an important tool to identify the direction of magmatic CO2 input into a basin system; this model also accounts for the higher CO2/3He and heavier δ13C(CO2) often found in intracrustal manifestations of magmatic gas (Griesshaber et al. 1992; Weinlich et al. 1999) compared with the values in pristine mantle samples (Javoy and Pineau 1991). • Diffusion experiments on other systems have been used to estimate the residence time of natural gas in a trapping structure (Kroos et al. 1992; Schlomer and Kroos 1997). The age of emplacement inferred from this study suggests that calculations of natural gas residence times based on these diffusion experiments seriously underestimate the storage efficiency of some trapping structures, and provide support for the viability of natural gas exploration in deeper, older, and therefore more unconventional, locations. The origin of 4He-associated N2 in the Hugoton-Panhandle giant gas field, USA. The most abundant non-hydrocarbon gas in sedimentary basins is nitrogen. In the USA, 10%, 3.5% and 1% of natural gases contain >25%, >50% and >90% by volume nitrogen respectively (Jenden and Kaplan 1989). Nevertheless, the dominant sources and mechanisms responsible for focusing and enrichment of nitrogen within natural gas fields are poorly constrained. This is in part due to the multiple sources of nitrogen in the subsurface including atmosphere-derived nitrogen dissolved in groundwater, nitrogen released from sedimentary organic matter, nitrogen released from metasediments during metamorphism and, in areas of magmatic activity, an igneous or mantle nitrogen origin. The overlapping range of nitrogen isotopic values for the respective systems has meant that nitrogen isotopes alone cannot be used to quantify the contribution of these different sources to natural gas systems. Nitrogen gas associated with high radiogenic 4He concentrations is particularly common (Gold and Held 1987; Jenden and Kaplan 1989; Jenden et al. 1988; Pierce et al. 1964; Poreda et al. 1986; Stilwell 1989; Hiyagon and Kennedy 1992; Hutcheon 1999). Because of the association of 4He, and therefore other crustal noble gases such as 21Ne* and 40Ar*, the noble gases are particularly appropriate for tracing the origin of He-associated nitrogen (N2*) (Ballentine and Sherwood Lollar 2002).

The Hugoton-Panhandle giant gas field is the case type example of a system containing N2* (Pierce et al. 1964). Extending 350 km across SW Kansas and the Oklahoma/Texas Panhandles this field contained more than 2.3×1012 m3 (STP) of recoverable gas, and produces from Permian carbonates between 400-900m depth on the south and western margins of the Anadarko basin. The isotopic compositions of the hydrocarbon gases across the entire field are indistinguishable, are hence co-genetic and

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reasonably originate from the Anadarko sedimentary basin. N2 concentrations throughout the field vary between 5-75%, averaging ~15%. The highest concentrations of nitrogen in the Texas Panhandle are found on the SSW margin of the field, on the side of the field furthest from the Anadarko basin hydrocarbon ‘kitchen’. In the Oklahoma and Kansas Hugoton, the highest nitrogen content is found to the north and west of the field, again on the opposite edge of the field to the Anadarko basin. The nitrogen is locally proportional to the 4He content, although the 4He/N2 systematically increases from 0.02 in the Kansas Hugoton to 0.077 in the Texas-Panhandle (Gold and Held 1987; Jenden et al. 1988; Pierce et al. 1964). Ballentine and Sherwood Lollar (2002) show that the nitrogen isotopic composition also changes systematically in this field, decreasing from δ15N(N2) = +9.4‰ in the Kansas Hugoton to δ15N(N2) = +2.7‰ in the Texas-Panhandle. 3He/4He, 21Ne/22Ne and 40 Ar/36Ar ratios enable noble gas contributions from mantle, crustal and groundwater sources to be resolved and quantified in the samples. Crustal radiogenic 4He/21Ne* and 4 He/40Ar* ratios show a 60% excess of 4He compared to predicted crustal production values, and are typical of noble gases released from the shallow crust (Ballentine and Burnard 2002, this volume). Although significant and resolvable mantle and groundwater-derived noble gases are present, mantle 3He/N2 and groundwater 36Ar/N2 values rule out significant mantle or atmosphere contributions to the gas field N2, which is crustal in origin.

4

He/N2

Figure 19. Plot of 20Ne/N2 vs. 4He/N2 in natural gases from the Giant Hugoton-Panhandle gas field in Texas-Oklahoma-Kansas, USA after Ballentine and Sherwood Lollar (2002). This natural gas field is the case-type system in which N2 content is related to 4He concentration. Most samples fall on a line indicating simple two-component mixing between one nitrogen component that is associated with both crustal 4He and groundwater-derived 20Ne (N2*) and another nitrogen component that has no resolvable association with any noble gases. Identifying one He/N2 end-member ratio enables the relative contribution of these two nitrogen components to any one sample to be calculated. The nitrogen isotopes also vary systematically with He/N2 and from the noble gas mixing relationship, the end-member nitrogen isotopic compositions can also be determined (see text).

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A plot of 4He/N2 vs. 20Ne/N2 shows that almost all samples lie on a simple mixing line between two crustal nitrogen components (Fig. 19). One N2 component is associated with the crustal 4He and groundwater-derived 20Ne (N2*). The other nitrogen component has no resolvable association with either crustal- or groundwater-derived noble gases. An end-member 4He/N2* = 0.077 (Pierce et al. 1964) is used to define the He-associated component and enables the relative contribution of non-He associated nitrogen to each sample to be calculated (Fig. 19). In turn this enables the δ15N(N2) for each end-member to be calculated, where δ15N(N2)=-3‰ and +13‰ for N2* and the non-He associated N2 respectively. The δ15N(N2*) value is not compatible with a crystalline or high grade metamorphic source and, similar to the 4He, probably originates from a shallow or low metamorphic grade source rock. 4He mass balance nevertheless requires a regional crustal source; its association with a resolvable magmatic 3He contribution pointing to a source to the recently active Sierra Grande uplift to the west of the gas field, in the opposite direction to the Anadarko basin hydrocarbon source which is unlikely to be a source of magmatic 3He. The ratio of radiogenic 4He to groundwater-derived 20Ne is almost constant throughout the entire system and clearly indicates a link between the crustal-derived 4He (and hence the N2*) and the groundwater system. To place perspective on the volume of groundwater constrained by the 20Ne mass balance this is equivalent to the water in a 100m thick static aquifer that covers three times the area of the Anadarko basin. This in no way suggests that this is the source of the groundwater (see below) but clearly shows that a regional groundwater system is involved and that this can provide the collection, transport and focusing mechanism for a diffuse crustal 4He and N2* flux. The N2*/20Ne ratio is three times lower than the gas field ‘saturation’ ratio (in the same way as 36 Ar/CH4) and shows that the N2* transport must be in the aqueous phase. The degassing mechanism of the groundwater can be accounted for by contact between the regional groundwater system and a pre-existing reservoir hydrocarbon gas phase. • The regional groundwater cover traps 4He and N2* released from shallow sediments and low grade metamorphic rocks during the thermal hiatus generated by the Sierra Grande uplift (source of 3He) some 200-300 km to the west of the HugotonPanhandle. • West-East migration of the groundwater (note the Anadarko basin is to the North and east of the gas field) containing the dissolved magmatic and crustal-derived gases contacts a pre-existing gas phase in the form of the Hugoton-Panhandle gas field resulting in degassing of the groundwater. • Noble gases have enabled the resolution of two different nitrogen sources and their respective isotopic compositions in this field and generated a simple model to account for the source, transport and relative timing of the 4He-associated gas input into the gas field. Sediment-derived atmospheric noble gases in the Elk Hills oil field, USA. The Elk Hills oil field is located in the southern San Joaquin valley, and is located about 30 km southeast of Bakersfield, California, USA. Production to date from the Elk Hills anticline has exceeded 2 billion barrels of oil from five producing intervals ranging from the ‘Dry Gas Zone’ (2-5 Ma) to the ‘Santos Oil Zone’ (30-35 Ma). Samples were collected from gas-oil separators and 4He, 36Ar, 22Ne, 84Kr and 132Xe abundance determined (Torgersen and Kennedy 1999). These workers did not tabulate the isotopic compositions, with the exception of 4He and 40Ar*, but report that the Ne, Kr and Xe isotopes are consistent with an atmosphere-derived source. Torgersen and Kennedy give the concentrations normalized to 36Ar and relate this to the air value to obtain a fractionation value where F(nNg)= {[nNg]/[36Ar]sample}/{[nNg]/[36Ar]air}. There is an increasing enrichment related to

Tracing Fluid Origin, Transport and Interaction in the Crust

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atomic number where F(22Ne) < F(84Kr) < F(132Xe). Although such enrichments have been previously noted in oil-related systems (Bosch and Mazor 1988; Hiyagon and Kennedy 1992; Pinti and Marty 1995), the Elk Hills study shows an average Xe enrichment factor of ~30, with the highest ratio 576 times the air ratio representing the highest 132Xe/36Ar ratio yet measured in a terrestrial fluid.

Figure 20. The Kr/Ar and Xe/Ar ratios normalized to air values to give F(84Kr) and F(132Xe) from the Elk Hills oil wells (California, USA) are plotted after Torgersen and Kennedy (1999). The filled triangles represent the composition of air and air-equilibrated water at 20°C (ASW). The open triangles represent the ratios predicted for oil-water equilibration at different oil/water volume ratios, and the shaded region labeled ‘Rayleigh Fractionation’ the range of values predicted after continuous gas loss from either the water or oil phase. The dashed line shows the weighted least squares fit through the Elk Hills data, forced through the ASW value. This data cannot be accounted for by solubility equilibrium or open system fractionation of groundwater-derived atmospheric noble gases. The data is explained by a trapped sedimentary source for atmospheric Kr and Xe that is released into the oil phase during oil formation and primary migration. The values measured in the Elk Hills samples can be compared with those measured in carbon rich extracts from cherts measured by Frick and Chang (1977).

F(84Kr) is strongly correlated with F(132Xe), but unlike earlier observations, such as the Paris Basin (Pinty and Marty 1995), the Elk Hills data cannot be attributed to fractionation from an oil-water-gas system (Fig. 20). Torgersen and Kennedy argue that the extreme values observed in the Elk Hills system are due to the preferential trapping of Xe > Kr by carbon rich sediments (the oil source rocks) which are subsequently released from the sediment during oil formation and primary migration. This argument is supported by observations of very high F(84Kr) and F(132Xe) values in a variety of carbon rich rocks including chert, thucolite and shale (Bogard et al. 1965; Frick and Chang 1977; Podosek et al 1980; Table 6). A model is proposed that considers mixing and dilution of the highly fractionated sediment-derived Kr and Xe with groundwater-derived species, and relates the decrease in F84Kr and F132Xe with an increase in 36Ar concentration in the hydrocarbons that is proportional to the groundwater/hydrocarbon ratio. This correlation is not yet well established.

580

•

Ballentine, Burgess & Marty

The identification of sediment-derived atmospheric Kr and Xe in some sample types means that the extension of solubility fractionation models (that assume a priori a groundwater origin for all atmospheric noble gases) to include Kr and Xe has to be assessed with caution on a case-by-case basis. 3

He/ 4He closure and 20Ne/ 36Ar fractionation in the Indus basin, Pakistan. The Indus basin is bound by the Pakistan foldbelt to the northwest and the Indian shield to the southeast and extends NE-SW for over 1200 km. The basin can be separated into Southern, Middle and Upper sub-basins. Rifting from Triassic times until collision with the Afghan blocks during the upper Cretaceous and Paleogene generated both the Middle and Southern Indus sub-basins. Hydrocarbon trapping structures may have been formed as early as the Cretaceous but most were formed by post collisional deformation during the Pliocene between 4.5-3.5 Ma. The sedimentary infill is dominantly marine in origin, with hydrocarbon source rocks identified mainly in Jurassic, Cretaceous and Eocene sequences. All potential source rocks have reached a high degree of maturity in the study area, with vitrinite reflectance values of between 0.85 and 0.93 Ro. There are no trapping structures producing oil and hydrocarbon reservoirs are dominated by thermogenic methane. A few fields in the middle Indus sub-basin in addition to hydrocarbon gas contain up to 70% CO2 and 23% N2. Battani et al. (2000) collected gas samples from the Middle and Southern Indus sub-basins and determined their composition, d13C(CO2, CH4, C2-C4), 4He, 20Ne, 36Ar, 3He/4He and 40Ar/36Ar. 3

He/4He ratios range between 0.009 and 0.056 Ra and are dominated by crustal radiogenic He. Battani et al. use this observation to rule out any significant magmatic contribution to either the N2 or CO2. They note that many extensional systems contain a resolvable magmatic He contribution (Oxburgh et al. 1986; Poreda et al. 1986; Ballentine et al. 1991), and in particular that old trapping structures can preserve a mantle 3He/4He signature for tens if not hundreds of million years (Ballentine et al. 1996; also see Ballentine et al. 2001). Battani et al (2000) argue that the lack of a resolvable magmatic 3 He/4He signature in the extensional Indus basin can only be explained if the basinal fluid system was open during rifting, and that no significant amounts of magmatic fluid have been preserved in the present day hydrocarbon traps. This is entirely consistent with the late development of trapping structures within the basin. Figure 21. Plot of 20Ne/36Ar vs. 1/36Ar for natural gases from the middle Indus subbasin, Pakistan (after Battani et al. 2000). Mixing is observed between two components, one with high 20Ne/36Ar and low 36Ar and the second with 20Ne/36Ar similar to unfractionated groundwater. Battani et al. have explored a variety of models to account for the high 20Ne/36Ar and conclude that the most viable mechanism is a two-stage process involving an oil phase. The first stage is Rayleigh fractionation of the residual noble gases in the groundwater after interacting with oil. The second stage is transfer of a small amount of the noble gases, accompanied by further fractionation, into the gas phase. Model conditions are discussed in the text.

Tracing Fluid Origin, Transport and Interaction in the Crust

581

20

Ne/36Ar ratios and 36Ar concentrations provides another important perspective on this system. In a plot of 20Ne/36Ar vs. 1/36Ar, there is clear mixing between a gas component that is characterized by low 36Ar and high 20Ne/36Ar values, and a second component that has higher 36Ar and groundwater 20Ne/36Ar values (Fig. 21). Battani et al interpret the low 20Ne/36Ar values to have an unfractionated groundwater source. The origin of the high 20Ne/36Ar values, up to 1.3 compared with a groundwater range of 0.13 to 0.19, is investigated in more detail. Other workers have also observed high 20Ne/36Ar values in both natural gases (e.g., Ballentine et al. 1991) and waters (Castro et al. 1998). Battani et al. note that a single stage equilibrium model for fractionation between groundwater and a gas (Bosch and Mazor 1988; Ballentine et al. 1991) cannot produce 20 Ne/36Ar values greater than ~0.6. This limit also applies for the gas phase during Rayleigh fractionation of a simple water-gas system. Battani et al. also investigate the possibility of a process of re-solution and effervescence. This would result in high 36Ar being correlated with high 20Ne/36Ar in the gas phase and can be discounted. Battani et al. (2000) argue that if groundwater is first equilibrated with an oil phase, the concentration of atmosphere-derived noble gases in the water phase will be reduced and the 20Ne/36Ar ratio increased due to the higher relative solubility of Ar in the oil phase. If this process is by a Rayleigh distillation process, the fractionation in the residual water can be extreme. Subsequent gas-water interaction can produce much more highly fractionated 20Ne/36Ar ratios than a single water-gas phase system, and the most highly fractionated values would correlate with the lowest 36Ar concentrations. Battani et al. model this process for a variety of conditions. They consider for example, the conditions required to account for the CO2-N2 rich gases in the middle Indus. The oil-water distillation process is taken to have left 30% of the 36Ar in the groundwater at ~120°C and at a depths of 3700m. Subsequent gas-water Rayleigh fractionation is modeled as leaving 90% of the 36Ar in the groundwater phase, at a temperature of 29°C and at a depth of 630 m (Fig. 21). There are clearly many variables in this modeling process that include the density of the oil and the salinity of the groundwater in addition to the parameters that have been varied to obtain a fit to the data set. In this respect it is not straightforward to assess the uniqueness of the model parameters used. Nevertheless, the model is sensitive to the depth of the gas water equilibration (pressure has a large effect on the water/gas volume ratio) and the model depths of 60° the cutoff rigidity drops below the minimum rigidity of cosmic ray particles existing within the solar system. Therefore at such high latitudes the cosmic ray flux remains constant. •

In the atmosphere cosmic rays are attenuated as a result of the interactions they undergo. The resulting decrease in cosmic ray flux is approximately exponential: N = N0 e–d/Λ

(2)

Here, N0 and N are the numbers of particles at the top of the atmosphere and at the location of the observer, respectively, d is the atmospheric depth (expressed in units of g/cm2), and Λ is the attenuation length or mean free path (in the same units). Λ is not a constant but changes with latitude, due to the modification of the cosmic ray energy spectrum depending on the cutoff rigidity, and is also slightly variable with altitude (e.g., Lal 1991). •

The flux direction of cosmic ray secondary particles in the lower atmosphere is distributed symmetrically around an intensity maximum in vertical (i.e., zenith) direction. The angular intensity distribution is given by J(θ,ϕ) = J0 sinm θ

for θ ≥ 0

(3a)

J(θ,ϕ) = 0

for θ < 0

(3b)

(e.g., Heidbreder et al. 1971), where J0 is the intensity in vertical direction, θ is the inclination angle measured from the horizontal, and ϕ is the azimuthal angle. For the exponent m a value of 2.3 is most widely used (e.g., Nishiizumi et al. 1989; Dunne et al. 1999). Other estimates are 3.5 ± 1.2 (Heidbreder et al. 1971) and 2.65 (Masarik et al. 2000). Scaling methods for latitude and altitude variation. To accommodate variations of cosmogenic nuclide production rates over the face of the globe, accurate scaling procedures are required. Early studies relied directly on observed distributions of cosmic ray intensities (e.g., Yokoyama et al. 1977; Porcelli et al. 1987; Brown et al. 1991; Staudacher and Allègre 1991). Based on the cutoff rigidity at the sampling location (e.g., Shea et al. 1987), they estimated the prevailing cosmic ray flux relative to some other coordinate position and applied an altitude correction according to Equation (2), with an attenuation length Λ appropriate for the respective latitude. Although this method is correct in principle, it has some disadvantages in practice. First, it can be quite complex searching the literature for applicable data and deriving a suitable scaling factor; in any case one will have to resort to approximations. Second, different people will use different

Noble Gases as Dating Tools for Surficial Processes

737

data and different approximations, therefore the methods used by various researchers will not be easily comparable. The introduction of an easily treatable formalism applicable to all altitudes and latitudes by Lal (1991) was thus soon accepted as a sort of standard method to convert production rates from one location to another. Lal’s method is based on third-order polynomials in altitude fitted to nuclear disintegration rates in the atmosphere (Lal and Peters 1967), which are given for geomagnetic latitudes from 0° to 60° with a 10° spacing. The polynomial coefficients are reproduced in Table 2. The production rate of any spallation-produced nuclide in a surface rock is expected to be proportional to the nuclear disintegration rate, so if the production rate at one location is known it can be scaled to any location. If Pn is the production rate of a nuclide at sea level and high latitudes (≥60°), the place usually taken for normalization, then P(λm,h) = N(λm,h) × Pn / 563.4

(4)

where P(λm,h) is the production rate at geomagnetic latitude λm and altitude h and N(λm,h) is the nuclear disintegration rate at the same location as calculated from Lal’s (1991) polynomials. Figure 2a displays the resulting altitude and latitude dependence. Lal’s (1991) scaling method has recently been challenged by Dunai (2000a), who claims that some of the approximations inherent in Lal’s procedure may lead to large systematic errors, especially for high altitudes and for latitudes around 30°. Among these approximations are the description of the geomagnetic field by an axial dipole and the extrapolation of attenuation path lengths from high to low altitudes. Dunai has proposed a different scaling method, in order to avoid such critical approximations. To include the non-dipole components, Dunai uses the geomagnetic field inclination instead of the geomagnetic latitude as the parameter describing the field and then derives the dependence of the neutron flux on latitude and altitude by fitting two separate fiveparameter sigmoidal functions to the observational data according to N 1030 (I) = Y + Λ(I) = y +

A [1 + exp((X − I) / B)]C

a [1 + exp(( x − I) / b )] c

(5) (6)

(Dunai 2000a, 2000b). Here, N1030(I) is the sea level (1030 g/cm2 atmospheric pressure) nuclear disintegration rate (or production rate) relative to I = 90°, Λ(I) is the attenuation path length in g/cm2, and I is the inclination in degrees. For the case of a dipole field, I can easily be calculated from the geomagnetic latitude λm using the relation tan I = 2 tan λm

(7)

whereas in other cases an average inclination value must be derived from paleomagnetic records. The parameters A, B, C, X, Y and a, b, c, x, y are given in Table 3. The nuclear disintegration rate (or production rate) at altitude h is then given by N(Δd,I) = N1030(I) × eΔd/Λ(I)

(8)

with Δd being the difference in atmospheric depth between sea level and altitude h, which is proportional to the difference in atmospheric pressure Δp (in SI units: Δd = Δp/g, where g = 9.80665 m/s2 is the standard gravitational acceleration). The conversion of altitude to atmospheric depth is described in the Appendix of Dunai (2000a). Although the calculation according to Dunai’s method is somewhat more complicated, it shares with Lal’s method the advantage of offering a consistent set of scaling parameters which enables direct comparison of results obtained by various workers from locations all over

738

Niedermann Table 2. Polynomial coefficients for the calculation of the nuclear disintegration rate N(λm,h) at geomagnetic latitude λm and altitude h in the atmosphere according to Lal (1991): N(λm,h) = a0 + a1h + a2h2 + a3h3, with h in km and N in g–1a–1.

λm [°]

a0

a1

a2

a3

0 10 20 30 40 50 60-90

330.7 337.9 382.1 469.3 525.6 571.1 563.4

255.9 252.1 272.1 394.6 505.4 588.1 621.8

98.43 111.0 132.5 97.76 142.0 170.9 177.3

20.50 20.73 24.83 47.20 58.87 76.12 78.91

25

25

Relative Production Rate P/P0

1.8

1.8

≥60° 40° 20

1.4

≥60° 50°

30°

20

0.6

0.9

30° 20°

0.6 0.0

0.3

0.6

0.9

0°

10

0°

5

5

Lal (1991) 0

40°

0°

15

20°

10

50°

30° 20°

30° 0.3

1.4

1.0

15 0.6 0.0

40°

40°

20° 0°

1.0

≥60°

≥60°

0

1

2

3

4

Dunai (2000a) 5

0

0

Altitude [km]

1

2

3

4

Altitude [km]

Figure 2. Dependence of cosmogenic nuclide production rates on altitude and latitude, as derived from the scaling methods of Lal (1991) and Dunai (2000a), respectively. Production rates are normalized to P0, the value at sea level and high latitudes. Though the curves look similar at first glance, differences between the two scaling methods become increasingly evident at higher altitudes. Near sea level, the largest differences occur at latitudes 20-40° as shown in the insets.

Table 3. Numerical values of the coefficients for the calculation of the nuclear disintegration rate N1030(I) and the attenuation length Λ(I) according to Equations (5) and (6), respectively (Dunai 2000a). Coefficients for N1030(I) A B C X Y

0.4450 4.1703 0.3350 62.698 0.5555

Coefficients for Λ(I) a b c x y

19.85 -5.430 3.590 62.050 129.55

5

Noble Gases as Dating Tools for Surficial Processes

739

the world. In Figure 2b, the altitude and latitude dependence of cosmogenic nuclide production according to Dunai (2000a) is shown and can be compared to that for Lal’s (1991) method. Although the curves are quite similar, distinct differences by more than 10% occur at high altitudes (> 3-4 km) and, for sea level (visible in the insets of Fig. 2), at 20-40° latitude. Desilets et al. (2001) claim that Dunai’s (2000a) scaling model is based on several false assumptions and thus does not provide an improvement over Lal’s (1991) method. In a reply, Dunai (2001a) gives reasons for his choices and maintains his view regarding the correctness of the procedure. It is not the purpose of this review to evaluate the arguments of either side nor to give recommendations as to which scaling method should be preferred in future. According to Desilets and Zreda (2001), both methods have shortcomings as they do not adequately appreciate the influence of effects such as energy-dependence of the nucleon attenuation length, energy sensitivity and background correction of instruments used for cosmic ray monitoring, parameters of the real geomagnetic field, and solar activity. However, no alternative scaling model is provided by these authors. Whichever method is preferred, it is important to remember that consistent sets of production rates and scaling methods should be used. It is, for example, not advisable to take an experimentally determined production rate which had been scaled to sea level and high latitude using Lal’s method and then scale it to the sampling location by Dunai’s method. For this reason sea level/high latitude production rates will be quoted according to both methods in the Production rates of cosmogenic nuclides section. Both the scaling procedures of Lal (1991) and Dunai (2000a) are strictly only valid for the spallation-produced component of cosmogenic nuclides. If production by muons cannot be neglected, their contribution must be scaled for altitude independently because of their higher attenuation path length. Lal (1991) offered explicit polynomial coefficients for the altitude and latitude dependence of the 10Be and 26Al production rates, which included a “muogenic” contribution of 15-17% at sea level (cf. Nishiizumi et al. 1989). However, according to Brown et al. (1995a) that contribution is much lower, only ~ 1-3%. Therefore it is not advisable to use the data of Lal’s (1991) Table 1 to scale 10Be and 26Al production rates, as the scaling error is probably smaller when the muonproduced component is completely neglected. Dunai (2000a) gives a modified form of Equation (8): N(Δd,I) = (1-α) N1030(I) × eΔd/Λ(I) + α N1030(I) × eΔd/Λµ

(9)

where α is the fraction of the nuclide produced by muons at sea level and Λµ = 247 g/cm2 (Lal 1988). Stone (2000) has pointed out an additional difficulty in scaling production rates. Strictly, they do not depend on altitude, but on air pressure, which is usually equivalent according to their relationship in the “standard atmosphere” (cf. Appendix of Dunai 2000a). However, the mean sea level air pressure is not completely uniform over the globe, but in some regions high-pressure (e.g., Siberia) or low-pressure systems (e.g., Iceland) persist. The most prominent deviations from the standard atmosphere occur in Antarctica, with mean ground level pressures 20-40 mbar lower than expected (Radok et al. 1996). Stone (2000) provides a formalism translating Lal’s (1991) scaling method to a function of pressure instead of altitude. He concludes that production rate variations are restricted to a few percent at most locations. In Antarctica, however, cosmogenic nuclide production may be 25-30% higher than implied by conventional scaling methods, which may have a significant impact on exposure histories of Antarctic rocks (see Glacier movement and ice sheet evolution in the Application examples section).

740

Niedermann

Effects of dip angle and shading. On an extended horizontal surface, cosmic ray particles are arriving from the whole upper half-space, i.e., from a solid angle of 2π. However if the surface is inclined, part of the flux is faded out, resulting in smaller cosmogenic nuclide production rates. Likewise, big objects on the horizon (such as nearby mountains) may block out a substantial portion of the cosmic ray flux. Dunne et al. (1999) have calculated scaling factors for both cases, assuming an angular distribution of the cosmic ray intensity according to Equations (3a) and (3b). For a surface of uniform slope with dip angle δ, the azimuth angle ϕ is defined as ϕ = 0 in the direction of maximum slope. The fraction of the total cosmic ray flux hitting the surface is then given by m+1 f(δ) = 2π

2π

∫

π/2

∫

(J/J 0 ) cosθ dθ dϕ

(10)

ϕ =0 θ=γ(ϕ,δ)

with the slope angle γ in direction ϕ inserted as γ(ϕ,δ) = arctan (cosϕ tanδ)

(11)

and J from Equations (3a), (3b). m is the exponent of the angular distribution (Eqn. 3a). Figure 3 depicts the dependence of the fractional cosmic ray flux on the dip angle. For the case of a horizontal surface shaded by a rectangular obstruction reaching up to an inclination angle θ0 and extending over an azimuth angle Δϕ, the remaining fraction of cosmic ray flux is calculated as

f(θ0 , Δϕ) = 1 −

Δϕ sin m+1 θ0 2π

(12)

(Dunne et al. 1999). The same formula can also be applied to a triangular obstruction having a baseline extension Δϕ and attaining a maximum inclination θT, if θ0 is replaced by θR = 0.62 θT – 0.00065 θT2

(13)

(for θR, θT in degrees and m = 2.3; see Dunne et al. 1999). The dependence of the remaining fraction of cosmic ray flux (“shielding factor”) on the inclinational and azimuthal extension of a rectangular obstruction is plotted in Figure 4. Temporal variation of cosmogenic nuclide production Solar activity dependence. Galactic cosmic rays reaching the solar system are modulated by solar magnetic fields, which reduce the intensity of the low-energy branch (up to ~ 10 GeV/nucleon). Since these magnetic fields vary with the 11-year solar cycle, the cosmic ray flux at the top of the Earth’s atmosphere is anticorrelated with indices of solar activity, such as the sunspot number. In consequence the cosmic ray neutron flux in the atmosphere and hence the production rates of cosmogenic nuclides change accordingly from solar minimum to maximum (e.g., Simpson 1983). Figure 5 shows that for a solar modulation parameter (the solar equivalent to the cutoff rigidity) between 300 and 900 MeV, production rates deviate by about ± 25% from the average at high latitudes (Masarik and Beer 1999). Near the geomagnetic equator, however, they remain almost constant because there the low-energy branch of primary cosmic ray particles (those which are most affected by solar modulation) is prevented from entering the atmosphere anyway due to the high cutoff rigidity.

Typical time spans relevant to cosmic ray exposure dating are at least tens to hundreds of solar cycles; therefore the dependence of production rates on the 11-year

Noble Gases as Dating Tools for Surficial Processes

741

Figure 3. Dependence of the cosmogenic nuclide production rate on the dip angle of an inclined surface, as calculated from Equation (10), with m = 2.3. P0 is the production rate on a flat horizontal surface.

Figure 4. Shielding factor S (= f(θ,Δϕ) from Eqn. 12) resulting from a cosmic-ray-blocking rectangular obstruction which extends over an azimuthal angle Δϕ and up to a zenith angle θ. m=2.3 was assumed for the exponent of the angular cosmic ray distribution. [Used by permission of the editor of Geomorphology, from Dunne et al. (1999), Fig. 1, p. 6.]

cycle can usually be neglected. One exception is the direct determination of cosmic ray production rates by exposing a target to cosmic ray irradiation for a few years and subsequently measuring the cosmogenic nuclides produced (Yokoyama et al. 1977; Graf et al. 1996; Nishiizumi et al. 1996; Brown et al. 2000). In this case the normalized production rate Pn can be estimated from the measured production rate Pm by comparing the average neutron counting rate (per hour or day) over one or several complete solar cycles Nav with the average counting rate during the exposure period: n Pn = Pm × N av × n (14) ∑ Ni i=1

where n is the number of exposure days and Ni is the counting rate on day i. Neutron counting rates should be obtained from nearby observatories. In Equation (14) the decay of a radionuclide during the exposure period has been neglected since it is only relevant for very short-lived nuclides such as 7Be (cf. Nishiizumi et al. 1996).

Figure 5. Latitudinal dependence of the production rates of 7Be (a), 10Be (b), 14C (c), and 36Cl (d) in the atmosphere for different solar modulation parameters, as modeled by Masarik and Beer (1999). Similar effects of solar activity can be expected for the in situ production of cosmogenic nuclides in rocks on the earth’s surface. [Used by permission of the American Geophysical Union, from Masarik and Beer (1999), J. Geophys. Res., Vol. 104, Fig. 8, p. 12108.]

742 Niedermann

Noble Gases as Dating Tools for Surficial Processes

743

The solar activity is also subject to variations on longer timescales. A prominent minimum of solar activity, during which almost no aurora and sunspots were observed, has been reported for the years 1645-1715 A.D., the so-called Maunder minimum (Eddy 1976). Other minima have been identified from 14C and 10Be records in tree rings and polar ice cores (e.g., Beer et al. 1988; Bard et al. 1997). To resolve such excursions from deviations caused by the geomagnetic field as discussed below, records from polar regions are best suited as the effects of solar modulation are maximal there whereas the geomagnetic field influence is minimal. Castagnoli and Lal (1980) have estimated that a solar modulation parameter of 100 MeV may describe the cosmic ray flux during the Maunder minimum and similar periods of quiet sun, corresponding to a doubling of cosmogenic nuclide production at high latitudes. Fortunately the time-integrated production rates over the whole exposure history of a rock are much less affected by such excursions. Secular variations of the geomagnetic pole position. The geomagnetic field undergoes secular variations, i.e., changes in the direction and intensity of its dipole and non-dipole components over periods of years to several millennia (e.g., Sternberg 1992). Such changes affect production rates of cosmogenic nuclides because they control the cutoff rigidity and thus the cosmic ray flux and energy spectrum. Three parameters are required to describe the geomagnetic field at a certain location, namely intensity, inclination, and declination. The latter two give the direction of the field vector; their evolution with time is quite well known for various places on Earth (e.g., Ohno and Hamano 1992, 1993). Averaging the magnetic field vectors all over the planet is expected to average out the non-dipole components and yield the positions of the geomagnetic poles. The movement of the north geomagnetic pole over the last 10 ka is plotted in Figure 6. It is qualitatively evident that for periods of more than a few thousand years, the geographic pole is a good approximation to the average geomagnetic pole position, i.e., geographic latitude can be used instead of geomagnetic latitude when scaling production rates. Sternberg (1996) estimates that the deviation is only ~ 1.5° for exposure times longer than several centuries. Therefore geographic latitude is the better choice in the majority of cosmogenic nuclide studies. Only for exposure ages shorter than ~ 1000 a is it recommended to estimate an average geomagnetic pole position from the data of Ohno and Hamano (1992, 1993) and calculate the average geomagnetic latitude from

sinλm = cosε cosλg cosΔϕg + sinε sinλg

(15)

where λm and λg are the geomagnetic and geographic latitudes of the site, ε is the geographic latitude of the average geomagnetic pole position, and Δϕg is the difference in geographic longitude of the site and the geomagnetic pole. The scaling method of Dunai (2000a) involves also non-dipolar components of the geomagnetic field. Although these components should cancel out when averaged over the whole Earth, they may yield substantial contributions at a specific location. Therefore Dunai based the scaling of production rates on inclination rather than geomagnetic latitude. Nevertheless, for time periods of > 10-20 ka, the time-averaged geomagnetic field at any location can be approximated by a geocentric axial dipole field (e.g., Merrill et al. 1998), and in these cases the geographic latitude may also be used with Dunai’s (2000a) scaling procedure, by converting inclination I to latitude λ according to Equation (7). For shorter exposures an average inclination should be estimated based on suitable paleomagnetic records from the vicinity of the sampling site. Long-term geomagnetic field intensity changes. Not only the direction of the geomagnetic field vectors, but also their intensity may change with time. Both paleomagnetic studies (e.g., Valet et al. 1998; Guyodo and Valet 1999; Juarez and Tauxe

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Figure 6. Location of the north geomagnetic pole from 10,000 a to present in 100 a increments, according to Ohno and Hamano (1992, 1993). Numbers denote years BP, arrows indicate the directions of the polar movement. It is qualitatively evident that for periods of the order of millennia, the geographic pole is a good approximation to the average geomagnetic pole position. [Used by permission of the American Geophysical Union, from Ohno and Hamano (1992), Geophys. Res. Lett., Vol. 19, Fig. 2, p. 1717.]

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2000) and investigations of cosmogenic nuclide production in earlier epochs (e.g., Frank et al. 1997; Plummer et al. 1997; Wagner et al. 2000) indicate variations of the dipole moment, which directly influence production rates. The relative paleointensities as derived from the 10Be deposition rate (Frank et al. 1997) and from the paleomagnetic record (Guyodo and Valet 1999; see Figure 7) in stacked deep-sea sediments agree remarkably well. However, there are some difficulties involved in the interpretation of these data. There has been a dispute as to the level of climatic influence in such data sets (e.g., Kok 1999; Frank 2000). Moreover, the pattern of temporal variation as shown in Figure 7 is not reproduced by all paleomagnetic data: Channell et al. (1997) did not observe the sharp increase in paleointensity from ~ 40 ka to the present at ODP site 983, and Goguitchaichvili et al. (1999) derive a geomagnetic dipole moment indistinguishable from the modern value for the Pliocene field, in contrast to the results of Juarez and Tauxe (2000). Although the magnitude of geomagnetic field intensity variations in the past may be disputed, it seems clear that variations did occur. The dependence of cosmogenic nuclide production rates on such variations has been estimated in several studies (e.g., Cerling and Craig 1994a; Shanahan and Zreda 2000). At high latitudes production rates remain constant because the cutoff rigidity is too low to influence the cosmic ray spectrum, whereas at low latitudes the marked changes in cutoff rigidity accompanying a change in the dipole moment result in corresponding modifications of production rates. Equation (1) describes the relation between cutoff rigidity, dipole strength, and geomagnetic latitude. In a modified form, it can be written as Pc = 14.9 GV × (Mt/M0) cos4 λm

(16)

Figure 7. Relative variations of the geomagnetic field intensity from 800 ka to present (dotted line) and their influence on instantaneous (thin solid line) and time-integrated (thick solid line) cosmogenic nuclide production rates at latitudes 0-10°, according to model calculations of Masarik et al. (2001). The geomagnetic field reconstruction relies on dendrochronologically derived Δ14C (Stuiver et al. 1998) for the last 10 ka and on paleomagnetic records from deep-sea sediments (Guyodo and Valet 1999) for the remainder of the period. The influence of geomagnetic field intensity on production rates decreases towards higher latitudes (cf. Fig. 8). [Used by permission of the editor of Geochim. Cosmochim. Acta, from Masarik et al. (2001), Fig. 1, p. 2998.]

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where M0 is the present dipole moment and Mt the dipole moment at an earlier time t. A variation of the dipole field is thus equivalent to a change in geomagnetic latitude, according to cos λm′ = (Mt/M0)1/4 cos λm

(17)

where λm′ is the geomagnetic latitude where the cutoff rigidity (and consequently the production rate) at dipole strength M0 is equal to that for λm at dipole strength Mt. Production rates at Mt can therefore be determined by scaling the sea level, high latitude production rates to latitude λm′ instead of λm. This method fails to work if (Mt/M0)1/4 > 1/cos λm, i.e., for high magnetic field intensities and low latitudes. Two recent studies have addressed the issue in more detail. Masarik et al. (2001) use model calculations of cosmic ray particle interactions with matter to estimate the relation between changes in cutoff rigidities and cosmogenic nuclide production rates. Their calculations are based on the Δ14C and paleomagnetic records of Stuiver et al. (1998) for the last 10 ka and Guyodo and Valet (1999) for the period 10-800 ka, respectively, and result in relatively modest variations of the time-integrated production rates. Figure 7 shows that at low latitudes (0-10°), the deviations from the present-day production rates are always < 1% during the last 10 ka, and less everywhere else. From ~ 5 to 40 ka, production rates increase steadily to ~ 10% above present-day values (at 0-10°) and remain ~ 9-12% higher than today back to 800 ka (Fig. 7). At latitudes of 30-40°, the deviations are always < 2.5%, and at > 40° production rates are virtually unaffected by variations of the geomagnetic field intensity. Dunai (2001b) has fitted his scaling parameters N1030 and Λ to cutoff rigidity instead of inclination in order to enable a full description of cosmic ray flux variations, including non-dipole components: A N1030 (Pc ) = Y + (18) C ⎡⎣1 + exp ( (X-Pc ) B ) ⎤⎦ a

Λ (Pc ) = y +

(19) c ⎡⎣1 + exp ( (x-Pc ) b ) ⎤⎦ where Pc is inserted in GV and the coefficients A, B, C, X, Y and a, b, c, x, y are compiled in Table 4. Dunai’s approach allows to calculate scaling factors for production rates at any place on the globe during any time in the past, provided that the geomagnetic field parameters are known for that time and place. Though Dunai’s (2001b) data yield similar qualitative production rate variations to those of Masarik et al. (2001) when

Table 4. Numerical values of the coefficients for the calculation of the nuclear disintegration rate N1030(Pc) and the attenuation length Λ(Pc) according to Equations (18) and (19), respectively (Dunai 2001b). Coefficients for N1030(Pc) A B C X Y

0.5221 -1.7211 0.3345 4.2822 0.4952

Coefficients for Λ(Pc) a b c x y

17.183 2.060 5.9164 2.2964 130.11

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evaluated based on the Guyodo and Valet (1999) record (Fig. 8), there is an obvious quantitative discrepancy. At low latitudes, time-integrated production rates reach a maximum of 20% above present-day values at ~140 and ~200 ka, compared to 12% according to Masarik et al. (2001). At mid-latitudes of 30 and 40°, Dunai’s (2001b) results imply maximum deviations of 15% and 7%, respectively, as compared to ~2.5% (Masarik et al. 2001). Furthermore, Dunai (2001b) obtains a marked altitude-dependence of production rate variations, whereas Masarik et al. (2001) claim that the corrections are virtually identical between sea level and 5000 m.

F 8 + caption

Figure 8 Time-integrated cosmogenic nuclide production rate at sea level, for various latitudes, relative to the present production rate at the same latitude according to the scaling method of Dunai (2001b). (a) shows the period 0-800 ka, (b) is an enlargement of the 0-50 ka range. While based of on the same paleomagnetic record (Guyodo and Valet 1999) for the period 10-800 ka, this model predicts substantially higher production reate vatations than that od Masarik et al. (2001; Fig. 7). [ Used by permission o fhte editor of Earth Planet. Sci. Lett., from Dunai (2001b), Fig. 6, p. 206.]

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At present, the issue of past variations of the geomagnetic field strength and how they control cosmogenic nuclide production rates is obviously not well enough understood and needs further investigation. Fortunately most cosmogenic nuclide studies in the past, including the majority of production rate determinations (see Experimental determinations of production rates section), have focused on mid- and high-latitude sites, where production rate variations are minor. However in future the accuracy of absolute exposure ages, particularly for locations at low latitudes, will critically depend on reliable correction factors for geomagnetic field variation. Cosmogenic nuclide production at depth and on eroding surfaces Depth dependence of cosmogenic nuclide production beneath a flat surface. So far, the influence of various parameters on cosmogenic nuclide production at the very surface of the Earth has been discussed, i.e., in the uppermost few centimeters of a surface rock. Upon penetrating into the rock, the intensity of cosmic rays is further attenuated exponentially (Eqn. 2), resulting in a corresponding decrease of production rates with depth. As a consequence of the angular distribution of incident cosmic ray particles, the depth dependence of production rates is not strictly exponential but obeys an incomplete gamma function (Dunne et al. 1999). However, that function can well be approximated by an exponential:

P(z) = P0 e–ρz/Λ

(20)

where P0 is the production rate at the surface, P(z) that at depth z (cm), ρ is the rock density (g/cm3), and Λ is the attenuation length (g/cm2). The value of Λ is lower than the attenuation length for the particle flux (Eqn. 2) by a factor of approximately 1.3 (Dunne et al. 1999), since particles incident at low angles must penetrate through more material to reach the same depth. Λ may also depend somewhat on target rock composition and on the nuclide that is produced, due to variable neutron absorption capabilities of different elements. Furthermore, as in the atmosphere Λ is expected to increase from high to low latitudes along with the cosmic ray energy spectrum getting harder. Experimental determinations of attenuation lengths with sufficient accuracy are scarce, but indeed the data from Antarctica (~ 150 g/cm2; Brown et al. 1992) are lower than those from New Mexico (~ 162 g/cm2; Nishiizumi et al. 1994), Réunion (~ 162 g/cm2; Sarda et al. 1993), or Hawaii (~ 170 g/cm2; Kurz 1986b). Calculated values range from 157 to 167 g/cm2 (Masarik and Reedy 1995). One peculiar feature in the calculations of Masarik and Reedy (1995) is the disturbance of the neutron flux near the air-surface interface. Figure 9 shows that in contrast to the exponential decrease on either side of the boundary, the pattern is flat to about 10-20 g/cm2 above and below the rock surface. That disturbance is caused by different neutron production and transport mechanisms in air and rock and the loss of neutrons from the rock to the atmosphere (O’Brien et al. 1978). There is a welcome side effect to this phenomenon in exposure age studies in that there is no need to correct for self-shielding in surface samples of less than a few centimeters size. Likewise, erosional loss by a few centimeters will not affect the cosmogenic nuclide production in the surface layer. On the other hand, the escape of neutrons from the outermost rock layers may reduce production rates in boulder samples relative to those in a flat surface (Masarik et al. 2000), by as much as 12% for the surface of a hemisphere with 1m radius. It must be cautioned that these results have so far not been tested experimentally. If the flat profile down to 10 g/cm2 is correct, production rates at depth would be some 6% higher than for an exponential decrease starting at the very surface. The exponential decrease according to Equation (20) is only valid for the spallationproduced component. Thermal neutron capture reactions have a distinct depth

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Flux (neutrons cm-2 s-1)

Distance from surface (g cm-2) Figure 9. Total cosmic ray neutron flux on either side of the air-surface interface as calculated by Masarik and Reedy (1995). Boundary effects produce a flat pattern for ~ 10-20 g/cm2 on both sides of the interface. [Used by permission of the editor of Earth Planet. Sci. Lett., from Masarik and Reedy (1995), Fig. 1, p. 387.]

dependence showing a maximum at some depth below the surface, ~30-50 g/cm2 (Liu et al. 1994). Since these reactions are only important for a few nuclides, in particular 36Cl, but not for cosmogenic He and Ne production (cf. Production mechanisms of cosmogenic nuclides in terrestrial rocks section above), they are not discussed further here. Reference is made to the investigations of Liu et al. (1994) and Phillips et al. (1996). For muon-induced cosmogenic nuclide production, a roughly exponential decrease with depth can again be assumed, with Pμ(z) = Pμ,0 e–ρz/Λμ

(21)

where Λμ ~ 1300 g/cm2 (Barbouti and Rastin 1983; Brown et al. 1995a), although this is only an approximation since the exact pattern of muon propagation and stopping rate depends much more strongly on the chemical composition of the target than in the case of neutrons, as is evident from the marked difference between the attenuation lengths in rock and air (247 g/cm2; Lal 1988). At depths greater than ~5000 g/cm2, Λμ increases further as the relative contribution of fast muon interactions increases (Bilokon et al. 1989; Heisinger and Nolte 2000). More accurate mathematical representations of the evolution of cosmogenic nuclide production rates with depth, including interactions with both stopped and fast muons, can be found in Stone et al. (1998) or Granger and Smith (2000), whereas Heisinger and Nolte (2000) have shown graphic illustrations based on their model calculations (cf. Fig. 1). Effective attenuation length beneath inclined or partially shaded surfaces. As the attenuation length Λ for cosmogenic nuclide production by spallation (Eqn. 20) is determined by the angular cosmic ray neutron distribution, it increases when part of the

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neutrons (those incident from shallow angles) are faded out due to shading or surface slope. In case of a rectangular obstruction extending over an azimuthal angle Δϕ and up to an inclination angle θ0, the effective attenuation length Λ* is given as 1 − (Δϕ / 2π)sin m + 2 θ0 Λ = Λ 1 − (Δϕ / 2π)sin m +1 θ0 *

(22)

(Dunne et al. 1999). m is the exponent of the angular distribution (Eqn. 3a). The influence of triangular obstructions on attenuation length is also described in Dunne et al. (1999). In case of a sloping surface, the mathematical treatment is rather complicated. Dunne et al. (1999) give “scaling factors” relating the production rate at depth z measured in a direction perpendicular to a sloped surface to that on a flat surface. Masarik et al. (2000), on the other hand, report effective attenuation lengths for 10Be production beneath surfaces of various slope angles, as obtained from model calculations. The agreement between both methods is quite satisfactory despite different angular distribution exponents used (m = 2.3 versus 2.65). Erosion. Erosion acts to continually or episodically remove layers of irradiated rock, thereby exposing at the surface material which was once buried and decreasing the depth of shielding of deeper material. Therefore, cosmogenic nuclide production in a sample exposed at the surface today was lower in earlier epochs. The shielding depth z(-t) of a rock sample at time t before present can be expressed as 0

z(-t) = z p + ∫ ε(t') dt'

(23)

-t

where ε(t) is the erosion rate and zp the present shielding depth. For the case of a constant erosion rate, one obtains z(-t) = zp + εt

(24)

and

P(z,-t) = P0 e

-ρ(z p +εt)/Λ

+ Pμ,0 e

-ρ(z p +εt)/Λμ

for the total production rate from spallation and negative muon capture (cf. Eqn. 20, 21). The latter term becomes important not only for deep samples, but also for surface samples where high erosion rates have prevailed for long periods of time. Derivation of exposure ages and erosion rates Simple exposure histories. The determination of the exposure age or erosion rate of a certain geomorphic feature is easiest in the case of a simple exposure history, i.e., if the rock studied was excavated from a completely shielded location by a brief process (such as a volcanic eruption) and then experienced cosmic ray irradiation at constant (or no) erosion without ever being moved. Temporal variations of the surface production rate due to solar activity or geomagnetic field changes should be negligible as well. In such a case, the temporal evolution of a cosmogenic nuclide concentration C in the rock is given by

dC(z,t)/dt = P(z,t) – λC(z,t)

(26)

where λ is the decay constant of a radionuclide (λ = 0 for stable nuclides). Defining t = 0 for the time of excavation and z0 the initial shielding depth, we obtain z(t) = z0 – εt (cf. Eqn. 24) and P(z,t) as in Equation (25), but replacing zp + εt by z0 – εt. Integration of (26)

(25)

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751

then yields the solution P0 e−ρ(z0 − εt)/Λ (1 − e− (λ+ρε/Λ)t ) λ + ρε/Λ Pμ,0 − ρ(z − εt)/Λμ − (λ+ρε/Λμ )t + e 0 (1 − e ) λ + ρε/Λ μ

C(z,t) = C(z,0) e− λt +

(27)

There are two important limiting cases for such simple exposure histories: The case of no erosion (ε = 0) and that of steady-state erosion. For ε = 0 Equation (27) reduces to C(z 0 ,t) = C(z 0 ,0) e − λt +

(

1 − e − λt − ρz /Λ P0 e−ρz0 /Λ + Pμ,0 e 0 μ λ

)

(28)

If the initial concentration C(z0,0) = 0, as can usually be assumed for radionuclides, and neglecting the muogenic production, the exposure age T is calculated from T= −

1 ⎛ C(z 0 ,T) λ ⎞ ln ⎜1 − ⎟ λ ⎝ P0 e −ρz0 /Λ ⎠

(29)

For λ = 0 (stable nuclides), the differential equation (26) reduces to dC/dt = P(z0) which has the solution C(z0,t) = C(z0,0) + P0 e–ρz0/Λ t + Pμ,0 e–ρz0/Λμ t

(30)

and therefore (for zero initial concentration or, equivalently, if only the cosmogenic component is considered)

T=

C(z 0 ,T)

P0 e −ρz0 /Λ + Pμ,0 e

− ρz0 /Λμ

(31)

In the case of steady-state erosion (cf. Lal 1991), the cosmogenic nuclide concentration is not determined by the exposure age T but by the erosion rate: When T >> 1/(λ + ρε/Λ) and T >> 1/(λ + ρε/Λμ), Equation (27) reduces to C(z,t) = C(z,0) e − λt +

Pμ,0 P0 − ρ(z − εt)/Λμ e −ρ(z0 − εt)/Λ + e 0 λ + ρε/Λ μ λ + ρε/Λ

(32)

It should be noted that due to the higher attenuation length, it takes longer for the muogenic component to reach erosion equilibrium, in particular for long-lived or stable nuclides (λ ≈ 0). For a sample on the present-day surface (z0 – εT = 0) and again assuming C(z0,0) = 0, we can nevertheless neglect the third term of (32) in first-order approximation since P0Λ is clearly higher than Pμ,0Λμ, though not by orders of magnitude (Pμ,0/P0 ≤ 0.03, Λμ/Λ ≈ 8). Using this approximation the steady-state erosion rate can be expressed as ε=

Λ ⎛ P0 ⎞ ⎜ − λ⎟ ρ⎝C ⎠

(33)

For a more exact solution including production by muons, a numerical approach or the use of modeled depth vs. concentration profiles (e.g., Heisinger et al. 1997; Heisinger and Nolte 2000) is required. Neither of the two limiting cases (no erosion or steady-state erosion) may be common in practice. The assumption of no erosion may be taken if less than a few

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centimeters of rock have been removed during the whole time of exposure, corresponding to a change in production rate by < 10%, which is typically well within the general precision of exposure dating. Using the same criterion, steady-state erosion is fulfilled when T > 3/(λ + ρε/Λ) or, for stable nuclides, T > 150 cm/ε (using ρ = 3 g/cm3 and Λ = 160 g/cm2). In all other cases the cosmogenic nuclide concentration is determined by a combination of exposure time and erosion rate, and Equations (29) or (31) and (33), respectively, give only the minimum exposure age and the maximum erosion rate compatible with the data. The temporal evolution of a cosmogenic nuclide concentration for various erosion rates is plotted in Figure 10. To derive the real exposure age (or the real erosion rate), independent information on either quantity is required. In principle, one possibility to achieve simultaneous information on both parameters is to study two (or more) cosmogenic nuclides in the same rock, yielding two equations such as (27) for the two unknowns. However, the method only works under certain conditions which have been documented in detail by Gillespie and Bierman (1995): First, the half-lives of the two nuclides must be sufficiently different. The ratio of the stable cosmogenic nuclides 3 He and 21Ne equals that of the production rates for any combination of ε and T, so the terms containing these quantities in Equation (27) cancel out, leaving only one equation for two unknowns. Second, one of the nuclides should be close to equilibrium between production and the combined effects of erosion and radioactive decay, while the other should still be growing; otherwise the precision of the determination will be poor. This is shown in Figure 11, for the pairs 10Be-26Al and 21Ne-26Al. Such plots are a common means to estimate both exposure age and erosion rate from the determination of two cosmogenic nuclides (e.g., Lal 1991; Nishiizumi et al. 1991a; Graf et al. 1991). The area between the two curves (“steady-state erosion island”; Lal 1991) comprises all the possible combinations of ε and T for the case of simple exposure histories.

Figure 10. Temporal evolution of the concentration of a stable cosmogenic nuclide in dependence of the erosion rate, for a production rate of 1 atom g–1 a–1. Expected concentrations in real samples can be calculated by multiplying the values from this plot with the actual production rate.

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Figure 11. Plots of the ratios 10Be/26Al and 21Ne/26Al, respectively, versus the 26 Al concentration for rocks with a simple exposure history. The lower solid lines represent the temporal evolution of these quantities in the absence of erosion, the upper solid lines in erosion equilibrium. Dotted lines indicate the temporal evolution for erosion rates of 10, 1, 0.1, and 0.01 m/Ma (1 m/Ma = 10–4 cm/a). The areas between the curves (“steadystate erosion island;” Lal 1991) comprise all possible combinations of exposure ages and erosion rates for simple exposure histories, data outside of these areas indicate complex exposure histories or experimental error. Curves are plotted for 10Be, 21 Ne, and 26Al production rates of 5.42, 19.0, and 35.2 atoms g–1 a–1, respectively (modified from Kubik et al. 1998; Niedermann 2000; see Table 6).

Complex exposure histories. Complex exposure histories are those which involve, for example, several stages of cosmogenic nuclide production at various depths, variable erosion rates, episodic burial, changes in altitude of exposure, etc. Such scenarios cannot usually be resolved by cosmogenic nuclide studies alone, but require independent information from other methods. Even if such information is not available, it is most important to recognize complex exposure histories in order to avoid gross mistakes in the interpretation of exposure age data. Evidence for complex exposure histories can be taken from plots such as Figure 11. Experimental data which do not fit into the allowed area of the plot clearly indicate that the assumptions on which Equation (27) is based are not met (provided that experimental errors can be excluded and contributions from noncosmogenic sources have been corrected for). For example, a data point left of the steadystate erosion island may result from a rock buried for some period of time, causing a change in the ratio of the two nuclides because the shorter-lived one decays more rapidly. Data points to the right of the allowed zone would imply exposure conditions under a higher production rate (e.g., at higher altitude). Ratios of short-lived to long-lived nuclide

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above the production ratio are not possible. Detailed discussions of complex exposure scenarios involving non-steady-state erosion can be found in Lal (1991), Gillespie and Bierman (1995), or Small et al. (1997). Complex exposure histories may also show up as a disturbed pattern of cosmogenic nuclide concentration versus depth. Particular care must be taken when dating soil or alluvial deposits, which often experience varying sedimentation rates, sudden burial, or bioturbation (i.e., soil mixing by living organisms) of the uppermost layers (e.g., Phillips et al. 1998; Braucher et al. 2000). In such studies it is extremely important that depth profiles are taken. Error considerations

Estimating the absolute precision of exposure ages and erosion rates deduced from cosmogenic nuclide studies is not easy. As detailed in the above sections, there are various factors controlling production rates even for the case of simple exposure histories, only some of which can be assessed in a strictly quantitative sense. The most important of these contributions are compiled in Table 5. For example, these data show Table 5. Compilation of important sources of error in surface exposure dating studies. The quoted uncertainties (2σ level) are only estimates for typical cases and may in reality be higher or lower depending on the special conditions. Additional error sources which cannot at all be quantified in any general manner include lacking information on erosion rate, uplift or subsidence, complex exposure histories, etc. Error source Measured nuclide concentration C Attenuation length Λ Shielding depth z Host rock density ρ

Typical individual uncertainty

Contribution to age uncertainty

5-20%

5-20%

10 g/cm2 5 cm 0.2 g/cm3

→ Shielding correction

Neglect of muogenic component

13% 20% 36% 2%

2% 10% 54% 7400%

Production rate Scaling method

10-20%

Scaling with elevation instead of air pressure Use of geographic latitude Non-dipole components of geomagnetic field Long-term solar activity variations Long-term variations of geomagnetic field

1% – 2-4% 1.5°

10-20% 10%1 2-10% 6-8% +15-30% 2%

10-25%

0-30%

± 400 MeV

10%?

± 4 × 1022 Am2

≤ 12% ≤ 20%

1

Remarks

z = 50 cm z = 100 cm z = 200 cm z = 0, sea level z = 100 cm z = 200 cm z = 500 cm at sea level, ≥ 60° Lal (1991) Dunai (2000a) outside Antarctica in Antarctica for T > 1 ka for T < 10-20 ka, depends on location time-integrated at high latitudes Masarik et al. (2001) Dunai (2001b)

Dunai (2000a) argues that Lal’s (1991) uncertainty estimate does not include all error sources and that 20% would be more realistic.

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the increasing influence of relatively small individual errors with increasing shielding depth, due to the exponential dependence. In consequence, erosion will also have a nontrivial effect on production rate uncertainties even if the value of the erosion rate is well constrained. Moreover, it is obvious that scaling of cosmogenic nuclide production rates in space and time remains a major source of error due to our lack of knowledge about relevant parameters, such as the local geomagnetic field, and their control on production rates. The usual way to circumvent trouble arising from such difficulties is to quote error limits which only include experimental uncertainties of the measurement but not those of the production rate determination, the scaling method, geomagnetic field variations, neglect of the muon-produced component, assumptions regarding simple or complex exposure histories, etc. This may not be a major problem as long as both authors and readers are aware of it and do not over-interpret the data. Also, it does not make much sense to include the (systematic) errors of production rate determinations or scaling when only results from one nuclide for a limited geographical area are compared. Whenever absolute ages or erosion rates are concerned however, such systematic errors must be taken into account. Bierman (1994) gives an excellent presentation on the limitations of the use of cosmogenic nuclides for geomorphic applications, although some of his error assessments may seem a little pessimistic. Systematic errors may be introduced in various stages of a cosmogenic nuclide study. Especially important to noble gas studies is the presence of non-cosmogenic components (see Discrimination of cosmogenic against trapped, radiogenic, and nucleogenic components in the Experimental issues section), which may lead to wrong conclusions if corrections for such components are based on faulty assumptions. Difficulties in the geomorphic setting may be another important error source. For example, snow or soil cover during part of a rock’s exposure is not always easy to detect but may have a substantial influence on cosmogenic nuclide production. Similar to erosion, neglect of these effects will yield too low exposure ages; it is therefore better to quote minimum exposure ages when they cannot be assessed. For studies of two cosmogenic nuclides with different half-lives (or one stable and one radioactive nuclide), I refer to the rigorous treatment of the precision of exposure age and erosion rate estimates by Gillespie and Bierman (1995). Though uncertainties resulting from cosmogenic nuclide studies may thus seem rather high when all error sources are considered, the unique possibilities of this method must be borne in mind (see Application examples). In many cases other techniques rely on much weaker assumptions or provide no quantitative answers at all. Furthermore, many error sources (e.g., those connected with production rate scaling) will certainly become considerably smaller in near future. PRODUCTION RATES OF COSMOGENIC NUCLIDES Experimental determinations of production rates

When the early studies of cosmogenic nuclides in terrestrial surface rocks were published, production rates were quite poorly known. Davis and Schaeffer (1955) and Phillips et al. (1986) estimated the 36Cl production rate from the cosmic ray neutron flux and the thermal neutron capture cross section of Cl. Srinivasan (1976) scaled the 126Xe production rate from the lunar surface to the Earth’s surface by considering the attenuation of cosmic rays all through the atmosphere. Kurz (1986a,b), Craig and Poreda (1986), and Porcelli et al. (1987) used the 3H production rate as calculated by Yokoyama et al. (1977) to estimate the 3He production rate. The problem with such estimates was a

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large influence of model calculations which were in part based on distant extrapolations or ill-constrained parameters, such as nuclear excitation functions for reactions involving neutrons (Lal 1988). Therefore, experimental determinations of cosmogenic nuclide production rates in rocks or minerals of various target element composition were urgently needed. Volcanic rocks are probably among the best-suited samples for such studies, since simple exposure histories are common, the influence of erosion can be judged from the preservation of surface textures, and exposure ages are identical to eruption ages, which can often be determined by conventional methods such as 14C or 40Ar-39Ar dating. Other possibilities to determine production rates experimentally include rocks exposed by glacial scouring, giant floods or landslides and artificial targets exposed for a few years at mountain altitudes. The following summary of experimental production rate determinations will focus on the noble gas isotopes 3He and 21Ne but will also include the radionuclides 10Be and 26Al, which have often been used in combined studies along with He and Ne. Production rates in basaltic rocks and their constituents. Volcanic rocks (mostly of basaltic composition) have provided the largest data set of experimental production rates, especially for 3He. The first such determination was that of Kurz (1986b), who measured the cosmogenic 3He (hereafter 3Hec) concentration in olivine phenocrysts from a radiocarbon-dated (28 ka) lava flow from Mauna Loa, Hawaii, and deduced a production rate of 97 atoms g-1 a-1 at sea level and 37°N geomagnetic latitude (Kurz 1987). That study was extended by Kurz et al. (1990) to a whole suite of Hawaiian lava flows, ranging in age from 600 to 14,000 a. The production rates obtained from that data set scatter by about a factor of 4, which can hardly be explained by real production rate variations. Dunai (2001b) has reevaluated these data and concludes that the association of samples to dated lava flows is unreliable in several cases. Cerling (1990), Poreda and Cerling (1992), and Cerling and Craig (1994b) studied olivine and pyroxene separates from various basalts in the western USA and France and derived a sea level, high latitude 3 He production rate of 115 ± 4 atoms g-1 a-1. In addition, Poreda and Cerling (1992) determined the (3He/21Ne)c ratio in olivines as a function of their Fo content (Fig. 12), providing the first experimental value for a 21Ne production rate: 45 ± 4 atoms g-1 a-1 in Fo81 olivine. The (3He/21Ne)c production ratio was also measured by Staudacher and Allègre (1991, 1993a) in various ultramafic nodules and olivine separates. They obtained values differing from each other and from that of Poreda and Cerling (1992). The scatter is probably related to difficulties in correcting for the magmatic He and Ne components, but may in part also be explained by variations in the composition of the minerals. Whereas the 3He production rate is rather insensitive to such variations (e.g., Eugster 1988) because 3He (and 3H) nuclei are sputtered off any target nucleus, the production rates of most other nuclides depend critically on nuclear excitation functions which may vary substantially for different target nuclides (cf. Production mechanisms of cosmogenic nuclides in terrestrial rocks section). For example, the Ne production rate from Mg is substantially higher than that from Al or Si, because of the large cross sections of the 24 Mg(n,α)21Ne and 25Mg(n,α)22Ne reactions (e.g., Hohenberg et al. 1978; Masarik and Reedy 1996; Leya et al. 2000).

More recently, the 3He production rate has been determined in 2450-7090 year old basalt flows from Oregon by Licciardi et al. (1999). These authors applied a correction for variations of the geomagnetic pole position (Ohno and Hamano 1992, 1993) to their production rate values, the difference between corrected and uncorrected data is < 5% in all cases. They also revised the earlier production rate determinations of Kurz (1987), Kurz et al. (1990), Cerling (1990), and Cerling and Craig (1994b), applying a consistent method for the conversion of radiocarbon ages to calendar ages (Stuiver et al. 1998), and derived a mean Holocene 3He production rate (normalized to sea level and high latitude)

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(21Ne/3He)c

Forsterite content (%) Figure 12. Compositional dependence of the (21Ne/3He)c ratio on the forsterite content of olivine, according to Poreda and Cerling (1992). Data around Fo80 are plotted enlarged in the inset. The solid line is a least squares fit through all the filled circles, with 95% confidence intervals (dotted lines). The dashed line gives the dependence expected by Lal (1991). The open circle is from Marti and Craig (1987). [Used by permission of the American Geophysical Union, from Poreda and Cerling (1992), Geophys. Res. Lett., Vol. 19, Fig. 2, p. 1865.]

of 119 atoms g-1 a-1. A minor mistake in the calculation of the weighted means, along with a new calibration of Kurz’s (1987) lava flow age, led J.M. Licciardi (pers. comm. 2001) to slightly adjust that value to 116 ± 4 atoms g-1 a-1. Whereas all the above production rate determinations were performed on rocks < 20 ka in age, Dunai and Wijbrans (2000) studied much older lava flows from Lanzarote, Canary Islands. In olivines from two flows with 40Ar-39Ar ages of 152 and 281 ka, they derived a normalized 3He production rate of 118 ± 11 atoms g-1 a-1, whereas a 1.35 Ma old flow yielded a nominally lower value. However, as the authors point out, erosion effects on that old surface can easily explain the 3He deficiency, and the production rate averaged over 1.35 Ma is also consistent with that derived from the younger flows. Interestingly, the 3He production rates derived by Licciardi et al. (1999) and Dunai and Wijbrans (2000) agree remarkably well, although the geomagnetic field intensity in the Holocene is considered to have been some 50% above the average of the preceding epochs (e.g., Guyodo and Valet 1996; Valet et al. 1998; Frank 2000; cf. Fig. 7). Therefore, as the Lanzarote calibration site is at relatively low latitude (29°N), one would expect a sizable effect on the production rate (Dunai 2001b; Masarik et al. 2001; Figs. 7 and 8). The lack of a resolvable difference may indicate that the influence of geomagnetic field intensity variations is not too large (cf. Long-term geomagnetic field intensity changes section). On the other hand, in an evaluation of the 3He production rate determinations of Kurz (1987), Kurz et al. (1990), Cerling and Craig (1994b), Licciardi et

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al. (1999), and Dunai and Wijbrans (2000) using his own method to accommodate reported geomagnetic field variations, Dunai (2001b) concludes that all these studies are consistent within error limit with a present-day 3He production rate of 103 ± 4 atoms g-1 a-1. The 3He production rates mentioned above were not all scaled to sea level and high latitudes by the same method, which may introduce systematic differences. Table 6 shows the published production rates scaled according to both methods widely used today, i.e., Lal (1991) and Dunai (2000a). No corrections for temporal variations of the geomagnetic field have, however, been applied to these data. Readers are encouraged to use consistent sets of production rates and scaling methods in their work to avoid unnecessary systematic errors. I would like to note that the good agreement between the 3 He production rates of Licciardi et al. (1999) and Dunai and Wijbrans (2000) is not affected by the different scaling methods used. Volcanic rocks dated by the 14C and 40Ar-39Ar methods have thus provided quite an extensive data set for the 3He production rate (predominantly determined in olivine separates), along with a few 21Ne data. In addition, the 36Cl production rate was determined in basalt lavas from the western USA (Zreda et al. 1991; Stone et al. 1996), in dependence of the K and Ca content. However, to my knowledge there have not been any studies of the 10Be and 26Al production rates in basalt, probably because of experimental difficulties to detect these radionuclides in such material. Production rates in quartz. Apart from basalts and their rock-forming minerals, most experimental determinations of cosmogenic nuclide production rates have dealt with quartz. Because of its simple chemical composition (SiO2) and its abundance in a wide variety of rock types, quartz is considered a very well-suited mineral for surface exposure dating. Nishiizumi et al. (1989) used quartz separates from glacially polished granitic rocks from the Sierra Nevada, California, to derive the production rates of 10Be and 26Al. The surfaces sampled had been exposed by the retreating glaciers of the Tioga period of the last ice age and showed evidence that at least several meters of rock had been removed by glacial scouring. Therefore simple exposure histories could be assumed for these rocks. Assuming an age of 11,000 a, which was based on several radiocarbon dates for the deglaciation of the Sierra Nevada, Nishiizumi et al. (1989) derived 10Be and 26Al production rates of 61.9 and 373.6 atoms g-1 a-1, respectively, at 3340 m altitude and 44°N geomagnetic latitude. A subset of the same samples was later used by Niedermann et al. (1994), who determined a 21Ne/26Al production ratio of 0.65 ± 0.11 in quartz, corresponding to a 21Ne production rate of 243 atoms g-1 a-1 at the given altitude and latitude. A 3He production rate was not established because 3He is not quantitatively retained in quartz (cf. Retentivity of minerals for cosmogenic He and Ne section below). A revised deglaciation age (~ 13,000 a; Clark et al. 1995) and inadequate assumptions used for scaling to sea level and high latitudes led the authors of these studies to reconsider their production rate estimates (Nishiizumi et al. 1996; Niedermann 2000). The values given in Table 6 are based on the revised figures and, in the case of 10Be and 26 Al, neglecting the production by muons whose contribution was overestimated in the studies of Nishiizumi et al. (1989, 1996) as shown by Brown et al. (1995a).

An independent production rate determination of 10Be and 26Al in quartz was carried out by Kubik et al. (1998). They sampled quartz veins from large boulders derived from the landslide deposits of Köfels, Austria, which had been dated at 9800 ± 100 a by 14C. The precision to which that age is known is clearly better than for the Sierra Nevada samples. If scaled to sea level and high latitudes by the same method, the Köfels production rates are 4-19% higher than those of Nishiizumi et al. (1989), yielding a 26 Al/10Be ratio of 6.52 versus 6.02. It is difficult to judge to what extent such differences

olivine olivine, pyroxene

Kurz et al. (1990)

Cerling (1990); Cerling and Craig (1994b) Licciardi et al. (1999) Dunai and Wijbrans (2000) Poreda and Cerling (1992) Staudacher and Allègre (1991)

SiO2 SiO2 SiO2 SiO2 SiO2

Nishiizumi et al. (1991a) Kubik et al. (1998) Nishiizumi et al. (1989)

Nishiizumi et al. (1991a) Kubik et al. (1998)

Be Al

26

10

Ne

21

He

3

Masarik and Reedy (1995)

Fo81 olivine SiO2 Fo81 olivine SiO2 SiO2 SiO2

SiO2

Staudacher and Allègre (1993a) Niedermann et al. (1994); Niedermann (2000) Nishiizumi et al. (1989)

For comparison:

Al

26

Be

10

Ne

21

olivine olivine Fo81 olivine ultramafic nodules, olivine ultramafic nodules, olivine SiO2

olivine

Kurz (1986b,1987)

He

3

Target rock or mineral

Reference

Nuclide

> 4000 9.8 13 > 4000 9.8

Antarctica Austria California Antarctica Austria

105 124 41.1 18.4 5.97 36.1

33.9 34.8

5.53 5.35 31.1

5.16

20.3

13 13

~32

113 99 46 ~80

112

109

123

105 124 41.1 18.4 5.97 36.1

34.4 35.2

5.61 5.42 29.5

4.90

19.0

~35

122 118 48 ~87

117

120

138

Scaling method Lal Dunai

~14-130

2.5-7.1 152-281 17.3 ~7-27

2.3-17.4

0.55-10.7

32.9

Age range (ka)

California

Western USA and France Oregon Canary Islands Utah Russia and Mongolia Russia, Mongolia, Réunion California

Hawaii

Hawaii

Calibration site

Model calculations

Revised age (Licciardi et al. 1999) Based on 3He production rates of 113/122 atoms g–1 a–1 Based on 3He production rates of 113/122 atoms g–1 a–1 Revised age (Clark et al. 1995; Niedermann 2000) Revised age (Clark et al. 1995); without muon contribution Saturation assumed Without muon contribution Revised age (Clark et al. 1995); without muon contribution Saturation assumed Without muon contribution

Revised age 32.94 ± 0.94 ka (Bard et al. 1998) Revised ages (Licciardi et al. 1999); individual data scatter by factor 4 Revised ages (Licciardi et al. 1999)

Remarks

Table 6. Production rates of the cosmogenic noble gas isotopes 3He and 21Ne and of the radionuclides 10Be and 26Al as determined in rocks and minerals of well-known exposure history. Most reported data are weighted means from several individual production rate determinations, which were all separately scaled to sea level and high latitudes by both the methods of Lal (1991) and Dunai (2000a). Temporal variations of the geomagnetic field and contributions from non-dipole components of the geomagnetic field were neglected to facilitate comparison of different data sets. Several data were revised according to more recent age calibrations, details are given in the remarks.

Noble Gases as Dating Tools for Surficial Processes   759

760

Niedermann

are caused by the scaling method (e.g., non-dipole components of the geomagnetic field), to variations in geomagnetic field intensity or solar activity during different periods of exposure, or just to experimental error. A determination of the 21Ne production rate in the Köfels samples, as planned in the Zurich lab (R. Wieler, pers. comm. 2001), will be most informative. Nishiizumi et al. (1991a) report 10Be and 26Al data from an Antarctic quartz in which the concentrations of both radionuclides were at saturation, corresponding to an exposure age of > 4 Ma and negligible erosion. In such a case Equation (27) reduces to C = P0 / λ (for a surface sample and neglecting the muogenic component), so the production rate can readily be calculated from the concentration and the decay constant. The resulting production rates compare quite well to those of Kubik et al. (1998), but are clearly higher than those of Nishiizumi et al. (1989), especially for 10Be (Table 6). Hudson et al. (1991) determined 21Ne in six of Nishiizumi et al.’s (1991a) quartz samples and deduced a minimum 21Ne production rate of 80 atoms g-1 a-1 at 1800m elevation and 77.6°S, converting to ~ 17 atoms g-1 a-1 at sea level. This value is consistent with that of Niedermann et al. (1994). These determinations did, however, not consider the exceptionally low air pressure in Antarctica, which increases production rates and leads to a shift of the “steady state erosion island” (cf. Fig. 11), possibly invalidating the argument of saturation concentration for 10Be (Stone 2000). Production rates in artificial targets. Cosmogenic nuclide production rates have also been determined in artificial targets exposed for a few years to cosmic ray irradiation. In an early study, Yokoyama et al. (1977) measured 22Na and 24Na in three metal targets (mainly Al) exposed in the Mont Blanc region, France. Of more relevance to surface exposure dating today are the experiments of Nishiizumi et al. (1996) and Graf et al. (1996), who exposed water and SiO2 targets, respectively, at mountain altitudes in Colorado, and that of Brown et al. (2000), who exposed water tanks at different altitudes in France. Nishiizumi et al. (1996) converted their results from H2O to SiO2 based on neutron irradiation experiments (Reedy et al. 1994) and argued that the average production rate over 10,000 years should be ~ 15% higher than that during the last four solar cycles due to geomagnetic field changes. In this way they achieved an agreement within 2% with the value obtained by Nishiizumi et al. (1989) in Sierra Nevada quartz (Table 6). However, a 15% difference between the present “instantaneous” production rate and that integrated over 10 ka is at odds with calculations of Dunai (2001b) and Masarik et al. (2001), who expect production rate variations of less than a few percent within that period of time. On the other hand, Brown et al. (2000) found an inconsistency between the neutron irradiation results of Reedy et al. (1994) and their observation of similar 3He/10Be production ratios in water and quartz, rendering the conversion of the 10 Be production rate from H2O to SiO2 uncertain. In their water targets, Brown et al. (2000) also determined the 3H/3He production ratio from oxygen as 0.32 ± 0.08.

The 21Ne production rate in a SiO2 target during exposure on Mt. Evans (4250 m) was 410 ± 60 atoms g-1 a-1 (Graf et al. 1996). As detailed by Niedermann (2000), this converts to a sea level, high latitude value of 17.7 atoms g-1 a-1 over four solar cycles if Dunai’s (2000a) scaling method is applied (or 21.6 atoms g-1 a-1 according to Lal 1991). The agreement with the production rate in Sierra Nevada quartz is well within experimental uncertainties and does not indicate a substantial production rate variation either. Production rates by muons. Experimental determinations of cosmogenic nuclide production by muons are scarce, and for noble gases even absent. In a depth profile from lateritic soil in the Congo, Brown et al. (1995a) were able to discern the muon-produced component of 10Be in quartz and constrain its contribution at the surface (300 m altitude)

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to between 1 and 3% of the total 10Be production, i.e., ~ 0.05-0.15 atoms g-1 a-1 scaled to sea level and high latitudes. A detailed study of 36Cl production in a 20 m marble profile from an Australian quarry was carried out by Stone et al. (1998). These authors derived a normalized 36Cl production rate by negative muon capture in calcite of 2.1 ± 0.4 atoms g-1 a-1, ~ 10% of the total production. The higher fraction of muon-produced 36Cl as compared to 10Be can be explained by a relatively high cross section for α emission after μ– capture, favoring the reaction 40Ca(μ–,α)36Cl. No experimental determinations of production rates by muons for other cosmogenic nuclides are available in the current literature. Therefore production models (e.g., Heisinger et al. 1997; Heisinger and Nolte 2000) should be used in all cases where production by muons might be relevant. Future needs. The data in Table 6 show the present state of production rate determinations in natural samples. There are quite a lot of consistent data for 3He in olivine, a few less consistent ones for 21Ne in olivine, and just one for 21Ne in SiO2. For 10 Be and 26Al, the three determinations in SiO2 agree within ~ 10-20%. One possible problem is the bias of production rate determinations to locations in the western USA and Hawaii, which may introduce a systematic error caused by the scaling method. In this respect, the Canary Island data of Dunai and Wijbrans (2000) are closer to the western hemisphere data when scaled by Dunai’s (2000a) method. On the other hand, the 10Be and 26Al production rates from California, Antarctica, and Austria scatter less if scaled by Lal’s (1991) procedure. However, different ages of calibration samples may also contribute to scatter of production rates. To investigate the reasons for such scatter, more determinations at various locations all over the globe are required. Ideally, these determinations should be made using samples of similar age. In addition, to better resolve temporal variations samples of sufficiently different ages from a limited area should be studied.

It is also evident from Table 6 that cosmogenic nuclide production rates should be determined for more compositions of target elements. Though not all minerals are likewise suited for surface exposure dating with 3He and 21Ne (see Experimental issues below), there are certainly more than just olivine and quartz. And finally, a determination of the contribution to 3He and 21Ne production from muon-induced reactions would be important to assess exposure histories of samples from shielded locations or where erosion is high. Production rates obtained by model calculations

In the early years of surface exposure dating production rate estimates of terrestrial cosmogenic nuclides relied mainly on models. Yokoyama et al. (1977) calculated production rates of various radionuclides in granite, basalt, and limestone, based on the flux and energy spectrum of cosmic ray particles in the atmosphere and on nuclear excitation functions, mainly for protons. Their 3H production rate, when scaled to sea level and high latitude, is ~ 60 atoms g-1 a-1, in good agreement with the 3He production rate established later (Table 6) if 3He is produced in equal parts directly and via 3H, although the latter assumption is questioned at least for O spallation by the results of Brown et al. (2000). The 10Be and 26Al production rates are not directly comparable since measured values are only available for SiO2. However, a value of < 2 atoms g-1 a-1 for 10 Be in granite (assumed SiO2 content 73%) is obviously much too low, whereas ~ 40 atoms g-1 a-1 for 26Al seem rather high. Lal (1991) estimated production rates for the noble gas isotopes 3He, 20,21,22Ne, and 36,38 Ar from the target elements O, Mg, Al, Si, Ca, and Fe. His 3He rates (47-83 atoms -1 -1 g a for O, Mg, Al, Si) are clearly below those determined in real samples (Table 6),

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Niedermann

and the 21Ne production rate in Si of 18 atoms (g Si)-1 a-1 is 2.5 times lower than that found by Niedermann et al. (1994). Lal (1991) was aware of the limited validity of his estimates which he ascribed mainly to the lack of applicable nuclear excitation functions. In this respect the work of Masarik and Reedy (1995) was a major step forward. These authors relied on cross sections as derived from measurements in extraterrestrial samples, where at sufficient shielding depths cosmogenic nuclide production is dominated by neutrons also. Indeed their production rates agree remarkably well with experimental values (Table 6). In a follow-up paper, Masarik and Reedy (1996) reported elemental coefficients for the production rates of 10Be, 14C, 26Al, 36Cl as well as 3He and 21 Ne, which are reproduced in Table 7. Using these coefficients it is possible to estimate production rates for any elemental composition of the target rock or mineral, i.e., also for minerals for which no experimental production rate determinations have been carried out. The good agreement of the “tested” production rates in quartz and olivine (Masarik and Reedy 1995) could suggest a high confidence in the elemental coefficients as well, at least for the major target elements. Another set of elemental production rates for the Ne isotopes (Schäfer et al. 1999), which is based on calculated neutron spectra in the centers of spherical meteorites, does however not agree very well with Masarik and Reedy’s (1996) data (see Table 7). Schäfer et al. (1999) argue that their coefficients provide a better consistency with experimental evidence from pyroxenes. EXPERIMENTAL ISSUES Retentivity of minerals for cosmogenic He and Ne

As for other geochronological methods (cf. Kelley 2002; Farley 2002), a prerequisite to successful surface exposure investigations is the quantitative retention of cosmogenic nuclides in the minerals studied. In general, it can be assumed that minerals with typically high noble gas concentrations from other sources (e.g., Carroll and Draper 1994), such as mantle or radiogenic gases, should also retain cosmogenic noble gases. However, a caveat to that contention is the distinct siting of different components within minerals. Trapped noble gases, especially those from crustal sources, are often concentrated in fluid inclusions, whereas the cosmogenic nuclei are produced within the crystal lattice. Unless there are cracks rendering the fluid inclusions leaky, diffusion loss from the lattice may be more severe than from inclusions. Because gases are preferentially partitioned in the Table 7. Elemental coefficients for the calculation of cosmogenic nuclide production rates (atoms g–1 a–1) in dependence of the chemical composition of the target mineral (Masarik and Reedy 1996). For example, P(26Al) is calculated as 225 [Al] + 77 [Si] + 0.15 [Fe], with element concentrations in weight fractions (g/g). For 21Ne, an alternative data set from Schäfer et al. (1999) is also reported. Nuclide 3

He

O 135

Na -

Mg 116

Al 107

Si 111

K

Ca

Ti

-

61

-

Fe 40

10

10.87

-

0.52

0.45

0.39

-

-

-

0.16

14

31.3

-

5.3

4.2

4.3

-

-

-

1.2

Be C

21

Ne

21

Ne

1

-

98

131

65

39

-

4

-

0.20

-

-

196

54

45

-

-

-

-

26

-

-

-

225

77

-

-

-

0.15

36

-

-

-

-

-

129

65

16

Al Cl

1

Schäfer et al. (1999)

0.9

Noble Gases as Dating Tools for Surficial Processes

763

fluid phase, they will tend to remain within fluid inclusions rather than diffuse through the crystal. Nevertheless, the minerals olivine and pyroxene, which are commonly used to investigate mantle noble gases, do retain cosmogenic He and Ne quantitatively, as implied by their low diffusion coefficients (Hart 1984; Trull et al. 1991) and demonstrated by ample data, such as identical concentrations of 3Hec in coexisting olivine and pyroxene grains (e.g., Craig and Poreda 1986; Cerling 1990) or the consistent 3He production rates discussed in the preceding section. Plagioclase is another mineral which has been tested for cosmogenic 3He and 21Ne retention. Cerling (1990) found virtually no 3 Hec; obviously this mineral is unsuited for 3He exposure dating. Poreda and Cerling (1992) concluded that the retention of 21Ne in volcanic plagioclase is satisfactory, based on constant ratios of 21Nec in plagioclase to 3Hec or 21Nec in coexisting olivine, independent of the age. They derived a normalized 21Ne production rate of 16.8 ± 1.7 atoms g-1 a-1 for Ab36 plagioclase, which seems, however, somewhat low. According to the formula of Masarik and Reedy (1996; cf. Table 7), 23.4 atoms g-1 a-1 21Ne are expected in Ab36 plagioclase, i.e., 40% more. In the case of Fo81 olivine, the agreement between Poreda and Cerling’s (1992) production rate and that calculated from Masarik and Reedy (1996) is very good: 45 ± 4 versus 44.4 atoms g-1 a-1. Therefore it is doubtful whether plagioclase indeed retains 21Ne quantitatively. Bruno et al. (1997) report diffusion losses of ~ 50% for plagioclase from Antarctic dolerites, confirming a poor retentivity of plagioclase for Ne as well. The extent to which 3Hec is lost from quartz by diffusion was a matter of debate in the early nineties (e.g., Cerling 1990; Graf et al. 1991; Trull et al. 1991; Brook and Kurz 1993). It can be concluded that 3Hec may be retained in quartz more or less quantitatively in favorable cases, e.g., in large grains and at prevailing low temperatures, but is more often lost to high percentages and can therefore not commonly be applied to exposure age studies. In contrast, 21Nec in quartz is obviously well-suited, in spite of relatively low degassing temperatures (Niedermann et al. 1993; see next section). This is indicated by a general consistency of 10Be, 26Al, and 21Ne exposure ages (e.g., Bruno et al. 1997; Schäfer et al. 1999; Hetzel et al. 2002a). Minerals other than those mentioned above have so far not been extensively used in cosmogenic noble gas studies. Judging from its generally high noble gas content (e.g., Carroll and Draper 1994), amphibole could be another candidate for quantitative He and Ne retention among relatively widespread minerals, but this supposition would have to be checked. Dunai and Roselieb (1996) reported a very high He retentivity of garnet and therefore expected a great potential for use in exposure age studies. However, high U and Th concentrations are quite common in garnet, so radiogenic He and nucleogenic Ne may interfere with the cosmogenic components (see next section). The same problem is expected for micas although they might as well retain Ne and possibly He. Farley et al. (2001) studied cosmogenic 3He in fluorapatite and expected good retention characteristics of this mineral, which is widely used in U/Th-He dating (cf. Farley 2002), but again radiogenic and nucleogenic components may be abundant. Discrimination of cosmogenic against trapped, radiogenic, and nucleogenic components

To allow for an accurate determination of the concentration of a cosmogenic noble gas isotope, noble gas components from other sources must be effectively discriminated against. These components include: •

trapped components, i.e., atmospheric, mantle, or crustal gases residing in the crystal lattice or in fluid inclusions. They may originate from solution in the magma,

764

Niedermann

incorporation of air or water during eruption, fluid interaction during metamorphic events, etc. •

radiogenic components produced by radioactive decay of U, Th, and 40K.

•

nucleogenic components produced by naturally occurring nuclear reactions, e.g., 6 Li(n,α)3H(β–)3He, 18O(α,n)21Ne, 19F(α,n)22Na(β+)22Ne, 24,25Mg(n,α)21,22Ne. In these reactions the α particles are derived from U and Th decay and the neutrons from other (α,n) reactions or from U fission.

Sample selection and preparation. Whereas cosmic ray irradiation is often the only important source of radionuclides in a rock, the stable noble gases are always mixtures of several components. To minimize uncertainties arising from decomposition of components, samples with high ratios of cosmogenic to non-cosmogenic gases should be used. High concentrations of cosmogenic isotopes, i.e., old samples taken from the surface at high altitude and latitude and where erosion is low, are of course most favorable, but choice in this respect is usually limited when specific questions are to be solved. Low concentrations of non-cosmogenic components are therefore essential.

There is no general recipe to avoid high concentrations of trapped noble gases. Depending on their source and the mechanism of incorporation, they may occupy various sites in a rock or mineral, but often they are concentrated in fluid inclusions. Therefore a low abundance of fluid inclusions in the minerals studied may be a useful criterion for sample selection. In addition, the sample may be crushed to small grain sizes before loading, so that gases are released from part of the fluid inclusions. At least for quartz, this method proved useful in reducing the amount of trapped Ne (Niedermann et al. 1994). On the other hand, care must be taken to avoid “irreversible” adsorption of atmospheric noble gases on fresh grain surfaces (e.g., Niedermann and Eugster 1992), which may not be a problem for Ne in quartz but might affect certain other minerals. Radiogenic and nucleogenic noble gas components are abundant in minerals with high concentrations of U, Th, and K, particularly when their age of formation is relatively high. Dating young lava flows, for which the mineral formation age and the surface exposure age are roughly identical, is therefore easier than studying landforms consisting of old rock which was exposed relatively recently. Typically low concentrations of U, Th, and K in olivine and pyroxene are another argument for studying those minerals, besides their good He and Ne retentivity. In quartz, these elements are also rare, but in practice there are often inclusions or intergrowths of U-rich minerals, such as biotites, especially in granitic host rocks. Clean quartz separates are therefore essential to avoid problems with interfering nucleogenic Ne components. Recent experience in Potsdam indicates that quartz clasts derived from very low-grade metasediments contain much less nucleogenic Ne than those from granites (Niedermann and Hermanns 1999), but on the other hand vein quartz may carry a lot of trapped gases in fluid inclusions (Hetzel et al. 2002a). Since α particles from U/Th decay have a range of 10-40 μm (e.g., Ziegler 1977), they may also be implanted from neighboring minerals into the quartz. Etching the quartz grains is therefore advisable to remove the surface layer. In addition to the quartz separation procedure of Kohl and Nishiizumi (1992), which was developed for 10Be and 26 Al analyses and includes etching by HF and HNO3, Niedermann et al. (1994) used a heavy liquid to obtain density fractions. The 2.66-2.69 g/cm3 fraction was least affected by nucleogenic Ne, whereas the heavy and light fractions were probably contaminated by U/Th-bearing mineral and fluid inclusions, respectively. Nucleogenic 3He produced by the reaction 6Li(n,α)3H(β–)3He is only important in minerals with high Li content. The 3He/4He ratio resulting from U/Th decay and 6Li neutron capture can be estimated from the chemical composition of the mineral and

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relevant nuclear parameters (Mamyrin and Tolstikhin 1984) and is, for common rock compositions, on the order of 10–8. In Li-rich minerals, such as hornblende or micas, substantially elevated 3He/4He ratios could however be established, rendering the distinction of a cosmogenic 3He component difficult. Therefore such minerals are probably not well suited for 3He surface exposure dating. Crushing versus melting. Those non-cosmogenic components which cannot be avoided by careful sample selection and preparation must be corrected for. Usually the first step in such a procedure is to calculate the excess of a noble gas isotope (e.g., 3He, 21 Ne) over the composition of the trapped gas: 3

Heex = [(3He/4He)m – (3He/4He)tr] × 4Hem

21

Neex = [(21Ne/20Ne)m – (21Ne/20Ne)tr] × 20Nem

(34a) (34b)

where indices ex, m, and tr mean excess, measured, and trapped. These equations are approximations assuming that 4Heex and 20Neex can be neglected. The isotope ratio in the trapped component must either be assumed or determined independently, if possible. Since (3He/4He)tr ratios may vary by several orders of magnitude in terrestrial samples, from ~ 1 × 10–8 to ~ 5 × 10–5, it is essential to use correct values. A reliable method to determine the (3He/4He)tr ratio in a sample is crushing in vacuo, which releases only gases stored in fluid inclusions but not those from the crystal lattice, where cosmogenic nuclides reside. This technique was already applied in the earliest studies of terrestrial cosmogenic He (Kurz 1986a,b; Craig and Poreda 1986) and yields consistent results under the prerequisite that the trapped He compositions in fluid inclusions and matrix are identical and that there is no contribution from radiogenic 4He produced in situ. These conditions are met at least for olivines from young volcanic rocks. In older rocks and in U/Th-rich minerals, a contribution from radiogenic 4He in the crystal lattice must be expected, which does not show up by crushing and requires a correction. Likewise, implantation of α particles may have a sizable effect if U/Th concentrations in neighboring minerals or the host rock are distinctly higher (cf. Dunai and Wijbrans 2000). Trapped Ne is usually assumed to have an atmospheric isotopic composition (21Ne/20Ne = 0.002959, 22Ne/20Ne = 0.1020; Eberhardt et al. 1965), but this is not always correct. In the mantle rocks studied by the Paris group (Staudacher and Allègre 1991, 1993a,b; Sarda et al. 1993), the presence of mantle Ne was demonstrated by elevated 20 Ne/22Ne ratios; these authors assumed mixtures of atmospheric and MORB-type Ne for the trapped component. Niedermann et al. (1994) found evidence for two distinct trapped Ne components in quartz from a Sierra Nevada granite, being somewhat fractionated from air in both directions. Hetzel et al. (2002a) applied in vacuo crushing to quartzite from the northeastern margin of the Tibetan Plateau and reported 21Ne/20Ne ratios up to 0.0039 (30% higher than atmospheric). Unlike He, for which part of the cosmogenic component is released along with the trapped one when crushing quartz (Brook and Kurz 1993), cosmogenic Ne is obviously retained. The non-atmospheric Ne in vein quartz was most probably trapped from crustal fluids. The crushing data are indeed representative of the trapped Ne composition, as shown by consistent 21Ne and 10Be ages, whereas assuming atmospheric trapped Ne yields erroneously high 21Ne ages (Hetzel et al. 2002a). Stepwise heating. Stepwise heating is another method for discriminating between cosmogenic and non-cosmogenic components. Although a quantitative separation of different components is seldom possible, distinct degassing characteristics may nevertheless yield significant information. The few papers which report stepwise heating data for cosmogenic He show the major release of 3Hec from mafic and ultramafic

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minerals (olivine, pyroxene) below ~ 900-1100°C (Kurz 1986a; Staudacher and Allègre 1991, 1993a; Sarda et al. 1993; Schäfer et al. 2000). Extraction temperatures for mantle He are somewhat higher in some samples, but similar in others, whereas radiogenic 4He is released concurrently with 3Hec (Kurz 1986a; Schäfer et al. 2000). The benefit of the stepwise heating technique for studies of cosmogenic helium is thus limited, at least in the minerals mentioned. Ne is different. In olivine and pyroxene, only minor amounts of cosmogenic Ne are degassed at temperatures ≤ 900°C, whereas the release of trapped Ne is relatively uniform over the whole temperature range (Staudacher and Allègre 1991, 1993a; Schäfer et al. 2000). A partial separation of trapped and cosmogenic Ne is therefore achievable by stepwise heating at ~ 900°C and 1600-1800°C, allowing for a more accurate determination of excess Ne due to enhanced 21Ne/20Ne and 22Ne/20Ne ratios in the steps with a reduced “background” of trapped Ne. A good resolution of different components is even more important for quartz, because the contribution of nucleogenic Ne is often higher than in pyroxenes and olivines. Niedermann et al. (1993) showed that in quartz ALH 85-4, separated from an Antarctic sandstone, > 97% of 21Nec is degassed at temperatures below 600°C. The main release of trapped Ne occurs at somewhat higher temperatures of ~ 400-800°C. These results are confirmed by studies of other quartz samples (Hudson et al. 1991; Phillips et al. 1998; Niedermann et al. 2001a), though it seems safer to increase the limit below which complete degassing of cosmogenic Ne occurs to 800°C. Figure 13 shows the 21Nec release pattern of six quartz samples. Despite some differences in the details, which are probably caused by distinct diffusion characteristics due to variations in the effective grain size (e.g., Trull and Kurz 1993), about 90% of 21Nec or more is degassed below

Figure 13. Release pattern of cosmogenic 21Ne for six quartz samples degassed by stepwise heating at 400°C, 600°C, 800°C, and ≥ 1000°C. Samples SLB 8-1, 6-1, 3-1, and 5-1 are from metamorphic and granitic host rocks (Sierra Laguna Blanca, Argentina; Niedermann et al. 2001a), Qt2 117cm is a soil sample consisting of sand-sized grains (Pajarito Plateau, New Mexico; Phillips et al. 1998), and ALH 85-4/1 is from a sandstone boulder (Allan Hills, Antarctica; Niedermann et al. 1993). Despite some differences in the details, it is obvious that virtually no cosmogenic Ne is left in these quartz samples above 800°C.

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600°C, and < 1.5% is left above 800°C in all samples. Since analytical uncertainties of Ne abundance determinations are typically ≥5%, any contributions from heating steps > 800°C can therefore safely be neglected, provided that the six samples plotted are representative for any quartz. However, in two quartz separates of Bruno et al. (1997), the fractions of excess 21Ne remaining at 800°C are 43% and 23%, respectively. There are no indications for a significant nucleogenic Ne component in these samples, so the excesses observed at high temperatures must be cosmogenic as well. There are two possible reasons for the unusual degassing pattern: Bruno et al. (1997) acknowledge that sample temperatures may have been considerably below those of the crucible because of relatively short extraction times. In addition, they used rather large grain sizes of 300-500 μm. If these grains were exceptionally free of internal fractures, which determine the effective grain size for diffusion (Trull and Kurz 1993), it is possible that indeed higher temperatures were required to extract 21Nec. In addition to reducing the trapped component, crushing the quartz grains to ~100 μm or less will most probably ensure complete release of cosmogenic Ne at 800°C. The completeness of 21Nec extraction at 800°C is important whenever nucleogenic Ne is present. Niedermann et al. (1994) found two distinct components of nucleogenic Ne in quartz. One of them was released along with cosmogenic Ne predominantly between 100 and 600°C and consisted only of 21Ne produced by the 18O(α,n) reaction. The other one, termed NeHT (for “high-temperature”), was mainly extracted above 800°C and was characterized by a 21Ne/22Ne ratio of ~ 1, close to that of cosmogenic Ne. Whereas the monoisotopic low-temperature component can easily be distinguished from cosmogenic Ne based on three-isotope systematics (see next section), the discrimination of Nec from NeHT depends critically on the distinct release characteristics. The two components can also be identified in granitic quartz from Argentina (Niedermann et al. 2001a). Probably the monoisotopic component is produced by 18O(α,n)21Ne reactions in the quartz crystal lattice, which would explain the similar release pattern to cosmogenic Ne. NeHT, however, may be located in solid or fluid inclusions, in which the presence of F along with U and Th enables 22Ne production through the reaction 19F(α,n)22Na(β+)22Ne. The higher temperatures required to extract NeHT would then be connected to the distinct retentivity of the inclusions. Significant contributions of nucleogenic 22Ne at temperatures < 600°C are reported in granitic quartz from Tibet by Schäfer et al. (2002). These authors suppose that microinclusions of biotite are the source of this component. It is indeed possible that reduced diffusion lengths in very small-grained biotites may decrease the temperature required for release of nucleogenic Ne. On the other hand, Schäfer et al. (2002) observed the nucleogenic signature also by in vacuo crushing, indicating that this component resides at least partly in fluid inclusions. Hence another explanation could be the presence of a crustal trapped component (see Crushing versus melting above) rather than in situ production. A better understanding of the relative importance of different components in various rock types will require more experimental experience as well as detailed petrographic observations, combined with other microanalytical techniques. Neon three-isotope systematics. For Ne, information about different components can also be extracted from three-isotope systematics. Figure 14 schematically shows the signatures of Ne components and their mixing lines with atmospheric Ne in three-isotope space. It is customary in cosmogenic Ne studies to use the 20Ne normalization, not 22Ne as in most other applications, because it facilitates the identification of 21Ne and 22Ne excesses. Moreover, the determination of 21Ne/20Ne is sometimes more accurate than that of 21Ne/22Ne because of the isobaric interference of CO2++ on m/e = 22. As cosmogenic Ne in a certain mineral is characterized by a unique 21Ne/22Ne ratio,

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Figure 14. Neon three-isotope diagram (22Ne/20Ne vs. 21Ne/20Ne) showing the compositions and trends of various Ne components. Mixtures of atmospheric and cosmogenic Ne plot on the spallation line, the slope of which has been experimentally determined for quartz and pyroxene (Table 8). Contributions from nucleogenic 21Ne and 22Ne are characterized by shifts in horizontal and vertical directions, respectively. The trends for crustal Ne (Kennedy et al. 1990) and MORB-type Ne (Sarda et al. 1988) are also given. The dotted line labeled mfl is the mass fractionation line.

any two-component mixture of atmospheric and cosmogenic Ne must lie on the “spallation line” in Figure 14. Deviations from that line indicate contributions from other components, such as nucleogenic Ne or non-atmospheric trapped Ne (be it mantle, crustal, or mass-fractionated air Ne). Of course, a position on the spallation line does not necessarily imply a cosmogenic origin, as fortuitous mixtures of nucleogenic 21Ne and 22 Ne (such as NeHT) or mantle and crustal Ne in the right proportions might also plot there. However, as long as there are not more than three components the contribution of each one can unequivocally be calculated. For example, if a three-component mixture of atmospheric Ne, cosmogenic Ne, and nucleogenic 21Ne is indicated with sufficient confidence, the 21Nec concentration is obtained from 21

Nec =

20

Ne m

( 21 Ne/ 20 Ne)c × m

⎡⎣( 22 Ne/ 20 Ne)air − ( 22 Ne/ 20 Ne) m ⎤⎦ ⎡⎣( 21 Ne/ 20 Ne)air − ( 21 Ne/ 20 Ne)c ⎤⎦

(35)

where m is the slope of the spallation line. Equation (35) corresponds to a horizontal shift of a data point in the three-isotope diagram to the spallation line and calculation of 21Neex for the shifted point according to (34b). Such a procedure is of course only applicable for data points lying to the right of the spallation line. The slope of the spallation line depends on the relative production ratios of cosmogenic 20Ne, 21Ne, and 22Ne and, therefore, on the mineral composition. It has been determined experimentally for quartz and pyroxene (Niedermann et al. 1993; Bruno et al. 1997; Phillips et al. 1998; Schäfer et al. 1999). The relevant data are compiled in Table 8. The values for (21Ne/20Ne)c were assumed based on observations in well-shielded lunar samples and meteorites and model calculations (Hohenberg et al. 1978; Schäfer et al.

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Table 8. Slope of the spallation line in a 22Ne/20Ne vs. 21Ne/20Ne three-isotope plot and Ne production ratios for quartz and pyroxene. The values for 22Ne/20Ne and 22Ne/21Ne are based on the experimental determination of the slope and an assumed 21Ne/20Ne ratio. Mineral

Slope

Quartz

1.120±0.021

Ne/20Ne

22

Ne/20Ne

22

Ne/21Ne

Reference

0.8±0.1 1.05±0.20 1 0.8±0.1

0.99±0.13 1.26±0.25 1 0.98±0.15

1.243±0.022

1.143±0.038

0.8±0.1

1.01±0.13

1.266±0.040

Schäfer et al. (1999)

1.055±0.017

1.05±0.20

1.21±0.23

1.15±0.03

Bruno et al. (1997)

1.069±0.035

1.10±0.20

1.27±0.24

1.159±0.040

Schäfer et al. (1999)

0.78

1.02

1.31

Hohenberg et al. (1978) 2

0.83

0.94

1.14

Schäfer et al. (1999)

0.91

0.79

0.86

Hohenberg et al. (1978) 2

1.10

1.11

1.01

Schäfer et al. (1999)

+0.05

1.10 −0.03 1.10±0.10 Pyroxene

21

+0.07

1.20 −0.04 1.22±0.10

Niedermann et al. (1993) Bruno et al. (1997) Phillips et al. (1998)

Model calculations: Quartz Pyroxene 21

20

1

Assumed Ne/ Ne too high, see text

2

Lunar surface, shielding depth 500 g/cm2

1999; Leya et al. 2000). The (21Ne/20Ne)c ratio of 1.05 ± 0.20 assumed by Bruno et al. (1997) is probably too high for quartz, since the production of 21Ne and 22Ne is only favored over that of 20Ne in Mg-containing minerals (such as pyroxene) due to large cross sections of the 24,25Mg(n,α)21,22Ne reactions. However the (22Ne/21Ne)c ratio is rather insensitive to that assumption. Likewise, the large error limits of the (21Ne/20Ne)c assumption have only minor influence on the precision of (22Ne/21Ne)c. Indeed, the agreement between individual determinations of the slope and the (22Ne/21Ne)c ratio is excellent, with error limits of only 2-3%. Any deviations from a two-component mixture of atmospheric and cosmogenic Ne in quartz or pyroxene can therefore clearly be recognized and the due corrections be applied. Cosmogenic noble gases as a nuisance

Even in a paper intended to boost the prospects of cosmogenic nuclides, their unpleasant sides should not remain unmentioned. Cosmogenic He and Ne components are not always welcome, because they may overprint other noble gas signatures and can even lead to erroneous conclusions when they are not recognized. A drastic example is the discovery of extremely high 3He/4He ratios up to 3.2 × 10–4 (230 RA) in South African diamonds by Ozima et al. (1983), prior to the first reports of cosmogenic 3He in terrestrial rocks. Based on these data the authors inferred that the initial terrestrial 3He/4He ratio was higher than that of planetary He and close to the solar value, and they discussed implications for the Earth’s accretion. However, as Lal et al. (1987) and McConville and Reynolds (1989) have shown, such high 3He/4He ratios are most probably due to cosmic ray irradiation during surface residence of the diamonds, which had been purchased from commercial sources. Today it has become clear that samples from the surface (or undocumented samples) are unsuitable for investigations of the isotopic compositions of mantle He and Ne, unless they are very young or were excavated only recently from well-shielded locations. Suboceanic rocks are of course not affected, but for continental samples, depending on

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the age and the altitude and latitude of the sampling location, a few meters of shielding are typically required to ensure that cosmogenic 3He and 21Ne will be low enough not to interfere with isotopic signatures of mantle, crustal, or nucleogenic components. Fortunately, since samples from well-shielded locations are not always easy to obtain, there is one possibility to extract information on trapped gases even from old surface samples, namely crushing in vacuo. As already mentioned, crushing only liberates gases from fluid inclusions, whereas the lattice-bound cosmogenic components are obviously retained (except for He in quartz; Brook and Kurz 1983). In any noble gas investigation dealing with continental rocks, possible contributions of cosmogenic isotopes have thus to be assessed. Despite careful sampling, problems cannot always be avoided. Matsumoto et al. (2000) studied mantle xenoliths from the Newer Volcanics, Australia, and found 3He/4He ratios up to 8 × 10–5 in some low temperature steps, although the rocks were sampled in active quarries where they were buried until recently. These authors report a striking decoupling of He and Ne in the xenoliths, with MORB-like 3He/4He and “nucleogenic” 21Ne/22Ne ratios, which they interpret in terms of mantle metasomatism. While Matsumoto et al. (2000) acknowledge the presence of cosmogenic He and Ne in a few of their samples, they dismiss it for most of them. However, as judged from the isotope systematics, I believe that the observed 21 Ne excesses may well be of cosmogenic origin. Noble gas data on Romanian mantle xenoliths obtained in Potsdam have revealed a similar component with high 3He/4He and high 21Ne/22Ne ratios (Althaus et al. 1998), despite 5-10 m of shielding by volcanic rocks. Though we earlier interpreted the extremely high 3He/4He ratios of > 30 RA as plumederived, we now consider the possibility whether muon-induced reactions might have produced enough cosmogenic 3He and 21Ne at such depths since eruption (~ 0.8 Ma), assuming reasonable erosion rates. Cosmogenic noble gases versus radionuclides

In the early years of terrestrial cosmogenic nuclide studies, the numbers of publications dealing with radionuclides and with noble gas isotopes, respectively, were more or less equal. However, later the radionuclide studies began to overwhelm, especially those with 10Be and 26Al. Among the reasons for such imbalance are experimental difficulties with noble gases as described above, first of all the interference of non-cosmogenic components. It is clear that radionuclides are superior in this respect, as contributions from such components are usually negligible or (as in the case of meteoric 10Be) can reliably be removed (cf. Kohl and Nishiizumi 1992), but noble gases have other advantages: •

As stable nuclides, the noble gases record cosmic ray irradiation at any time in the past. This offers possibilities to investigate exposure histories which are longer than the limit of a few million years imposed by the half-life of 10Be, the longest-lived radionuclide commonly used (see Glacier movement and ice sheet evolution below). Equivalently, very low erosion rates can only be quantified with noble gases. Even for exposure ages which are still within reach of the 10Be method, results get inaccurate and depend strongly on assumptions for the production rate when saturation levels are approached (e.g., Stone 2000). Moreover, complex exposure histories involving prior surface residence followed by burial on a timescale of several million years can only be identified using cosmogenic noble gases along with radionuclides. It is even possible to study “paleo-exposures” (Libarkin et al. 2002), i.e., surface exposures which occurred many million years ago in rock which was later covered permanently.

•

Typically, the sample size needed for noble gas studies is much smaller than for

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radionuclide analyses. Whereas 0.5-1 g of quartz is usually sufficient to determine 21 Nec concentrations on the 106 atoms/g level with < 10% uncertainty (2σ), some 20-30 g of quartz are required to achieve a similar accuracy for 10Be and 26Al. Therefore noble gases may be the better choice when suitable material is scarce. •

So far, the 10Be and 26Al methods have been successfully applied in quartz only. In other minerals, the separation of meteoric 10Be has not reliably been attained, and the high concentrations of stable 27Al render the detection of 26Al impossible (e.g., Gosse and Phillips 2001). 36Cl can be determined in more minerals, but the interpretation of results is often difficult due to various production mechanisms and uncertain production rates (Gosse and Phillips 2001), and relatively few labs have used it for surface exposure dating. The noble gases provide an alternative for mafic minerals and may work with other minerals also (Retentivity of minerals for cosmogenic He and Ne section).

•

The production rate of 3He is much higher than for any other cosmogenic nuclide. In minerals where it is quantitatively retained, 3He thus provides the chance to date exceptionally young surfaces. For example, we have derived an age of 13+−21 13 years for a sample from the 1993 lava flow of Láscar volcano (Chile) taken at 4540 m altitude a few months after eruption (Niedermann et al. 2001b), which illustrates the potential of 3He on timescales also relevant to archeological studies. New developments in noble gas mass spectrometry, such as the “compressor ion source” which improves the mass spectrometer sensitivity for He and Ne by two orders of magnitude (Baur 1999), may further increase the precision of 3He (as well as 21Ne) determinations in the future.

For many applications a combined study of cosmogenic noble gases and radionuclides will be useful, for example in order to identify and evaluate complex exposure histories. Clearly, both techniques should be further developed without concentrating too much on just one of them. APPLICATION EXAMPLES

In the last part of this review, I will present an overview of relevant work on terrestrial in situ produced cosmogenic nuclides. The intention is to show the broad range of possible applications and refer interested readers to the cited literature, without trying to provide detailed recapitulations of the various investigations nor claiming for completeness. Many of the cited studies were carried out with radionuclides; however there is no reason why similar studies should not be possible with noble gases as well. Dating of lava flows

Lava flows were among the first surface features to be studied by cosmogenic nuclides. In many cases they are indeed ideal subjects, because the conditions for simple exposure histories are obviously met, and concentrations of radiogenic/nucleogenic isotopes are generally low since the rocks are only as old as their surface residence. The only major difficulty in dating lava flows is the influence of erosion, as was recognized even in the earliest studies (Craig and Poreda 1986; Kurz 1986a,b). Erosion effects tend to be more severe in humid climates (e.g., Hawaii) and for old lava flows. Cerling (1990) dated four lava flows at Owens Valley, California, and concluded that the 3He ages of the younger flows (~ 13 and ~ 57 ka) were reasonable, while flows with K-Ar ages of > 2 Ma yielded 3He ages too low by an order of magnitude. Erosion rates of 2-3 m/Ma or a parttime cover are required to account for the difference. Anthony and Poths (1992) compared their 3He ages for lavas in the Potrillo volcanic field, New Mexico, with age estimates based on K-Ar and accumulation of calcium carbonate in soils and found

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consistent results, except for one set of K-Ar data. Linking their findings with compositional trends, they were able to establish a time sequence of magma evolution in individual eruptive centers and the volcanic field overall. Staudacher and Allègre (1993b) measured cosmogenic 3He and 21Ne in rocks from two eruption phases of Piton de la Fournaise volcano (Réunion) and derived an age of 23.8 ± 2.0 ka for the collapse of its second caldera. These examples illustrate the potential of cosmogenic nuclides for dating of lavas and volcanic eruptions, especially when other methods fail due to the young age of volcanic rocks or lack of 14C-datable material. Glacier movement and ice sheet evolution

Another important field of in situ cosmogenic nuclide studies is the history of glacier advance and retreat or the evolution of polar ice sheets. Such questions are directly linked to global climate variations. Many papers have therefore dealt with glacial landforms on the Antarctic continent, trying to gain information regarding their stability under changing climatic conditions. Apart from their potential significance to social or political issues, such studies are particularly suited for the method because exceptionally high exposure ages of several million years and exceptionally low erosion rates (~ 0.1-1m/Ma) promote accurate determinations of cosmogenic nuclide concentrations. This was shown in the earliest studies of Antarctic rocks (Nishiizumi et al. 1986, 1991a; Brown et al. 1991; Graf et al. 1991). One possible problem in the interpretation of Antarctic exposure histories is the low air pressure over that continent, which causes production rates substantially higher than expected as recognized only recently (Stone 2000). In this respect the stable noble gas isotopes provide more robust data than radionuclides having concentrations close to the saturation level (Stone 2000). In the following summary of relevant work it must be remembered that the reported exposure ages, which do not take account of the pressure effect, may be too high by ~ 25%. Investigations have concentrated on the Dry Valleys in East Antarctica, for obvious reasons because the overwhelming majority of the continent is permanently covered by ice. Brook et al. (1993) studied boulders from a sequence of Taylor Glacier moraines and found significant scatter of ages within a single deposit. Moraine boulders are certainly not ideal samples for surface exposure dating because of frequent complex exposure histories (e.g., pre-exposure prior to glacier transport, past soil cover, boulders of an older deposit exposed between younger material, etc.). Nevertheless Brook et al. (1993) could place constraints on regional uplift rates and ice thickness variability within a timeframe of 2-3 Ma, and other studies also show that equivocal results can largely be avoided by careful sampling. Brook et al. (1995) dated clasts from drift sheets deposited by advances of the Ross Sea ice and again got widely scattered results. However, in this case the authors argued that the scatter may indicate several distinct drift depositions connected with episodes of ice sheet grounding, because the relatively young ages (8-106 ka for the younger drift) are less likely to be affected by erosion or soil cover and pre-exposure can in most cases be ruled out as the material is derived from the sea floor. Ivy-Ochs et al. (1995), Bruno et al. (1997), Schäfer et al. (1999), Summerfield et al. (1999), and Van der Wateren et al. (1999) determined 10Be, 3He, and 21Ne minimum exposure ages of erratic boulders, glacial sediments, and bedrock samples from various locations in the Dry Valleys region and the Transantarctic Mountains. These data include the highest exposure ages for terrestrial rocks reported to date, 10.08 ± 0.24 Ma for a dolerite boulder from Mount Fleming (Schäfer et al. 1999) and 11.2 ± 0.2 Ma for a bedrock sample from Daniels Range (Van der Wateren et al. 1999). Such high exposure ages can only be resolved with the stable noble gas isotopes, whereas radionuclides are at secular equilibrium between production and decay. Moreover, erosion rates have to be exceptionally low, < 6 cm/Ma in the case of the Mount Fleming dolerite. These data

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prove the extreme stability of the Dry Valleys landscape since at least Miocene time and exclude any significant climate variations during the Pliocene warm period. They support models of a stable East Antarctic Ice Sheet and limit uplift rates of the Transantarctic Mountains to ~ 100 m/Ma. Schäfer et al. (2000) investigated the sublimation rate of a remnant ice body covered by a layer of till in the Dry Valleys region. Based on the exposure ages of two erratic boulders, one on top of the till and the other one sticking in the ice, they derived a sublimation rate of a few m/Ma at most and a minimum age of the remnant ice of 3 Ma. Ackert et al. (1999) measured 3He and 36Cl in moraine boulders deposited by a high-stand of the West Antarctic Ice Sheet and found evidence against a significant contribution to the eustatic sea level rise at 11 ka from West Antarctic Ice Sheet meltwater. Such studies have undoubtedly demonstrated the relevance of cosmogenic nuclide studies in Antarctica to global climate issues. A few cosmogenic nuclide papers have dealt with glaciation history in temperate regions. 36Cl dating of moraine boulders at Bloody Canyon, California, enabled Phillips et al. (1990) to establish the chronology of glacial sequences in the eastern Sierra Nevada and correlate them to global ice volume and the marine oxygen isotope record. In a similar study, Fernandez Mosquera et al. (2000) examined the sequence of deglaciation in the northwestern Iberian Peninsula (Spain and Portugal), based on cosmogenic 21Ne in moraine boulders and bedrock. Bierman et al. (1999) sampled bedrock outcrops in Minnesota and Baffin Island, at the southern and northern paleo-margins of the Laurentide Ice Sheet, respectively. Based on 26Al/10Be ratios which were lower than expected for a simple exposure history, these authors proposed a model involving periods of temporary burial and concluded that the surfaces were at least 0.5 Ma old, despite individual exposure age data of only ~ 100 ka. An important conclusion from that work is the fact that the burial periods would have remained undetected if only one nuclide (or only stable nuclides) had been measured. The extent of Pleistocene glaciations in the Himalayas and the Tibetan Plateau has been the subject of several recent studies. The paleoclimate in that region my have significant consequences for the South Asian monsoon system and even for global climate. Owen et al. (1999) and Phillips et al. (2000) dated glacial deposits in Northern Pakistan with cosmogenic radionuclides and arrived at different conclusions regarding the timing of glacial advances. Schäfer et al. (2002) applied 10Be, 26Al, and 21Ne dating of erratic boulders to two locations in Central and Eastern Tibet. They found surprisingly old ages of 160-170 ka for moraines in the Central Plateau, excluding a major advance of glaciers in that area ever since. Extensive glaciations of the Tibetan Plateau, which could have substantially influenced the monsoon system or even triggered the onset of glacial advances in the whole northern hemisphere, are obviously not consistent with these data. Rates of erosion and soil accumulation

Studies of erosion rates have been another important application for in situ produced cosmogenic nuclides, since quantitative data from other methods are scarce. The earlier papers focused mainly on exposure ages and provided erosion rate estimates only as a byproduct. Craig and Poreda (1986), Kurz (1986b), and Nishiizumi et al. (1990) derived erosion rates of 8-12 m/Ma for lavas from Haleakala volcano, Maui (Hawaii), by comparing the nominal surface exposure ages with established lava ages. Sarda et al. (1993) estimated a somewhat lower erosion rate in a similar climate at Piton de la Fournaise volcano, Réunion. These authors pointed out that erosion rates derived from river loads or from neotectonic studies are often much higher, up to 1000 m/Ma, most probably because samples for cosmogenic studies are typically taken from relatively stable locations and are therefore not representative for vast areas where erosion is

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dominated by river incision. A similar conclusion was reached by Small et al. (1997), who assessed erosion rates of alpine bedrock summits in the western USA and inferred that the mean value of 7.6 ± 3.9 m/Ma was similar to erosion rates in other environments, excluding very arid regions. Extremely low erosion rates were reported for the Antarctic Dry Valleys, as already mentioned in the preceding section. Values on the order of 0.1 m/Ma or less are obviously typical for the cold and arid climate conditions prevailing there (Nishiizumi et al. 1991a; Graf et al. 1991; Ivy-Ochs et al. 1995; Bruno et al. 1997; Schäfer et al. 1999; Summerfield et al. 1999). Even for samples taken from slopes inclined by 36-38°, Summerfield et al. (1999) found erosion rates ≤ 1m/Ma. In sharp contrast, locations affected by glacier incision may have suffered downcutting rates of some 1000m/Ma (Van der Wateren et al. 1999). In warmer climates under arid conditions, denudation is about an order of magnitude faster than in the extremely stable Dry Valleys surfaces. Bierman and Turner (1995) deduced a rate of ~ 0.7 m/Ma for the tops of granitic inselbergs in South Australia. Stone and Vasconcelos (2000) determined lowest rates of 1-2 m/Ma for the highest surface settings near Mt Isa (Queensland), whereas 2-5 m/Ma were typical at lower elevations. These authors also pointed out the dependence of erosion rates on rock types, being two to five times lower in silicate rocks than in limestone. Fleming et al. (1999) and Cockburn et al. (2000) used erosion rate determinations with cosmogenic radionuclides to assess the rates of escarpment retreat at the passive continental margins of eastern South Africa and Namibia, respectively. In both cases the retreat rates were shown to be much too low to be compatible with an escarpment origin at the continental margin during the time of continent break-up. Brown et al. (1995b) and Bierman and Steig (1996) developed a method to estimate mean erosion rates for entire river basins from cosmogenic nuclide concentrations in river sediment. Under the assumptions that production rates within the catchment area do not vary much, that cosmogenic nuclide concentrations are in erosion equilibrium, and that sediments are not stored for long times on the valley floor, the basin-wide erosion rate is given by ε =

P0 Λ C ρ

(36)

where symbols are as in Equations (26)-(33) and the brackets indicate average values. Brown et al. (1995b) demonstrated the suitability of their method by comparing the derived value of 43 m/Ma for the Icacos river basin, Puerto Rico, with erosion rates determined from 10Be concentrations in samples from ridge crests and hill slopes and from river-load mass balance studies. In a subsequent investigation Brown et al. (1998a) applied the technique to assess the anthropogenic influence on denudation in another Puerto Rican watershed affected by agricultural exploitation. Schaller et al. (2001) studied the catchments of four middle European rivers with drainage areas from 3000 to 43,000 km2 and found that erosion rates based on cosmogenic nuclides were higher by a factor of 1.5-4 than those obtained from river-load gauging. They proposed several possible causes for the difference and concluded that cosmogenic nuclides may provide better estimates of long-term erosion rates over timescales of some 104 years. Granger et al. (2001) assessed the dependence of erosion rates on slope gradient by studying 10Be and 26Al in bedrock and river sediment from granitic terrain in the northern Sierra Nevada, California. Their results demonstrate that erosion is controlled by the extent of bedrock outcropping, due to much faster weathering of soil-covered rock compared to bare rock, which may in some cases lead to higher erosion rates on valley bottoms and

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gentle footslopes than on steep mountain slopes. Soil accumulation on a stable surface may, in principle, be regarded as negative erosion. However, there are some complications. One of these is the presence of cosmogenic nuclides inherited from earlier exposure on hillslopes or during transport, which will vary among different sediment clasts. As shown by Anderson et al. (1996) and Repka et al. (1997), such inheritance can be treated by amalgamating a statistically adequate number of clasts and by comparing measured depth profiles of cosmogenic nuclide concentrations with those expected for zero inheritance. There may be several types of depth profiles, depending on the rate and persistence of soil accretion (short-term deposition or steady accumulation) and the role of bioturbation within the soil. Explicit treatments of the theoretical background as well as application examples are given by Phillips et al. (1998) and Braucher et al. (2000). Rates of tectonic uplift

Cosmogenic nuclides can be used in two ways to estimate rates of tectonic uplift. For samples which were exposed long enough, such as those in the Antarctic Dry Valleys (see Glacier movement and ice sheet evolution section), the mere concentration of a cosmogenic nuclide provides constraints on the uplift rate because of the dependence of production rates on altitude. A maximum uplift rate is obtained under the assumption of steady uplift from sea level to the present elevation (cf. Bruno et al. 1997). This method has been applied to limit uplift in the Transantarctic Mountains to < 170 m/Ma (Brook et al. 1993; Ivy-Ochs et al. 1995; Bruno et al. 1997), as opposed to certain studies suggesting much higher rates up to 1000 m/Ma. The second method utilizes the relationship between river incision and uplift. Rivers in tectonically active regions compensate for uplift by downcutting. In the course of this process, river benches (straths) or fluvial terraces are created, which are later abandoned as incision continues. Therefore the time since abandonment of the strath or terrace, along with its elevation above the present valley floor, provides a measure for tectonic uplift, provided that a climatic control on terrace formation and abandonment can be ruled out. Burbank et al. (1996) and Leland et al. (1998) investigated incision of the Indus river in the northwestern Himalayas. They derived incision rates between 2 and 12 km/Ma (or mm/a) and argued for an equilibrium between incision and bedrock uplift in this region, because otherwise the longitudinal river profile would have been severely perturbed. In other tectonic settings the direct equation of incision and regional uplift may not apply as well. However, on a more local scale, the relative shifts in altitude caused by tectonic fault scarps can be assessed quite safely. Brown et al. (1998b) used 10Be to date alluvial fans on the southern margin of the Tien Shan mountains (China), which were offset by tectonic faults. Combining their ages with the vertical offsets, they deduced vertical slip rates of 0.4-1.1 km/Ma. In a similar study at the northeastern margin of the Tibetan plateau, Hetzel et al. (2002b) found slip rates of 0.3-0.5 km/Ma for the Yumen thrust fault north of the Qilian Shan mountains, an order of magnitude less than assumed earlier based on the assumption of Holocene ages for the tectonically offset alluvial fan surfaces. Earthquakes and landslides

Alluvial fans cut by tectonic faults can also yield information on the frequency of earthquakes. An offset debris flow fan in Owens Valley (California) was dated with 10Be by Bierman et al. (1995). Based on boulder ages from different parts of the flow and the record of three faulting events, the authors deduced earthquake recurrence intervals of 5800-8000 years for that location. Zreda and Noller (1998) measured 36Cl in a bedrock fault scarp at Hebgen Lake (Montana) and were able to extract the ages of six prehistoric

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earthquakes dating from 24 to 0.4 ka. Hermanns et al. (2001) determined 21Ne exposure ages of eight giant landslides and one tectonically uplifted terrace at Sierra Laguna Blanca (Argentina). Based on landslide ages of 150-430 ka, these authors argued for tectonic oversteepening of the mountain front generating repeated collapses, probably triggered by earthquakes larger than magnitude 7.5, whereas a relation to wetter climatic conditions was deemed unlikely. The cessation of landsliding after ~ 150 ka was interpreted as marking thrust-front migration towards the western piedmont, consistent with uplift of the terrace dated at ~ 85-130 ka. Further applications

A few more examples are finally presented to illustrate the range of applications of terrestrial cosmogenic nuclide studies. On the moon, cosmogenic nuclides have been a common means of dating meteorite impact craters (e.g., Eugster et al. 1977). A similar study on a terrestrial impact crater, namely Meteor Crater in Arizona (also known as Barringer crater or Canyon Diablo), was carried out by Nishiizumi et al. (1991b) and Phillips et al. (1991), using 10Be, 26Al, and 36 Cl as well as rock-varnish 14C. They obtained consistent ages of 49.2 and 49.7 ka, respectively, in agreement with thermoluminescence studies but a factor of two higher than earlier estimates based on soil development and 14C in lake sediments on the crater floor. Wells et al. (1995) tested several competing models of desert stone pavement formation by comparing the 3He exposure ages of pavement clasts with those of basalt outcrops from which they were derived. They obtained identical ages, which they judged as clear evidence that clasts had remained at the surface continuously and were not concentrated randomly by processes such as surface runoff or upward migration within soils. Small et al. (1999) developed a method to estimate the rate of regolith production from bedrock based on cosmogenic nuclide concentrations and a mass balance model. They found regolith production rates of 7-20 m/Ma on alpine hillslopes in the Wind River Range, Wyoming, about a factor of 2 higher than bedrock erosion in similar environments, perhaps due to increased frost weathering in the water-bearing regoliths. Cerling et al. (1999) determined 3He exposure ages of boulders collected from four different debris flow fans at a tributary river mouth in Grand Canyon, Arizona. Together with radiocarbon dates obtained for the younger surfaces, they were able to establish a chronology for the episodic debris flows and draw conclusions regarding the relation between magnitudes and recurrence intervals of debris flows at the studied location. Future prospects

Besides the various themes which have already been covered by cosmogenic nuclide studies as shown in the above sections, new fields may become amenable to the method in near future. I can only give a very limited outlook based on recent case studies. Three investigations have tried to apply cosmogenic nuclides to issues of paleontology or archeology. Boaretto et al. (2000) made an effort to date cave sediments and flint tools from a prehistoric site in Israel. Although they could not derive the ages of the investigated layers due to experimental difficulties, their concept might be promising in the future. Ivy-Ochs et al. (2001) dated chert flakes from Egypt, which were produced by prehistoric men when making flintstone tools. Although they found reasonable ages of ~ 300 ka, the question what proportion of the exposure took place before the cobbles were worked on could not yet be answered. Farley et al. (2001) observed cosmogenic 3He

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in fossil tooth enamel fluorapatite. Provided that He is quantitatively retained in tooth enamel, which the authors expect based on the retention of radiogenic 4He, cosmogenic He could provide a new tool for dating fossil remains. Libarkin et al. (2002) measured cosmogenic 21Ne in quartz from the 28 Ma old Fish Canyon Tuff (Colorado), which is covered by later-emplaced tuff and sediment layers. Comparing the 21Nec concentration with the time interval during which Fish Canyon Tuff was exposed at the surface, which is constrained by 40Ar/39Ar dating of the two tuff layers, these authors discuss implications for paleoerosion rates and paleoelevation in that area of the Colorado Rocky Mountains. Libarkin et al.’s (2002) study shows a novel perspective especially for applications of cosmogenic noble gas isotopes, as radionuclides do not record such ancient exposures. Renne et al. (2001) have presented the first report of cosmogenic Ar in terrestrial rocks. Though the background of atmospheric Ar is usually higher than for Ne, a substantial production rate in Ca-rich rocks (e.g., Hohenberg et al. 1978; Lal 1991) makes its detection feasible. In fluorapatite, fluorite, sphene, and plagioclase samples from Antarctica and Namibia, they found elevated 38Ar/36Ar ratios up to 0.289 by total fusion and 0.364 by stepwise heating (atmosphere: 0.188). The 38Ar/36Ar ratios in Antarctic apatites correlate with 3Hec concentrations, confirming the cosmogenic nature of the 38Ar excesses. However, estimates for the production rates of 38Ar and 36Ar are not consistent for apatite and fluorite samples and also disagree with calculations of Lal (1991). The Ar production rate systematics are complicated by 36Ar production through decay of the cosmogenic radionuclide 36Cl (half life 3×105a), which is produced in part by neutron capture of 35Cl (Table 1). Nevertheless, cosmogenic Ar may provide a valuable additional tool for surface exposure dating once production rates are better constrained, as it will be applicable in minerals which do not retain He and Ne quantitatively and in particular in Ca-rich rocks such as limestones. In the future, investigations of terrestrial cosmogenic nuclides produced in situ in rock material will certainly become more and more widespread. As shown in the preceding sections, applications of the method span many fields of Earth sciences and beyond. Accuracy of results is expected to improve as both experimental methods and knowledge of the background, such as production rates, are refined. The use of other minerals and additional nuclides will further expand the range of treatable subjects. ACKNOWLEDGMENTS

I thank Joe Licciardi for unpublished data and Ralf Hetzel and Rainer Wieler for comments on the manuscript. Reviews by Ken Farley, Simon Kelley, William Phillips, and Jörg Schäfer were very helpful to better achieve this article’s purpose to become a survey of the current state of the art in cosmic ray exposure dating using noble gases. REFERENCES Ackert RP Jr, Barclay DJ, Borns HW Jr, Calkin PE, Kurz MD, Fastook JL, Steig EJ (1999) Measurements of past ice sheet elevations in interior west Antarctica. Science 286:276-280 Althaus T, Niedermann S, Erzinger J (1998) Noble gases in ultramafic mantle xenoliths of the Persani Mountains, Transylvanian Basin, Romania. Mineral Mag 62A:43-44 Anderson RS, Repka JL, Dick GS (1996) Explicit treatment of inheritance in dating depositional surfaces using in situ 10Be and 26Al. Geology 24:47-51 Anthony EY, Poths J (1992) 3He surface exposure dating and its implications for magma evolution in the Potrillo volcanic field, Rio Grande Rift, New Mexico, USA. Geochim Cosmochim Acta 56:4105-4108 Barbouti AI, Rastin BC (1983) A study of the absolute intensity of muons at sea level and under various thicknesses of absorber. J Phys G: Nucl Phys 9:1577-1595

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Kurz MD (1986a) Cosmogenic helium in a terrestrial igneous rock. Nature 320:435-439 Kurz MD (1986b) In situ production of terrestrial cosmogenic helium and some applications to geochronology. Geochim Cosmochim Acta 50:2855-2862 Kurz MD (1987) Erratum. Geochim Cosmochim Acta 51:1019 Kurz MD, Colodner D, Trull TW, Moore RB, O’Brien K (1990) Cosmic ray exposure dating with in situ produced cosmogenic 3He: results from young Hawaiian lava flows. Earth Planet Sci Lett 97:177-189 Lal D (1988) In situ-produced cosmogenic isotopes in terrestrial rocks. Ann Rev Earth Planet Sci 16: 355-388 Lal D (1991) Cosmic ray labeling of erosion surfaces: in situ nuclide production rates and erosion models. Earth Planet Sci Lett 104:424-439 Lal D, Nishiizumi K, Klein J, Middleton R, Craig H (1987) Cosmogenic 10Be in Zaire alluvial diamonds: implications for 3He contents of diamonds. Nature 328:139-141 Lal D, Peters B (1967) Cosmic ray produced radioactivity on the earth. In Handbook of Physics, 46/2. Springer, Berlin, p 551-612 Leland J, Reid MR, Burbank DW, Finkel R, Caffee M (1998) Incision and differential bedrock uplift along the Indus River near Nanga Parbat, Pakistan Himalaya, from 10Be and 26Al exposure age dating of bedrock straths. Earth Planet Sci Lett 154:93-107 Leya I, Lange H-J, Neumann S, Wieler R, Michel R (2000) The production of cosmogenic nuclides in stony meteoroids by galactic cosmic-ray particles. Meteorit Planet Sci 35:259-286 Libarkin JC, Quade J, Chase CG, Poths J, McIntosh W (2002) Measurement of ancient cosmogenic 21Ne in quartz from the 28 Ma Fish Canyon Tuff, Colorado. Chem Geol 186:199-213 Licciardi JM, Kurz MD, Clark PU, Brook EJ (1999) Calibration of cosmogenic 3He production rates from Holocene lava flows in Oregon, USA, and effects of the Earth’s magnetic field. Earth Planet Sci Lett 172:261-271 Liu B, Phillips FM, Fabryka-Martin JT, Fowler MM, Stone WD (1994) Cosmogenic 36Cl accumulation in unstable landforms 1. Effects of the thermal neutron distribution. Water Resour Res 30:3115-3125 Mamyrin BA, Tolstikhin IN (1984) Helium Isotopes in Nature. Elsevier Science Publishers, Amsterdam Marti K, Craig H (1987) Cosmic-ray-produced neon and helium in the summit lavas of Maui. Nature 325:335-337 Marti K, Eberhardt P, Geiss J (1966) Spallation, fission, and neutron capture anomalies in meteoritic krypton and xenon. Z Naturforschung 21a:398-413 Masarik J, Beer J (1999) Simulation of particle fluxes and cosmogenic nuclide production in the Earth’s atmosphere. J Geophys Res 104:12099-12111 Masarik J, Frank M, Schäfer JM, Wieler R (2001) Correction of in situ cosmogenic nuclide production rates for geomagnetic field intensity variations during the past 800,000 years. Geochim Cosmochim Acta 65:2995-3003 Masarik J, Kollar D, Vanya S (2000) Numerical simulation of in situ production of cosmogenic nuclides: Effects of irradiation geometry. Nucl Instr Meth Phys Res B172:786-789 Masarik J, Reedy RC (1995) Terrestrial cosmogenic-nuclide production systematics calculated from numerical simulations. Earth Planet Sci Lett 136:381-395 Masarik J, Reedy RC (1996) Monte Carlo simulation of in situ produced cosmogenic nuclides. Radiocarbon 38:163-164 Matsumoto T, Honda M, McDougall I, O’Reilly SY, Norman M, Yaxley G (2000) Noble gases in pyroxenites and metasomatised peridotites from the Newer Volcanics, southeastern Australia: implications for mantle metasomatism. Chem Geol 168:49-73 McConville P, Reynolds JH (1989) Cosmogenic helium and volatile-rich fluid in Sierra Leone alluvial diamonds. Geochim Cosmochim Acta 53:2365-2375 Merrill RT, McElhinny MW, McFadden PL (1998) The Magnetic Field of the Earth. Academic Press, San Diego Niedermann S (2000) The 21Ne production rate in quartz revisited. Earth Planet Sci Lett 183:361-364 Niedermann S, Althaus T, Hahne K (2001b) A minimum age for Llullaillaco south flow from cosmogenic 3 He: much older than 19th century. In III South American Symposium on Isotope Geology, Extended Abstracts Volume (CD), Sociedad Geológica de Chile, Santiago, Chile, p 56-59 Niedermann S, Eugster O (1992) Noble gases in lunar anorthositic rocks 60018 and 65315: Acquisition of terrestrial krypton and xenon indicating an irreversible adsorption process. Geochim Cosmochim Acta 56:493-509 Niedermann S, Graf T, Kim JS, Kohl CP, Marti K, Nishiizumi K (1994) Cosmic-ray-produced 21Ne in terrestrial quartz: the neon inventory of Sierra Nevada quartz separates. Earth Planet Sci Lett 125: 341-355 Niedermann S, Graf T, Marti K (1993) Mass spectrometric identification of cosmic-ray-produced neon in terrestrial rocks with multiple neon components. Earth Planet Sci Lett 118:65-73

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Niedermann S, Hermanns RL (1999) The chronology of giant landslides at Sierra Laguna Blanca, Argentina, as deduced from cosmogenic 21Ne. European Geophysical Society, Geophys Res Abstr 1:43 Niedermann S, Hermanns RL, Strecker MR (2001a) 21Ne surface exposure dating of giant landslides at Sierra Laguna Blanca, Argentina: Evidence for tectonic control on slope oversteepening. In III South American Symposium on Isotope Geology, Extended Abstracts Volume (CD), Sociedad Geológica de Chile, Santiago, Chile, p 413-416 Nishiizumi K, Finkel RC, Caffee MC, Southon JR, Kohl CP, Arnold JR, Olinger CT, Poths J, Klein J (1994) Cosmogenic production of 10Be and 26Al on the surface of the earth and underground. Proc 8th Intl Conf Geochron, Cosmochron and Isotope Geol, U S Geol Surv Circular 1107:234 Nishiizumi K, Finkel RC, Klein J, Kohl CP (1996) Cosmogenic production of 7Be and 10Be in water targets. J Geophys Res 101:22225-22232 Nishiizumi K, Klein J, Middleton R, Craig H (1990) Cosmogenic 10Be, 26Al, and 3He in olivine from Maui lavas. Earth Planet Sci Lett 98:263-266 Nishiizumi K, Kohl CP, Arnold JR, Klein J, Fink D, Middleton R (1991a) Cosmic ray produced 10Be and 26 Al in Antarctic rocks: exposure and erosion history. Earth Planet Sci Lett 104:440-454 Nishiizumi K, Kohl CP, Shoemaker EM, Arnold JR, Klein J, Fink D, Middleton R (1991b) In situ 10Be26 Al exposure ages at Meteor Crater, Arizona. Geochim Cosmochim Acta 55:2699-2703 Nishiizumi K, Lal D, Klein J, Middleton R, Arnold JR (1986) Production of 10Be and 26Al by cosmic rays in terrestrial quartz in situ and implications for erosion rates. Nature 319:134-136 Nishiizumi K, Winterer EL, Kohl CP, Klein J, Middleton R, Lal D, Arnold JR (1989) Cosmic ray production rates of 10Be and 26Al in quartz from glacially polished rocks. J Geophys Res 94: 17907-17915 O’Brien K, Sandmeier HA, Hansen GE, Campbell JE (1978) Cosmic ray induced neutron background sources and fluxes for geometries of air over water, ground, iron, and aluminum. J Geophys Res 83:114-120 Ohno M, Hamano Y (1992) Geomagnetic poles over the last 10,000 years. Geophys Res Lett 19:1715-1718 Ohno M, Hamano Y (1993) Global analysis of the geomagnetic field: Time variation of the dipole moment and the geomagnetic pole in the Holocene. J Geomag Geoelectr 45:1455-1466 Owen LA, Caffee M, Finkel RC, Gualtieri L, Spencer JQ, Richards B (1999) Timing of multiple glaciations throughout the Himalayas. Geol Soc Am Abstr Progr 31:A141 Ozima M, Zashu S, Nitoh O (1983) 3He/4He ratio, noble gas abundance and K-Ar dating of diamonds – An attempt to search for the records of early terrestrial history. Geochim Cosmochim Acta 47:2217-2224 Paneth FA, Reasbeck P, Mayne KI (1952) Helium 3 content and age of meteorites. Geochim Cosmochim Acta 2:300-303 Phillips FM, Leavy BD, Jannik NO, Elmore D, Kubik PW (1986) The accumulation of cosmogenic chlorine-36 in rocks: a method for surface exposure dating. Science 231:41-43 Phillips FM, Zreda MG, Flinsch MR, Elmore D, Sharma P (1996) A reevaluation of cosmogenic 36Cl production rates in terrestrial rocks. Geophys Res Lett 23:949-952 Phillips FM, Zreda MG, Smith SS, Elmore D, Kubik PW, Dorn RI, Roddy DJ (1991) Age and geomorphic history of Meteor Crater, Arizona, from cosmogenic 36Cl and 14C in rock varnish. Geochim Cosmochim Acta 55:2695-2698 Phillips FM, Zreda MG, Smith SS, Elmore D, Kubik PW, Sharma P (1990) Cosmogenic chlorine-36 chronology for glacial deposits at Bloody Canyon, eastern Sierra Nevada. Science 248:1529-1532 Phillips WM, McDonald EV, Reneau SL, Poths J (1998) Dating soils and alluvium with cosmogenic 21Ne depth profiles: case studies from the Pajarito Plateau, New Mexico, USA. Earth Planet Sci Lett 160:209-223 Phillips WM, Sloan VF, Shroder JF Jr, Sharma P, Clarke ML, Rendell HM (2000) Asynchronous glaciation at Nanga Parbat, northwestern Himalaya Mountains, Pakistan. Geology 28:431-434 Plummer MA, Phillips FM, Fabryka-Martin J, Turin HJ, Wigand PE, Sharma P (1997) Chlorine-36 in fossil rat urine: An archive of cosmogenic nuclide deposition during the past 40,000 years. Science 277:538-541 Porcelli DR, Stone JOH, O’Nions RK (1987) Enhanced 3He/4He ratios and cosmogenic helium in ultramafic xenoliths. Chem Geol 64:25-33 Poreda RJ, Cerling TE (1992) Cosmogenic neon in recent lavas from the western United States. Geophys Res Lett 19:1863-1866 Radok U, Allison I, Wendler G (1996) Atmospheric surface pressure over the interior of Antarctica. Antarct Sci 8:209-217 Reedy RC, Nishiizumi K, Lal D, Arnold JR, Englert PAJ, Klein J, Middleton R, Jull AJT, Donahue DJ (1994) Simulations of in situ cosmogenic nuclide production. Nucl Instr Meth Phys Res B92:297-300 Renne PR, Farley KA, Becker TA, Sharp WD (2001) Terrestrial cosmogenic argon. Earth Planet Sci Lett 188:435-440

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Repka JL, Anderson RS, Finkel RC (1997) Cosmogenic dating of fluvial terraces, Fremont River, Utah. Earth Planet Sci Lett 152:59-73 Sarda P, Staudacher T, Allègre CJ (1988) Neon isotopes in submarine basalts. Earth Planet Sci Lett 91: 73-88 Sarda P, Staudacher T, Allègre CJ, Lecomte A (1993) Cosmogenic neon and helium at Réunion: measurement of erosion rate. Earth Planet Sci Lett 119:405-417 Schäfer JM, Baur H, Denton GH, Ivy-Ochs S, Marchant DR, Schlüchter C, Wieler R (2000) The oldest ice on Earth in Beacon Valley, Antarctica: new evidence from surface exposure dating. Earth Planet Sci Lett 179:91-99 Schäfer JM, Ivy-Ochs S, Wieler R, Leya I, Baur H, Denton GH, Schlüchter C (1999) Cosmogenic noble gas studies in the oldest landscape on Earth: surface exposure ages of the Dry Valleys, Antarctica. Earth Planet Sci Lett 167:215-226 Schäfer JM, Tschudi S, Zhao Z, Wu X, Ivy-Ochs S, Wieler R, Baur H, Kubik PW, Schlüchter C (2002) The limited influence of glaciations in Tibet on global climate over the past 170000 yr. Earth Planet Sci Lett 194:287-297 Schaller M, von Blanckenburg F, Hovius N, Kubik PW (2001) Large-scale erosion rates from in situproduced cosmogenic nuclides in European river sediments. Earth Planet Sci Lett 188:441-458 Shanahan TM, Zreda M (2000) Chronology of Quaternary glaciations in East Africa. Earth Planet Sci Lett 177:23-42 Shea MA, Smart DF, Gentile LC (1987) Estimating cosmic ray vertical cutoff rigidities as a function of the McIlwain L-parameter for different epochs of the geomagnetic field. Phys Earth Planet Inter 48: 200-205 Simpson JA (1983) Elemental and isotopic composition of the galactic cosmic rays. Ann Rev Nucl Part Sci 33:323-381 Small EE, Anderson RS, Hancock GS (1999) Estimates of the rate of regolith production using 10Be and 26 Al from an alpine hillslope. Geomorphology 27:131-150 Small EE, Anderson RS, Repka JL, Finkel R (1997) Erosion rates of alpine bedrock summit surfaces deduced from in situ 10Be and 26Al. Earth Planet Sci Lett 150:413-425 Srinivasan B (1976) Barites: anomalous xenon from spallation and neutron-induced reactions. Earth Planet Sci Lett 31:129-141 Staudacher T, Allègre CJ (1991) Cosmogenic neon in ultramafic nodules from Asia and in quartzite from Antarctica. Earth Planet Sci Lett 106:87-102 Staudacher T, Allègre CJ (1993a) The cosmic ray produced 3He/21Ne ratio in ultramafic rocks. Geophys Res Lett 20:1075-1078 Staudacher T, Allègre CJ (1993b) Ages of the second caldera of Piton de la Fournaise volcano (Réunion) determined by cosmic ray produced 3He and 21Ne. Earth Planet Sci Lett 119:395-404 Sternberg RS (1992) Radiocarbon fluctuations and the geomagnetic field. In Taylor RE et al. (eds) Radiocarbon After Four Decades. Springer, New York, p 93-116 Sternberg R (1996) Workshop on secular variations in the rates of production of cosmogenic nuclides on Earth: Paleomagnetic averages of geomagnetic latitude. Radiocarbon 38:169-170 Stone J, Vasconcelos P (2000) Studies of geomorphic rates and processes with cosmogenic isotopes— examples from Australia. J Conf Abstr 5:961 Stone JO (2000) Air pressure and cosmogenic isotope production. J Geophys Res 105:23753-23759 Stone JO, Allan GL, Fifield LK, Cresswell RG (1996) Cosmogenic chlorine-36 from calcium spallation. Geochim Cosmochim Acta 60:679-692 Stone JO, Evans JM, Fifield LK, Allan GL, Cresswell RG (1998) Cosmogenic chlorine-36 production in calcite by muons. Geochim Cosmochim Acta 62:433-454 Stuiver M, Reimer PJ, Bard E, Beck JW, Burr GS, Hughen KA, Kromer B, McCormac G, van der Plicht J, Spurk M (1998) INTCAL98 radiocarbon age calibration, 24,000-0 cal BP. Radiocarbon 40:1041-1083 Summerfield MA, Stuart FM, Cockburn HAP, Sugden DE, Denton GH, Dunai T, Marchant DR (1999) Long-term rates of denudation in the Dry Valleys, Transantarctic Mountains, southern Victoria Land, Antarctica based on in situ produced cosmogenic 21Ne. Geomorphology 27:113-129 Trull TW, Kurz MD (1993) Experimental measurements of 3He and 4He mobility in olivine and clinopyroxene at magmatic temperatures. Geochim Cosmochim Acta 57:1313-1324 Trull TW, Kurz MD, Jenkins WJ (1991) Diffusion of cosmogenic 3He in olivine and quartz: implications for surface exposure dating. Earth Planet Sci Lett 103:241-256 Valet J-P, Tric E, Herrero-Bervera E, Meynadier L, Lockwood JP (1998) Absolute paleointensity from Hawaiian lavas younger than 35 ka. Earth Planet Sci Lett 161:19-32

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Van der Wateren FM, Dunai TJ, Van Balen RT, Klas W, Verbers ALLM, Passchier S, Herpers U (1999) Contrasting Neogene denudation histories of different structural regions in the Transantarctic Mountains rift flank constrained by cosmogenic isotope measurements. Global Planet Change 23: 145-172 Wagner G, Masarik J, Beer J, Baumgartner S, Imboden D, Kubik PW, Synal H-A, Suter M (2000) Reconstruction of the geomagnetic field between 20 and 60 kyr BP from cosmogenic radionuclides in the GRIP ice core. Nucl Instr Meth Phys Res B172:597-604 Wells SG, McFadden LD, Poths J, Olinger CT (1995) Cosmogenic 3He surface-exposure dating of stone pavements: Implications for landscape evolution in deserts. Geology 23:613-616 Wieler R (2002) Cosmic-ray-produced noble gases in meteorites. Rev Mineral Geochem 47:125-170 Yokoyama Y, Reyss J-L, Guichard F (1977) Production of radionuclides by cosmic rays at mountain altitudes. Earth Planet Sci Lett 36:44-50 Ziegler JF (1977) Helium: Stopping powers and ranges in all elemental matter. Pergamon, New York Zreda MG, Noller JS (1998) Ages of prehistoric earthquakes revealed by cosmogenic chlorine-36 in a bedrock fault scarp at Hebgen Lake. Science 282:1097-1099 Zreda MG, Phillips FM, Elmore D, Kubik PW, Sharma P, Dorn RI (1991) Cosmogenic chlorine-36 production rates in terrestrial rocks. Earth Planet Sci Lett 105:94-109

17

K-Ar and Ar-Ar Dating Simon Kelley Department of Earth Sciences The Open University Milton Keynes MK7 6AA, United Kingdom [email protected]

INTRODUCTION — A BIT OF HISTORY The aim of this chapter is to present the K-Ar and Ar-Ar dating techniques in the context of noble gas studies, since there are already several recent texts on K-Ar and ArAr dating (Dickin 1995; McDougall and Harrison 1999). The focus of this section will be aspects of argon transport and storage in the crust, which affect K-Ar and Ar-Ar dating including Ar-loss from minerals by diffusion and Ar-gain by minerals or “excess argon.” The K-Ar dating technique was one of the earliest isotope dating techniques, developed soon after the discovery of radioactive potassium, and provided an important adjunct to U-Pb and U-He dating techniques. The ease of measurement and ideal half-life (1250 million years; see Table 2 below), for dating geological events has made this the most popular of isotopic dating techniques. Aldrich and Nier (1948) first demonstrated that 40Ar was the product of the decay of 40K, and soon after K-Ar ages were being measured in several laboratories most often using an absolute method such as a McLeod gauge to measure argon concentrations. The first published K-Ar results by such a technique were those of Smits and Gentner (1950) who analyzed sylvite from the Buggingen Oligocene evaporite deposits, obtaining an age of 20 million years. Mass spectrometers, which simultaneously measured very small amounts of gas, and the isotope ratios necessary to make corrections for atmospheric contamination, quickly replaced manometric techniques. Crucially the use of static vacuum techniques, pioneered by John Reynolds at the University of California-Berkeley, meant that mass spectrometers were sufficiently sensitive to analyse the small amounts of gas released from common rocks and minerals. Although the earliest mass spectrometers were built ‘in house’, the introduction of the commercially available MS10 (Farrar et al. 1964), a small 180° metal mass spectrometer built for leak testing, made K-Ar dating generally available. Complete descriptions of early K-Ar development and techniques can be found in Schaeffer and Zähringer (1966) and Dalrymple and Lanphere (1969). Although Thorbjorn Sigurgeirsson proposed the principles of Ar-Ar dating in an unpublished Icelandic laboratory report in 1962, he never succeeded in publishing or testing the idea. The Ar-Ar dating technique as it is practised today originated in the noble gas laboratory of John Reynolds in Berkeley, where Craig Merrihue and Grenville Turner were working on neutron irradiated meteorite samples using the I-Xe dating technique. Merrihue recognised that a 39Ar signal seen in the chart recorder traces was the result of neutron irradiation and published the idea in an abstract (Merrihue 1965). The publication of Merrihue and Turner (1966) saw the birth of the Ar-Ar dating technique. Written by Turner after the untimely death of Merrihue, this paper unusually describes a fully formed isotope dating technique (compare this with the slow emergence of the full K-Ar technique), possibly because the Berkeley Laboratory had been recording the full traces of all noble gases for some time, allowing Merrihue and Turner rapid access to a considerable database of measurements. The advantage of the Ar-Ar technique is that potassium and argon are effectively measured simultaneously on the same aliquot of sample, providing greater internal precision and also the ability to analyse very small and 1529-6466/02/0047-0017$05.00

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heterogeneous samples. Ar-Ar dating proved to be an ideal technique for dating meteorites because it made the best use of the extremely limited number of samples and also provided thermal histories. Indeed when lunar samples were returned from the Apollo 11 mission, Ar-Ar provided a crucial dating technique. Some samples were dated using K-Ar and yielded ages in the broad range 3 to 4 Ga, testifying to the antiquity of the lunar surface, although this much had been estimated from crater densities. In contrast, the Ar-Ar dating technique provided a wealth of precise ages and thermal histories. Using very small samples, Grenville Turner was able to unravel the crystallization histories, thermal histories during post-eruption heating and the cosmic ray exposure histories in a classic series of papers (Turner 1970b,c; 1971b, 1972). Turner applied quantitative diffusion concepts to stepwise argon release and recovered information from partially outgassed samples, establishing techniques and protocols that are still used to interpret stepwise heating Ar-Ar spectra today. This work and much of the early history of the Ar-Ar dating technique are set out in detail in McDougall and Harrison (1999). Although this chapter describes both K-Ar and Ar-Ar techniques, it should be noted that K-Ar dating is now important in only limited situations including standardization (i.e., first principles dating of standards), dating fine grained clay samples, dating young basalts and obtaining dates in rapid turnaround times. Ar-Ar dating is now used in a very wide range of geological applications, dating samples as old as lunar basalts and primitive meteorites, and volcanic rocks erupted only 2000 years ago. Ar-Ar dating has been applied to many areas of Earth Sciences for dating igneous, metamorphic and sedimentary events. In recent years the introduction of laser techniques for single spot and laser heating analysis has widened the range of applications for Ar-Ar dating and the introduction of more sophisticated models for stepwise heating continue to provide ever more detailed thermal histories from K-feldspars. THE K-AR AND AR-AR DATING METHODS Introduction Both K-Ar and Ar-Ar dating techniques are based upon the decay of a naturally occurring isotope of potassium, 40K to an isotope of argon, 40Ar (Fig. 1). The decay of 40 K is by a branching process; 10.48% of 40K decays to 40Ar by β+ decay (Beckinsale and Gale 1969, also proposed gamma-less electron capture decay but this has never been verified), followed by γ decay to the ground state, and by electron capture direct to the ground state, and 89.52% decays to 40Ca by β- to the ground state (Fig. 1). 40K-40Ca dating using the more common branch is also possible (e.g., Marshall and DePaolo 1982), but its application is generally restricted to old potassium-rich rocks since 40Ca is the most abundant naturally occurring isotope (96.94%), making the small amounts of radiogenically produced 40Ca very difficult to measure. Argon, in contrast, is a rare trace element and radiogenically produced 40Ar generally exceeds the levels of trapped 40Ar (although this is not always the case—see later). The naturally occurring isotopes of argon are measured by mass spectrometry for K-Ar dating (36Ar, 38Ar and 40Ar). The 36 Ar/38Ar ratio is almost constant (see Table 1), although cosmogenic 38Ar can be detected in some Ca-rich samples (Renne et al. 2001). Absolute argon concentrations, required for the K-Ar technique, are measured as a ratio against a known amount of 38Ar tracer gas. Mass spectrometry for Ar-Ar dating requires only isotope ratios between naturally occurring isotopes and also reactor produced isotopes 39Ar and 37Ar which have half lives of 269 years and 34.95 days respectively. As we shall see later, the irradiation procedure produces not only the radioactive isotopes but also small amounts of stable isotopes of argon, and it is thus important to measure all argon masses precisely by mass spectrometry in order to correct for neutron-induced interferences.

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Figure 1. Branching diagram showing the decay scheme for 40K, showing decay to 40Ar and 40Ca (after McDougall and Harrison 1999).

Table 1. Naturally occurring iso- topes of argon and potassium. Isotope

Abundance (%)

40

Ar Ar 36 Ar

99.600 0.632 0.336

39

93.2581 0.01167 6.7302

38

K K 41 K 40

After Steiger and Jäger (1977).

The essential difference between K-Ar and Ar-Ar dating techniques lies in the measurement of potassium. In K-Ar dating, potassium is measured generally using flame photometry, atomic absorption spectroscopy, or isotope dilution and Ar isotope measurements are made on a separate aliquot of the mineral or rock sample. In Ar-Ar dating, as the name suggests, potassium is measured by the transmutation of 39K to 39Ar by neutron bombardment and the age calculated on the basis of the ratio of argon isotopes.

Assumptions The “date” measured by both K-Ar and Ar-Ar techniques reflects the time since radiogenic argon produced by decay of 40K, became trapped in the mineral or rock. This may be the “age” of the rock or the most recent cooling event and in some samples may even reflect an integrated cooling age for a range of sub-grains. However, like all isotopedating techniques, there is a set of assumptions that must be valid if the number measured is to be interpreted as the age of a geological event: 1. The decay of the parent nuclide, potassium must be independent of its physical state. This is the standard assumption that must be valid for any isotope dating technique. 2. The 40K/K ratio must be a constant at any given time (Table 1). Most recently

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(Humayun and Clayton 1995a; Humayun and Clayton 1995b) measured a range of samples and found less than 0.05% variation in the 39K/41K ratio, even in samples where previous studies had measured some variation. 3. All radiogenic 40Ar measured in the sample results from 40K decay. The occasional presence of contaminating 40Ar from various sources can make determining the actual radiogenic content difficult but these are not strictly speaking radiogenic argon (see below). 4. Corrections can be made for any non-radiogenic argon. This is a simple procedure in terrestrial samples where there is generally some contaminating argon from the atmosphere (0.934% argon), but with a constant 40Ar/36Ar ratio of 295.5 (Table 1). Such corrections are less simple in extra-terrestrial samples where the initial 40 Ar/36Ar ratios are not constant, and are generally achieved using an isochron plot. Cosmogenic contributions are considered elsewhere in this volume (Wieler 2002; Niedermann 2002). 5. The sample, whether mineral or whole rock, must have remained a closed system since the event being dated. This includes gain or loss of either argon or potassium. This assumption is sometimes invalid, particularly in systems with complex geological and thermal histories. However, Ar-Ar stepwise heating and laser spot techniques can often be used to extract thermal history information from partially opened systems, taking advantage of the manner and extent of argon loss. Table 2. Decay constants for K-Ar and Ar-Ar dating. After Steiger and Jäger (1977). Decay by β40K→40Ar by electron capture and γ 40K→40Ar by electron capture combined value 40K→40Ca

Decay factor λβλe λ'e λ = λβ- + λec + λ'ec present day 40K/K

Value 4.962 × 10-10 a-1 0.572 × 10-10 a-1 0.0088 × 10-10 a-1 5.543 × 10-10 a-1 0.0001167

CALCULATING K-AR AND AR-AR AGES The age equation for the K-Ar isotope system is: t=

1 ⎡ λ ln ⎢1 + λ ⎣ λ e + λ' e

40

Ar * ⎤ ⎥ K ⎦

40

(1)

where t is the time since closure, λ is the total decay of 40K, and (λe + λ′e) is the partial decay constant for 40Ar (Begemann et al. 2001) (Table 2). 40Ar*/40K is the ratio of radiogenic daughter product (shown conventionally as 40Ar* to distinguish it from atmospheric 40Ar) to the parent 40K. Since there is no common natural fractionation of potassium isotopes (Humayun and Clayton 1995a,b), the modern ratio of 40K/K is a constant (Table 1), and thus measurement of potassium and argon concentrations together with isotope ratios of Ar, enable an age to be calculated. The Ar-Ar technique, first described by Merrihue and Turner (1966), is based on the same decay scheme as K-Ar, but instead of measurement on a separate aliquot, potassium is measured by creating 39Ar from 39K by neutron bombardment in a nuclear reactor. The reaction induced is: 39 19

K(n, p) 39 18 Ar

(2)

The ratio of 39K to 40K is effectively constant (see above) and thus the critical 40Ar*/40K

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ratio is proportional to the ratio of the two argon isotopes 40Ar/39Ar. Although 39Ar is radioactive, decaying with a half-life of 269 years, this effect is small for the period between irradiation and analysis (generally less than 6 months) and is easily corrected for. Mitchell (1968) showed that the number of 39Ar atoms formed during irradiation can be described by the equation: Ar = 39 K Δ∫ ϕ (ε) σ (ε)d(ε)

39

(3)

where 39K is the number of atoms, Δ is the duration of the irradiation, ϕ(ε) is the neutron flux density at energy ε, and σ(ε) is the neutron capture cross section of 39K for neutrons of energy ε for the neutron in/proton out reaction shown in Equation (2). Rearranging Equation (1) in terms of 40Ar* yields: Ar * = 40 K

40

λe + λ'e λt (e ) − 1 λ

[

]

(4)

Combining Equations (3) and (4) for a sample of age t yields: 40

Ar * = 39 Ar

[

]

(e λt ) −1 K λ e + λ' e 1 39 K λ ΔT ∫ ϕ(ε)σ(ε)d(ε) 40

(5)

This can be simplified by defining a dimensionless irradiation-related parameter, J, as follows: J=

39 40

K λ ΔT ∫ ϕ(ε )σ(ε)d(ε ) K λ e + λ' e

(6)

The J value is determined by using mineral standards of known age to monitor the neutron flux. Substituting Equation (6) into Equation (5) and rearranging, yields the Ar-Ar age equation: t=

40 1 ⎡ Ar * ⎤ ln ⎢1+ J 39 λ ⎣ Ar ⎥⎦

(7)

The ratio of the two isotopes of argon, naturally produced radiogenic 40Ar and reactor-produced 39Ar is thus proportional to the age of the sample. For terrestrial samples, the 40Ar peak measured in the mass spectrometer most often has two components (neglecting the 40ArK interference reaction), radiogenic and atmospheric. The 40 Ar/36Ar ratio of the atmosphere was determined by IUGS convention as 295.5 (Table 1; Steiger and Jäger 1977), though Nier determined a value of 296 (Nier 1950). When the 40 Ar/36Ar ratio of contaminating argon components is >295.5, the extra argon is known as extraneous argon. The term extraneous argon includes both excess and inherited argon following the terminology of Dalrymple and Lanphere (1969) and McDougall and Harrison (1999). Excess argon is the component of argon incorporated into samples by processes other than in situ decay, generally from a fluid or melt at the grain boundary. Inherited argon results from the incorporation of older material in a sample, such as for example grains of sand caught up in an ignimbrite eruption. However, in the simple case, assuming that all the non-radiogenic argon is atmospheric, the daughter/parent ratio (40Ar*/39Ar) can be determined from the equation: 40

Ar*/39Ar = [40Ar*/39Ar]m – 295.5[36Ar/39Ar]m

(8)

where subscript m denotes the measured ratio. This equation is always a simplification; in most terrestrial samples surface contamination ensures that some atmospheric argon is present, though fluids at depth rarely have atmospheric ratios (see below). However, in

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extraterrestrial samples atmospheric argon is recognised as a modern contamination or the result of weathering. In this light, it might seem strange to assume that all contaminating argon in terrestrial samples has an atmospheric isotope ratio, given that many trap argon at depth, not in equilibrium with the atmosphere. This puzzle will be discussed in some detail in the later section on argon diffusion and solubility. Sample irradiation for Ar-Ar dating induces not only Reaction (2) but also a series of interfering reactions caused by neutron bombardment of potassium, calcium, chlorine and argon. The complete series of interfering reactions is detailed in Table 3, but most have low production rates relative to the reaction in Equation (3) and can be ignored. The most important reactions are those involving calcium and potassium. The corrections are generally small, though they are critical for samples younger than 1 Ma when the interfering reactions producing 40Ar from 40K are important, and for samples with Ca/K > 10, when reactions producing 36Ar and 39Ar from isotopes of Ca become important. The magnitude of the interference from these reactions varies with the irradiation time and neutron flux energy spectrum. The range of measured interference factors for many of the world's reactors are listed in McDougall and Harrison (1999). The 42Ca(n,α)39Ar and 40Ca(n,nα)36Ar production ratios do not vary a great deal, because they are caused by fast neutrons and the energy spectrum of fast neutrons in most reactors is fairly similar. The far larger variation in the interference in the 40K(n,p)40Ar reaction is caused by its higher sensitivity to the ratio of fast to thermal neutrons in the reactor. This ratio varies between reactors and also between different irradiation positions within a reactor. In fact samples are often shielded with cadmium foil to reduce the thermal neutron flux and lower the efficiency of the 40K(n,p)40Ar reaction. The precise correction factors can be determined by irradiating pure salts of Ca and K (often CaF2, KCl and K2SO4). An additional correction must also be made for the decay of 37Ar (produced by neutron bombardment of calcium) which has a half-life of 34.95±0.08 days (Renne and Norman 2001). The short half-life of 37Ar means that all Ca-rich samples must be analyzed within about 6 months of irradiation otherwise the precision determining the original 37Ar concentrations may be affected, compromising the corrections to 36 Ar and 39Ar for Ca irradiation. Another factor affecting the accuracy of Ar-Ar dating is 39Ar recoil. This effect is crucial when studying very fine scale argon distributions or fine grained minerals such as clays, but 39Ar recoil from mineral surfaces can also affect high precision dating. Turner and Cadogan (1974), calculated the likely distances of 39Ar recoil during irradiation to be a mean of 0.08 μm, a study which was refined by Onstott et al. (1995) and measured directly by Villa (1997). The effects are most obviously detected in measurements of fine grained clays (e.g., Foland et al. 1983), but are commonly cited as a reason for variable ages produced from altered minerals (Lo and Onstott 1989), and basaltic rocks (e.g., Baksi 1994; Feraud and Courtillot 1994). The Ar-Ar technique is able to achieve higher levels of internal precision than K-Ar dating since it does not depend upon separate absolute measurements but instead requires only the ratios of Ar isotopes and can achieve precision of better than 0.25%. However, external precision and accuracy are affected by the uncertainty in the age of mineral standards, as we will see in the following section. In order to achieve optimum precision in the mass spectrometric measurements, the neutron flux (which affects the magnitude of the J value) must be carefully selected. The flux must be sufficient for precise measurement of 39Ar and a 40Ar*/39Ar ratio within the dynamic range of the mass spectrometer (generally less than 100 for good precision). Further, at higher flux levels the interfering reactions on Ca and K also become more important, degrading the precision and accuracy with which the 40Ar*/39Ar ratio may be determined. Therefore, for

40K(n,d)39Ar

42Ca(n,nα)38Ar

42Ca(n,α)39Ar 43Ca(n,nα)39Ar

38Ar

39Ar

40K(n,p)40Ar 41K(n,d)40Ar

43Ca(n,α)40Ar 44Ca(n,nα)40Ar

41K(n,α)38Cl→β-→38Ar

39K(n,d)38Ar

40Ar(n,d)39Cl→β-→39Ar

38Ar(n,γ))39Ar

40Ar(n,nd)38Cl→β-→38Ar

36Ar(n,γ)37Ar

Ar

37Cl(n,γ)38Cl→β-→38Ar

35Cl(n,γ)36Cl→β-→36Ar

Cl

The terminology (a,b) used here refers to nuclear reactions taking place during irradiation where a is the incident particle and b is the resulting emission. The terms are n=neutron, p = proton, d = deutron, α=and alpha particle,γ = a gamma particle and β- = a positron.

40Ar

39K(n,p)39Ar

40Ca(n,α)37Ar

37Ar

39K(n,nd)37Ar

40Ca(n,nα)36Ar

36Ar

K

Ca

Argon isotope

Important interfering reactions are shown in bold type. The main 39Ar-producing reaction is shown with a single border.

Table 3. Interfering reactions on Ca, K, Ar and Cl.

3

K-Ar and Ar-Ar Dating 791

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each sample there is an optimum flux level and given that many samples are irradiated together, each package sent for irradiation is a compromise. Turner (1971a) calculated the fields for optimum J value, and correspondingly integrated neutron flux, which were upgraded by McDougall and Harrison (1999) in the light of higher sensitivity, higher resolution mass spectrometers (Fig. 2).

Figure 2. A Figure for optimizing irradiation parameters, taking account of age and Ca/K. The irradiation parameter is plotted against Age (Ma) and zones of optimum irradiation level are highlighted (after McDougall and Harrison 1999; Turner 1971a).

The availability of five argon isotopes provided by the Ar-Ar technique facilitates isotope correlation plots, the most common of which is the three isotope plot 36Ar/40Ar vs. 39Ar/40Ar (Fig. 3). Samples containing a mixture of radiogenic and atmospheric Ar plot along a line with negative slope between the 39Ar/40Ar ratio representing the age and the atmospheric 36Ar/40Ar ratio of 0.003384 (= 1/295.5) (Fig. 3a). The correlation plot also allows Ar-Ar ages to be calculated for samples with contamination other than modern air, since the age can equally be determined from lines passing through the 36 Ar/40Ar axis at values other than the atmospheric ratio (Fig. 3b). However, a mixture of contaminating phases with more than one isotope composition in a sample results in a scatter of points not defining a line, and no age can be calculated. In many cases atmospheric ‘blank’ argon released from furnaces during heating is the only detected contaminating argon component. In cases where the contaminating argon is not homogeneous, physical techniques such as stepped heating, in vacuo crushing and laser spot dating have been used to separate components (see below). The values of constants and estimation of errors As the internal precision of Ar-Ar ages has improved over the years, the following have been the focus of debate: 1. The commonly accepted values for the K decay constants (Steiger and Jäger 1977) 2. The inter-laboratory and inter-standard calibration of Ar-Ar ages.

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Figure 3. (a) An argon isotope correlation diagram, showing a correlation between atmospheric and radiogenic argon components which form an isochron. Any pure 40Ar component would lie at the origin and thus any excess argon component tends to pull the point B towards the origin. (b) Two samples of amphibole analyzed by laser spot technique. The upper line intercepts within errors of atmospheric argon, the lower line yields a very similar age yet intercepts at a much lower 36Ar/40Ar ratio and contains excess argon.

K-Ar can be regarded as an absolute dating technique, dependent only on the value of the decay constant, and calibration of 38Ar spike. However, all Ar-Ar ages are derived relative to the age of mineral standards, which are irradiated at the same time as the sample. The external precision of Ar-Ar ages is thus limited by the external precision of the age of the mineral standard as determined by the K-Ar method. The most widely used international standards are the hornblendes Hb3gr and MMHb1; biotites GA1550, GHC305 and B4B, muscovite B4M, and sanidines from the Fish Canyon Tuff, Taylor Creek and Alder Creek (widely accepted ages for these standards are found in McDougall and Harrison 1999). Many other pure mineral samples are used as internal standards and several have been proposed as international standards but are not mentioned here since they are not in wide use. The advantage of using mineral standards is that they are freely available but since they are natural, series errors can be introduced if the various standards are not inter-calibrated. This has been an area of particular controversy in recent years, somewhat masking the improvements in internal precision. Fish Canyon Tuff is a prime example of the problems that are faced by those attempting to achieve accurate, and high precision Ar-Ar ages. Fish Canyon sanidine was proposed as an international standard by Cebula et al. (1986), who reported an age of 27.79 Ma but this was relative to another standard, MMHb1 with an age of 518.9 Ma (Alexander et al. 1978). When the age of the MMHb1 was revised to 520.4 Ma (Samson and Alexander 1987), the age of Fish Canyon sanidine became 27.84 Ma, though some workers used a value of 27.55 Ma, based on a different value for the age of MMHb1. In 1994, Renne et al. determined the age of Fish Canyon sanidine to be 28.03 Ma, an age later confirmed by

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cross calibration (Renne et al. 1998b) with biotite standard GA1550. However, subsequent to 1994, many workers continued to use the value 27.84 Ma possibly because this yielded ages in agreement with the spline-fitted magneto-stratigraphic timescale, and in particular, the Cretaceous/Tertiary boundary of 65.0 Ma. Using the value recommended by Renne et al. (1998b) yields an age of 65.4 Ma for tektites from the K/T boundary. In addition, Lanphere and Baadsgaard (2001), maintain that a value of 27.51 Ma is the best age for Fish Canyon sanidine, based on Rb/Sr and U/Pb dates on Fish Canyon Tuff. The problem with this approach is that it requires cross calibration between dating methods, something that is even more fraught with problems. The decay constant of 87Rb is no better constrained than 40K (see below) and meteorite cross calibrations indicate ages may be as much as 2% too low (Renne 2000). In addition, the 27.52±0.09 Ma bulk U/Pb age of Lanphere and Baadsgaard (2001) differs strongly from the U/Pb age of 28.476±0.064 Ma obtained by Schmitz and Bowring (2001) on single grain and small multi-grain fractions of zircon, which confirmed an earlier determination of 28.41±0.05 Ma by Oberli et al. (1990). The work to improve the inter-calibration of standards has been accompanied by parallel discussions of the accuracy and precision of accepted decay constants of several important parent isotopes including 40K (e.g., Begemann et al. 2001). It is notable that the decay constant quoted in most of the physical sciences literature is not the same as the one generally accepted for K-Ar and Ar-Ar dating. In 1977, based mostly on work by Beckinsale and Gale (1969) and Garner et al. (1975), Steiger and Jäger (1977) recommended the use of the decay constants in Table 2. However, Endt and Van der Leun (1973) later compiled the same data as Beckinsale and Gale (1969) to produce different results mainly by different selection criteria and statistical techniques. In a recent summary of 40K decay constants, Audi et al. (1997) report a total decay constant of 5.428 × 10-10 a-1, as previously cited in the nuclear physics literature, a branching ratio of 89.28% and 40K/K = 1.17 × 10-4. Min et al. (2000) showed that a better correlation could be achieved between Ar-Ar and U-Pb ages using the decay constants of Endt and Van der Leun (1973) together with modern physical constants resulting in a total decay constant of 5.463±0.054 × 10-10 a-1 which corresponds to a half-life of 1269±13 Ma. Renne (2000) further demonstrated that in one of the oldest rapidly cooled meteorites, called Acapulco, nuclear physics decay constants (Audi et al. 1997) produced Ar-Ar ages within errors of U-Pb ages, in line with the petrologic interpretation of these samples as rapidly cooled, whereas the existing Ar-Ar ages indicated slow cooling. Note that the interpretation of this important result is controversial (Trieloff et al. 2001; Renne 2001) and final resolution may await a more complete characterization of the thermal history of the Acapulco meteorite. (UTh)/He studies of Acapulco phosphates (Min et al. 2000) appear to validate the rapid cooling which supports the Audi et al. (1997) constants, particularly for the β decay (Min et al. 2001). Reading this section may leave those new to K-Ar and Ar-Ar dating bemused by the current controversy over decay constants and the inter-calibration of standards in such a mature isotope dating technique. It must be emphasised that this controversy has arisen recently as the precision on age determinations has improved, and attempts are made to correlate with U-Pb ages, where decay constants are better constrained. Ages in the literature other than those cited above are still calculated using the decay constants recommended by Steiger and Jäger (1977), and Renne et al. (1998b) represents the most precise inter-calibration of standards combined with a precise K-Ar date on the GA1550 biotite standard using isotope dilution for the K analysis. Analytical errors on the Ar-Ar age have generally been calculated using the simple error propagation of Dalrymple et al. (1981):

K-Ar and Ar-Ar Dating σ 2t ≈

J 2σ 2R + F 2 σ 2J λ2 (1 + FJ) 2

795 (9)

where σt is the error on the age, J is the value from Equation (6), F is the 40Ar*/39Ar ratio, λ is the combined decay constant, σR is the error on the 40Ar*/39Ar ratio and σJ is the error on the J value. A more complete numerical error analysis is given by Scaillet (2000). Errors on the age generally include the error in determining the J value, but not the external errors on determining the K-Ar age of the standard (e.g., GA1550) used to determine the neutron flux or the decay constants. However, when comparing Ar-Ar ages with ages determined by other isotope techniques such as U-Pb, these errors must be considered, something that is particularly important in dating short lived events such as terrestrial meteorite impacts or flood basalts. Renne et al. (1998a) present discussion of this problem and equations for calculating full external errors and appropriate errors when ages for samples measured against two different standards where those standards have been inter-calibrated. ARGON DIFFUSION AND SOLUBILITY K-Ar and Ar-Ar dating are isotope dating techniques, not isotope tracer techniques, but the trace element chemistry of argon plays an integral role in the assumptions and interpretation of K-Ar and Ar-Ar data. The use of the Ar-Ar dating technique to investigate cooling histories and understand extraneous argon are premised on an understanding of argon diffusion rates and argon solubility in hydrous fluids, melts and minerals. Below we will see how diffusion data provide the link between K-Ar or Ar-Ar age and the thermal history of the sample being investigated. We will also see how the assumption that all radiogenic daughter product (40Ar*) in a mineral results from the decay of 40K is commonly contravened. In order to understand how this can happen, we need to consider partition and solubility of argon in fluid, melts and minerals. Argon diffusion (and its use to determine thermal histories) A full exposition of diffusion theory and diffusion mechanisms is beyond the scope of this chapter but the reader is referred to McDougall and Harrison (1999). Many observations show that argon diffusion rates in natural and laboratory experiments follow an Arrhenius relationship. There are however, important departures where fast track diffusion dominates in nature (Kramar et al. 2001; Reddy et al. 2001b) and laboratory analysis of hydrous minerals (Gaber et al. 1988; Lee 1993; Lo et al. 2000). Even in natural cases which might earlier have been identified as volume diffusion effects, careful compositional control shows that phase mixing can mimic argon loss profiles (e.g., Onstott and Peacock 1987; Wartho 1995). In such cases the data can not easily be inverted to produce thermal histories. Having considered cases in which lattice or ‘volume’ diffusion does not explain the data, it should be noted that lattice diffusion seems to dominate many natural systems, in argon and other noble gases (cf. Farley 2002, this volume). Lattice diffusion follows a second order diffusion mechanism (McDougall and Harrison 1999) with an Arrhenius relationship given by the equation: D = D0 e

⎛ −E ⎞ ⎝ RT ⎠

(10) 2 -1

where D is the diffusion coefficient (m s ), D0 is the pre-exponential factor (the theoretical diffusion constant at infinite temperature or a measure of ‘conduction’ in the mineral), E is the activation energy (J), R is the gas constant (J K-1 mol-1) and T is temperature (K). This logarithmic relationship between diffusion rate and temperature means that at high temperatures, argon can diffuse through a mineral lattice and escape

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into the grain boundary network much faster than it is produced by decay of 40K. At low temperatures, diffusion rates are so slow that the daughter atoms are quantitatively retained in the mineral lattice (Fig. 4). There is also a temperature range of partial retention similar to the helium partial retention zone (HePRZ) of (U-Th)/He dating (Farley 2002, this volume), though this is generally quite narrow and represents a small length of time relative to the age of the sample. However the partial retention zone may become significant in slowly cooled metamorphic rocks (Dodson 1986; Lister and Baldwin 1996; Wheeler 1996). In early work, K-Ar ages of minerals such as biotite and muscovite were regarded as recording instantaneous closing temperatures and thus times in the cooling thermal history of the rock and closure temperatures were largely qualitative (e.g., Purdy and Jäger 1976), but Dodson (1973), provided an important advance by quantifying the closure temperature effect. By making a simplifying assumption, Dodson, showed that the closure temperature of a mineral could be described by the equation: E ⎛ A R Tc 2 D0 / r 2 ⎞ = ln ⎜ ⎟ RTc ⎝ E dT / dt ⎠

(11)

Figure 4. (a) A Temperature-time (Tt) curve for a cooling metamorphic terrain, in which the cooling rate is linear in 1/T. (b) Although the cooling rate changes slowly, argon retention within a mineral, following the Tt path changes rapidly between rapid exchange to the grain boundary and no exchange. This sudden flip from open to closed gives rise to the closure temperature concept.

where E and D0 are the activation energy and pre-exponential factor for argon diffusion, Tc is the closure temperature, R is the gas constant, r is the effective grain radius for argon diffusion (see discussion of diffusion domains below), dT/dt is the cooling rate at closure and A is a geometrical factor describing the variation of diffusion in the mineral (A = 55 for spherical geometry or equal diffusion in all three dimensions, A = 27 for cylindrical or diffusion dominantly in only two dimensions, and A = 8.7 for planar or dominantly one dimensional diffusion). The main simplifying assumption made by Dodson (1973) in order to produce this simple equation is that the cooling path of the rocks is linear in 1/T. Although metamorphic cooling is commonly not linear in 1/T, the

K-Ar and Ar-Ar Dating

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results are quite robust unless rocks are cooling very slowly. At cooling rates slower than around 5°C per million years, inner zones of grains close significantly earlier than the outer zones, leading to a closure temperature profile within the grain. Dodson (1986) took account of this slow cooling effect by the addition of an expression for the closure temperature as a function of position within the crystal. The function is tabulated in McDougall and Harrison (1999) and still suitable for rapid hand calculation. In more complex cases the diffusion equations can be solved numerically by finite difference, and programs such as DIFFARG (Wheeler 1996) allow the user to input any thermal history, limited only by the power of the computer. DIFFARG is suitable for calculations on single grains, but Lovera et al. (1997) have provided a technique for analysing multiple sub-grains or ‘domains’ by modeling 39Ar release during laboratory experiments (see below).

Figure 5. Laser profiles showing the effects of cracks and defects in minerals. (a) Biotite in an argon loss experiment, showing an indentation in the age contours resulting from a crack (Onstott et al. 1991). (b) Muscovite exhibiting young ages along an obvious fracture, shown left to right just below the center of the grain (Hames and Cheney 1997). (c) Biotite from a slowly cooled terrain again showing an indentation resulting from a crack (Hodges et al. 1994). (d) K-feldspar from a deformed sandstone, note that old ages are preserved in the center of this boudinaged grain but close to the boudinage and in the smaller fragment, only younger ages were measured (Reddy et al. 2001).

The discussion above considered only simple whole grain or sub-grain diffusion when lattice diffusion was the transporting mechanism. In recent years, the advent of laser analysis has demonstrated just how variable the actual within grain patterns can be (Fig. 5) (e.g., Hodges et al. 1994; Kelley and Turner 1991; Kramar et al. 2001; Lee et al.

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1990; Phillips and Onstott 1988; Scaillet et al. 1990; Villa et al. 1996). Even before laser analysis had illustrated the patterns produced, bulk mineral analysis had shown that the relationship between grain size and argon loss broke down at larger grain sizes. Both natural (Layer et al. 1989; Wright et al. 1991) and laboratory (Grove and Harrison 1996) bulk heating experiments showed that the closure temperature did not increase as biotite grain size increased beyond around 150 μm, other bulk stepwise heating work determined effective radii up to 340 μm (Copeland et al. 1987), and yet age gradients have been measured in even larger grains (Hames and Cheney 1997). It seems likely that the effective diffusion radii are actually the result of variable cracking and defect density of the mica (Kramar et al. 2001; Mulch et al. 2001), particularly biotite which is the physically weaker of the two most commonly analyzed micas (Dahl 1996). Muscovite grains seem to retain metamorphic cooling age gradients even after erosion, transport and deposition in later sediments (e.g., Sherlock 2001). Lee (1995) has proposed a model for fast-track diffusion to explain the effects of combined lattice diffusion and diffusion along fast diffusion pathways through the lattice such as defects. Lee proposed that the combined diffusion could be modeled as two parallel diffusion mechanisms with argon atoms partitioning between the two. The mathematical model produces realistic release patterns, but does not currently take account of the distances between fast track pathways and the time taken for atoms to reach one (cf. Arnaud and Kelley 1995). Future development of the fast track model may provide very fruitful avenues for research. Argon solubility (and the causes of extraneous argon) One of the fundamental assumptions of K-Ar dating (assumption 3, above) is that after correcting for atmospheric argon, all 40Ar in the sample is the result of the in situ decay of 40K, an assumption which is not always valid. However, the amounts of extraneous argon (see earlier definition) are small and generally remain undetected. Perhaps the best-known examples that contravene this assumption are the glassy rims to MORB basalts which retain significant quantities of argon derived from the upper mantle (Graham 2002, this volume). The solubility of noble gases in hydrous fluids and silicate melts have been studied extensively (Carroll and Draper 1994) but the data on mineral solubility and thus mineral fluid partition coefficients are less well known. Only in recent times have techniques been developed which can address the very low concentrations of noble gases in mineral lattices (e.g., Brooker et al. 1998; Chamorro et al. 2002; Wartho et al. 1999). The data which do exist imply very low partition coefficients between minerals and hydrous fluids and between minerals and melts. This explains how assumption 3 (above) can sometimes be invalidated and yet the dating technique still yields accurate, though often less precise, ages (Kelley 2002). Early reports of excess argon covered the whole range of minerals and rocks (e.g., Dalrymple and Moore 1968; Damon and Kulp 1958; Livingston et al. 1967; Lovering and Richards 1964; Pearson et al. 1966) and most importantly, fluid inclusions (Rama et al. 1965). The step heating technique has been quite successful in discriminating against low concentrations of homogeneously distributed excess argon (e.g., Renne et al. 1997) which can be plotted on an isochron diagram (e.g., Heizler and Harrison 1988; Roddick 1978). In fact many published Ar-Ar ages contain small amounts of excess argon reflected in an initial 40Ar/36Ar ratio that is within a few percent of the atmospheric ratio. Such determinations yield precise results when the ratio of the contaminating component is close to that of atmospheric argon. However, as the ratio of the contaminant increases and the small 36Ar peak becomes more difficult to measure, and the possibility of obtaining a precise age is quickly compromised. In extreme cases, excess argon may be undetected (e.g., Arnaud and Kelley 1995; Foland 1983; Pankhurst et al. 1973; Sherlock and Arnaud

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799

1999). Further, the initial ratio correction only works when the isotope ratio within the samples is homogeneous, in cases of heterogeneous excess argon, a spread of data makes precise age determination impossible (e.g., Cumbest et al. 1994; Pickles et al. 1997). The development of new analytical techniques for Ar-Ar dating has led to several advances in our understanding of excess argon. Ar-Ar stepped heating provides a physical technique to separate and analyse phases within individual samples as a result of their different breakdown temperatures (e.g., Belluso et al. 2000). Stepwise heating also produces decrepitation of fluid inclusions at low temperatures resulting in the high initial ages commonly observed in release spectra. Stepwise heating has also demonstrated excess argon diffusion into grain boundaries (Harrison and McDougall 1981). One feature of excess argon commonly associated with low potassium rocks and minerals such as plagioclase, amphibole and clinopyroxene is the saddle or 'U'-shaped Ar-Ar stepped heating release spectrum (Harrison and McDougall 1981; Lanphere and Dalrymple 1976; Wartho et al. 1996) (Fig. 6). Several explanations have been offered for this release pattern, first described by Dalrymple et al. (1975) and Lanphere and Dalrymple (1976) in samples from dykes intruding Pre-Cambrian rocks in Liberia. The initial Ar release yields old apparent ages, which decrease with progressive 39Ar release, approaching the true age of the sample, and finally returning to older ages towards the end of argon release. Although arguments have been made for excess argon incorporation via special diffusion mechanisms such as anion diffusion (Harrison and McDougall 1981), recent data seem to demonstrate that the most likely candidates are fluid inclusions, released at low temperature, and melt or solid inclusions, released at high temperature during mineral breakdown (Boven et al. 2001; Esser et al. 1997) (Fig. 7).

Figure 6. Excess argon in amphibole producing ‘U’ or ‘saddle’ shaped release spectrum. Dark spectra are original grains, light spectra are the same samples after acid leaching which reduced the excess argon component at both low and high ends of the release (Wartho et al. 1996).

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Figure 7. Saddle shaped spectra in plagioclase caused by excess argon dissolved in inclusions. (a) A zero age volcanic plagioclase with high melt inclusion content. (Esser et al. 1997). (b) A plutonic plagioclase showing a complex saddle shaped release pattern, probably the result of argon release from mixed phases (Boven et al. 2001).

Similar patterns have been seen in K-feldspars and these also probably relate to the interplay of fluid inclusion decrepitation at low temperature and excess argon in large sub-grains released at high temperature (Foster et al. 1990; Harrison et al. 1994; Zeitler and FitzGerald 1986). In vacuo crushing has also been used to study the close correspondence between excess argon and saline crustal fluids trapped in fluid inclusions in quartz (Kelley et al. 1986; Turner and Bannon 1992; Turner et al. 1993) and K-feldspar (Burgess et al. 1992; Harrison et al. 1993, 1994; Turner and Bannon 1992). Finally, in situ laser spot extraction techniques have provided a method of investigating excess argon distributions within minerals, demonstrating close correlation between excess argon, and composition signatures inherited from the protolith in ultra-high-pressure UHP terrains (e.g., Giorgis et al. 2000; Sherlock and Kelley 2002), and demonstrating diffusion of excess argon through the mineral lattice (Fig. 8) (e.g., Lee et al. 1990; Pickles et al. 1997; Reddy et al. 1996). Excess argon is most commonly found in metamorphic rocks, and is less common in volcanic systems where outgassing to the atmosphere provides a release mechanism. Excess argon is also particularly common in hydrothermal systems associated with large

K-Ar and Ar-Ar Dating

801

granite intrusions (Kelley et al. 1986; Kendrick et al. 2001a,b; Turner et al. 1993), shear zones in ancient metamorphic terrains (Allen and Stubbs 1982; Smith et al. 1994; Vance et al. 1998), contact metamorphic aureoles (Harrison and McDougall 1980) and high and ultra-high-pressure metamorphic rocks (Arnaud and Kelley 1995; Inger et al. 1996; Li et al. 1994; Ruffet et al. 1997; Scaillet 1996; Scaillet et al. 1990, 1992; Sherlock et al. 1999; Sherlock and Kelley 2002).

Figure 8. Excess argon diffusing into a biotite grain from the Seconda Zona Diorito Kinzigitica (IIDK) in the western Alps, the true cooling age of the biotites samples is around 40 Ma (from Pickles et al. 1997).

Figure 9. Comparison diagram of solubilities for hydrous fluids, melts and minerals based on data in Kelley (2002).

The importance of argon geochemistry is demonstrated clearly in Figure 9, which shows that argon is a highly incompatible trace element, strongly favoring partition from minerals into grain boundary fluids in metamorphic rocks, or from crystals into melts or melts into bubbles in magmatic systems, leaving minerals as the most highly depleted part of any system. Small amounts of excess argon may often be present but the concentrations are so low as to be masked by the larger radiogenic component. The corollary to this hypothesis is that more cases of excess argon are detected in low potassium minerals and the excess Ar is very often present in quartz (e.g., Kelley et al. 1986; Kendrick et al. 2001b; Rama et al. 1965; Turner and Bannon 1992; Vance et al. 1998). The highly incompatible nature of the argon in solid/melt and solid/fluid systems makes the fluids and melts effectively “infinite reservoirs” for radiogenic argon in these systems. More importantly however, the fact that some experimental data are available means that it is possible to model simple systems and define the limits of K-Ar and Ar-Ar dating. Although mineral data are sparse, reliable data now exist for olivine, phlogopite, clinopyroxene and K-feldspar (Brooker et al. 1998; Roselieb et al. 1999; Wartho et al.

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1999). Emerging work on quartz, plagioclase, leucite and other minerals may add to this database (Roselieb et al. 1999; Wartho et al. unpublished data). In the following sections we will explore how the data for K-feldspar can be used to explore how natural systems might work. Excess argon has been detected in both fluid-rich (open system) environments such as shear zones or hydrothermal systems and in fluid-poor (closed system) environments such as granulites or high-pressure metamorphic rocks. In fluid-rich environments, a fluid with a high concentration of excess argon is in contact with minerals above their closure temperature, and significant quantities of excess argon partition into the minerals, or just as likely, become incorporated as fluid inclusions (e.g., Cumbest et al. 1994; Kendrick et al. 2001a,b). In fluid-poor environments such as dry granulites or high- and ultra-highpressure metamorphic rocks, fluids may only be present as transient phases in restricted zones. In such rocks, fluids do not travel significant distances (Philippot and Rumble 2000) and thus transport of argon along the grain boundaries might be as little as a few centimetres over millions of years as demonstrated for oxygen in high-pressure and ultrahigh-pressure rocks (Philippot and Rumble 2000; Foland 1979). Radiogenic argon produced in potassium bearing minerals above their closure temperatures would accumulate in the grain boundary network and quickly reach levels where significant quantities partition into the minerals (e.g., Baxter et al. 2002; Foland 1979; Kelley and Wartho 2000). The two different occurrences of excess argon, in open and closed systems, are mirrored in the argon contents of quartz (fluid inclusions). Where quartz has been analyzed, excess argon concentrations signal excess argon in other minerals in fluidrich environments (e.g., Vance et al. 1998; Kendrick et al. 2001a, b) though in fluid-poor environments, they exhibit lower excess argon contents than co-existing micas (Arnaud and Kelley 1995; Sherlock et al. 1999; Sherlock and Kelley 2002). Although these two modes of occurrence are very different, both can be successfully modeled by considering argon as a trace element partitioning between fluid and solid using the measured argon solubility. K-feldspar is a good example for a simple model since some solubility data are available (Wartho et al. 1999; Wartho and Kelley unpubl. data) and it is one of the most common potassium bearing minerals used in K-Ar and ArAr dating (cf. McDougall and Harrison 1999). Excess argon in open systems. Argon solubility in saline waters has been measured precisely only at low temperatures (Smith and Kennedy 1983; Kipfer et al. 2002, this volume), but argon solubility at high temperatures can be estimated by extrapolating the salinity data in proportion with the high-temperature solubility data. Such extrapolation indicates that argon solubility in pure and saline grain boundary fluids up to 300°C (peak closure temperature for K-feldspar grains) lies in the range 25 to 100 ppm bar-1 atm-1. The large range reflects uncertainty in these values but they serve to constrain the model. The solubility of argon in K-feldspar is 0.66 ppb bar-1 atm-1 (data derived from Wartho et al. 1999) thus the DAr for the K-feldspar grain boundary fluid system lies in the range 6.6 × 10-6 to 2.6 × 10-5. Figure 10 illustrates concentrations of excess argon in K-feldspar (expressed as the increase in age they would cause in an orthoclase K-feldspar) in equilibrium with variable concentrations of argon in the corresponding grain boundary fluid. The shaded area indicates concentrations of (atmospheric) argon found in near surface ground waters and some deeper waters (Smith and Kennedy 1983; Torgersen et al. 1989; Kipfer et al. 2002, this volume). More extreme concentrations of excess argon are found in hydrothermal fluids and in fluid inclusions which range up to 21 ppm (Burgess et al. 1992; Harrison et al. 1993; Harrison et al. 1994; Kelley et al. 1986; Turner and Wang 1992). If such concentrations of excess argon were introduced into grain boundary fluids in the model, they would increase the apparent ages of K-feldspar

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samples by less than 0.6 Ma. Even if 100 ppm of excess argon was introduced into the model fluid, K-feldspar ages would rise by only 4 Ma. Therefore in the huge majority of cases, the level of excess argon in K-feldspar will be 1 to 2 orders of magnitude below detection limits (around 0.1 Ma in a 100-Ma K-feldspar). Even when excess argon is present in fluids at the highest levels measured in the upper crust, excess argon in Kfeldspar would be a minor component in most analyzed samples. This observation corroborates the many measurements showing that excess argon is uncommon in Kfeldspar, much of the excess argon which is detected in K-feldspar is confined to fluid inclusions (Burgess et al. 1992; Harrison et al. 1994).

Figure 10. A plot of age excess vs 40Ar concentration in the grain boundary fluid for an open system. Solid lines are those derived from the mineral/fluid partition coefficients (6.6 × 10-6 to 2.6 × 10-5) for K-feldspar, dashed lines are those for another mineral with higher partition coefficients such as biotite. The shaded area indicates concentrations of (atmospheric) argon found in near surface ground waters and some deeper waters (Smith and Kennedy 1983; Torgersen et al. 1989; Kipfer et al. 2002, this volume). The vertical line indicates the current limit of excess argon concentration found in crustal fluids (Harrison et al. 1994).

The model confirms how robust the Ar-Ar system is when applied to K-feldspar, but minerals with higher mineral/fluid partition coefficients such as biotite will behave differently in a similar system. Note that biotite has been shown to yield ages over 100 Ma older than the expected age (e.g., Brewer 1969; Smith et al. 1994) and several studies describe excess ages in biotite but not in co-existing muscovite (e.g., Roddick et al. 1980). Roddick et al. (1980) argued that this reflected greater solubility of argon in biotite, an observation corroborated by crystal chemical (ionic porosity) arguments that argon solubility in biotite should be higher than muscovite (Dahl 1996). The dashed lines on Figure 10 illustrate the behavior of a mineral with a partition coefficient of 1 × 10-3 and similar radiogenic argon concentrations in grain boundary fluids to those seen in fluid inclusions. Such a mineral would commonly yield ages of the order of 1 Ma older than the true closure age but under extreme excess argon conditions, might yield ages more than 100 Ma older than the true closure age (compared with 4 Ma for K-feldspar in the same fluid). Where fluids are derived from basement rocks in fluid-rich regimes such as orogenic thrust belts, excess argon is unsurprisingly common (Brewer 1969; Kelley 1988; Smith et al. 1994; Reddy et al. 1996; Vance et al. 1998). The widespread influx of fluids in such regimes is illustrated by the occurrence of excess argon over broad areas (e.g., Brewer et al. 1969; Smith et al. 1994).

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Excess argon in closed systems. Modeling excess argon in a closed system is less well constrained since there are no measured argon solubilities in hydrous fluids at high temperatures and even the presence of a fluid on the grain boundaries of such rocks may amount to an absorbed OH-layer or a CO2-rich fluid. Further, any fluids which were present may have been isolated in pores or as grain edge tubules (Holness 1997) and are likely to have existed only transiently. However, the model illustrates how excess argon develops in a system without an infinite reservoir of hydrous fluid or melt. The same Kfeldspar/fluid system used in the previous section can also be used to investigate the behavior of argon in fluid-poor systems. In a closed system, the controlling factors on the distribution of excess argon between K-feldspar and fluid are temperature, fluid salinity, the volume fraction of fluid (as a ratio of the total volume of the rock), and potassium content or K-feldspar content of the whole rock. In fluid-rich systems such as porous sandstones in a basin environment, porosity might reach several percent, but in metamorphic rocks, porosity generally decreases with increasing grade to as low as 0.01% (10-4 vol fraction) in dry systems (Holness 1997). In order to model a closed system, zero radiogenic and excess argon concentrations are initially assumed in both fluid and minerals. In natural systems there may be pre-existing radiogenic argon in minerals or fluid inclusions, particularly when the rocks have a polymetamorphic history and the boundary between excess argon and inherited argon becomes blurred in these cases (e.g., Li et al. 1994). This model system uses the same range of K-feldspar/fluid partition coefficients (6.6 × 10-6 to 2.6 × 10-5) to account for salinity variations and was run for rocks with 1 to 100% K-feldspar. Argon is assumed to be even more incompatible in any other mineral phases within the rock. As the model runs, excess argon builds up in K-feldspar as radiogenic argon is produced by 40K decay. Unlike the open system, the fluid “reservoir” is finite and after a while radiogenic argon increases to the level where partition back into the K-feldspar becomes detectible. The fractional age excess has been calculated for fluid filled porosity of 1%, 0.1% and 0.01%. Figure 11 shows that samples with greater than 0.1% porosity do not generate significantly old ages in K-feldspar. If the porosity is 0.1%, excess argon in the same Kfeldspar would yield ages as much as 1.5 Ma too old in a 100 Ma rock with 100% Kfeldspar rock, but only 0.03 to 0.8 Ma in the same age rock with 5 to 30% K-feldspar (common to crustal rocks such as granite). In all probability this would still be below detection levels. Only in the most fluid-poor systems with 0.01% porosity, does the system start to exhibit detectable excess argon with up to 8 Ma excess argon in a rock with 30% 100 Ma old K-feldspar, and even in this case only the most K-feldspar-rich rocks containing very saline fluids will produce detectable excess argon. Furthermore, in this model it was assumed that the other minerals in the rock exhibited lower argon solubility but if another mineral such as biotite (with higher partition coefficient) is present, excess argon concentrations in K-feldspar quickly drop below detection levels even in the most fluid-poor terrains (e.g., Foland 1979). K-feldspar ages measured in eclogite terrains which exhibit closed system excess argon in phengite (Arnaud and Kelley 1995), sometimes reveal high-temperature excess argon, but this may also result from plagioclase or inclusions outgassing during the cycle-heating experiment (Arnaud and Kelley 1995; Boven et al. 2001). It is not clear how rigorously this model can be applied to the most fluid-poor rocks since the grain boundary fluid phase in dry systems such as eclogites and granulites may be as little as a layer of OH- molecules at the grain boundaries. However, the model serves to illustrate how a closed system can explain phenomena observed in highpressure and ultra-high-pressure terrains, particularly the occurrence of excess argon in phengite when it is the predominant potassium-bearing mineral in the rock. In several UHP terrains, it has been noted that excess argon is more prevalent in rocks with

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protracted histories or old protoliths (Li et al. 1994; Arnaud and Kelley 1995; Inger et al. 1996; Giorgis et al. 2000). Although much of this ‘inherited argon’ will partition into the limited fluid phase, excess argon is more likely in rocks with older protoliths. Another prediction of a closed system model is that potassium-rich rocks will contain greater concentrations of excess argon than potassium-poor rocks, even when only one potassium bearing mineral phase is present (e.g., Sherlock and Kelley 2002). Ironically, in fluidpoor high-pressure rocks it would be more appropriate to measure ages from low-K metabasalts or meta-sandstones rather than mica-rich schists or K-feldspar-rich rocks. Figure 11. A plot of age excess vs. 40 Ar concentration in the grain boundary fluid for a closed system. Solid lines are those derived from the mineral/fluid partition coefficients (6.6 × 10-6 to 2.6 × 10-5) for Kfeldspar in a rock with 1% fluid. Dashed lines are those for the same rock with 0.1% fluid and the dashed/dotted lines are those for a rock with 0.01% fluid. The shaded area represents the range of rock compositions from which K-feldspar is commonly analyzed (5 to 50%).

In summary, the presence of excess argon in most crustal fluids means that assumption No. 3 of the K-Ar and Ar-Ar may never be strictly valid, but in the majority of cases, the concentration of excess argon in minerals is swamped by the in situ radiogenic component and is not a factor in determining an accurate age. Excess argon in fluid and melt inclusions. There is considerable overlap between the studies of excess argon in minerals and crustal noble gas studies (Ballentine and Burnard 2002; Ballentine and Marty 2002, both in this volume) in studies of fluid and melt inclusions. The recent publications of Kendrick et al. (2001a, b) demonstrate just how close the two sets of studies have become. Fluid inclusions and melt inclusions are commonly important sources of excess argon in minerals analyzed for K-Ar and Ar-Ar dating, particularly in low potassium minerals such as amphibole and plagioclase. The simple reason for the importance of fluid and melt inclusions is illustrated by Figure 11. Melt inclusions incorporated from a magma containing excess argon will contain ~100 times more argon than minerals crystallising from the same melt, and fluid inclusions in equilibrium may contain as much as 10,000 times the excess argon concentration (by weight) of the mineral structure. Inclusions provide some of the most intractable analytical problems in K-Ar and Ar-Ar dating although they can also occasionally be used to advantage in preserving mineral ages (e.g., Kelley et al. 1997; van der Bogaard and Schirnick 1995). Cumbest et al. (1994) showed how the interaction of metamorphic fluids could result in different generations of excess argon in fluid inclusions and Harrison et al. (1994) showed that fluid inclusions were present in many K-feldspar samples but repeated cycles of step heating could reduce and correct for their effects. Melt inclusions are common in mineral separates used to date volcanic and hyperbyssal rocks, but have received little attention. However, the effects of melt inclusions from a historic flow on Mount Erebus were studied by Esser et al. (1997) showing that their presence resulted in saddle-shaped age spectra during stepped 39Ar release (Fig. 7a).

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The applications cited here focus upon the Ar-Ar dating technique, simply because most recent developments have utilized the Ar-Ar dating technique. The aim of this section is to highlight applications which illustrate the breadth of problems which have been addressed using Ar-Ar dating. Thermochronology The diffusion of radiogenic noble gas daughter products provides several routes for determining not only the chronology of events but also thermal histories (e.g., Swindle 2002; Farley 2002, both in this volume). One of the most common applications of Ar-Ar geochronology has been dating mineral grains or mineral separates in order to study cooling and uplift in metamorphic terrains. Often in the past, these ages were combined with U-Pb and Rb-Sr ages in order to produce cooling histories using mineral closure temperatures based upon laboratory diffusion determinations (e.g., Foland 1974; Grove and Harrison 1996) or field estimates (e.g., Kirschner et al. 1996; Purdy and Jäger 1976). However, after the success of unravelling the thermal histories of lunar samples (see Turner 1977 for a full review), and the many studies on thermal histories of meteorites (Turner 1969, 1970a), step heating was applied to terrestrial samples. As Turner (op cit) showed, the diffusion parameters of a mineral can be recovered from the release of 39Ar during the step heating procedure, and combined with age information to deduce a thermal history. The results of applying this technique to terrestrial samples have been mixed. While step heating can discriminate against contamination by mixed phases, thermal histories are rarely extracted from hydrous minerals since the minerals become unstable long before they melt, releasing argon not only from outer layers but along cleavages parted by explosive loss of water in biotite (Gaber et al. 1988) or via breakdown reactions in the case of amphibole (Lee 1993; Lee et al. 1991; Wartho 1995; Wartho et al. 1991) and biotite (Lo et al. 2000). However this aspect has recently been exploited (Belluso et al. 2000), using the preferential breakdown of different amphibole compositions in an attempt to date the formation of separate generations. Feldspars do not break down during short term in vacuo heating experiments prior to melting since they are anhydrous, and thus do not suffer the same effects as hydrous minerals. However, the complex sub-solidus reactions and exsolution in plagioclase feldspar means that Ar-Ar step heating generally reflects mixed phases, and interpretation may be difficult (e.g., Boven et al. 2001). Stepwise heating of K-feldspar on the other hand, provides an opportunity to explore thermal histories without the need to resort to other isotope systems or even other minerals. The microstructure of K-feldspar is relatively simple in volcanic rocks, but in plutonic and basement rocks, progressive episodes of reaction during cooling result in complex microtexture of intergrowths and exsolution (Parsons et al. 1999). Plutonic and basement feldspars are thus composed of many variably sized and shaped diffusion domains within single grains of K-feldspar, some as small as a micron. The bulk closure temperature of large K-feldspar grains would be around 300°C if they acted as whole grains (cf. Foland 1974) but in fact the ‘effective’ closure temperature is closer to 150-200°C, and this is the reason why they were regarded as “leaky” by early workers. Several workers suggested that argon loss from K-feldspars might be the result of variable sub-grain sizes, but it was not until Lovera et al. (1989, 1991) that this effect was quantified using cycled heating steps designed to preferentially outgas different diffusion domain sizes, rather than the standard monotonic temperature increase normally used in step heating. Lovera et al. (1989, 1991) also formulated a mathematical model using the deviation of the 39Ar release from a simple Arrhenius relationship, to monitor argon diffusion in domains of differing sizes. The individual step Ar-Ar ages were then combined with derived diffusion information to

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produce a continuous thermal history over a range of as much as 200°C. This was a considerable advance over a single meaningless closure temperature. In their initial work, Lovera et al. used an age spectrum to monitor the cooling of the Chain of Ponds Pluton in northwestern Maine (Fig. 12), showing that the age spectrum could not be modeled using a single diffusion domain but that a reasonably good fit could be achieved using three domains of different sizes (i.e., diffusion radii) and specific volume fractions.

Figure 12. Cycle heating results for a K-feldspar sample from the Chain of Ponds pluton (Lovera et al. 1989), showing the fit of modeled argon release for a single diffusion domain lines and the improved fit which can be obtained by using three diffusion domains. The lower figure shows the departure of argon diffusion from a simple Arrhenius relationship.

A great deal of discussion of multi-domain modeling (MDD) has ensued mainly surrounding the diffusion mechanisms (Arnaud and Kelley 1997; Harrison et al. 1991; Lovera et al. 1989, 1991, 1993, 1997; McDougall and Harrison 1999; Villa 1994; Wartho et al. 1999) and the physical reality of the diffusion domain sizes used in the model (Arnaud and Eide 2000; Burgess et al. 1992; FitzGerald and Harrison 1993; Lovera et al. 1989, 1991, 1993; Parsons et al. 1999; Reddy et al. 2001a). The model has also evolved considerably since its inception with the use of an increasing number of discrete diffusion domain sizes (e.g., Mock et al. 1999a,b), evocation of multiple activation energies (Harrison et al. 1991) which was later discontinued, a realization that fluid inclusions played an important role in apparent excess argon seen at low temperatures (Turner and Wang 1992; Burgess et al. 1992; Harrison et al. 1993, 1994), and more recently focus on the low-temperature release to determine an activation energy used throughout (Lovera et al. 1997) since higher temperatures may be affected by pre-melting and other artifacts. Monte Carlo inversion techniques have also been used to define domain sizes and thermal histories (Warnock and Zeitler 1998; Lovera et al. 1997) and there has been discussion of the effects of recoil in very small domains upon the release patterns and their interpretation (Onstott et al. 1995; Villa 1997). If the MDD technique had produced no useful information, the controversy and constantly evolving modeling techniques would have dissuaded workers from applying it to geological problems. However, as McDougall and Harrison (1999) document, the cycle-heating MDD model technique has resulted in reasonable thermal histories in a

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wide range of geological situations. Moreover, McDougall and Harrison (1999) set out a full set of testable assumptions and techniques for accepting or rejecting data. In summary, the cycle-heating technique must be undertaken with great care and awareness of the potential advantages and pitfalls. The cycle-heating MDD modeling technique offers the chance to extract thermal information from bulk K-feldspar separates, but what of hydrous minerals such as micas and amphiboles? In fact, laser spot dating has succeeded in imaging age profiles in micas and amphiboles where step heating failed to produce thermal histories (e.g., Kelley and Turner 1991; Lee et al. 1991; Phillips and Onstott 1988; Scaillet et al. 1990), demonstrating that hydrous minerals do retain thermal information (Fig. 13). Laser age profiles have been used to address a range of geological problems including; heat loss from igneous intrusions (Kelley and Turner 1991; Wartho et al. 2001); slow cooling in metamorphic terrains (Hames and Cheney 1997); ingress of excess argon (Scaillet et al. 1990; Scaillet 1996; Pickles et al. 1997; Ruffet et al. 1991, 1997; Sherlock et al. 1999; Giorgis et al. 2000) and the relationship between deformation and Ar-Ar ages (Reddy et al. 1996, 1997, 2001). UV lasers have been used to measure experimentally produced argon diffusion profiles at a spatial resolution of as little as 1 micron (Arnaud and Kelley 1997; Wartho et al. 1999) and measure argon crystal/melt partition coefficients (Brooker et al. 1998; Chamorro et al. 2002). The advantage of such studies is that, unlike step heating, they provide images of the argon distribution within grains, showing the effect of sub-grain boundaries and compositional effects. The disadvantage is that spatial resolution and the range of suitable samples are limited by the gas required for precise isotope measurement—in other words, older samples and larger grains. Figure 13. Ar-Ar age profiles from micas and amphiboles showing Ar loss or gain by diffusion which has been modeled to recover the thermal history of the sample. (a),(b) show two-amphibole grains from a rock close to the Duluth Gabbro. The age profiles show diffusion domains defined by the presence of biotites in the outer zone of the grains (Note the Ca/K monitored by 37 Ar/39Ar on the right). Both domains yielded the same time-integrated diffusion parameters (Kelley and Turner 1991). (a) A small domain within the grain has suffered 78% argon loss. (b) A larger domain showing only 43% Ar loss. (c) Ar-loss profiles from biotite grains close to a sill on the Isle of Mull, Scotland, outgassed to varying extents dependent upon distance from two sills, one sill was 2.7 m wide, the other was 6 m wide. The variable outgassing was used to produce a magma flow history for the sill (Wartho et al. 2001).

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Dating young volcanic eruptions K-Ar and Ar-Ar have long been used to date young (Pre-historic) volcanic eruptions (cf. Curtis in Schaeffer and Zähringer 1966) but in recent years, the advent of low blank furnaces and laser extraction techniques, particularly the CO2 laser, have pushed the boundaries. It remains unclear whether laser extraction or furnace extraction is the superior technique for very young volcanic samples. Lasers offer slightly lower blanks, though the quantities of sample can be too large for focussed laser extraction (>0.5 g in some cases), but furnaces offer better temperature control. In a study of zero age samples Esser et al. (1997) showed that anorthoclase samples suffered small amount of excess argon which were better discriminated using the furnace extraction technique. There are particular analytical challenges offered by dating young volcanic samples, including precise calibration of the mass spectrometer since many samples contain only small amounts of radiogenic argon. Contamination and control of the sample quality is also particularly difficult since sample sizes are often 0.1-1 g. Finally, melting large quantities of sample also produces large quantities of H2O, CO2 and other gases released from minerals, and requires larger getters to clean the gas prior to mass spectrometric measurement of isotope ratios. The ages of the youngest volcanic rocks dated by the Ar-Ar technique have now reached the historic realm. For example, Renne et al. (1997) have measured the age of the Vesuvius eruption observed and recorded by Pliny in 79 AD. Other work has focussed upon dating human evolution (Deino et al. 1998; Ludwig and Renne 2000; Renne et al. 1997), the relationship between volcanic eruptions and recent climate change (Ton-That et al. 2001), glacial advances (Singer et al. 2000; Wilch et al. 1999), and precise calibration of young magnetic reversals (e.g., Singer and Pringle 1996). High-precision ages on altered basalts One of the reasons why precision of the standards and problems of the decay factors used for Ar-Ar dating came to light was the push to achieve high precision ages for huge flood basalt provinces, which have been linked by some workers to climate change and global mass extinctions. A great deal of effort has gone into analytical techniques and protocols in order to obtain high precision ages for flood basalts (e.g., Foland et al. 1993; Hofmann et al. 2000; Min et al. 2000; Renne 1995; Renne and Basu 1991; Renne et al. 1992, 1995, 1996a,b; Storey et al. 1995; Duncan et al. 1997) this work also contributed strongly to the problems concerning accuracy and precision, decay constants and international standards as a result. Precision (excluding decay and standard errors), even on samples with low potassium are commonly less than 0.5%, sometimes as low as 0.25%. In addition, analytical techniques involving acid clean-up have been developed to analyse extremely difficult samples from the sea floor (Pringle et al. 1991). Dating low-temperature processes Both the K-Ar and Ar-Ar techniques have been used to date low-temperature phases such as clay minerals and in particular the mineral illite. The problem for Ar-Ar dating of such phases is that recoil of 39Ar during neutron irradiation can compromise the measured age. In fine grained phases such as clays, this can cause up to 30% 39Ar loss by recoil and resulting 40Ar*/39Ar ratios, calculated without the recoiled 39Ar are 30% too high. This problem has been overcome to some extent by encapsulating clay samples in evacuated silica vials, breaking the vials after irradiation, and measuring the amount of recoiled 39Ar (e.g., Foland et al. 1983; Smith et al. 1993). More recently, Onstott et al. (1997) showed that the process could be miniaturized, using micro-ampoules and opening them with a UV laser. The technique has been utilized (Dong et al. 1995, 1997a,b) to investigate variations in argon retention properties of clays. Other low-temperature K-bearing

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minerals such as cryptomelane and alunite have been used to date surface processes (Vasconcelos 1999), including the formation of ore gangue (Itaya et al. 1996; Vasconcelos 1999) and cave deposits (Polyak et al. 1998). Ar-Ar dating of authigenic K-feldspar overgrowths has also been investigated using dissolution (Hearn et al. 1987), stepped heating (Mahon et al. 1998) and laser extraction (e.g., Girard and Onstott 1991) with varying success. Recent work using a UV laser extraction technique provides further evidence that such overgrowths can provide meaningful ages (Hagen et al. 2001). Another growing field of work has been dating detrital minerals in sediments as a provenance tool, but also in order to study the sedimentary systems supplying basins and uplift or thermal history of the source. Clauer (1981) and Adams and Kelley (1998) have shown that while biotite is often quickly altered in the sedimentary environment, white micas survive and retain ages even in second cycle sediments (those which were originally deposited and then re-excavated during later basin inversions; e.g., Sherlock 2001). K-feldspar also retains pre-erosion cooling ages (e.g., Copeland et al. 1990) but the lower closure temperatures mean that it has more commonly been used for thermal analysis of sedimentary basins (e.g., Harrison and Be 1983). Unique samples The Ar-Ar dating technique depends only upon two isotopes of argon, and thus it is possible to analyse individual particles such as lunar soil grains (Burgess and Turner 1998), terrestrial sand grains (Kelley and Bluck 1989), and glass spherules (Culler et al. 2000). Ar-Ar dating has also successfully been used to date small amounts of heterogeneous samples, such as pseudotachylytes in fault zones (Kelley et al. 1994; Magloughlin et al. 2001; Muller et al. 2001), and even terrestrial meteorite impacts (Spray et al. 1995). Finally, in its earliest days the Ar-Ar dating was called upon to extract a great deal of temporal and thermal information from unique and precious moon rocks (Turner 1977; Swindle 2002, this volume), and this is still the case as demonstrated by the recent dating of the unique ALH84001 meteorite sample from Mars (Turner et al. 1997). ACKNOWLEDGMENTS The author was greatly assisted in writing this chapter by discussion with Sarah Sherlock. John Taylor kindly drew the figures. Paul Renne, Nicolas Arnaud, Matt Heizler and Mike Cosca provided detailed and insightful reviews which greatly improved the text. Rainer Wieler spotted the many mistakes and typos. REFERENCES Adams CJ, Kelley SP (1998) Provenance of Permian-Triassic and Ordovician metagreywake terranes in New Zealand: Evidence from 40Ar/39Ar dating of detrital micas. Geol Soc Am Bull 110:422-332 Aldrich LT, Nier AO (1948) Argon 40 in potassium minerals. Phys Rev 74:876-877 Alexander ECJ, Mickelson GM, Lanphere MA (1978) MMhb-1: A new 40Ar-39Ar dating standard. U S Geol Surv Open-file Rept 70-701:6-8 Allen AR, Stubbs D (1982) An 40Ar/39Ar study of a polymetamorphic complex in the Arunta Block, central Australia. Contrib Mineral Petrol 79:319-332 Arnaud NO, Kelley SP (1995) Evidence for widespread excess argon during high-pressure metamorphism in the Dora Maira (Western Alps, Italy), using an ultra-violet laser ablation microprobe 40Ar/39Ar technique. Contrib Mineral Petrol 121:1-11 Arnaud NO, Kelley SP (1997) Argon behavior in gem-quality orthoclase from Madagascar: Experiments and some consequences for Ar-40/Ar-39 geochronology. Geochim Cosmochim Acta 61:3227-3255 Arnaud NO, Eide EA (2000) Brecciation-related argon redistribution in alkali feldspars: An in naturo crushing study. Geochim Cosmochim Acta 64:3201-3215

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Renne PR, Deino AL, Walter RC, Turrin BD, Swisher CC, Becker TA, Curtis GH, Sharp WD, Jaouni AR (1994) Intercalibration of astronomical and radioisotopic time. Geology 22:783-786 Renne PR, Zhang ZC, Richards MA, Black MT, Basu AR (1995) Synchrony and causal relations between Permian-Triassic boundary crises and Siberian flood volcanism. Science 269:1413-1416 Renne PR, Glen JM, Milner SC, Duncan AR (1996a) Age of Etendeka flood volcanism and associated intrusions in southwestern Africa. Geology 24:659-662 Renne PR, Deckart K, Ernesto M, Feraud G, Piccirillo EM (1996b) Age of the Ponta Grossa dike swarm (Brazil), and implications to Parana flood volcanism. Earth Planet Sci Lett 144:199-211 Renne PR, Sharp WD, Deino AL, Orsi G, Civetta L (1997) Ar-40/Ar-39 dating into the historical realm: Calibration against Pliny the Younger. Science 277:1279-1280 Renne PR, Karner DB, Ludwig KR (1998a) Radioisotope dating—Absolute ages aren't exactly. Science 282:1840-1841 Renne PR, Swisher CC, Deino AL, Karner DB, Owens TL, DePaolo DJ (1998b) Inter-calibration of standards, absolute ages and uncertainties in 40Ar/39Ar dating. Chem Geol 145:117-152 Renne PR, Farley KA, Becker TA, Sharp WD (2001) Terrestrial cosmogenic argon. Earth Planet Sci Lett 188:435-440 Roddick JC (1978) The application of isochron diagrams in 40Ar/39Ar dating: A discussion. Earth Planet Sci Lett 41:233-244 Roddick JC, Cliff RA, Rex DC (1980) The evolution of excess argon in Alpine biotites—A 40Ar-39Ar analysis. Earth Planet Sci Lett 48:185-208 Roselieb K, Wartho J-A, Buttner H, Jambon A, Kelley SP (1999) Solubility and diffusivity of noble gases in synthetic phlogopite: A UV LAMP investigation. 4:368 Ruffet G, Gruau G, Ballèvre M, Féraud G, Philipot P (1997) Rb-Sr and 40Ar-39Ar laser probe dating of high-pressure phengites from the Sesia Zone (Western Alps): Underscoring of excess argon and new age constraints on high-pressure metamorphism. Chem Geol 141:1-18 Samson SD, Alexander EC (1987) Calibration of the interlaboratory 40Ar-39Ar dating standard—MMhb-1. Chem Geol (Isotop Geosci Sect) 66:27-34 Scaillet S (1996) Excess 40Ar transport scale and mechanism in high-pressure phentites: A case study from an eclogitised metabasite of the Dora-Maira nappe, western Alps. Geochim Cosmochim Acta 60: 1075-1090 Scaillet S (2000) Numerical Error Analysis in Ar-40/Ar-39 Dating. Chem Geol 162:269-298 Scaillet S, Feraud G, Lagabrielle Y, Ballevre M, Ruffet G (1990) Ar-40/Ar-39 laser-probe dating by step heating and spot fusion of phengites from the Dora-Maira nappe of the Western Alps, Italy. Geology 18:741-744 Scaillet S, Féraud G, Ballèvre M, Amouric M (1992) Mg/Fe and [(Mg,Fe)Si-Al2] compositional control on argon behavior in high-pressure white micas: A 40Ar/39Ar continuous laser-probe study from the DoraMaira nappe of the internal western Alps, Italy. Geochim Cosmochim Acta 56:2851-2872 Schaeffer OA, Zähringer J. (1966) Potassium Argon Dating. Springer-Verlag, New York, 234 p Schmitz MD, Bowring SA (2001) U-Pb zircon and titanite systematics of the Fish Canyon Tuff: an assessment of high-precision U-Pb geochronology and its application to young volcanic rocks. Geochim Cosmochim Acta 65:2571-2587 Sherlock SC (2001) Two-stage erosion and deposition in a continental margin setting: a 40Ar/39Ar laserprobe study of offshore detrital white micas in the Norwegian Sea. J Geol Soc London 158: 793-800 Sherlock SC, Arnaud NO (1999) Flat plateau and impossible isochrons: Apparent 40Ar-39Ar geochronology in a high-pressure terrain. Geochim Cosmochim Acta 63:2835-2838 Sherlock SC and Kelley SP (2002) Excess argon evolution in HP-LT rocks: A UVLAMP study of phengite and K-free minerals, NW Turkey. Chem Geol (in press) Sherlock SC, Kelley SP, Inger S, Harris NBW, Okay A (1999) 40Ar-39Ar and Rb-Sr geochronology of highpressure metamorphism and exhumation history of the Tavsanli Zone, NW Turkey. Contrib Mineral Petrol 137:46-58 Singer B, Hildreth W, Vincze Y (2000) Ar-40/Ar-39 evidence for early deglaciation of the central Chilean Andes. Geophys Res Lett 27:1663-1666 Singer BS, Pringle MS (1996) Age and duration of the Matuyama-Brunhes geomagnetic polarity reversal from Ar-40/Ar-39 incremental heating analyses of lavas. Earth Planet Sci Lett 139:47-61 Smith PE, Evensen NM, York D (1993) First successful 40Ar-39Ar dating of glauconites: Argon recoil in single grains of cryptocrystalline material. Geology 21:41-44 Smith PE, York D, Easton RM, Özdemir Ö, Layer PW (1994) A laser 40Ar/39Ar study of minerals across the Grenville Front: Investigations of reproducible excess Ar patterns. Can J Earth Sci 31:808-817 Smith SP, Kennedy BM (1983) The Solubility of noble gases in water and in NaCl brine. Geochim Cosmochim Acta 47:503-515

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Smits F, Gentner W (1950) Argonbestimmummungen an Kalium-Mineralien I. Bestimmungen an tertiären Kalisalzen. Geochim Cosmochim Acta 1:22-27 Spray JG, Kelley SP, Reimold WU (1995) Laser probe argon40/argon39 dating of coesite- and stishovitebearing pseudotachylytes and the age fo the Vredefort impact event. Meteoritics 30:335-343 Steiger RJ, Jäger E (1977) Subcommission on geochronology: Convention on the use of decay constants in geo- and cosmocchronology. Earth Planet Sci Lett 36:359-362 Storey M, Mahoney JJ, Saunders AD, Duncan RA, Kelley SP, Coffin MF (1995) Timing of hot spot-related volcanism and the breakup of Madagascar and India. Science 267:852-855ds Swindle TD (2002) Noble gases in the moon and meteorites—Radiogenic components and early volatile chronologies. Rev Mineral Geochem 47:101-124 Ton-That T, Singer B, Paterne M (2001) Ar-40/Ar-39 dating of latest Pleistocene (41 Ka) marine tephra in the Mediterranean Sea: implications for global climate records. Earth Planet Sci Lett 184:645-658 Torgersen T, B.M. K, Hiagon H, Chiou KY, J.H. R, Clarke WB (1989) Argon accumulation and the crustal degassing flux of 40Ar in the Great Atresian Basin, Australia. Earth Planet Sci Lett 92:43-56 Trieloff M, Jessberger EK, Fieni C (2001) Comment on “Ar-40/Ar-39 age of plagioclase from Acapulco meteorite and the problem of systematic errors in cosmochronology” by Paul R. Renne. Earth Planet Sci Lett 190:263-265 Turner G (1969) Thermal histories of meteorites by the 39Ar-40Ar method. In Millman PM (ed) Meteorite Research. D. Reidel Publishing, Dordrecht, The Netherlands, p 407-417 Turner G (1970a) Thermal histories of meteorites. In Runcorn SK (ed) Paleogeophysics. Academic Press, London, p 491-502 Turner G (1970b) 40Ar-39Ar age determination of lunar rock 12013. Earth Plant Sci Lett 9:177-180 Turner G (1970c) Argon-40/argon-39 dating of lunar rock samples. Science 167:466-468 Turner G (1971a) 40Ar-39Ar dating: The optimization of irradiation parameters. Earth Planet Sci Lett 10:227-234 Turner G (1971b) 40Ar-39Ar ages from the lunar maria. Earth Planet Sci Lett 11:169-191 Turner G (1972) 40Ar-39Ar age and cosmic ray irradiation history of the Apollo 15 anorthosite 15415. Earth Planet Sci Lett 14:169-175 Turner G (1977) Potassium-argon chronology of the moon. Phys Chem Earth 10:145-195 Turner G, Cadogan PH (1974) Possible effects of 39Ar recoil in 40Ar-39Ar dating. Geochim Cosmochim Acta 5:1601-1615 Turner G, Bannon MP (1992) Argon isotope geochemistry of inclusion fluids from granite-associated mineral veins in southwest and northeast England. Geochim Cosmochim Acta 56:227-243 Turner G, Wang S (1992) Excess argon, crustal fluids and apparent isochrons from crushing K-feldspar. Earth Planet Sci Lett 110:193-211 Turner G, Burnard P, Ford JL, Gilmour JD, Lyon IC, Stuart FM (1993) Tracing fluid sources and interactions. Phil Trans Roy Soc London Ser A 344:127-140 Turner G, Knott SF, Ash RD, Gilmour JD (1997) Ar-Ar chronology of the Martian meteorite ALH84001: Evidence for the timing of the early bombardment of Mars. Geochim Cosmochim Acta 61:3835-3850 van der Bogaard P, Schirnick C (1995) 40Ar/39Ar laser-probe ages of Bishop Tuff quartz phenocrysts substantiate long-lived silicic magma chamber at Long Valley, United States. Geology 23:759-762 Vance D, Ayres M, Kelley SP, Harris NBW (1998) The thermal response of a metamorphic belt to extension: constraints from laser Ar data on metamorphic micas. Earth Planet Sci Lett 162:153-164 Vasconcelos PM (1999) K-Ar and Ar-40/Ar-39 geochronology of weathering processes. Ann Rev Earth Planet Sci 27:183-229 Villa IM (1994) Multipath Ar transport in K-feldspar deduced from isothermal heating experiments. Earth Planet Sci Lett 122:393-401 Villa IM (1997) Direct determination of 39Ar recoil distance. Geochim Cosmochim Acta 61:689-691 Villa IM, Grobéty B, Kelley SP, Trigila R, Wieler R (1996) Assessing Ar transport paths and mechanisms in the McClure Mountains hornblende. Contrib Mineral Petrol 126:67-80 Warnock AC, Zeitler PK (1998) Ar-40/Ar-39 thermochronometry of K-feldspar from the KTB borehole, Germany. Earth Planet Sci Lett 158:67-79 Wartho JA (1995) Apparent Argon Diffusive Loss Ar-40/Ar-39 Age Spectra in Amphiboles. Earth Planet Sci Lett 134:393-407 Wartho JA, Dodson MH, Rex DC, Guise PG (1991) Mechanisms of Ar release from Himalayan metamorphic hornblende. Am Mineral 76:1446-1448 Wartho J-A, Rex DD, Guise PG (1996) Excess argon in amphiboles linked to greenschist facies alteration in Kamila amphibolite belt, Kohistan island arc system, northern Pakiston: Insights from 40Ar/39Ar step-heating and acid leaching experiments. Geol Mag 133:595-609

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Wartho JA, Kelley SP, Brooker RA, Carroll MR, Villa IM, Lee MR (1999) Direct measurement of Ar diffusion profiles in a gem-duality Madagascar K-feldspar using the ultra-violet laser ablation microprobe (UVLAMP). Earth Planet Sci Lett 170:141-153 Wartho JA, Kelley SP, Blake S (2001) Magma flow regimes in sills deduced from Ar isotope systematics of host rocks. J Geophys Res-Solid Earth 106:4017-4035 Wieler R (2002) Cosmic-ray-produced noble gases in meteorites. Rev Mineral Geochem 47:125-170 Wheeler J (1996) A program for simulating argon diffusion profiles in minerals. Computers Geosciences 28:919-929 Wilch TI, McIntosh WC, Dunbar NW (1999) Late Quaternary volcanic activity in Marie Byrd Land: Potential Ar-40/Ar-39-dated time horizons in West Antarctic ice and marine cores. Geol Soc Am Bull 111:1563-1580 Wright N, Layer PW, York D (1991) New insights into thermal history from single grain Ar-40/Ar-39 analysis of biotite. Earth Planet Sci Lett 104:70-79 Zeitler PK, FitzGerald JD (1986) Saddle-shaped 40Ar/39Ar age spectra from young, microstructurally complex potassium feldspars. Geochim Cosmochim Acta 50:1185-1199

18

(U-Th)/He Dating: Techniques, Calibrations, and Applications Kenneth A. Farley Division of Geological and Planetary Sciences California Institute of Technology Pasadena, California 91125 [email protected]

INTRODUCTION The possibility of dating minerals by the accumulation of 4He from U and Th decay has been recognized for many years (e.g., Strutt 1905), but in the century since the idea was first conceived, the method has rarely been applied successfully. After several investigations of (U-Th)/He dating of various minerals (e.g., Damon and Kulp 1957; Fanale and Kulp 1962; Damon and Green 1963; Turekian et al. 1970; Bender 1973; Leventhal 1975; Ferreira et al. 1975) the technique was essentially abandoned as yielding unreliable and usually low ages, presumably as a result of diffusive He loss possibly associated with radiation damage. In 1987, Zeitler and coworkers rekindled interest in the method by proposing that in the case of apatite, He ages might be meaningfully interpreted as ages of cooling through very low temperatures. Laboratory diffusion data presented by these authors indicated a closure temperature of about 100ºC, a value supported by more recent studies (Lippolt et al. 1994; Wolf et al. 1996b; Warnock et al. 1997). Consistent with this interpretation Wolf et al. (1996a) found that apatite He ages increase systematically with sample elevation in a mountain range, as expected for exhumation-induced cooling through a low closure temperature. Based on the strength of these results and additional laboratory (Farley 2000) and natural (Warnock et al. 1997; House et al. 1999; Stockli et al. 2000) constraints on He diffusivity, recent attention has focused on applications of apatite He thermochronometry. There is also renewed interest in He dating of other U- and Th-bearing minerals both for dating mineral formation and for thermochronometry. For example, Lippolt and coworkers have undertaken detailed studies of He diffusion and dating of various phases, most notably hematite formed in hydrothermal systems (Lippolt and Weigel 1988; Wernicke and Lippolt 1992; Lippolt et al. 1993; Wernicke and Lippolt 1994a,b). Here I present an overview of recent techniques, calibrations, and applications of the (U-Th)/He dating method; Hurley (1954) provides an excellent summary of earlier work in this field. Much of this paper focuses on apatite, because the He behavior and requisite analytical techniques are better established for this phase than for other target minerals, such as zircon and titanite. Similarly, much of this paper concerns He diffusivity behavior required for thermochronometric applications, yet recent work is also considering applications to direct dating, for example, of young tephras (Farley et al. 2001). Rationale A wide variety of dating techniques are already well developed and widely applied. Given the methodological complexity, the limited analytical precision, and the rather small observational base presently available, it seems appropriate to consider why development and application of (U-Th)/He dating has received renewed attention. Two immediate answers come to mind. First, many materials are inappropriate for dating by existing techniques, most obviously because they are too poor in parent or radiogenic daughter isotope or too rich in non-in situ produced daughter. Thus a new technique with different requirements is potentially advantageous. This is indeed true for the (U-Th)/He 1529-6466/02/0047-0018$05.00

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system: the sensitivity for measurement of U, Th, and He is extremely high and the background of “excess” He is low, so both young materials and those with only a trace of U and Th are potentially datable. As a result young volcanic rocks lacking sanidine for Ar/Ar dating are a potential target of the (U-Th)/He method (e.g., Graham et al. 1987; Farley et al. 2002). At most this is likely to be a niche market. Greater interest in (U-Th)/He dating arises from the fact that He ages of various minerals can be used to delineate the cooling history of rocks through a temperature range that is only partially accessed by existing dating techniques. There is a large family of minerals potentially suitable for He thermochronometry, only a few of which have been explored. Figure 1 presents an overview of estimated He closure temperatures of several minerals (described in detail below) compared with closure temperatures of other thermochronometric methods. He dating complements existing techniques, and of particular interest is the apatite helium method, which is sensitive to temperatures substantially lower than any other method. Geologic applications which illustrate the unique uses of (U-Th)/He dating are presented at the end of this chapter. TECHNICAL ASPECTS He ingrowth 4

He nuclei (α particles) are produced by the series decay of 238U, 235U, and 232Th and by α decay of 147Sm. In essentially all minerals the overwhelming majority of radiogenic helium derives from actinide decay, so the ingrowth equation is: He = 8 238U (eλ 238 t −1) + 7 (238U/137.88) (eλ 235 t −1) + 6 232Th (eλ 232 t −1) (1)

Figure 1. Nominal closure temperatures of various thermochronometers showing how He systems (z) complement existing techniques („: Ar; Δ: fission track). Systems are simply ordered by closure temperature on the Y-axis. With the exception of apatite and titanite, the He closure temperatures are not yet well-known. Data sources: apatite He (Farley 2000); apatite fission track (Gallagher et al. 1998); hematite He (Wernicke et al. 1994); K feldspar multi-diffusion domain Ar (MDD) and other Ar methods (McDougall and Harrison 1999 and references therein); zircon fission track (Yamada et al. 1995); titanite fission track (Coyle and Wagner 1995), garnet He (Dunai and Roselieb 1996).

(U-Th)/He Dating

821

where 4He, U, and Th refer to present-day amounts, t is the accumulation time or He age, and λ is the decay constant (λ238 = 1.551 × 10–10 yr-1, λ235 = 9.849 × 10–10 yr-1, λ232 = 4.948 × 10–11 yr-1). The coefficients preceding the U and Th abundances account for the multiple α particles emitted within each of the decay series, and the factor of (1/137.88) is the present day 235U/238U ratio. This equation assumes secular equilibrium among all daughters in the decay chain, a condition guaranteed for crystals formed more than ~350 kyr prior to the onset of He accumulation. For most applications this assumption is valid, but in certain cases the effects of secular disequilibrium must be considered (see below). Equation (1) assumes the absence of initial 4He in the crystal being dated, and this is in general a good assumption. For example, while atmospheric Ar frequently accounts for a substantial fraction of the 40Ar in a K/Ar or Ar/Ar analysis (Kelley 2002, this volume), the concentration of He in the atmosphere is so low (5 ppm vs. ~1% for 40Ar) that trapped atmospheric He is unlikely to be important. In some cases fluid inclusions may carry crustal or mantle helium, but for U,Th-rich minerals like apatite, zircon, and titanite, the He concentration of such fluids and/or the inclusion density would have to be high to affect He ages except when the He ages are very low. Nevertheless some workers believe they have detected fluid-inclusion-hosted helium in a few apatites (Lippolt et al. 1994; Stockli et al. 2000). The presence of helium “inherited” from some prior history, for example due to incomplete degassing of a crystal stoped into a magma chamber, is unlikely given the high diffusivity of He in most solids. However, beryl and other cyclosilicates sometimes harbor very high concentrations of “excess” helium from a poorly understood source (e.g., Toyoda and Ozima 1988). This helium probably enters the cyclosilicates through the large central channels of these minerals, making them unsuitable for He dating. A final potential source of excess He is solution from surrounding fluids into grain interiors (see Kelley 2002, this volume, for a discussion of the analogous problem in Ar geochronology). At present neither laboratory data (such as He solubilities in minerals at relevant temperatures) nor sufficient age data are available to quantitatively evaluate the significance of this phenomenon. DIFFUSION BEHAVIOR Knowledge of the He-retention characteristics of the phase being dated is critical for correct interpretation of (U-Th)/He data. For example, it is known that He is not retained under Earth-surface conditions in quartz (Trull et al. 1991), sanidine, and muscovite (Lippolt and Weigel 1988), so these phases have little obvious potential for any type of helium dating. However, helium is thought to be retained at the Earth’s surface in olivine (Trull et al. 1991), pyroxene (Lippolt and Weigel 1988), amphibole (Lippolt and Weigel 1988), garnet (Dunai and Roselieb 1996), non-metamict zircon (Hurley 1952; Damon and Kulp 1957), non-metamict titanite (Hurley 1952; Reiners and Farley 1999), apatite (Zeitler et al. 1987), allanite (Wolf 1997), magnetite (Fanale and Kulp 1962), hematite (Wernicke and Lippolt 1994a,b; Bahr et al. 1994) and submarine basaltic glass (Graham et al. 1987). Several of these phases, especially olivine and pyroxene, have been used extensively for cosmogenic 3He studies, so He retention is robustly established (Niedermann 2002, this volume). In others the conditions for retention (temperature, grain size, degree of radiation damage) have not been explored. While in the case of rapidly cooled rocks at the Earth’s surface (e.g., volcanics) demonstration of quantitative retention at ~25ºC is sufficient to successfully apply He dating, thermochronometry of slowly cooled rocks requires precise knowledge of how diffusivity scales with temperature. Typically, laboratory experiments are used to constrain the parameters of the Arrhenius relationship:

Farley

822 D D o − E a /RT = e a2 a2

(2)

where D is the diffusivity, Do the diffusivity at infinite temperature, Ea the activation energy, R the gas constant, T the Kelvin temperature, and a the diffusion domain radius (Fechtig and Kalbitzer 1966). If this relationship is obeyed, then measurements of lnD/a2 as a function of reciprocal temperature will plot on a straight line with intercept lnDo/a2 and slope -Ea/R. If the measurements do not plot on a straight line, more complex behavior such as multiple diffusion mechanisms or domain sizes may be involved. It is important to note that such laboratory measurements may not apply under natural conditions. For example, diffusion coefficients are commonly measured at temperatures far higher than are relevant in nature, so large and potentially inaccurate extrapolations are often necessary. Similarly, some minerals undergo chemical or structural transformations and possibly defect annealing during vacuum heating; extrapolation of laboratory data from these modified phases to natural conditions may lead to erroneous predictions. Substantial effort is required to measure the He diffusivity parameters in a given phase and to determine how those parameters vary with mineral characteristics such as grain size and shape, chemical composition, and defect and/or radiation damage density. Only apatite has been studied in detail, but limited He diffusivity information is available on other minerals as well, as discussed below. Apatite Zeitler et al. (1987) initiated interest in He thermochronometry by demonstrating that apatite has a closure temperature of about 100ºC (here and elsewhere closure temperatures are referenced to a cooling rate of 10ºC/Myr (Dodson 1973; Kelley 2002, this volume). More recent efforts (Lippolt et al. 1994; Wolf et al. 1996b; Warnock et al. 1997; Farley 2000) confirm this approximate closure temperature, and in addition suggest: 1) He diffusion from Durango apatite (Young et al. 1969, a common “standard” apatite), as well as a variety of other apatites, obeys an Arrhenius relationship (Eqn. 2; illustrated in Fig. 2), suggesting that He diffusion from apatite is a singlemechanism thermally activated volume diffusion process, at least at temperatures