Oxygen in the Solar System 9781501508509, 9780939950805

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Frontispiece (on facing page): This diagram is a compilation of oxygen isotopic analytical data for meteoritic and lunar samples, collected since about 1973 in the University of Chicago lab of Robert N. Clayton and Toshiko Mayeda. 8lsO and 8170 give the deviations, in parts per 1000 (permil; %o), of the ratios ls0/160 and 170/160, respectively, in samples relative to Standard Mean Ocean Water (SMOW). The diagram is, in effect, an isotopic map of solar system bodies. All terrestrial and lunar samples plot on the so-called terrestrial fractionation line (TF), which has a slope of ~'/2, along which the variation can be explained by simple mass-dependent, physicochemical processes such as evaporation, condensation, and igneous crystallization. Meteoritic samples mostly do not plot on this line, and either plot on separate slope-'A lines parallel to but displaced from TF, or else along linear arrays having slopes closer to 1. The Carbonaceous Chondrite Anhydrous Mineral (CCAM) line is the dominant slope-1 array, whose extreme 160-rich end is defined by calcium-, aluminum-rich inclusions (CAIs). The CCAM line requires non-mass-dependent isotopic effects or mixing of an additional 160-rich component. Some CAIs disperse to the right of the CCAM line owing to melt volatilization that resulted in massdependent isotopic fractionation (e.g., the FUN line).

Reviews in Mineralogy and Geochemistry, Volume 68

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1529-6466

978-0-939950-80-5

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The Solar System in Clayton-Mayeda Space —

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CAIs (bulk) CO & CV chondrites CI chondrites CM chondrites CV3 chondrules LL chondrites L chondrites H chondrites OC chondrules Moon Aubrites Brachinites Silicates in irons Primitive achondrites Ureilites Mesosiderites Pallasites HEDs SNCs

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DEDICATION TO ROBERT N. CLAYTON

It is most fitting that this volume be dedicated to Professor Robert N. Clayton who, among geochemists and cosmochemists, could easily wear the name "Mr. Oxygen." His 1973 discovery (with coauthors Larry Grossman and Toshiko Mayeda), that calcium-, aluminumrich inclusions (CAIs) in the Allende meteorite have oxygen isotonic (specifically, 160-rich) compositions that cannot be explained by simple physical processes such as condensation or evaporation, led to the more general discovery of isotope anomalies in CAIs. Working with his long-time associate Tosh Mayeda and many other colleagues over the years, Bob demonstrated the power of 3-isotope oxygen measurements to fingerprint different meteorite classes and, by extension, their different parent bodies. Although the original interpretation of the 160rich signature in CAIs and other early solar system materials as being presolar in origin has since been abandoned even by Bob himself, the intense interest generated by the 1973 work revolutionized cosmochemistry in a way that continues to this day. Perhaps one way of putting into perspective the giant stature of Bob Clayton within the planetary science and isotope chemistry community is to note the following observation. Isotope geochemists (a group that includes many of us) tend to be a disputatious lot who will readily disagree with one another at meetings over just about anything large or small. Yet Bob is one person with whom most of this community will only rarely disagree: he very commonly is allowed to have The Last Word. In our experience, it is a rare thing for anyone to be so revered. Bob has already received many high honors, not least being the National Medal of Science, election to the U.S. National Academy of Sciences and the Royal society of London, and the Leonard Medal of the Meteoritical Society. The significance of dedicating this volume to him, although seemingly small in comparison with those other exalted tributes, is that this comes directly from a large number of Bob's fiiends and colleagues with whom he has worked and collaborated over many years. Thank you, Bob, for your many great accomplishments and for being an inspiration to us all. The authors and editors of this volume are proud to dedicate it in your honor, and hope that it lives up to your high standards.

OXYGEN IN THE SOLAR SYSTEM 68

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FROM THE SERIES EDITOR This volume was jointly published by the Mineralogical Society of America (MSA) and the Lunar and Planetary Institute (LPI). Such huge undertakings rarely go with out hitting a few "bumps in the road," yet Glenn MacPherson and his team prevailed and produced a finished product all could be proud of. Steve Simon, the behind-the-scenes guy, made my job exceptionally easy and I thank him for all his hard work! Plus, he created the index for this volume—a rarity in this series yet greatly appreciated! You can learn more about this volume and the "Oxygen Inititive" in the Introduction written by Glenn MacPherson. Any supplemental material and errata (if any) can be found at the MSA website www. minsocam.org. A searchable, electronic version of this volume can be found on the GeoScienceWorld website www.geoscienceworld.org. Todi T. P-osso, Series Editor West Richland, Washington January 2008

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OXYGEN IN THE SOLAR SYSTEM 68

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TABLE OF CONTENTS Introduction Glenn J. MacPherson 1

Oxygen Isotopes in the Early Solar System A Historical Perspective Robert N. Clayton 5 5 6 8 9 9 10 10 11 12 12 12

ABSTRACT BEFORE ALLENDE AFTER ALLENDE FUN C Als OXYGEN ISOTOPES IN PRESOLAR GRAINS CHEMICAL ISOTOPI ! EFFECTS PHOTOCHEMICAL EFFECTS INTERNAL ASTEROIDAL PROCESSES NITROGEN CONCLUSIONS ACKNOWLEDGMENT REFERENCES

3

Abundance, Notation, and Fractionation of Light Stable Isotopes Robert E. Criss, James Farquhar

ABSTRACT INTRODUCTION

15 15 vii

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ISOTOPIC ABUNDANCES AND ATOMIC WEIGHTS NOTATION Isotope ratios 5-values Isotopic fractionation factor Big delta and related approximations Capital delta Capital delta prime and delta prime Material balance COMMONLY-USED DIAGRAMS 5 - 5 plot Big A and Cap A plots Three-isotope plot ISOTOPIC FRACTIONATION PROCESSES Mass-dependent fractionation Kinetic processes Non-mass-dependent fractionations CONCLUSIONS REFERENCES

T1

16 18 18 19 19 19 20 20 21 21 21 21 22 24 24 26 27 28 29

Nucleosynthesis and Chemical Evolution of Oxygen Bradley S. Meyer, Larry R. Nittler, Ann N. Nguyen, Scott Messenger

ABSTRACT INTRODUCTION NUCLEOSYNTHESIS OF THE ISOTOPES OF OXYGEN Production of oxygen in mainline stellar burning stages Analysis of the oxygen yields from massive stars Low-mass stars Novae and Type la supernovae CHEMICAL EVOLUTION OF THE ISOTOPES OF OXYGEN OXYGEN IN PRES OL AR GRAINS Oxygen in carbonaceous grains Presolar oxide and silicate grains CONCLUDING REMARKS ACKNOWLEDGMENTS REFERENCES

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31 31 32 32 36 38 41 41 45 46 47 50 51 51

Oxygen in the Solar System - Table of Contents

3

Oxygen in the Interstellar Medium Adam G. Jensen, F. Markwick-Kemper, Theodore P. Snow

ABSTRACT INTRODUCTION Phases in the interstellar medium Forms of oxygen in the interstellar medium OXYGEN IN I III! GAS PHASE Measurements of gas-phase oxygen Isotope measurements from gas-phase oxygen and carbon monoxide Inferring gas-phase depletions of oxygen OXYGEN IN INTERSTELLAR DUST Solar System silicates Silicates in circumstellar environments of young stars Dust properties in the interstellar medium Dust production by evolved stars CONSISTENCY BETWEEN GAS AND SOLID PHASES Abundance and depletion constraints Transitions between the solid and gas phase in the interstellar medium SUMMARY REFERENCES

O

55 55 56 56 56 56 60 61 63 63 63 64 65 66 66 67 68 68

Oxygen in the Sun Andrew M. Davis, Ko Hashizume, Marc Chaussidon, Trevor R. Ireland, Carlos Allende Prieto, David L. Lambert

ABSTRACT INTRODUCTION THE SOLAR PHOTOSPHERIC ABUNDANCE OF OXYGEN OXYGEN ISOTOPIC COMPOSITION OF THE SUN Predictions of the isotopic composition of the Sun Spectroscopic constraints on the oxygen isotopic composition of the Sun Identification of the solar isotopic composition trapped in lunar samples Oxygen isotopic composition of the solar wind: direct measurements Summary of solar oxygen isotopic composition ACKNOWLEDGMENTS REFERENCES

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73 74 74 77 77 78 79 87 88 89 89

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Redox Conditions in the Solar Nebula: Observational, Experimental, and Theoretical Constraints Lawrence Grossman, John R. Beckett, Alexei V. Fedkin, Steven B. Simon, Fred J. Ciesla

ABSTRACT INTRODUCTION OXYGEN FUGACITY DURING CRYSTALLIZATION OF REFRACTORY INCLUSION Ml I I S Experimental technique Results Thermochemistry Selection of fassaite-melilite pairs Oxygen barometry THE OXIDATION STATE OF IRON IN ORDINARY CHONDRITES The problem Radial transport processes Vertical transport processes Relationship between f0l of cosmic gases and abundances of C, O and H Condensation of fayalitic olivine Change of FeO/(FeO + MgO) during chondrule melting REDOX CONDITIONS INFERRED FROM OTHER IRON-BEARING NEBULAR MATERIALS Amoeboid olivine aggregates Metal grains in CH chondrites FORMATION CONDITIONS OF ENSTATITE CHONDRITES Mineralogy of EH3 enstatite chondrites Condensation at high C/O ratio Bulk chemical compositions of EH enstatite chondrites Condensation of EH enstatite chondrites Formation conditions of EH3 enstatite chondrites CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

O

93 94 94 94 96 99 103 105 109 109 Ill 112 114 115 124 126 126 127 127 127 127 130 130 134 135 136 136

Oxygen Isotopes of Chondritic Components Hisayoshi Yurimoto, Alexander N. Krot, Byeon-Gak Choi, Jerome Aleon, Takuya Kunihiro, Adrian J. Brearley

ABSTRACT INTRODUCTION CHONDRITES AND THEIR COMPONENTS x

141 142 144

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OXYGEN ISOTOPIC COMPOSITIONS OF SECONDARY PHASES OXYGEN ISOTOPIC COMPOSITIONS OF REFRACTORY INCLUSIONS Alteration and secondary minerals of fine grained CAIs (FGIs) ORIGINAL OXYGEN ISOTOPIC DISTRIBUTION OF FGIs FGIs in primitive O chondrites FGIs in primitive E chondrites FGIs in CO 3.0 chondrites FGIs in CR chondrites FGIs in CM chondrites FGIs in CV chondrites FGIs in CH chondrites FGIs in CB chondrites Chondrule-bearing FGI Summary of oxygen isotopie characteristics of FGIs Alteration and secondary minerals of amoeboid olivine aggregates (AOAs) ORIGINAL OXYGEN ISOTOPIC DISTRIBUTION OF AOAs AOAs in CV chondrites AOAs in CO, CR, Acfer 094 and CM chondrites Summary of oxygen isotopie characteristics of AOAs OXYGEN ISOTOPIC DISTRIBUTION OF COARSE-GRAINED CAIs (CGIs) 7R-19-1, a compact Type A CGI E49, a compact Type A CGI SS-02, a Type B2 CGI IT A 1 01. a Type B2 CGI 1623-2, a compact Type A CGI V2-01, a fluffy Type A CGI Chondrule-bearing CGIs Summary of oxygen isotopie characteristics of CGIs OXYGEN ISOTOPIC COMPOSITIONS OF CHONDRULES Chondrules in CH chondrites Chondrules in CR chondrites Chondrules in CB chondrites Chondrules in CO and Acfer 094 chondrites Chondrules in CV chondrites Chondrules in E chondrites Chondrules in ordinary chondrites Refractory inclusion-bearing chondrules Summary of oxygen isotopie characteristics of chondrules OXYGEN ISOTOPIC COMPOSITIONS OF MATRIX Existence of submicron silicate grains with extreme non-solar oxygen isotopie compositions Oxygen isotopie heterogeneity of matrix Summary of oxygen isotopie characteristics of matrix IMPLICATIONS FOR ASTROPHYSICAL SETTING OF CHONDRITIC COMPONENT FORMATION ACKNOWLEDGMENTS REFERENCES

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145 149 149 151 151 151 151 152 154 154 155 155 155 156 156 158 158 159 161 161 163 163 165 166 166 167 168 169 169 170 170 170 171 172 172 172 172 174 175 176 177 179 179 181 182

Oxygen in the Solar System - Table of Contents

Mass-independent Oxygen Isotope Variation in the Solar Nebula Edward D. Young, Kyoshi Kuramoto, Rudolph A. Marcus, Hisayoshi Yurimoto, Stein B. Jacobsen ABSTRACT 187 INTRODUCTION 188 GALACTIC OXYGEN ISOTOPE EVOLUTION-A NON-CHEMICAL PATH TO M ASS INDEPENDENCE 189 Testing the hypothesis - the oxygen isotopic composition of the Sun 191 CHEMICAL MASS-INDEPENDENT OXYGEN ISOTOPE FRACTIONATION 192 The MIF in ozone formation 193 Conditions for a chemical MIF in the formation of CAIs 195 A possible chemical mechanism for MIF in CAIs 195 Consequences of chemical mechanism for MIF in the early water 196 Testing the hypothesis: experiment to test gas phase MIF at high temperature 197 PHOTOCHEMICAL MASS-INDEPENDENT OXYGEN ISOTOPE FRACTIONATION: CO SELF-SHIELDING 198 CO photodissociation and self-shielding 198 Astronomical observations of oxygen isotope fractionation by CO self-shielding... 200 The pivotal role of H 2 0 201 CO self-shielding at the inner annulus of the solar circumstellar disk 203 CO self-shielding at the surfaces of the solar circumstellar disk 204 CO self-shielding in molecular clouds and inheritance in the Solar System 210 Testing the hypotheses: predictions of the CO self-shielding models 212 SUMMARY 213 REFERENCES 214

Oxygen and Other Volatiles in the Giant Planets and their Satellites Michael H. Wong, Jonathan I. Lunine, Sushil K. Atreya, Torrence Johnson, Paul R. Mahajfy, Tobias C. Owen, Thérèse Encrenaz ABSTRACT INTRODUCTION Oxygen-based insights from the outer planets and their moons The protosolar abundances MEASURING OXYGEN IN JUPITER'S ATMOSPHERE Structure of the cloud layers Galileo Probe Mass Spectrometer water mixing ratio measurements The probe entry site: A 5-jam hot spot Spectroscopic measurements of Jovian water xii

219 220 220 221 222 222 223 225 226

Oxygen in the Solar System - Table of Contents Lightning on Jupiter Oxygen isotopes in Jupiter Summary of Jovian oxygen OUTER PLANET VOLATILE GASES Oxygen and other heavy element enrichments in Jupiter Volatile enrichments in the other outer planets OXYGEN IN OUTER PLANET SATELLITES Jupiter's satellites Saturn's satellites Outer Solar System satellites and Kuiper Belt Objects FORMATION OF THE OUTER PLANETS Volatile enrichment by icy planetesimals Volatile enrichment by carbonaceous planetesimals Volatile enrichment by disk evolution CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

I I

227 228 228 229 229 231 232 234 234 235 236 237 238 239 240 241 241

Oxygen in Comets and Interplanetary Dust Particles Scott A. Sandford, Scott Messenger, Michael DiSanti, Lindsay Keller, Kathrin Altwegg

ABSTRACT INTRODUCTION THE CHEMICAL FORM OF OXYGEN IN THE INTERSTELLAR MEDIUM, "COMETARY" INTERPLANETARY DUST PARTICLES, AND COMETS Oxygen carried by carbonaceous materials in the interstellar medium, meteorites, cosmic dust, and cometary samples Direct detection of oxygen-bearing volatiles in comets The oxygen-bearing minerals in "cometary" IDPs and samples from comet 8IP/Wild 2 OXYGEN ISOTOPES IN THE INTERSTELLAR MEDIUM, COMETS, COMETARY SAMPLES, AND "COMETARY" IDPS Oxygen isotopes in interstellar materials In situ measurement of the oxygen isotopes in the volatile material of comet Halley Oxygen isotopic compositions of meteorites, "cometary" IDPs and samples from comet 8IP/Wild FUTURE IN SITU MEASUREMENTS OF ISOTOPIC RATIOS IN COMETS CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

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247 247 249 249 253 258 260 261 261 262 264 265 265 265

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Oxygen and Asteroids Thomas H. Burbine, Andrew S. Rivkin, Sarah K. Noble, Thais Mothe-Diniz, William F. Bottke, Timothy J. McCoy, M. Darby Dyar, Cristina A. Thomas

ABSTRACT INTRODUCTION DYNAMICAL STRUCTURE OF THE ASTEROID BELT ASTRONOMICAL TECHNIQUES Brightness Reflectance spectroscopy Spectral data Corrections Interaction of photons with a surface ABSORPTION BANDS Electronic absorption features Vibrational absorption features SPACE WEATHERING Effect of space weathering on reflectance spectra Space weathering environment of asteroids Experimental studies Evidence of space weathering on asteroids Implications for visible/near-IR remote sensing ORDINARY CHONDRITES, LODRANITES/ACAPULCOITES, AND UREILITES DETERMINING MINERAL CHEMISTRIES Determining the ratio of olivine to pyroxene Modified Gaussian Modeling ASTEROID TAXONOMY A-types C-complex D- and P-types E- and Xe-types K- and L-types M-types O-types Q-types R-types S-complex T-types V- types HELIOCENTRIC DISTRIBUTIONS OF TAXONOMIC CLASSES IN THE MAIN BELT DISTRIBUTION OF HYDRATED ASTEROIDS IN THE MAIN BELT NEAR-EARTH ASTEROIDS SPACECRAFT MISSIONS COLLISIONAL AND DYNAMICAL EVOLUTION OF ASTEROIDS xiv

273 273 275 276 276 276 278 279 281 281 281 288 291 291 293 293 294 294 295 296 296 297 297 301 304 306 307 308 308 310 310 310 310 312 312 314 321 323 324 326

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DELIVERY Ol METEOROIDS TO EARTH THE EFFECTS OF PLANETARY EMBRYOS AND RADIAL MIXING IN I III! MAIN BELT COULD IRON METEORITES HAVE COME FROM THE TERRESTRIAL PLANET REGION? SUMMARY ACKNOWLEDGMENTS REFERENCES

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327 329 330 331 331 331

Oxygen Isotopes in Asteroidal Materials Ian A. Franchi

ABSTRACT INTRODUCTION ORDINARY CHONDRITES Introduction Ordinary chondrites - whole-rock Ordinary chondrites - components R CHONDRITES ENSTATITE METEORITES Introduction EH and EL chondrites Aubrites CARBONACEOUS CHONDRITES Introduction CV chondrites CK chondrites CO chondrites CM chondrites CI chondrites CR chondrites CH chondrites CB chondrites PRIMITIVE ACHONDRITES Introduction Acapulcoites and lodranites Brachinites Winonaites Ureilites BASALTIC ACHONDRITES Introduction Howardites, eucrites and diogenites Angrites Basaltic inclusions IRONS AND STONY-IRONS Introduction

345 346 349 349 349 351 356 358 358 358 360 361 361 361 365 366 368 371 372 373 374 375 375 376 377 377 378 379 379 380 381 382 382 382 xv

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IAB Complex IIAB III! IIIAB IVA Mesosiderites Pallasites Ungrouped irons CONCLUSIONS REFERENCES

383 384 384 385 386 387 387 388 389 390

Oxygen Isotopie Composition and Chemical Correlations in Meteorites and the Terrestrial Planets David W. Mittlefehldt, Robert N. Clayton, Michael J. Drake, Kevin Righter ABSTRACT INTRODUCTION BACKGROUND Nebular element fractionations Oxygen isotope anomalies Mechanisms of non-mass-dependent isotope fractionation CHONDRITIC METEORITES Micro- and meso-scale correlations Correlations among chondrite groups UREILITES TERRESTRIAL PLANETS SUMMARY ACKNOWLEDGMENTS REFERENCES

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Record of Low-Temperature Alteration in Asteroids Michael E. Zolensky, Alexander N. Krot, Gretchen Benedix

ABSTRACT INTRODUCTION C, P AND D ASTEROIDS - CARBONACEOUS CHONDRITES Aqueous activity on the CI parent asteroid(s) and its oxygen isotope record Aqueous activity on the CM parent asteroid(s) and its oxygen isotope record Oxygen isotopic compositions of secondary minerals in the ungrouped carbonaceous chondrite Tagish Lake Aqueous alteration of CR chondrites and their oxygen isotope record Hydrous and anhydrous alteration of CV chondrites and their oxygen isotope records Low-temperature aqueous alteration of CO chondrites Veritas asteroids - hydrous chondritic interplanetary dust particles S ASTEROIDS - ORDINARY AND R CHONDRITES Aqueous alteration of ordinary chondrites and its oxygen isotope record Aqueous alteration of R-chondrites and its oxygen isotope record M AND E ASTEROIDS - INCLUDING ENSTATITE CHONDRITES OXYGEN ISOTOPIC COMPOSITION OF ASTEROIDAL WATER AND EVOLUTION OF OXYGEN ISOTOPIC COMPOSITION OF THE INNER PROTOPLANETARY DISK SUMMARY AND FUTURE WORK ACKNOWLEDGMENTS REFERENCES

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429 429 430 430 434 437 438 439 448 448 451 451 452 452

454 455 456 456

The Oxygen Cycle of the Terrestrial Planets: Insights into the Processing and History of Oxygen in Surface Environments James Farquhar, David T. Johnston

ABSTRACT INTRODUCTION ISOTOPIC VARIATIONS AMONG TERRESTRIAL MATERIALS Historical account of oxygen isotopic variations of terrestrial reservoirs Molecular oxygen Ozone Other oxygen-bearing atmospheric species with nonzero A 17 0 Multiply substituted molecular species EVOLUTION OF OXYGEN IN EARTH'S SURFACE ENVIRONMENTS Planetary processing of oxygen OBSERVATIONS RELEVANT TO THE EVOLUTION OF OXYGEN IN THE ATMOSPHERE AND OCEANS xvii

463 463 464 465 471 471 473 473 474 475 476

Oxygen in the Solar System - Table of Contents Hypotheses about the levels of oxygen in Earth's early environments The transition from a low-oxygen atmosphere to a high oxygen atmosphere Into the Paleoproterozoic and Mesoproterozoic Oxygen and Proterozoic carbon cycle Oxygen concentration variations since the end of the Proterozoic Conceptual model for oxygenation of Earth surface environments M \ Y FRONTIERS CONCLUDING STATEMENTS ACKNOWLEDGMENTS REFERENCES

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Redox Conditions on Small Bodies, the Moon and Mars Meenakshi

ABSTRACT INTRODUCTION SMALL BODIES Brachinites and other primitive achondrites Ureilites Aubrites Angrites Eucrites THE M O O N MARS OTHER TERRESTRIAL PLANETS SUMMARY A N D CONCLUSIONS ACKNOWLEDGMENTS REFERENCES

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476 480 481 482 483 484 485 486 487 487

Wadhwa 493 493 494 494 495 496 496 496 497 499 503 505 506 506

Terrestrial Oxygen Isotope Variations and Their Implications for Planetary Lithospheres Robert E. Criss

ABSTRACT INTRODUCTION OXYGEN ISOTOPE GEOCHEMISTRY OF TERRESTRIAL ROCKS Earth's primordial 5 l s O value Oxygen isotope variations of terrestrial rocks ISOTOPIC FRACTIONATION PROCESSES Isotopic fractionation factors Fractional crystallization and AFC Real magmas Subsolidus fractionation processes xviii

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OXYGEN ISOTOPE ZONATION AND HETEROGENEITY IN PLANETARY IITHOSPIIIRIS Processes producing ls O zonation Processes producing ls O heterogeneity Bulk ls O composition of the continents Isotopic changes over geologic time CONCLUSIONS REFERENCES

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519 519 521 523 524 525 525

Basalts as Probes of Planetary Interior Redox State Christopher D. K. Herd

ABSTRACT INTRODUCTION THE OXIDATION STATE OF THE EARTH'S MANTLE The lower mantle The upper mantle OXYBAROMETERS APPLICABLE TO BASALTIC ROCKS Oxygen fugacity from mineral equilibria Multivalent trace elements Oxygen fugacity from multivalent trace elements THE BASALT-MANTLE SOURCE REDOX RELATIONSHIP Is basalt oxygen fugacity reflective of the redox state of its mantle source? Implications for understanding the redox states of planetary interiors ACKNOWLEDGMENTS REFERENCES

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527 527 528 530 531 533 535 541 541 546 546 548 549 549

Rheological Consequences of Redox State Stephen Mackwell

ABSTRACT INTRODUCTION DEFORMATION OF OLIVINE Olivine single crystal studies Olivine aggregate studies How does oxygen fugacity affect creep of olivine? DEFORMATION OF OTHER SILICATES How does oxygen fugacity affect creep of other silicates? SUMMARY REFERENCES

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Oxygen in the Solar System - Table of Contents

Appendix: Meteorites - A Brief Tutorial David W. Mittlefehldt ABSTRACT INTRODUCTION CHONDRITES Carbonaceous chondrites Ordinary chondrites Enstatite chondrites Rumuruti-like and Kakangari-like chondrites ACHONDRITES Acapulcoite-lodranite clan Winonaites and silicate inclusions from IAB (and possibly IIICD) irons Angrites Aubrites Brachinites Howardite-eucrite-diogenite clan Ureilites IRONS Magmatic iron meteorite groups Non-magmatic iron meteorite groups STONY IRONS Main-group and Eagle Station grouplet pallasites Mesosiderites ACKNOWLEDGMENTS REFERENCES

571 571 572 574 575 576 576 576 577 579 579 579 580 580 581 581 583 585 585 585 586 587 587

Subject Index

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Meteorite Index

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 1-3, 2008 Copyright © Mineralogical Society of America

Introduction Glenn J. MacPherson Smithsonian Institution PO Box 37012, MRC 119 Washington, D.C. 20013-7012, U.S.A. macphers @ si. edu

Oxygen (O) Atomic No. = 8 Atomic Wt. = 15.9994 Solar System Abundance (relative to 106 Si atoms)1 = 1.413 X 107 Three stable isotopes:

16

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Relative isotopic abundances on Earth2: 99.762%

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O, mO

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17

O, 0.200%

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O

th

First isolated as an element in the late 18 century by Scheele and Priestley (independently), oxygen is the third most abundant element in the universe. It is abundant on Earth as a colorless elemental gas in the atmosphere, in combination with hydrogen in water, and in combination with silicon and other metals in the silicates and oxides that make up Earth's crust and mantle. Oxygen is chemically very electronegative, and only a few metals (mainly, the platinum group metals, gold, silver, mercury, and copper) persist in their elemental form on Earth's surface in the presence of the oxygenated atmosphere. The abundance of elemental oxygen in Earth's atmosphere is more or less steady-state, being produced and sustained by the action of photo synthetic plants against the constant removal by physical and biogenic oxidation processes.

Hydrogen may be the most abundant element in the universe, but in science and in nature oxygen has an importance that is disproportionate to its abundance. Human beings tend to take it for granted because it is all around us and we breathe it, but consider the fact that oxygen is so reactive that in a planetary setting it is largely unstable in its elemental state. Were it not for the constant activity of photosynthetic plants and a minor amount of photo dissociation in the upper atmosphere, we would not have an oxygen-bearing atmosphere and we would not be here. Equally, the most important compound of oxygen is water, without which life (in the sense that we know it) could not exist. The role of water in virtually all geologic processes is profound, from formation of ore deposits to igneous petrogenesis to metamorphism to erosion and sedimentation. In planetary science, oxygen has a dual importance. First and foremost is its critical role in so many fundamental Solar System processes. The very nature of the terrestrial planets in our own Solar System would be much different had the oxygen to carbon ratio in the early solar nebula been somewhat lower than it was, because elements such as calcium and iron and 1 2

Abundance from Lodders (2003) From Böhlke et al. (2005)

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MacPherson

titanium would have been locked up during condensation as carbides, sulfides and nitrides and even (in the case of silicon) partly as metals rather than silicates and oxides. Equally, the role of water ice in the evolution of our Solar System is important in the early accretion and growth of the giant planets and especially Jupiter, which exerted a major control over how most of the other planets formed. On a smaller scale, oxygen plays a critical role in the diverse kinds of physical evolution of large rocky planets, because the internal oxidation state strongly influences the formation and evolution of the core, mantle and crust of differentiated planets such as the Earth. Consider that basaltic volcanism may be a nearly universal phenomenon among the evolved terrestrial planets, yet there are basalts and basalts. The basalts of Earth (mostly), Earth's Moon, Vesta (as represented by the HED meteorites) and Mars are all broadly tholeiitic and yet very different from one another, and one of the primary differences is in their relative oxidation states (for that matter, consider the differences between tholeiitic and calc-alkaline magma series on Earth). But there is another way that oxygen has proven to be hugely important in planetary science, and that is as a critical scientific clue to processes and conditions and even sources of materials. Understanding the formation and evolution of our Solar System involves reconstructing processes and events that occurred more than 4.5 Ga ago, and for which the only contemporary examples are occurring hundreds of light years away. It is a detective story in which most of the clues come from the laboratory analysis of the products of those ancient processes and events, especially those that have been preserved nearly unchanged since their formation at the Solar System's birth: meteorites; comets; and interplanetary dust particles. For example, the oxidation state of diverse early Solar System materials ranges from highly oxidized (ferric iron) to so reducing that some silicon exists in the metallic state and refractory lithophile elements such as calcium exist occur in sulfides rather than in silicates or carbonates. These variations reflect highly different environments that existed in different places and at different times. Even more crucial has been the use of oxygen 3-isotope variations, which began almost accidentally in 1973 with an attempt to do oxygen isotope thermometry on high-temperature solar nebula grains (Ca-, Al-rich inclusions) but ended with the remarkable discovery (see Clayton 2008) of non-mass-dependent oxygen isotope variations in hightemperature materials from the earliest Solar System. The presolar nebula was found to be very heterogeneous in its isotopic composition, and virtually every different planet and asteroid for which we have samples has a unique oxygen-isotopic fingerprint. The idea for this book originated with Jim Papike, who suggested the idea of a study initiative (and, ultimately, a published volume) focused on the element that is so critically important in so many ways to planetary science. He recognized that oxygen is such a constant theme through all aspects of planetary science that the proposed initiative would serve to bring together scientists from a wide range of disciplines for the kind of cross-cutting dialogue that occurs all too rarely these days. In this sense the Oxygen Initiative is modeled on the Basaltic Volcanism Study Project, which culminated in what remains to this day a hugely important reference volume (Basaltic Volcanism Study Project 1981). After obtaining community input and feedback, primarily through the Curation and Analysis Planning Team for Extraterrestrial Materials (CAPTEM) and the Management Operations Working Group for NASA's Cosmochemistry Program, a team of scientists was assembled who would serve as chapter writing leads, and the initiative was formally proposed to and accepted by the Lunar and Planetary Institute (LPI; Dr. Stephen Mackwell, Director) for sponsorship. A formal proposal was then submitted to and approved by the Mineralogical Society of America to publish the resulting volume in the Reviews in Mineralogy and Geochemistry (RiMG) series. Three open workshops were held as preludes to the book: Oxygen in the Terrestrial Planets, held in Santa Fe, NM July 20-23, 2004; Oxygen in Asteroids and Meteorites, held in Flagstaff, AZ June 2-3, 2005; and Oxygen in Earliest Solar System Materials and Processes (and including the outer planets and comets), held in Gatlinburg, TN September 19-22, 2005. The workshops were each organized around a small number of sessions (typically 4-6), each focusing on a

Introduction

3

particular topic and consisting of invited talks, shorter contributed talks, and ample time for discussion after each talk. In all of the meetings, the extended discussion periods were lively and animated, often bubbling over into the breaks and later social events. As a consequence of the cross-cutting approach, the final book spans a wide range of fields relating to oxygen, from the stellar nucleosynthesis of oxygen, to its occurrence in the interstellar medium, to the oxidation and isotopic record preserved in 4.56 Ga grains formed at the Solar System's birth, to its abundance and speciation in planets large and small, to its role in the petrologic and physical evolution of the terrestrial planets. Thanks are due to many people and organizations, without whose help and support neither the workshops nor this volume would have happened. Dr. Steve Mackwell immediately recognized the importance of the initiative and pledged LPI logistical and financial support for the workshops and publication. In particular, Sue McCown and Kimberly Taylor of the LPI provided terrific logistical and on-site support for the three workshops. Financial support for the workshops and this book was also provided by NASA's Cosmochemistry Program and its Discipline Scientist, Dr. David Lindstrom. I particularly wish to acknowledge the invaluable contributions my co-editors, Dave (Duck) Mittlefehldt and John Jones, and of the associate editors who handled papers coming out of the third workshop on Earliest Solar System Materials and processes: Drs. Andy Davis, Sasha Krot, Larry Nittler, Ed Scott, Sara Russell, and Ed Young. Dr. Steve Simon undertook what probably was the most Herculean task of all, acting as technical editor for all of the book chapters to ensure that journal style and internal consistency were maintained. He did this efficiently, cheerfully, and well. RiMG Series Editor Jodi Rosso patiently and expertly turned the edited chapters into final camera-ready copy. Finally, in the background but always present, was Jim Papike himself keeping us all on target and on schedule (and kicking some posteriors where posterior-kicking was needed). This volume is a testimony to Jim's vision, and it would not have happened without him. Basaltic Volcanism Study Project (1981) Basaltic Volcanism on the Terrestrial Planets, Pergamon Press, New York Bohlke JK, de Laeter JR, De Bievre P, Hidaka H, Peiser HS, Rosman KJR, Taylor PDP (2005) Isotopic compositions of the elements, 2001. J Phys Chem Ref Data 34:57-67 Clayton RN (2008) Oxygen isotopes in the early Solar System — A historical perspective. Rev Mineral Geochem 68:5-14 Lodders K (2003) Solar system abundances and condensation temperatures of the elements. Astrophys J 591:1220-1247

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 5-14, 2008 Copyright © Mineralogical Society of America

Oxygen Isotopes in the Early Solar System — A Historical Perspective Robert N. Clayton Enrico Fermi Institute, Department of Chemistry, Sciences Department of the Geophysical University of Chicago Chicago, Illinois 60637, U.S.A. r-clayton @ uchicago. edu

ABSTRACT The first suggestion for the use of oxygen isotopes in cosmochemistry was that of H. C. Urey and colleagues in 1934, but appropriate instrumentation had not yet been developed. The modern era of oxygen isotope cosmochemistry began with the study of Apollo lunar samples in 1969 and of Allende refractory inclusions in 1973. The large (>5%) variations in 1 7 0 / 1 6 0 and 1 8 0 / 1 6 0 ratios, and small variations in 1 7 0 / 1 8 0 were first interpreted as nucleosynthetic effects, but are now recognized to be the result of chemical processes early in Solar System history. Thus oxygen isotopes provide natural tracers for processes of formation of solid bodies in the inner Solar System. In particular, oxygen isotopes are very useful in recognizing genetic associations among meteorite groups. They also have been valuable in the study of parent body processes, such as metamorphism and aqueous alteration. There is conjecture that the ultimate cause of the oxygen isotope effects may be isotope-selective photodissociation of CO, which will be tested by isotopic measurement of solar oxygen and nitrogen collected in the NASA Genesis mission.

BEFORE ALLENDE Measurements of the oxygen isotopic compositions of extraterrestrial materials have provided unique insights into processes of formation of our Solar System (Clayton 2003). Since the three stable isotopes of oxygen ( l e O, 1 7 0, l s O) are synthesized by different nuclear processes in different astrophysical sites, any incomplete homogenization in the interstellar medium could result in variations in the relative abundances of the isotopes in samples with diverse origins. In a remarkably prescient paper, Manian et al. (1934) attempted to observe such variations by measurement of 1 8 0/ 1 6 0 ratios in meteorites that may have originated beyond the Solar System, as judged from their apparent hyperbolic orbits. Their search failed for two major reasons: (1) the absence of any "extra-solar rocks" in the size range of meteorites; and (2) the inadequacy of mass spectrometers at that time. In fact, the mass spectrum shown by Manian et al. has such low resolution that it does not even show the existence of 1 7 0, quantitative determination of which is essential for recognition of extra-solar oxygen. Analogous measurements can be carried out today with modern mass spectrometers on micrometer-sized grains of oxides and silicates, found within meteorites, and do indeed yield variations in 1 8 0/ 1 6 0 and 1 7 0/ 1 6 0 by factors of ten or more (Nittler et al. 1994), which are interpreted as residual heterogeneities of nucleosynthetic origin (to be discussed below). It was anticipated by Brown (1947) that meteorites might contain daughter products of radionuclides with half-lives in the range 10 3 -10 9 yr, yielding measurable quantities of the daughter isotopes in some phases. Reynolds (1960) found excess 129Xe from decay of the 1529-6466/08/0068-0002$05.00

DOI: 10.2138/rmg.2008.68.2

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extinct radionuclide 129I (t1/2 = 17 Ma) in the Richardton chondrite, thus indicating a short time-interval between a nucleosynthetic process and formation of the Solar System. This observation was followed by the identification by Black (1972) of neon-E, a component within carbonaceous chondrites, consisting of almost pure 22 Ne. Black stated: "... 'E' is not indigenous to the solar mix, but rather is to be assigned an extra-solar system origin". Neon-E may be produced directly in a supernova explosion or indirectly as a decay product of 22 Na (t1/2 = 2.6 yr) (Nichols et al. 1994). In light of the background information described above, it would be expected that a renewed search for "isotopic anomalies" in oxygen, measuring all three stable isotopes, would have been initiated. Such was not the case: oxygen isotope studies of meteorites in the 1960s and early 1970s continued to be based only on 1 8 0/ 1 6 0 variations, interpreted as mass-dependent fractionation effects (Taylor et al. 1965; Onuma et al. 1972a). Mass-dependent isotopic fractionation in light elements, such as hydrogen, carbon, nitrogen, oxygen, and sulfur, was put on a firm theoretical foundation by the classic work of Urey (1947). In his tabulations of vibrational frequencies and isotopic partition functions for oxygen-bearing compounds, Urey considered only l e O and l s O, tacitly implying that information from 1 7 0 was redundant and unnecessary. Almost all lightelement isotope effects in terrestrial geochemistry follow the principles described by Urey, and, to this day, almost all terrestrial oxygen isotope studies are based only on 1 8 0/ 1 6 0 variations. In some extraterrestrial applications, such as the study of oxygen isotope variations in lunar rocks, the approach based on terrestrial experience has worked well (Onuma et al. 1970; Taylor and Epstein 1970), although there are some clear exceptions, such as the virtual absence of deuterium in the hydrogen of implanted solar wind (Epstein and Taylor 1970). It was later observed that the 15 N/ 14 N ratio of solar wind nitrogen implanted in lunar soils varied by more than 30%, a range larger than can be explained by known mass-dependent effects (Thiemens and Clayton 1980), which could imply some specific solar process.

AFTER ALLENDE In 1969, laboratories around the world were gearing up for study of the first lunar samples, due in autumn of that year. On February 8, 1969, the remarkable CV3 chondrite, Allende, fell in Chihuahua State, Mexico, providing more than two tons of primitive meteorite for detailed chemical, mineralogical, and petrographic study. The fall was perfectly timed to contribute to L. Grossman's doctoral research at Yale University (Grossman 1972). Of special interest to him were the refractory, calcium-aluminum-rich inclusions (CAIs; Fig. 1), which were interpreted as being primary condensates from a hot solar gas. This suggested a novel application of oxygenisotope thermometry: a determination of the condensation temperature and the composition of the nebular gas (Onuma et al. 1972b) by detailed isotopic measurements of the individual minerals that make up the CAIs, predominantly spinel, pyroxene, melilite, plagioclase and occasionally olivine. As was standard practice at the time, mass spectrometry was done with C 0 2 as the sample gas, with mass 46/mass 44 used for measurement of 1 8 0/ 1 6 0, and mass 45/ mass 44 for measurement of 13C/12C. In the chemical preparation of C0 2 , oxygen was liberated from the sample as 0 2 , by reaction with BrF 5 , and was reacted with hot graphite to make C0 2 . Hence, the carbon isotope ratio should be constant for all samples. However, given the huge variations found for 1 8 0/ 1 6 0 in CAI minerals, we also found a linear correlation between 46/44 variations and 45/44 variations, which we recognized were due to variations in 1 7 0/ 1 6 0, since 17 0 also contributes a small amount to the ion beam at m/e = 45 ( 1 6 0 1 2 C 1 7 0). Such a correlation is expected in mass-dependent fractionations (Craig 1957), but our observations gave a slope that was twice the Craig value, showing that the magnitude of the 1 7 0/ 1 6 0 variations was equal to that of 1 8 0/ 1 6 0 variations, rather than one-half of the 1 8 0/ 1 6 0 variations, as is known for mass-dependent fractionation. These observations were the basis for the first paper on "oxygen

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Figure 1. A type B 2 CAI from Allende, showing the roughly spherical shape and igneous texture indicative of crystallization from a melt. In this false-color B S E image, dark grey is melilite, light grey is clinopyroxene, medium grey is anorthite and black is spinel (MacPherson and Grossman 1981). In a very similar Allende inclusion, A13S4, melilite has 5 l s O and 5 1 7 0 near zero, whereas pyroxene and spinel have 5 1 8 0 and 5 1 7 0 near -40%o, even for spinels entirely enclosed within melilite (Clayton et al. 1977).

isotope anomalies" (Clayton et al. 1973). Subsequent oxygen isotope analyses have been done by measuring 0 2 + , using 0 2 as the sample gas in gas-source spectrometers, or by measuring ions by Secondary Ion Mass Spectrometry ( S I M S ) . In the 1970's, it was shown that the dominant oxygen isotopic pattern in primitive extraterrestrial materials at the sub-centimeter scale (chondrules and CAIs) is a linear array with slope near 1 on a "three-isotope" graph of 5 1 7 0 versus 5 l s O (e.g., Clayton et al. 1977). Such an array corresponds to a nearly constant 1 7 0 / l s 0 ratio, with variable amounts o f l e O (Fig. 2). This could be a mixing line between two end-members, one enriched in l e O , the other depleted in l e O. The simplest explanation for the meteoritic mixing line, or C C A M (carbonaceous chondrite anhydrous minerals), was the addition of a component enriched in l e O by stellar nucleosynthesis. In the limit o f pure " l e O-nuggets", these would have to constitute about 5 % o f those mineral fractions that were the richest in l e O , and thus should be readily identified. This extreme case was clearly not supported by observation, and was replaced by the postulate of a component that was enriched by only a few percent in l e O , but still reflecting a nucleosynthetic origin (Clayton et al. 1977). One of the earliest observations of the oxygen isotope distribution in Allende CAIs was the recognition of a mineralogical control: the greatest 1 6 0-enrichments occurred in spinel and clinopyroxene, and the smallest enrichments in melilite, plagioclase, and secondary garnet and feldspathoids (Clayton et al. 1977). This was attributed to a two-stage process, with l e O-rich minerals retaining their primary isotopic compositions, and l e O-poor phases reflecting some kind of secondary exchange, probably with a nebular gas.

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5180 (%»rel. SMOW) Figure 2. An oxygen three-isotope plot showing the CCAM (carbonaceous chondrite anhydrous mineral) line defined by minerals from Allende CAIs, and, for reference, the terrestrial fractionation line (TF) defined by terrestrial rocks, minerals and waters. In early interpretations of the CCAM line, the compositions at the upper end were considered "normal" and others were called "anomalous." In the self-shielding interpretation, the solar composition lies at the lower end of the CCAM line, and other compositions reflect exchange with a photochemically processed gas.

One obvious test for a nucleosynthetic origin of l e O excesses is a search for correlated nuclear effects in other low-mass elements with which oxygen is chemically bound, such as magnesium and silicon. This test was performed on a set of typical Allende CAIs (Mittlefehldt et al. 2008), in which the fluorination reaction on each CAI yielded simultaneously 0 2 and SiF4, for oxygen and silicon isotope measurements. The oxygen data all followed the slope-1 CCAM trend, and the silicon data (829Si vs. 830Si) all followed a slope-1/2 mass-dependent fractionation trend. The oxygen and silicon isotope data were uncorrelated, showing that two separate processes were involved, and in particular, casting serious doubt on a nucleosynthetic origin for the oxygen isotope effects.

FUN CAIs Early in the isotopic studies of Allende CAIs, two inclusions were found in which unusually large mass-dependent heavy-isotope enrichment was observed in magnesium (Wasserburg et al. 1977) and silicon (Clayton et al. 1978). These CAIs, unremarkable chemically and mineralogically, also have apparent nucleosynthetic isotope anomalies in all elements that have been studied (e.g. McCulloch and Wasserburg 1978; Lee et al. 1978). They were dubbed "FUN" inclusions, for "fractionated" and "unidentified nuclear" effects (Wasserburg et al. 1977). Their oxygen isotopic patterns differ from those of most CAIs, in that an intermediate stage of mass-dependent evaporation appears to have occurred between the time of primary crystallization and the secondary isotope exchange with a more le O-poor reservoir (Clayton and Mayeda 1977; Lee et al. 1980; Clayton et al. 1984). The connection between an evaporation

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event and the presence of nuclear anomalies remains unknown. The FUN inclusions are also devoid of radiogenic 26Mg (MacPherson 2003). There are about six known FUN CAIs, all but one of which were known by 1984; only one, Vigarano 1623-5, has been discovered in the last 20 years (Davis et al. 1991), in spite of the isotopic analysis of hundreds of CAIs.

OXYGEN ISOTOPES IN PRESOLAR GRAINS The terms "presolar grains" and "stardust" refer to mineral grains that acquired their chemical and isotopic characteristics before the Sun and Solar System formed, and were subsequently incorporated into meteorite parent bodies (Anders and Zinner 1993). They are recognizable by their large differences in isotopic compositions in most elements with respect to Solar System materials. In most cases, the abundances of presolar grains in meteorites are at the level of tens of ppm, so their influence on the bulk isotopic compositions of meteorites is very small. Thus, although extremely interesting in their own right, they are not responsible for the isotopic variations discussed above. "Anomalous" isotopic variations in the noble gases due to presolar grains are indeed measurable in bulk meteorites, however; this is the property that led to the discovery of presolar grains (Lewis et al. 1987). Presolar grains condensed in cool atmospheres of earlier generations of stars, and their notable isotopic abundances result from nuclear processes in the parent stars or in previous generations of stars (galactic chemical evolution). The chemical separation techniques used by the Anders group favor isolation of carbon-rich phases, notably graphite, diamond, and silicon carbide (Anders and Zinner 1993), most of which were derived from carbon-rich stars of the Asymptotic Giant Branch (AGB). Subsequently, search techniques involving oxygen-isotope mapping have revealed presolar oxides (spinel, corundum) (Hutcheon et al. 1994) and silicates (Nguyen and Zinner 2004). The oxygen-rich grains have been assigned to four classes according to their oxygen isotopic compositions on a three-isotope plot (Nittler et al. 1997). Grains in all classes may have been derived from oxygen-rich red giant stars at various stages of stellar evolution. Grains enriched only in le O, once believed to be present in meteoritic CAIs, are exceedingly rare (Nittler et al. 1998).

CHEMICAL ISOTOPE EFFECTS In 1983, a new process was discovered that could reproduce the slope-1 array for oxygen isotopes by a purely chemical means, generically known as a non-mass-dependent (NMD) isotopic fractionation (Thiemens and Heidenreich 1983). This phenomenon was first recognized in gas phase laboratory synthesis of ozone (0 3 ) from oxygen (0 2 ), with heavyisotope enrichment in the product ozone. The authors interpreted this result as being due to isotope-selective photodissociation of 0 2 . Thiemens and Heidenreich (1983) included CO in a list of molecules that might exhibit this isotopic self-shielding effect (see below), and Navon and Wasserburg (1985) discussed the advantages of CO over 0 2 for the origin of oxygen isotope variations in meteorites. Since neither 0 2 nor 0 3 is expected to have been a major constituent of the hydrogen-rich solar nebula, it has been important to determine the underlying cause of the NMD chemical effect, to assess its generality and applicability to other molecular species. Although a number of systems have shown departures from classical mass-dependent isotope effects, only the 0 2 - 0 3 system has produced an extended slope-1 relationship. A theoretical interpretation of the NMD effect, based on symmetry-dependent kinetics, has been presented (Gao and Marcus 2002), and Marcus (2004) has attempted to extend these ideas to include the formation of le Oenriched minerals, as are seen in CAIs. Proposals to account for NMD effects in meteorites by such chemical processes have assumed that the oxygen isotopic composition of the nebular gas, and of the Sun, was similar to the terrestrial composition, and that CAIs are "anomalous"

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and require a special mechanism for their formation. In the self-shielding model, the solar composition is taken to be that of the CAIs, so that the terrestrial (and inner Solar System) composition is "anomalous" and requires a special mechanism. Determination of the isotopic composition of the solar wind, to be obtained by analysis of samples returned by the Genesis mission and assumed to be representative of the isotopic composition of the Sun, should allow us to distinguish between these proposed scenarios.

PHOTOCHEMICAL EFFECTS Isotope effects associated with the phenomenon of photochemical self-shielding have been discussed in the astrophysical literature since the early 1980s (Bally and Langer 1982; van Dishoeck and Black 1988). The isotope effects of self-shielding are based on the photodissociation of gaseous carbon monoxide, the most abundant oxygen-containing molecule in the Galaxy. This process occurs by predissociation, in which absorption of an ultraviolet photon (X = 90-100 nm) excites a molecule to a short-lived state which then dissociates into ground-state C and O atoms. The excitation occurs in narrow lines that are wavelengthspecific for the various isotopologues of CO, due to the mass-dependence of their vibrational energies. Because of its much greater abundance, the 12 C le O absorption line becomes optically thick (saturated absorption), while the lines of the rare isotopic species remain optically thin (unsaturated absorption). Thus the interior of a molecular cloud or solar nebula undergoes preferential dissociation of 13 C le O, 12 C 17 0 and 12 C ls O relative to 12 C le O, enhancing the local abundances of atomic 17 O and l s O, which can then be incorporated into other molecules, such as H 2 0, and eventually into mineral grains. The photochemical process has been modelled quantitatively for molecular clouds (Warin et al. 1996) and the predicted deficits of molecular 12 17 C 0 and 12 C ls O have been observed by ultraviolet spectroscopy (Sheffer et al. 2002). The self-shielding effect predicts that all phases produced from the gas enriched in the heavier, rare isotopes will themselves be enriched in these isotopes. As a consequence, the isotopic composition of the initial bulk cloud or nebula must be at least as le O-rich as the 16 0-richest minerals. For the Solar System, this implies that the Sun should have an oxygen isotopic composition similar to that of the spinels in CAIs, i.e., at the lower end of the mixing line in Figure 2 (Clayton 2002). Analysis of samples returned by NASA's Genesis mission will provide an oxygen isotopic analysis of the solar wind, thus providing a direct test of the selfshielding proposal. Of the three classes of proposals for the origin of the meteoritic oxygen isotope variation: nucleosynthetic; NMD chemical; and photochemical, none has been disproved, and each has adherents. The nucleosynthetic scenario led to predictions of correlated nucleosynthetic effects in magnesium and silicon, and to expectations of very le O-rich presolar grains. Neither has been found. The NMD chemical process clearly does occur in the Earth's atmosphere, but a viable extension to solar nebular conditions has not yet been found. The self-shielding mechanism has been demonstrated to occur in cold, low-density molecular clouds, but also has not yet led to a comprehensive model for production of isotopically "anomalous" solids in the Solar System.

INTERNAL ASTEROIDAL PROCESSES Both the chemical NMD processes and the photochemical processes inherently involve the gas phase in the solar nebula, with different degrees of interaction resulting in final solid products (chondrules, CAIs, planetesimals) with differing l e O excesses or deficits, expressed in terms of A 1 7 0 = 8 1 7 0 - 0.52 8 l s O. Thus an asteroid or rocky planet has a characteristic value of A 17 0, inherited from the nebula, but remaining constant for subsequent internal processes,

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such as metamorphism or melting. For example, for Earth, A 1 7 0 = 0 (by definition), but for Mars (SNC meteorites) A 17 0 = +0.32 (Franchi et al. 1999); for the HED (howardite-eucritediogenite) parent body A 1 7 0 = -0.22 (Wiechert et al. 2004); for equilibrated H-chondrites A 1 7 0 = +0.73; and for equilibrated L-chondrites A 1 7 0 = +1.07 (Clayton et al. 1991). Igneous and metamorphic processes within a parent body occur at constant A 17 0, following massdependent fractionation patterns, just as they do on Earth. Inter-mineral isotopic fractionations can, therefore, be used for isotopic thermometry, as was done for metamorphosed chondrites (Onuma et al. 1972a) and for lunar igneous rocks (Onuma et al. 1970). The use of oxygen isotope abundances to trace aqueous alteration processes in carbonaceous chondrites illustrates the additional power of the three-isotope system. It is likely that water accreted to the carbonaceous chondrites as ice, which almost certainly had A 17 0 different from that of the anhydrous silicates and other minerals. For example, in Murchison (CM2), A 1 7 0 for anhydrous minerals is -5.2%o, whereas A 17 0 for phyllosilicates is -1.9%o, implying interaction with an aqueous phase with A 17 0 more positive than -1,9%c. The data can be interpreted either in terms of simple closed-system mass balance (Clayton and Mayeda 1999) or in terms of a more complicated model including fluid flow (Young et al. 1999). These models assume the accretion of water as ice, which reacts with olivine and pyroxene upon melting. The aqueous alteration conditions are more akin to those in terrestrial sea-floor weathering than to terrestrial hydrothermal processes. An important advance in isotopic thermometry of carbonates has been made by John Eiler and colleagues (Ghosh et al. 2006), which allows estimation of precipitation temperatures without independent knowledge of the oxygen isotopic composition of the fluid phase. Application of this technique to carbonates in carbonaceous chondrites yields temperatures from 0 to 40 °C (Guo et al. 2007).

NITROGEN Another long-standing problem in Solar System isotope studies is the very large range of variations observed in the nitrogen isotopic composition. Because of the close similarity between the N 2 and CO molecules (isoelectronic, isobaric), solution of one isotope problem (oxygen or nitrogen) may aid in solution of the other. Of course, nitrogen has only two stable isotopes, so tests for mass-dependence cannot be made, and the only observable variable is the 15 N/14N ratio, or 8 15 N. However, nitrogen from meteorites exhibits a range of more than a factor of two in 15N/14N, strongly suggesting that some process beyond ordinary mass-dependent fractionation has occurred. Furthermore, the observed isotopic variations, although systematic from one meteorite group to another, have not shown correlations with any other property. In many ways, the history of the development of nitrogen isotope studies of the Solar System parallels that for oxygen isotopes. "Mainstream" presolar SiC grains are enriched in 14N by a factor of 10 or more, due to operation of the CNO nuclear cycle in their source stars (Anders and Zinner 1993), but concentrations of these grains in meteorites are too low to account for the observed whole-rock isotopic variations. The large variation in 15N/14N in nitrogen implanted in lunar soils was first interpreted as evidence for a secular increase in the ratio due to nuclear reactions near the surface of the Sun (Kerridge 1975). Geiss and Bochsler (1982) showed that such nuclear processes were quantitatively inadequate to account for the isotopic variations in implanted lunar nitrogen. Models in the late 1980s assumed two implanted components, of either lunar or solar origin (Kerridge 1989). Hashizume et al. (2000) used ion microprobe techniques on lunar ilmenite grains to distinguish implanted nitrogen of solar origin (correlated with deuterium-free hydrogen) from implanted nitrogen of non-solar origin, which they labelled "planetary," without identifying a specific origin. They concluded that lunar soil contains only one solar component, with 8 15 N in the range - 2 5 0 to

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-300%o. A composition of about -300%o has also been inferred for nitrogen in ammonia in Jupiter's atmosphere (Owen et al. 2001). This interpretation leaves unanswered the question of the origin, presumably within the Solar System, of the huge range of nitrogen isotopic compositions among various reservoirs. Owen et al. (2001) and especially Terzieva and Herbst (2000) considered the possibility of large chemical isotopic fractionations among nitrogen compounds in cold clouds. The latter authors concluded that such processes could not account for 15N enhancements by factors greater than 1.5, as has been seen in one interplanetary dust particle (Messenger et al. 1996), whereas the observed meteoritic enhancement extends to a factor of 3 (Prombo and Clayton 1985).

CONCLUSIONS If we accept the generalization by Suess (1965) that the initial solar nebula was isotopically well-homogenized, then we must find late-stage processes, operating within the Solar System, to account for the observed large isotopic differences in oxygen and nitrogen between the Sun and the inner Solar System. The process of isotopic self-shielding in the photodissociation of CO and N 2 may satisfy this requirement; the molecules have very similar dissociation energies and both have an abundant lighter isotope, and one or two rare heavier isotopes. The apparent similarity in nitrogen isotope ratios between the Sun and Jupiter would imply that the photochemical isotopic effects were limited to matter now found in the inner Solar System, effectively ruling out processes that affected the entire solar nebula (Yurimoto and Kuramoto 2004; Lyons and Young 2005). Measurements of the solar wind isotopic compositions of both elements in samples returned by NASA's Genesis mission will go a long way toward testing these hypotheses.

ACKNOWLEDGMENT This research has been supported over many years by grants from the U.S. National Science Foundation and from NASA. The most recent support is from NASA grant NAG513165. G.J. MacPherson and M.H. Thiemens provided helpful reviews.

REFERENCES Anders E, Zinner E (1993) Interstellar grains in primitive meteorites: diamond, silicon carbide, and graphite. Meteoritics 28:490-514 Bally J, Langer WD (1982) Isotope-selective photodestruction of carbon monoxide. Astrophys J 255:143-148 Black DC (1972) Trapped helium, neon and argon isotopic variations in meteorites — II Carbonaceous meteorites. Geochim Cosmochim Acta 36:377-394 Brown H (1947) An experimental method for the estimation of the age of the elements. Phys Rev 72:348 Clayton RN (2002) Self-shielding in the solar nebula. Nature 415:860-861 Clayton RN (2003) Oxygen isotopes in the solar system. Space Sci Rev 106:19-32 Clayton RN, Grossman L, Mayeda TK (1973) A component of primitive nuclear composition in carbonaceous meteorites. Science 182:485-498 Clayton RN, MacPherson GJ, Hutcheon ID, Davis AM, Grossman L, Mayeda TK, Molini-Velsko C, Allen JM (1984) Two forsterite-bearing FUN inclusions in the Allende meteorite. Geochim Cosmochim Acta 48:535-548 Clayton RN, Mayeda TK (1977) Correlated oxygen and magnesium isotopic anomalies in Allende inclusions: I. Oxygen. Geophys Res Lett 4:295-298 Clayton RN, Mayeda TK (1999) Oxygen isotope studies of carbonaceous chondrites. Geochim Cosmochim Acta 63:2089-2104 Clayton RN, Mayeda TK, Epstein S (1978) Isotopic fractionation of silicon in Allende inclusions. Proc Lunar Planet Sci Conf 9th: 1267-1278 Clayton RN, Mayeda TK, Goswami JN, Olsen EJ (1991) Oxygen isotope studies of ordinary chondrites. Geochim Cosmochim Acta 55:2317-2337

Historical Perspective: Oxygen Isotopes in the Early Solar System

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Clayton RN, Onuma N, Grossman L, Mayeda TK (1977) Distribution of the presolar component in Allende and other carbonaceous chondrites. Earth Planet Sci Lett 34:209-224 Craig H (1957) Isotopic standards for carbon and oxygen and correction factors for mass-spectrometric analysis of carbon dioxide. Geochim Cosmochim Acta 12:133-149 Davis AM, Clayton RN, Mayeda TK, Sylvester PJ, Grossman L, Hinton RW, Laughlin JR (1991) Melt solidification and late-stage evaporation of a FUN inclusion from the Vigarano C3V chondrite. Geochim Cosmochim Acta 55:621-637 Epstein S, Taylor HP Jr. (1970) The concentration and isotopic composition of hydrogen, carbon and silicon in Apollo 11 lunar rocks and minerals. Proc Apollo 11 Lunar Sci Conf, 1085-1096 Franchi IA, Wright IP, Sexton AS, Pillinger CT (1999) The oxygen isotopic composition of Earth and Mars. Meteorit Planet Sci 34:657-661 Gao YO, Marcus RA (2002) On the theory of the strange and unconventional isotope effects in ozone formation. JChemPhys 116:137-154 Geiss J, Bochsler P (1982) Nitrogen isotopes in the solar system. Geochim Cosmochim Acta 46:529-548 Ghosh P, Adkins J, Affeck H, Balta B, Guo W, Schauble EA, Schrag D, Eiler JM (2006) 1 3 C- l s O bonds in carbonate minerals: a new kind of paleothermometer. Geochim Cosmochim Acta 70:1439-1450 Grossman L (1972) Condensation, chondrites and planets. Ph.D. Dissertation, Yale University, New Haven, Connecticut Guo W, Perronnet M, Zolensky ME, Eiler JM (2007) Temperatures of aqueous alteration on carbonaceous chondrite parent bodies. Meteorit Planet Sci 42:A61 Hashizume K, Chaussidon M, Marty B, Robert F (2000) Solar wind record on the Moon: deciphering presolar from planetary nitrogen. Science 290:1142-1145 Hutcheon ID, Huss GB, Fahey AJ, Wasserburg GJ (1994) Extreme 26 Mg and 1 7 0 enrichments in an Orgueil corundum: identification of a presolar oxide grain. Astrophys J 425:L97-L100 Kerridge JF (1975) Solar nitrogen: evidence for a secular increase in the ratio of nitrogen-15 to nitrogen-14. Science 188:162-164 Kerridge J F (1989) What has caused the secular increase in solar nitrogen-15? Science 245:480-486 Lee T, Mayeda TK, Clayton RN (1980) Oxygen isotopic anomalies in Allende inclusion HAL. Geophys Res Lett 7:493-496 Lewis RS, Tang M, Wacker JF, Anders E, Steel E (1987) Interstellar diamonds in meteorites. Nature 326:160162 Lyons JR, Young ED (2005) CO self-shielding as the origin of oxygen isotope anomalies in the early solar nebula. Nature 435:317-320 MacPherson GJ (2003) Calcium-aluminum-rich inclusions in chondritic meteorites. In: Treatise on Geochemistry, vol. 1. Davis A (ed) Elsevier, Oxford UK, p 201-246 MacPherson GJ, Grossman L (1981) A once-molten, coarse-grained Ca-rich inclusion in Allende. Earth Planet Sci Lett 52:16-24 Manian SH, Urey HC, Bleakney W (1934) An investigation of the relative abundance of the oxygen isotopes 0 1 6 : 0 1 8 in stone meteorites. J Am Chem Soc 56:2601-2609 Marcus RA (2004) Mass-independent isotope effect in the earliest processed solids in the solar system: A possible chemical mechanism. J Chem Phys 121:8201-8211 Messenger S, Keller LP, Thomas KL, Walker RM (1996) Nitrogen petrography in two 15N-rich IDPs. Meteorit Planet Sci31:A88 Mittlefehldt DW, Clayton RN, Drake MJ, Righter K (2008) Oxygen isotopic composition and chemical correlations in meteorites and the terrestrial planets. Rev Mineral Geochem 68:399-428 Navon O, Wasserburg GJ (1985) Self-shielding in 0 2 — a possible explanation for oxygen isotopic anomalies in meteorites? Earth Planet Sci Lett 73:1-16 Nguyen A, Zinner E (2004) Discovery of ancient silicate stardust in a meteorite. Science 303:1496-1499 Nichols RH Jr., Kehm K, Brazzle R, Amari S, Hohenberg CM, Lewis RS (1994) Ne, C, N, O, Mg and Si isotopes in single interstellar graphite grains: Multiple stellar sources for Neon E (L). Meteoritics 29:510-511 Nittler LR, Alexander CMO'D, Gao X, Walker RM, Zinner E K (1994) Interstellar oxide grains from the Tieschitz ordinary chondrite. Nature 370:443-446 Nittler LR, Alexander CMO'D, Gao X, Walker RM, Zinner E K (1997) Stellar sapphires: the properties and origins of presolar A1 2 0 3 in meteorites. Astrophys J 482:475^195 Nittler LR, Alexander CMO'D, Wang J (1998) Meteoritic oxide grain from supernova found. Nature 393:222 Onuma N, Clayton RN, Mayeda T K (1970) Oxygen isotope fractionation between minerals and an estimate of the temperatures of formation. Science 167:536-538 Onuma N, Clayton RN, Mayeda T K (1972a) Oxygen isotope temperatures of "equilibrated" ordinary chondrites. Geochim Cosmochim Acta 36:157-168 Onuma N, Clayton RN, Mayeda T K (1972b) Oxygen isotope cosmothermometer. Geochim Cosmochim Acta 36:169-188

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Owen T, Mahaffy PR, Niemann HB, Atreya S, Wong M (2001) Protosolar nitrogen. Astrophys J 5 5 3 : L 7 7 - L 7 9 Prombo CA, Clayton RN (1985) A striking isotope anomaly in the Bencubbin and Weatherford meteorites. Science 230:935-937 Reynolds JH (1960) Determination of the age of the elements. Phys Rev Lett 4:8-10 Sheffer Y, Lambert DL, Federman S R (2002) Ultraviolet detection of interstellar 1 2 C n O and the CO isotopomeric ratios toward X-Persei. Astrophys J 5 7 4 : L 1 7 1 - L 1 7 4 Suess HE (1965) Chemical evidence bearing on the origin of the solar system. Ann Rev Astron Astrophys 3:217-234 Taylor HP Jr., Epstein S (1970) 0 1 8 / 0 1 6 ratios of Apollo 11 lunar rocks and minerals. Proc Apollo 11 Lunar Sci Conf, 1613-1626 Taylor HP Jr., Duke M B , Silver LT, Epstein S (1965) Oxygen isotope studies of minerals in stony meteorites. Geochim Cosmochim Acta 29:489-512 Terzieva R, Herbst E (2000) The possibility of nitrogen isotopic fractionation in interstellar clouds. Mon Not Royal Astron Soc 317:563-568 Thiemens MH, Clayton RN (1980) Ancient solar wind in lunar microbreccias. Earth Planet Sci Lett 47:34-42 Thiemens MH, Heidenreich J E III (1983) The mass-independent fractionation of oxygen — a novel isotope effect and its possible cosmochemical implications. Science 2 1 9 : 1 0 7 3 - 1 0 7 5 Urey HC (1947) The thermodynamics of isotopic substances. J Chem Soc (London), 562-581 van DishoeckEF, Black JH (1988) The photodissociation and chemistry of interstellar CO. Astrophys J 334:771802 Warin S, Benayoun J J , Viala Y P (1996) Photodissociation and rotational excitation of interstellar CO. Astron Astrophys 308:533-564 Wasserburg GJ, LeeT, Papanastassiou DA (1977) Correlated O andMg isotopic anomalies inAllende inclusions: II Magnesium. Geophys Res Lett 4 : 2 9 9 - 3 0 2 Wiechert UH, Halliday AN, Palme H, Rumble D (2004) Oxygen isotope evidence for rapid mixing of the HED meteorite parent body. Earth Planet Sci Lett 2 2 1 : 3 7 3 - 3 8 2 Young ED, Ash RD, England P, Rumble D (1999) Fluid flow in chondritic parent bodies: Deciphering the compositions of planetesimals. Science 2 8 6 : 1 3 3 1 - 1 3 3 5 Yurimoto H, Kuramoto K (2004) Molecular cloud origin for the oxygen isotope heterogeneity in the solar system. Science 3 0 5 : 1 7 6 3 - 1 7 6 6

3

Reviews in Mineralogy & Geochemistry Vol. 68, pp. 15-30, 2008 Copyright © Mineralogical Society of America

Abundance, Notation, and Fractionation of Light Stable Isotopes Robert E. Criss Washington University St. Louis, Missouri 63130, U.S.A. criss @ wustl. edu

James Farquhar University of Maryland College Park, Maryland 20742, U.S.A. jfarquha @essic. umd. edu

ABSTRACT Stable isotopes have become an essential tool to characterize and understand terrestrial and extraterrestrial matter. This chapter will briefly review the abundances of important light stable isotopes, demonstrate the link between abundance and atomic weight, introduce the notations and diagrams that are commonly used to report isotopic measurements, describe and partially explain the types of fractionation effects known to occur in nature, and direct the reader to more comprehensive sources of information on each subject. The special techniques needed to make accurate isotopic measurements gave rise to special notation for reporting stable isotope data, and these notations in turn gave rise to special diagrams that emphasize compositional differences and facilitate interpretation. Fundamental definitions are the isotope ratio R, representing the ratio of the abundance of a heavy isotope to that of a lighter, typically much more common isotope, and the isotopic fractionation factor a , representing the quotient RA/RB of the isotope ratios of two substances A and B. Under equilibrium conditions, l n a can theoretically vary linearly with 1 IT at low temperatures or with 1 IT 2 at high temperatures, forming the basis for a standard graph. For practical reasons the ratio R is difficult to measure and inconvenient to report, so stable isotope abundances are usually reported as delta values (5-values) that describe their deviations from a defined "standard" material. Thus, the most important diagram for data interpretation is the "5-5 plot" where the 5-values of two coexisting phases are simply plotted against each other. In systems where two different heavy isotopes exist, two different delta values may be defined, each normalizing the abundance of one of the heavy isotopes to the common light isotope. In such cases, a very important diagram called the "three isotope" plot involves simply plotting these two different 5-values against each other for a given material, and the slopes of data arrays on such graphs can be used to distinguish ordinary "mass-dependent" fractionation (MDF) effects from "non-mass-dependent" fractionations (NMF). Numerous algebraic convolutions of the above definitions have been made, providing special definitions that can elucidate different phenomena. The processes that govern isotope distribution have become progressively better understood, yet recent studies show that these processes are more diverse than anticipated only ten years ago.

INTRODUCTION Isotopic variations in light elements, particularly hydrogen, carbon, nitrogen, oxygen and sulfur (HCNOS), provide key information on planetary formation and on the evolution and interactions of their lithosphères, atmospheres and hydrospheres. The atomic abundances of 1529-6466/08/0068-0003505.00

DOI: 10.2138/rmg.2008.68.3

16

Criss & Farquhar

the isotopes of interest are only a few percent or less, and their absolute concentrations are difficult to accurately measure given their typically small natural variations. Accordingly, a "delta" (8) notation was devised long ago to report isotopic abundances as per mil (%o), or parts per thousand, deviations from defined isotopic standards, as this notation exploits comparisons that can be precisely made with Nier-type mass spectrometers modified for this purpose (McKinney et al. 1950). In addition, a notation termed "Big-delta" (A) is commonly used both to report the simple difference in 8-values between two phases, and a similar notation, termed "Cap-delta," is used to report the deviation of samples from a defined reference line on a "three isotope plot," so attention to detail is required. In systems having multiple phases or components, the isotopes may "fractionate," or differ in relative abundance, among those parts. The isotopic fractionation factor is analogous to a distribution coefficient, and it may reflect either equilibrium or disequilibrium processes or both. Normal fractionation processes depend critically on the vibrational frequencies, and hence on isotopic masses, so these effects are called "mass-dependent fractionations," or MDF. Normal equilibrium fractionations among simple gases may be calculated from statistical thermodynamics (Urey 1947; Richet et al. 1977), but estimates for complex substances and for kinetic processes are more difficult to make. More recently, several "non-mass-dependent fractionation" (NMF) processes have been found to occur among rarefied gas molecules, which in special cases can be transferred to other molecules and even to solid phases. Processes that can produce NMF are discussed in detail in the chapter by Young et al. (2008), and include photo-dissociation in cases where the symmetry or atmospheric optical opacity differs for a given molecule and its various, isotopically-substituted forms ("isotopologues," Fig. 1). Isotopomers, molecules with the same isotopic composition, but with a different arrangement (e.g., the linear molecules 14 N 15 N le O and 15 N 14 N le O) are also subject to small differences in their chemical behavior and have become a focus of recent research. This chapter will briefly review the abundances of important light stable isotopes, demonstrate the link between isotopic abundances and elemental atomic weight, introduce the notations and diagrams that are commonly used to report isotopic measurements, describe and partially explain the types of fractionation effects known to occur in nature, and direct the reader to more comprehensive sources of information on each subject. Subsequent chapters in the volume will apply this framework to several different isotope systems in diverse natural settings.

ISOTOPIC ABUNDANCES AND ATOMIC WEIGHTS Each chemical element is an assemblage of different types of atoms, or nuclides, that share the same number "Z" of protons but can differ in the number "N" of neutrons. The mass number "A," representing the simple sum of Z plus N, is an integer that is close to the actual, non-integral atomic weight of the nuclide of interest. Isotopes of different elements are signified by a standard notation that is the chemical symbol for the element preceded by a superscript indicating the mass number. For example, 12C or "carbon-12" is the most common carbon atom, 13C is a stable isotope called "carbon-13," and 14C, the most important radioisotope, is commonly called "carbon-14" or "radiocarbon." All of these carbon atoms have six protons, as carbon is the sixth element in the periodic table, but they respectively have 6, 7, and 8 neutrons, a difference that affects the mass as well as the nuclear character and stability of each nuclide. Atomic weights are reported in atomic mass units, abbreviated "amu" or sometimes simply "u". The amu is defined as exactly 1/12 of the mass of the carbon-12 nuclide, itself defined to have a rest mass of 12.000000 amu. Atoms are tiny, so the amu is a small unit, only 1.66054 x 10~27 kg. Nuclide masses are routinely reported to great precision, because isotopes

Abundance,

Notation,

Fractionation

of Light Stable Isotopes

17

STABLE ISOTOPOLOGUES OF WATER 16

16

'H

17

0

0

17

'H

18

0

0

V

1 8.01 0 6 amu A b = 9 9 7 , 3 0 0 ppm 1 9 . 0 1 4 8 amu A b = 4 0 0 ppm P°=23.641 torr P°=23.756 torr

16

W

17

0

1

o

19.01 6 8 amu P°=22.01 torr

A b = 3 0 0 ppm

16

2

2

h.16O

2 0 . 0 2 3 1 amu A b = 0 . 0 2 2 ppm P°=20.54 torr

18

0

W

Ab=.12ppm

o

2 1 . 0 2 1 1 amu

18

0

2

h217o

2 1 . 0 2 7 4 amu A b = 1 0" 5 ppm

A b = 2 , 0 0 0 ppm

P°=23.535 torr

hzh17o

17

0

0

2 0 . 0 1 4 8 amu

0

2 0 . 0 2 1 1 amu

8

0

A b = 0.6 ppm

0

h18O

2 2 . 0 2 7 4 amu

A b . = 5 * l 0"" ppm

Figure 1. The water molecule has nine stable isotologues, all with different molecular weights and relative abundances (Ab), as indicated. Physical properties also vary among these molecules, as demonstrated by the vapor pressures at 25 °C (P°) that are provided or estimated for the most important isotopologues.

of different elements can share identical mass n u m b e r s (e.g., 4 0 K , 4 0 A r , 4 0 Ca), but in detail their masses differ slightly. In addition, tiny mass differences a m o n g nuclides can translate into t r e m e n d o u s energy differences according to Einstein's formula, E = mc 2 . T h e masses and relative abundances of the stable H C N O S isotopes are given in Table 1. In addition, the approximate atomic weight of each element, representing the weighted average of the constituent isotopes in an average sample, is given in bold. T h e latter is the n u m b e r reported in an ordinary periodic table, and is simply the sum of the atomic abundance (AZ>;) of each isotope multiplied by its atomic mass (Wtj), i.e.: Element Atomic W e i g h t = ^ j A b j W t j

(1)

i

For example, the atomic weight of 12.011 amu for the element carbon is calculated as follows: Carbon A t o m i c Weight = 0 . 9 8 9 x 1 2 . 0 0 0 0 + 0 . 0 1 1 x 1 3 . 0 0 3 3 5

(2)

18

Criss &

Farquhar

Table 1. Stable isotopes of the HCNOS elements. Element

Atomic Weight (amu)

Isotope

Hydrogen (Z=l)

Abundance (atom %)

1.0079 'H (Protium)

1.007825

99.985

2

2.014102

0.015

H (D, or Deuterium)

Carbon (1=6)

12.011 12

C

13

C

Nitrogen (1=7)

12.00000

98.90

13.00335

1.10

14.0067 14N

14.003074

99.63

15N

15.000109

0.37

Oxygen (Z=8)

15.9994 160

15.994915

99.76

17

16.999131

0.04

17.999160

0.20

o 18Q Sulfur (Z=16)

32.07 32

s

31.9720705

95.02

33

s

32.9714583

0.75

34S

33.9678665

4.21

36S

35.9670808

0.02

(Source: Walker et al. 1989)

Note that the atomic weights of the elements are not intrinsic, as they vary slightly from sample to sample depending upon the exact atomic proportions of the constituent isotopes. In contrast, the atomic weights of the individual isotopes are intrinsic and invariant.

NOTATION Isotope ratios The isotope ratio R, here representing the simple quotient of an isotope of interest normalized to the most abundant nuclide of a given element (e.g., 2 H/ 1 H; 13 C/ 12 C; 1 8 0/ 1 6 0, etc.) is the fundamental variable of stable isotope geochemistry. Note the convention that R depicts the quotient of the heavy isotope abundance over the light isotope abundance, which is invariably used for the HCNOS isotopes, and is recommended for stable isotope studies of many other elements that are now being studied (Johnson et al. 2004). The ratio R is the variable of choice for understanding the processes that control isotope distribution; moreover, many needless approximations are avoided if R is used as the starting point in mathematical derivations. However, for numerous reasons R is not a convenient means for reporting natural variations of most light stable isotopes. First, Table 1 shows that the most common nuclide of each of the HCNOS elements constitutes more than 95% of the constituent atoms, so R is a small number, typically between 0.05 and 0.0001. Second, in most natural samples, the variations among the stable isotopes are rather small, and are difficult to measure as absolute ratios.

Abundance,

Notation,

Fractionation

of Light Stable

Isotopes

19

8-values For all of the above reasons, a special method was developed long ago to report variations in the abundances of light stable isotopes. This method cleverly exploits the fact that a mass spectrometer can be used to determine the difference between the isotope ratios of two substances much more precisely than the absolute ratio of each individual substance. In particular, McKinney et al. (1950) invented "8-values" to report the abundance ratio of a measured sample, R„ as the normalized deviation from the abundance ratio, Rsid, of a defined isotopic standard, such that: Ô =1000

/ R

, ~R

X

Rstd

(3)

Ordinarily, the delta symbol is followed not by a subscript x but rather by notation indicating the isotope ratio of interest, i.e., 8D or 8 2 H for the 2H/1H ratio, 8 13 C for the 13C/12C ratio, etc. The factor of 1000 magnifies these small deviations into convenient numbers, called per mil (%o) differences, where l%c is one tenth of one percent. For most natural samples, the variations in isotope abundances of the HCNOS elements are a few per mil to tens of per mil, but larger variations occur, particularly for hydrogen or for certain extraterrestrial materials. Mass spectrometers routinely measure isotopic differences to better than ± l%c for H and better than ± 0.1%c for CNOS. The 8 notation requires comparison to a defined isotopic standard for each element of interest. Samples of many different substances have been used as isotopic standards over the years, especially for oxygen. O'Neil (1986a) provides a concise summary of the most important HCNOS isotopic standards and the means to interrelate them. Isotopic fractionation factor The isotopic fractionation factor " a " represents the partitioning of isotopes between two phases, and is analogous to a geochemical partition coefficient. This factor is most commonly used to represent the theoretical equilibrium condition between the phases, but fractionation factors may be used to quantify a non-equilibrium condition or process, or simply used to represent the measured isotopic difference between the phases. The fractionation factor is directly defined as the simple quotient of the isotope ratios R A and RB of two coexisting phases A and B: a A . B = R A /RB

(4)

Equation (3) can be used to translate this definition in terms of 8 values: O-a-B = (1000 + 8 a ) / (1000 + 8 b )

(5)

The value calculated for a is independent of the isotopic standard chosen to report the 8 values of samples A and B in Equation (5), but of course both samples must be reported as per mil deviations from the same standard. Because the 8-values of most natural samples differ by only a few per mil or so, most isotopic fractionation factors are close to unity, typically between 0.95 and 1.10. Big delta and related approximations The symbol A, called "Big delta," is commonly used to represent the simple isotopic difference between two phases A and B: AA-B

= 8 a - 8b

(6)

Big delta is commonly called the "isotopic fractionation" between A and B, because of the following approximations that can be directly derived from Equation (5) in the typical case where a is close to unity:

20

Criss & Farquhar 1000 (aA-B - 1 ) ~ 8 a - SB

(7a)

1000 In aA-B « 5 a - 8 b

(7b)

or Thus, the simple difference in the 8-values of two coexisting substances can be conveniently related to the effective fractionation factor to reasonably good accuracy, and rapid estimates of expected isotopic differences can be made for a given fractionation factor. Of course, use of Equation (5) will avoid any inaccuracies. Capital delta The symbol A a , called "Capital delta" is used to describe the difference between the isotopic composition of a substance (A) and a reference "mass-dependent" fractionation line, usually the terrestrial fractionation line (see below). As originally defined, "Capital delta" is similar to "Big delta" because it represented the simple difference between the "8-values" for the rarest isotopes in a given sample and the associated reference value on the identified massdependent fractionation (MDF) line. The MDF line for oxygen can be approximated by (e.g., Clayton et al. 1973): §mfl » 0 . 5 2 5 ^

(8a)

and one of the corresponding definitions for "Capital delta" is given as: ^a

=

^a - 0 . 5 8 ^

(8b)

Another definition for "Capital delta" is formulated using an alternative definition of the MDF line for oxygen: §MFL= 1000X
15N + 4 He quickly depletes the initial l s O and involves O in the main CNO cycling. As time progresses in the calculation, the 1 7 0 abundance rises due to proton capture onto l e O and achieves a maximum enrichment of about a factor of ten relative to its initial abundance. After about 104 years, however, the 1 7 0 abundance attains a steady state as destruction via 17 0(p, 4 He) 14 N balances the production from l e O. Finally, after 106 years, the full CNO bi-cycle achieves a steady state. The ! H converts into 4 He and the oxygen isotopes convert into 14N. After complete CNO burning, 4 He and 14N are enriched while 12C and 16.17.lsO are all depleted (Fig. 1). Nevertheless, if the CNO cycling is not complete or the burning happens at a lower temperature, the matter may be enriched in 1 7 0 and depleted in 16>180, as is the case in Figure IB at a time of 106 yr. This is typically the case in the envelopes of stars that have experienced dredge up (mixing) of matter from a hydrogen burning shell. Helium burning. After a star has burned its hydrogen into helium, the next available fuel is the 4 He. Due to the lack of stable isotopes with mass number five and eight, the helium burning proceeds via the triple-alpha process, which may be viewed as the reaction 4 He + 4 He + 4 He —> 12 C + y. Helium burning typically occurs at temperatures of ~l-3xl0 8 K and densities near 1000 g/cm 3 . Figure 2 shows the evolution of mass fractions relative to solar values in a helium burning calculation at a temperature of 2.5x10 s K and a density of 1000 g/cm 3 and using as the initial abundances the final yields from the previous hydrogen burning calculation. In the initial stages of helium burning, the abundant 14N captures 4 He to produce 18F, which decays to l s O, thereby strongly enriching l s O relative to l e O and 1 7 0. The l s O itself then captures another 4 He, which depletes the l s O and creates 22 Ne. At the high temperature of the calculation in Figure 2, this conversion of 14N into 22 Ne occurs within the first 0.1 yr. As this is occurring, the triple-alpha reaction is also converting 4 He into 12C. Beginning at about one year, the 12C becomes sufficiently abundant that it can capture 4 He to become le O. This strongly enriches l e O relative to the other oxygen isotopes. The 12 C/ le O ratio resulting from helium burning determines the nature of the subsequent carbon burning in the star, which, in turn, determines the whole subsequent shell structure of the star (e.g., El Eid et al. 2004). For this reason, the reaction 12 C( 4 He,y) le O, which is not yet fully characterized in the laboratory, is the subject of intense experimental study. As a final point, it is worth noting that the reaction l e O + 4 He —> 20 Ne + y does not occur at helium-burning temperatures because of the lack of an appropriate resonance in 20Ne. A dominant mechanism for two nuclei to react is to proceed through formation of a compound nucleus, which is, in effect, a resonance in the scattering state of the two interacting nuclei. The compound nucleus subsequently breaks apart into different nuclei or de-excites to the product nucleus ground state. For example, when 12C and 4 He combine, they can form an excited state (that is, a resonance) of l e O if the energy and the spin and parity changes in the interaction are appropriate. Most of the time, the excited l e O nucleus will break apart into 4 He and 12C again. In some cases, however, the excited l e O decays by photon emission to the ground state. In the case of 20 Ne, the typical interaction energies of l e O and 4 He in helium burning are appropriate to form the 4.969 MeV state in 20 Ne, which would make an ideal compound nucleus, but the spin and parity changes are not right, so the reaction does not occur (e.g., Clayton 1968). Alpha capture therefore ceases at l e O, and, as a consequence l e O is, in fact, the dominant product of helium burning. From Figure 2, it is clear that there is a phase in helium burning in which l s O is enhanced while l e O and 1 7 0 are depleted. As the burning progresses, however, the l s O converts to 22 Ne

Nucleosynthesis

& Chemical Evolution of Oxygen

35

Figure 2. Evolution of the mass fractions of the indicated species during helium burning at a temperature of 2 . 5 x l 0 8 K and a density of l.OxlO 3 g/cm 3 .

and l e O becomes the dominant oxygen isotope. Because the l e O so greatly dominates the other oxygen isotopes after helium burning, we will henceforth neglect these minor isotopes in the subsequent discussion. Carbon burning. The next stellar burning stage is carbon burning. Carbon burning typically occurs at temperatures near 9x10 s K and densities near 105 g/cm 3 . It predominantly converts 12 C into 20 Ne and 24 Mg via the reactions 12C + 12C H> 20 Ne + 4 He and 12C + 12C H> 24 Mg + y. At carbon-burning temperatures, there are resonances in the 16O(4He,y)20Ne reaction available that allow the reaction to proceed efficiently. This means that the l e O left after helium burning is depleted somewhat during carbon burning due to capture of 4 He produced by the main carbonburning reactions. Neon burning. After carbon burning, a star burns neon. This effectively occurs in a twostep process. The first reaction is 20 Ne + y —> l e O + 4 He. This reaction is endothermic. The 4 He produced, however, then captures on another 20 Ne to produce 24 Mg. This latter reaction is sufficiently exothermic that the net reaction, 20 Ne + 20 Ne —> l e O + 24 Mg, is exothermic. Clearly, neon burning replenishes some of the l e O depleted during carbon burning.

36

Meyer, Nittler, Nguyen,

Messenger

Oxygen and silicon burning. Once neon burning is complete, the next burning stage is that of oxygen. The dominant reaction is l e O + l e O —> 28Si + 4 He. This phase clearly depletes the oxygen. After oxygen burning, the star will burn 28Si into 56Fe. The effective reaction is 28 Si + 28Si —> 56Ni + y followed by decay of 56Ni to 56Fe; however, the burning actually occurs by burning through quasi-equilibrium clusters (e.g., Woosley et al. 1973). Little oxygen is produced in this 28Si burning (e.g., Bodansky et al. 1968). Stellar explosion. Stars that are able to evolve all the way to silicon burning develop iron cores. At that point, the nuclei are in the state with the highest nuclear binding; hence, any rearrangement of their abundances cannot release energy. This means that the star can no longer burn fuel to maintain its pressure against gravity and the stellar core collapses homologously. The inner core collapses subsonically. Once it attains nuclear matter densities, the collapse halts and the core bounces. The outer core, however, collapses supersonically and therefore does not receive the signal to stop collapsing. It crashes onto the already collapsed inner core. This highly non-adiabatic effect generates a shock wave that propagates out through the outer layers of the star. As the shock passes through the stellar layers, it heats and compresses them. This causes further burning (known as explosive burning). The energy imparted to the outer layers by the shock eventually expels them and injects them into interstellar medium. This dramatic stellar death is known as a supernova. In particular, it is a "core-collapse" supernova since the energy powering the explosion arose from the release of gravitational binding energy in the collapse of the stellar core. For a review, see, for example, Woosley et al. (2002). The modification of the oxygen isotopes due to explosive burning is largely confined to the inner layers of the star. In particular, explosive oxygen burning will deplete the l e O in the le O-rich layers and explosive neon burning will enhance the l e O in the 20Ne-rich layers. By the time the shock reaches the 12C-rich and 4 He-rich regions, however, the post-shock temperatures and densities do not lead to significant changes in the pre-supernova oxygen abundances. Analysis of the oxygen yields from massive stars Armed with an understanding of the nucleosynthesis in the mainline burning stages of stellar evolution, we may now analyze the yields from massive stars—those with masses greater than about ten solar masses. Such stars are able to burn their nuclei all the way to iron. Stars with mass less than about ten solar masses develop cores that are supported by electron degeneracy, which halts their evolution before the advanced nuclear burning stages can occur. For a discussion of degeneracy, see the section on low-mass stars below. To analyze the yields from massive stars, we consider a single model. While this is only one particular model, its yields are fairly representative of the ejecta from any star more than ten times the mass of the Sun. Chemical evolution models using detailed stellar model yields have long shown that the Galaxy's supply of l e O is dominantly from massive stars (e.g., Tinsley 1980). Because l e O so dominates the oxygen abundance, massive stars are thus nearly the sole contributors to the Galaxy's inventory of this important element even though they compose only a few percent of the stars born in any stellar generation. Massive stars also dominate the galactic production of l s O (e.g., Timmes et al. 1995). Low-mass stars only dominate the galactic synthesis of the low-abundance isotope 17 0. The model presented here is that of Meyer (2005), using the stellar evolution code of The et al. (2000) and El Eid et al. (2004) but with coupled nuclear burning and convection using routines written by Jordan et al. (2005), for a star that began with a mass 25 times greater than that of the Sun. By the end of its life, it had lost 3.4 solar masses of material due to stellar winds. The model ran through silicon burning. The collapse of the core and the explosion of the star were then simulated by numerically increasing the energy in the core of the stellar model. This led to the generation of an outwardly propagating shock wave in the model. Shock propagation and the concomitant explosive burning were tracked with the appropriate computer codes.

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The final yields (after stellar explosion) of the isotopes of oxygen relative to their initial abundances are shown in Figure 3. The abscissa is the so-called Lagrangian mass coordinate, M r . The value of M r indicates the mass contained within a given spherical shell in the star. For example, M r = 0 is the center of the star. M r = M, where M is the total mass of the star, is the star's surface. M r = 10 solar masses locates a spherical shell that contains 10 solar masses inside. The Lagrangian mass coordinate is convenient because, although the radius of a shell expands or contracts as the star evolves, the mass it contains does not. As Figure 3 shows, the abundances of the oxygen isotopes vary dramatically inside the stellar ejecta. Particularly noteworthy is the fact that abundances are uniform over certain mass ranges. This indicates mixing within the stellar regions. For example, the abundances of le O, 17 0, and l s O are uniform from M r ~7.8 to 21.6 solar masses. This represents the outer envelope of the star, which stretches from what was the hydrogen-burning shell to the stellar surface. The strong convection in the envelope prior to the explosion homogenizes the abundances. Similarly, the helium-burning shell in the star extends from about M r ~ 5.5 to 7.3 solar masses. It is also worth noting that stellar burning proceeds through a sequence of core and shell burning. In particular, a given burning stage commences in the center of the star since the temperature is generally highest there. If the energy release is strong enough, the burning drives convection. This draws fuel in from throughout the convective core. Once the burning is complete, the fuel for that burning stage is exhausted throughout the convective core. The star will then contract until the next burning stage begins. Shell burning occurs when the region outside the formerly convective core contracts and heats to the point that the appropriate nuclear fuel can ignite. With these preliminaries in mind, we may trace a star's evolution from the oxygen isotopes in the star. It is convenient to consider major zones in the stellar ejecta and use the terminology from Meyer et al. (1995) to label these zones by the most dominant elemental abundances, working our way inward from the surface of the star. The H envelope stretches from M r ~ 7.8 to 21.6 solar masses. As the star evolves past the main sequence phase of its life, convection begins in the envelope and reaches down to matter that had experienced partial burning of hydrogen 10 4

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Figure 3. Final mass fractions (after stellar explosion) of the isotopes of oxygen relative to solar in a onedimensional stellar model of an initially 25 solar mass star as a function of Lagrangian mass coordinate M r . The final stellar mass was 21.6 solar masses, and the abundances are uniform from M r = 7.8 to 21.6 solar masses (the surface).

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by the CNO cycle. As a consequence, the abundances throughout the envelope look similar to the early stages of CNO burning in Figure 1, with modest enrichments in 14N and 1 7 0. It is also noteworthy that mass loss begins as envelope convection commences and carries away some products of CNO burning prior to the explosion. The He/N zone stretches from M r ~ 7.3 to 7.8 solar masses. It is a thin, radiative (nonconvective) shell that was originally part of the convective hydrogen-burning core. Since it largely completed hydrogen burning at the time of stellar explosion, it converted most of its oxygen into 14N, in analogy with the late stages of the single-zone calculation shown in Figure 1, in which all three oxygen isotopes are depleted relative to their initial abundances. The He/C zone stretches from M r ~ 5.5 to 7.3 solar masses. This is the part of the star that had experienced convective core hydrogen burning and was burning helium in a convective shell at the time of stellar explosion. In this zone, the l s O is enriched, and the abundances resemble those in the single-zone calculation in the early stages of the burning. It is worth emphasizing that, because of the convective burning, a 14N atom that captures a 4 He atom in the burning region is likely to mix out into the non-burning parts of the shell before capturing another 4 He. This means that the l s O has to cycle back into the burning region before capturing another 4 He to become 22Ne, which lengthens the lifetime of the l s O compared to that in the single-zone calculation. This, in turn, explains why the convective helium shell can be enriched in both 12C and ls O, even though Figure 2 might suggest that 12C builds up only after the l s O is declining. The O/C zone contains material from M r ~4.5 to 5.5 solar masses. This is matter that was part of the convective helium-burning core but did not partake in convective carbon shell burning. Since this matter completed helium burning, the oxygen here is dominated by le O. Nevertheless, in this particular stellar model, some 1 7 0 and l s O are present due to non-convective burning in this region just prior to the stellar explosion. The O/Ne zone, ranging from M r ~3 to 4.5 solar masses, is a region of the star that experienced convective carbon shell burning. As discussed above, at the temperatures of carbon burning, l e O can capture 4 He nuclei released by the carbon burning reactions. This thereby depletes the l e O left over from the previous convective core helium burning. As is evident from Figure 3, however, this depletion is fairly slight, and the l e O still strongly dominates the oxygen abundances. In the present stellar model, the O/Si zone ranges from M r ~ 1.8 to 3 solar masses. This is the region of the star that experienced neon shell burning. As discussed above, the effective neonburning reaction is 20Ne + 20Ne —> 24Mg + le O, which increases the l e O abundance slightly. Finally, inside M r ~ 1.8 solar masses in the present model, the star burns its l e O into 28Si and heavier isotopes both in pre-supernova and supernova nucleosynthesis. These regions of the stellar ejecta, the Si/S and Ni zones, are devoid of any oxygen, except for trace amounts produced during the stellar explosion. In broad summary, the ejecta from a massive star is characterized by an 17 0-rich envelope, which contains most of the stellar ejecta, a narrow helium shell enriched in ls O, and the inner regions, which are strongly enriched in le O. It is important to emphasize that, although it is true that massive stars can eject several solar masses of le O, the massive star ejects all three O isotopes, and the bulk ejecta are not necessarily le O-rich. Low-mass stars Apart from mass loss during pre-supernova evolution, which can be significant, massive stars tend to lose all their mass in a single event, namely, the supernova explosion. In contrast, low-mass stars, those with masses less than roughly eight solar masses, lose their mass in periodic episodes during their post-main sequence phases. It is thus necessary to follow the evolution of the star in detail to understand the ejected oxygen yields. We are brief in our

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discussion; for more detailed reviews of the evolution of, and mass loss from, low-mass stars, please see Busso et al. (1999), Herwig (2005), or Straniero et al. (2006). Before discussing the evolution of low-mass stars, it is important to consider the concept of degeneracy in gases. The principal subatomic components of stars are electrons, protons, and neutrons, which all have intrinsic angular momentum quantum number ("spin") Vi, which means they are fermions. By the Pauli Exclusion Principle, no two fermions can occupy the same quantum state. This is not a problem at low densities when there are a vastly greater number of states energetically available than electrons present. When the density increases, however, the number of energetically available states does not greatly exceed the number of electrons. This is the condition of degeneracy, and the result is that electrons may find themselves trapped in highenergy states because lower-energy states are already occupied by other electrons. Because these high-energy electrons carry momentum, they contribute significantly to the pressure of the gas. The remarkable consequence is that even a zero-temperature degenerate gas can have a high pressure. To put this in perspective, consider the pressure at the center of the Sun. The temperature ( I ) = 1.5xl0 7 Kelvins (K) and the mass density is about 150 g/cm 3 , which corresponds to a particle density of about n = 2 x l 0 2 5 particles/cm3 in the fully ionized plasma. This gives rise to a gas pressure P = w k B r = about 4 x l 0 1 6 dynes/cm2, where k B is Boltzmann's constant. There are essentially no degenerate electrons in the center of the Sun, so the degeneracy pressure is zero. In the evolved core of a low-mass star, by contrast, the temperature is roughly 10 s K and the mass density is roughly 10 6 g/cm 3 . The gas pressure is thus about P = 8 x l 0 2 1 dynes/cm2. At the same conditions, however, the degeneracy pressure is about 2 x l 0 2 2 dynes/cm2. The degeneracy pressure dominates that due to the thermal motions of the gas and can be great enough to support the star against contraction due to gravity. A final introductory point regards burning under degenerate conditions. When nuclear burning occurs, nuclear binding energy is released, which raises the temperature of the local environment. Under non-degenerate conditions, the pressure rises. This causes gas in the local environment to expand, and thus cool. This feedback mechanism can thus lead to steady, stable burning. Under degenerate conditions, however, the pressure is dominated by the degenerate electrons and is thus largely insensitive to the temperature. When burning occurs, the temperature rises but the pressure does not. The rising temperature causes the nuclear burning to proceed at a more rapid rate, which leads to a further increase in temperature and even faster burning. This is a runaway condition that stops only when the temperature is so high that the degeneracy is lifted and the normal gas pressure again dominates. Such violent burning can lead to nova outbursts or to Type la supernova explosions (see below). The typical evolution of a low-mass star begins with core hydrogen burning during main sequence evolution. Once the hydrogen is exhausted, the center of the star contracts, and hydrogen shell burning commences on top of the hydrogen-exhausted core. The star begins to ascend the giant branch in the Hertzsprung-Russell diagram. As the star ascends the giant branch, it grows in radius, and the outer envelope expands and cools. This increases the opacity, and the envelope becomes convective. The convective envelope grows in extent and eventually "dredges up" material that had experienced CNO processing (the "first dredge up"). This enriches the envelope material in 1 7 0 and depletes it slightly in l e O and l s O. The stellar core continues to contract until helium burning ignites. This occurs either under electron-degenerate conditions as a "flash" for stars less than roughly 2 solar masses or under non-degenerate conditions and, hence, more quietly for stars of mass greater than ~2 solar masses. After the core has exhausted its helium, it contracts again and ascends the "asymptotic" giant branch (AGB). At this time, stars more massive than ~4 solar masses experience "second dredge up," which brings products of hydrogen shell and helium core burning to the surface. The CNO burning products tend to dominate so that, again, the 1 7 0 is enriched.

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By this point, the star's structure consists of an inert, degenerate C/O core, surrounded by a helium-burning shell, surrounded by a hydrogen-burning shell, surrounded by the stellar envelope. The hydrogen- and helium-burning shells burn alternately in a complicated choreography. Most of the time, hydrogen shell burning occurs quiescently. Eventually, however, the temperature and density rise in the helium-rich region between the C/O core and the hydrogen shell. The helium ignites in a thermal pulse, which drives convection within the helium-rich zone and extinguishes the hydrogen-burning shell. The convective envelope reaches down into matter that has experienced helium burning and dredges it up ("third dredge up") to the surface, thereby enriching it in the helium-burning products, including l e O and l s O, as well as 1 7 0 from the hydrogen shell. Once the helium burning has ceased, the newly produced carbon and oxygen settle onto the core, and the hydrogen-burning shell reignites. The thermal pulses help drive the periodic mass loss from the AGB stars. As a typical AGB star progresses through multiple thermal pulses, the C/O ratio in the envelope increases as the helium shell nucleosynthesis and third dredge up tend to preferentially add 12C over l e O. Because the oxygen abundance initially dominated that of carbon, the star began with a C/O abundance ratio less than one. The preferential addition of 12C increases the C/O ratio until it becomes a "carbon" star (C/O > 1). Eventually the pulsations drive off most the star's envelope, and only a degenerate C/O white dwarf star with a thin hydrogen or helium atmosphere is left behind. The thin atmosphere is due to the remnants of the star's original envelope, or possibly to accretion from the interstellar medium. So-called "naked" white dwarf stars that show no surface hydrogen or helium have also been observed (Werner et al. 2004). Naked white dwarf stars have clearly lost their entire envelopes. In white dwarf stars, degeneracy pressure is sufficient to hold the star up against gravity, even at zero temperature, so these burnt-out cinders slowly radiate and cool with time. It is interesting that, given good estimates of bare white dwarf cooling rates, one may estimate the age of the galactic disk from the populations of white dwarf stars as a function of their luminosity (e.g., Oswalt etal. 1996). While this picture of stellar evolution is largely successful in explaining surface abundances of low-mass stars, certain observational puzzles present themselves. First, the surface abundance ratio of 12C/13C observed for many red-giant branch stars is lower than predicted by the models. This suggests extra mixing below the conventional convective envelope. Such mixing would bring envelope material into the vicinity of the H-burning shell and allow for some nuclear processing, and then transport it back to the surface. This has been termed "cool bottom processing" (Boothroyd et al. 1995). In addition to helping to explain the carbon abundance puzzle in low-mass red-giant branch stars, cool bottom processing can also explain the low 18Q/16Q r a t j 0 j n cer ^ain AGB star atmospheres and presolar oxide grains (Wasserburg et al. 1995; Boothroyd and Sackmann 1999; Nollett et al. 2003) and the carbon and nitrogen isotopic compositions of many presolar SiC grains (Huss et al. 1997; Nollett et al. 2003). Another puzzle is the large surface 1 7 0 enrichments in the handful of observations of J type carbon stars which cannot be explained by standard low-mass-star evolution. The answer may lie in so-called "hot bottom burning". In this scenario, the convective envelope of a star with mass greater than ~5 solar masses extends down to the H-burning shell so that the envelope material itself experiences H burning, heavily depleting l s O and enriching 1 7 0 at the stellar surface (e.g., Boothroyd et al. 1995; Lattanzio et al. 1997). As we shall see in the section "Oxygen in Presolar Grains," one generally invokes hot bottom burning and cool bottom processing to explain the oxygen isotopic ratios in certain presolar oxide and silicate grains as well as in stellar atmospheres. For completeness, it is worth noting that stars with masses in the range of ~8-10 solar masses do not develop degenerate C/O cores; they proceed to core carbon burning. They then eject their envelopes in a manner similar to lower mass stars and leave behind O/Ne/Mg white dwarf stars. As we have seen, the pressure in high-mass stars (stars with mass greater than about

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41

ten times that of the Sun) is never dominated by electron degeneracy. These stars therefore proceed all the way through silicon burning.

Novae and Type la supernovae Novae are thermonuclear explosions that occur when material from a binary stellar companion (main sequence or red giant star) is gradually accreted onto a white dwarf. As the density increases in the outer layers of the white dwarf due to ongoing accretion, nuclear reactions begin to occur. Because these reactions take place under degenerate conditions, a relatively brief thermonuclear runaway occurs, resulting in the optical outburst we know as a "nova." A small amount of processed accreted matter and even some underlying white dwarf material is ejected. Then the system settles down to accrete more material until the next nova outburst. Although there are many uncertainties in nova modeling, there is reasonably good agreement between nova nucleosynthesis calculations and observed elemental abundances in novae. Novae are not believed to be significant contributors to the galactic budget of most elements; however, they are probably an important source of some rare isotopes, including 13C, 15 N, and of interest here, 1 7 0. Calculations predict ejecta with 1 7 0/ 1 6 0 ratios 25-2000 times the solar ratio (José et al. 2004). The ejecta are also typically enriched in l s O. Nevertheless, He burning in massive stars is the dominant source of the galactic l s O, and novae contribute only a small fraction to the galactic inventory. In some white dwarf-stellar companion binary systems, the central density in the accreting white dwarf can build up to such high temperatures that carbon-burning reactions can ignite. Like the nova outburst, such burning occurs under degenerate conditions but, given the higher temperature sensitivity of carbon-burning reactions compared to those of the CNO bi-cycle relevant for novae, this burning is much more violent. The result is the complete thermonuclear disruption of the white dwarf in what is most likely a Type la supernova. Because of the high temperatures attained in a Type la event, much of the matter reaches nuclear statistical equilibrium, which means that such supernovae are major contributors to the Galaxy's supply of irongroup isotopes. Nuclear statistical equilibrium does not favor oxygen, however, so the inner parts of Type la's do not synthesize much oxygen. In the outer layers of Type la supernovae for which the burning front is subsonic, considerable l e O can be produced, principally by explosive carbon and neon burning. Nevertheless, given the relative infrequency of la events, they probably only contribute at best a few percent to the Galaxy's inventory of oxygen (e.g., Thielemann et al. 1986). Confirmation of this basic picture comes from the fact that very old, low-metallicity stars show O/Fe ratios greater than the solar value (e.g., McWilliam 1997). Such stars formed early from gas that inherited the ejecta from massive stars that lived and died in our Galaxy's first few million years. Massive stars produce a higher ratio of oxygen ( l e O) to iron than that present in the Solar System, which explains the high O/Fe ratio in low-metallicity stars. Type la supernovae, on the other hand, did not start occurring until the first white dwarf stars formed, many tens to hundreds of millions of years after the Galaxy formed. Since Type la supernovae produce much iron with little oxygen, their ejecta lowered the O/Fe ratio in the interstellar medium from its high value in the early Galaxy to the solar value by the time of the Sun's birth.

CHEMICAL EVOLUTION OF THE ISOTOPES OF OXYGEN A fundamental distinction in nucleosynthesis theory is between primary and secondary isotopes, and this distinction has important consequences for the evolution of the oxygen abundances in the Galaxy. A primary isotope is one that can be produced from a star initially composed only of hydrogen and helium. A secondary isotope is one that can only be made from pre-existing seed nuclei. These pre-existing seed nuclei come from previous generations of stars, hence the name secondary.

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The oxygen isotopes provide clear examples of this distinction. l e O is a classic example of a primary isotope. A star composed initially only of the principal products of Big-Bang nucleosynthesis ( ! H and 4 He) could produce l e O. It would do this by converting its initial ! H into 4 He via the PP burning chains. When the 4 He subsequently burns by the triple-alpha process, the dominant products would be 12C (another primary isotope) and l e O. By contrast, 17 0 and l s O are both secondary isotopes, because a star initially composed only of ! H and 4 He could not form them. 1 7 0 is made by proton-capture on l e O. This occurs during hydrogen burning, and a star initially composed only of ! H and 4 He would have no l e O during its hydrogen-burning stage to capture a proton. Similarly, l s O traces its origin back to 4 He capture by 14N in the early stages of helium burning. As evident from previous discussions, CNO burning converts most of the star's initial C, N, and O isotopes into 14N because of the slowness of the 14 N(p,y) ls O reaction. A star composed initially only of ! H and 4 He would not have the C, N, or O isotopes necessary to convert into 14N and, hence, would produce no l s O. Figure 4 demonstrates the primary vs. secondary nature of the oxygen isotopes. Shown are the results from a suite of stellar evolution and nucleosynthesis calculations by Woosley and Weaver (1995) for a range of initial stellar mass and metallicity. In chemical evolution studies, metallicity is an astronomical term that refers to the mass fraction of isotopes heavier than helium. Early astronomical studies of galactic abundances and chemical evolution only distinguished between hydrogen, helium, and everything else, and the respective mass fractions were X (hydrogen), Y (helium), and Z (the "metals"). In the Solar System, the mass fraction of "metals" is Z®~0.02, with l e O accounting for half of that component. As Figure 4A shows, massive stars with a metallicity 1/10,000 as large as that in the Solar System eject as much l e O as do stars with solar or even twice solar metallicity. This confirms le O's status as a primary isotope because it shows that the l e O yield from a star is essentially independent of the star's initial abundance of elements other than hydrogen or helium. The general trend for all metallicities is that the larger the mass of the star, the more l e O it ejects. This reflects the fact that larger stars have larger helium-burning cores; thus, they produce more l e O. The exception to this trend is that some stars in the suite of models with mass greater than 30 times that of the Sun produce less than their lower-mass peers. This is not a nucleosynthetic effect; rather, it is a consequence of the fact that these models experienced significant fallback during their explosion and formed black holes. The black holes swallowed up much of the l e O, which reduced the ejected yield. Figures 4B and 4C show the corresponding ejected yields from the suite of stellar models for 1 7 0 and l s O. From these plots it is clear that the larger the initial metallicity of the star, the greater the production of 1 7 0 and l s O. It is also interesting that, for a given initial metallicity, the yield is only weakly sensitive to initial stellar mass. As we previously saw, the ejected 1 7 0 and l s O are produced in shell burning. The size of convective hydrogen and helium burning cores is strongly dependent on the stellar mass. By contrast, the sizes of convective hydrogen and helium shells are less dependent on the stellar mass. This explains the weaker dependence of 17 - 18 0 on initial stellar mass than that of le O. An important caveat is that the rates for the key proton-capture reactions 1 7 0(p,a) 1 4 N and 17 0(p,y) 18 F have been revised since the calculations shown in Figures 4B and 4C were run. In particular, the preferred values for these reaction rates have increased (Blackmon et al. 1995; Angulo et al. 1999). These increases cause greater destruction of 1 7 0 during CNO burning, thereby leading to a decrease in the expected yield of 1 7 0 from massive stars. In particular, massive star models using the new 1 7 0 proton-capture rates typically show nearly ten-fold reductions in the yield of 1 7 0 (e.g., Rauscher et al. 2002). This, in turn, suggests that other sites, particularly AGB stars with hot bottom burning and novae play a significant, if not dominant, role in the synthesis of this isotope. Nevertheless, as in the massive star models, the production of 1 7 0 is secondary in these other sites (e.g., Romano and Matteucci 2003).

Nucleosynthesis

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With these basics in mind, the galactic evolution of the abundances of the oxygen isotopes is now easy to understand. The high entropy of the Universe allowed primordial nucleosynthesis to produce only ! H, 2 H, 3 He, 4 He, and trace amounts of lithium, beryllium, and boron from the initial soup of neutrons, protons, and electrons. The initial composition of the first generation of stars, then, was devoid of C, N, and O, and as a consequence, produced little 1 7 0 and l s O but normal amounts of l e O. As the Galaxy evolved, the succeeding generations of stars formed from interstellar media that had initial compositions enriched in "metals" from previous generations. These stars were then able to produce increasingly large amounts of 1 7 0 and l s O. This behavior is evident in Figure 5, which shows a one-zone model of the evolution of the oxygen abundances in the Galaxy. Like Figure 4, Figure 5 was produced from the online CUGCE Tool at http://webnucleo.org (for a description, see Meyer et al. 2001). The model followed a single zone in the Galaxy, used the instantaneous recycling approximation, and employed Clayton's family of analytic galactic infall models (Clayton 1984). Figure 5 shows the mass fraction of oxygen isotopes in the interstellar gas, normalized to their mass fraction at the time of Solar System formation in the model. The behavior of l e O is distinctly different from that of 1 7 0 and l s O in the figure. Because l e O is a primary isotope, each generation of stars ejects about the same number of grams of l e O per gram of mass going into stars; thus, the gas becomes enriched in l e O in a linear fashion. In contrast, 1 7 0 and l s O are secondary isotopes; therefore, the number of grams of these isotopes ejected per gram going into stars increases with each generation. This gives rise to the quadratic evolution seen in Figure 5. At the time of Solar System formation, the oxygen isotopes had all reached their solar values in the Galaxy. The galactic chemical evolution presented in Figure 5 is a simplification, given the fact that the model contains only a single zone and uses the instantaneous recycling approximation—the real Galaxy is an inhomogeneous mix of different interstellar phases and stars that have finite lifetimes. Nevertheless, the general trend of increasing 1 7 0 and l s O relative to l e O in the evolution

Oxygen Evolution

t(Gyr) Figure 5. Mass fraction of the isotopes in the interstellar gas in a chemical evolution model as a function of time. The mass fractions are normalized to their values at the time of Solar System formation (roughly 9.5 Ga after Galaxy formation in the present model). On the scale of the plot, t=0 corresponds to the point in time when the Galaxy starts forming. Build up of the mass of the Galaxy occurs over ~2-3 Ga after t=0 by infall of intergalactic gas.

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of the Galaxy is quite evident from astronomical data. In particular, stars that formed recently in the local stellar neighborhood show a lower 1 6 0/ 1 7 0 ratio than the solar value, which indicates the ratio has declined over the last 4.5 Ga, as expected from galactic chemical evolution (e.g, Wilson and Rood 1994). As for l s O, the 1 6 0/ 1 8 0 ratio decreases inward from the Sun's position in the Galaxy. Since the inner part of the Galaxy is thought to be more chemically evolved than the outer regions, this supports a secondary nature of l s O. Despite this agreement between observations and galactic chemical evolution expectations for oxygen, there are two significant puzzles. The first is that the solar 1 6 0/ 1 8 0 ratio is actually less than that in today's interstellar medium in the local solar neighborhood (Wilson and Rood 1994). This is not expected, because the 1 6 0/ 1 8 0 ratio in the Galaxy should decline with time. The second puzzle is that the 1 8 0/ 1 7 0 ratio in today's interstellar medium is about 3.5, which is significantly less than the solar value of 5.2. From Figure 5, one would expect this ratio to be little changed since the time of the Sun's formation. A proposed solution to the " l s O puzzle" is that the Sun formed in an association of highmass stars, a so-called "OB association". In such a stellar association, the high-mass stars evolve, explode, and self-enrich the cluster. Since the high-mass stars preferentially eject l e O and l s O, the 1 8 0/ 1 6 0 and 1 8 0/ 1 7 0 ratios can grow over the course of a ~10-20 Myr period of the association's evolution (e.g., Prantzos et al. 1996). This means that the Sun would have formed with higher values for these ratios than the ambient interstellar medium from which the entire association itself formed. D. D. Clayton has proposed an alternative scenario. It is based on his idea (Clayton 2003) for origin of the silicon isotopes in presolar silicon carbide grains (see section "Oxygen in Presolar Grains") in which the absorption of a metal-poor satellite galaxy (like the Magellanic Clouds) by the Milky Way initiated a burst of star formation. The number of AGB stars formed during this starburst exceeded that formed by normal galactic star formation, so the former contributed the bulk of the SiC in this period. Since the parent composition of these AGB stars was a mix of the evolved Milky Way Galaxy and the metal-poor satellite, the silicon isotopic compositions of the SiC grains from these stars essentially lie on a mixing line between these two initial compositions. In this scenario, the Sun itself formed with a composition relatively close to the metal-poor satellite; thus, most SiC grains have a silicon isotopic composition that looks more evolved than solar. This scenario has interesting implications for the l s O puzzle (Clayton 2004). In particular, immediately after the merger, high-mass stars quickly evolved and enriched the interstellar medium with l e O and l s O but not 1 7 0, which had to wait for the low-mass stars to evolve. The Sun formed during this period of enriched l e O and l s O, with an 1 8 0/ 1 7 0 ratio of 5.2. Over the course of the last 4.5 Ga, however, low-mass stars formed during the starburst returned their 17 0-enriched matter and lowered the interstellar 1 8 0/ 1 7 0 ratio to its present value of 3.5. Clayton (2004) points out that a consequence of this scenario would be a correlation between 18 0/ 1 6 0 and 1 7 0/ 1 6 0 and the 30 Si/ 28 Si in presolar SiC grains. Oxygen and silicon isotopes in presolar silicates from low-mass stars would also be expected to show these correlations. The current presolar silicate data set (see section "Oxygen in Presolar Grains") does not show an unambiguous correlation between Si and O isotopes, but there are some technical difficulties with such measurements (Nguyen et al. 2007). While details of this scenario need to be worked out, it is a nice illustration of oxygen at the nexus of nuclear physics, stellar evolution, galactic astronomy, and presolar grains.

OXYGEN IN PRESOLAR GRAINS Presolar grains are rare and small (nm to 10 )im) mineral grains recovered from primitive meteorites and interplanetary dust particles (IDPs; 90% of which are believed to have formed in C-rich AGB stars and only 1% from supernovae. Some recent studies of cold dust in supernova remnants suggest that supernovae are much more prodigious dust producers than previously believed (e.g., Dunne et al. 2003). However, other studies have questioned these results (Dwek 2004; Krause et al. 2004) and the presolar grain evidence indicates that supernovae were relatively minor contributors to dust in the Galaxy at the time that the Solar System formed. Finally, we note that presolar grains are identified by virtue of having highly anomalous isotopic compositions relative to the range of materials known to have formed in the Solar System. However, this is essentially an operational definition, and analytical uncertainty plays a significant role in deciding whether any given grain is demonstrably a stellar condensate. That is, searches for presolar oxides and/or silicates generally involve the measurement of large numbers of grains, and grains whose isotopic compositions are significantly different (e.g., 3a) from the rest are defined to be presolar. Figure 7 shows the O-isotopic compositions (expressed as 8-values; see the caption or Criss and Farquhar (2008) for the definition) of many of the presolar grains. Also shown as small ellipses near the origin is the range of O isotopes measured with high precision in meteoritic materials. The grey shaded ellipse illustrates schematically the region of the plot for which analytical uncertainty on small grain measurements has excluded identification of presolar grains. Almost certainly, as analytical methods improve, this region will get smaller and grains with less extreme composition will be defined as presolar. However, it is possible that some presolar grains have isotopic compositions similar to the solar values. One source of such grains may have been other young stars in the molecular cloud from which the Solar System formed. Young stars are observed to eject prodigious amounts of materials in bipolar outflows, and the Sun's stellar neighbors should have had isotopic compositions close to solar values. In any case, we wish to emphasize that establishing the true limits (if any) of presolar grain isotopic compositions is an important topic for future research.

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1000 800

° 0 F

O

OoQ

oA* S

0

-200

. O OOcP °o »a

A _ 8o°

°

„

o %

°o

-400 —I

-400

I — I —I

i

-200 5

i A m L_ _i

1 8

0

0

/ 1 6

0

i

i

_i

200

i

i_

400

(0/oo)

Figure 7. O isotopie ratios of presolar oxides and silicates, expressed as 5-values: S'O = 10 3 x[('0/ 16 0) gIaill / ('0/ 16 0) TelTestlia i — 1 ]. Dotted lines indicate solar isotopie ratios (assumed to be terrestrial). Solid ellipses near origin indicate range of high-precision isotopie measurements of Solar System-derived meteoritic materials. Grey shaded ellipse indicates region in which analytical errors preclude identification of presolar grains. See Fig. 6 for data sources.

CONCLUDING REMARKS We have reviewed in detail the nucleosynthesis and chemical evolution of the O isotopes in the Galaxy. Although the relative abundances of these isotopes are affected by many nuclear processes occurring in many types of stars, it is clear that most l e O and l s O atoms in the Universe were synthesized in massive stars, whereas a m a j o r fraction of the 1 7 0 probably f o r m e d in lower-mass stars and nova ejecta. This basic picture is confirmed by astronomical observations of stars and molecular clouds and by the ability of nucleosynthesis and galactic chemical evolution models to roughly reproduce the composition of the Solar System. Nonetheless, there are significant puzzles, especially the unusual l s O / l e O ratio of the Solar System, and many important remaining uncertainties in the details of the stellar models. In this regard, presolar grains provide particular promise for improving our understanding of the origin and evolution of O in the Galaxy. Advances in the sensitivity and spatial resolution of isotopic imaging capabilities of S I M S instruments have m a d e it easier to find and study presolar oxides and silicates in situ, especially those of submicron size. A s w e have seen above, such advances have already significantly enhanced our k n o w l e d g e of stellar evolution by constraining mixing in supernova ejecta and cool bottom processing and hot b o t t o m burning in red giant branch stars. A n d w e anticipate study of oxygen-rich presolar matter will continue to reward u s with insights into issues such as chemistry in supernova debris (for example, low-density graphite), chemical evolution (for example, the possible origin of the G r o u p 4 oxide grains), and transit times of dust through the interstellar m e d i u m and conditions in the S u n ' s parent molecular cloud (for example, presolar silicates in IDPs). Such rich rewards are only fitting for oxygen, the king of the strictly stellar synthesized elements.

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ACKNOWLEDGMENTS The authors would like to thank the organizers, particularly Glenn MacPherson, for the stimulating 2005 Oxygen in the Early Solar System workshop in Gatlinburg. BM, LRN and SRM all acknowledge NASA's Cosmochemistry program for financial support of this work.

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 55-72, 2008 Copyright © Mineralogical Society of America

Oxygen in the Interstellar Medium Adam G. Jensen Center

for Astrophysics and Space Astronomy, University of Colorado Boulder, Colorado 80309-0389, U.S.A. (Present address: Goddard Space Flight Center, Code 665, Greenbelt, Maryland 20771, U.S.A) Adam. G Jensen @ nasa.gov

F. Markwick-Kemper Jodrell

Bank Centre for Astrophysics University of Manchester M13 9PL, Manchester, United Kingdom [email protected] University P.O. Box 400325,

of Virginia, Department Charlottesville Virginia,

of Astronomy, 22904-4325,

U.S.A.

Theodore P. Snow Center

for Astrophysics and Space University of Colorado Boulder, Colorado 80309-0389, tsnow @ casa. Colorado, edu

Astronomy U.S.A.

ABSTRACT The oxygen that is observed in the Solar System today is a remnant of the interstellar oxygen that was in the dense molecular cloud that collapsed to form the Solar System. While the chemical evolution of the Galaxy has progressed since then, processes in the interstellar medium (ISM) that involve oxygen are relevant to the origins of oxygen in the Solar System. Oxygen in the ISM can be found as neutral or ionized atomic gas and as a constituent of molecular gas, volatile ices, and refractory minerals in dust, with the dominant state depending on the specific environment. The gas-phase abundance of atomic oxygen is well-known in the diffuse ISM that fills most of the Galaxy's volume, but the state of oxygen in denser environments is poorly understood. The ISM abundances of isotopes of oxygen other than 1 6 0 cannot be easily determined due to observational constraints. Oxygen in interstellar dust is primarily found in the form of silicates that are created in evolved stars and then ejected into the ISM before being incorporated into the formation of new solar systems. Some of the important unknowns concerning oxygen in the ISM include the "cosmic" (i.e., total) abundance of oxygen, the abundance of oxygen in dust, and the details of dust grain processing in the ISM.

INTRODUCTION T h e o x y g e n f o u n d in t o d a y ' s S o l a r S y s t e m o r i g i n a t e d in t h e i n t e r s t e l l a r m e d i u m ; it w a s o n c e p a r t of t h e d e n s e m o l e c u l a r c l o u d that c o l l a p s e d to f o r m t h e S u n a n d t h e S o l a r S y s t e m . T h e S o l a r S y s t e m h a s p r e s e r v e d a s a m p l e of interstellar m e d i u m m a t e r i a l t h a t w a s a v a i l a b l e ~ 5 b i l l i o n y e a r s ago, b u t t h e c h e m i c a l e v o l u t i o n of t h e G a l a x y h a s p r o g r e s s e d s i n c e then. It is, h o w e v e r , 1529-6466/08/0068-0005$05.00

DOI: 10.2138/rmg.2008.68.5

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worthwhile to study the processes that take place in the interstellar medium today to obtain clues to the origins of oxygen in the Solar System. Oxygen is an important component in both gas-phase interstellar chemistry and models of solid-state interstellar material (dust and ices). The dominant form of oxygen in the interstellar medium (ISM) varies greatly between different environments. We discuss the abundance and evolution of oxygen in the ISM, both in gas-phase atomic and molecular forms and as a component of oxygen-rich solid-state components. Phases in the interstellar medium Based on gas density, astronomers distinguish several phases in the interstellar dust and gas reservoir (Snow and McCall 2006). The diffuse atomic regions take up approximately ~90% of the available volume, but only ~1% of the mass, in the interstellar medium, while the dense molecular clouds contain the bulk of the mass (~90%), mostly in molecular form, in only ~1% of the volume of the ISM. The remaining intermediate-density material, which accounts for about 9% of the total mass in 9% of the total volume in the ISM, can be divided into diffuse molecular clouds and translucent clouds, based on their density and extinction1, and whether or not carbon is ionized, neutral, or in the form of CO (Snow and McCall 2006). Phase transitions between these phases occur, and dust and gas can cycle between the diffuse, translucent and dense phases. Material can only leave this eternal cycle in star formation processes when it becomes part of a star or a planetary system. Replenishment of the interstellar medium occurs through stellar ejecta; nucleosynthesis in stellar interiors may have altered the original compositions of stars, and the stellar ejecta often no longer resemble the original ISM material. In addition, the conditions in stellar winds are often favorable for additional chemical and physical processing of the stellar ejecta. Oxygen is found in all phases in the interstellar medium, and can be present in gas phase species (atomic or molecular), as well as in solid-state components (dust or ices). Overall, the interstellar medium is oxygen-rich and enough oxygen is available to drive an oxygen-rich chemistry. Forms of oxygen in the interstellar medium The dominant form of oxygen in the ISM varies with the type of gas cloud. In the densest molecular clouds, oxygen can be found in molecules such as CO, OH, and H 2 0. Over a very wide range of densities, including both diffuse molecular clouds and diffuse atomic clouds, oxygen is found in its neutral form (O I). In the hot intercloud medium, oxygen is found in various ionized forms, with the most easily detectable and diagnostically important (though not the most abundant) form being O VI. In dust (found to varying degrees in all phases of the ISM) oxygen can be found in silicates such as Mg2(i_x)Fe2rSi04, oxides such as FeO, and ices such as H 2 0. Measuring the amount of oxygen in each form presents its own unique challenges. In the next section we summarize the methods of observing oxygen in the gas phase, significant results, and implications for the amount of oxygen in solid-state forms.

O X Y G E N IN T H E GAS PHASE Measurements of gas-phase oxygen Neutral atomic oxygen. Neutral atomic oxygen (O I) has most of its resonance transitions in the far-UV (below 1100 A), though two of the most useful lines are found at 1302 A and 1356 A. A wide range of instruments have been used to undertake studies of interstellar O I, The degree to which radiation from a background source is extinguished. This is usually expressed as I =I0 with I0 the intensity of the background source, I the measured intensity and the optical depth x as a measure of extinction. 1

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from Copernicus (e.g., York et al. 1983; Keenan et al. 1985) to the Goddard High Resolution Spectrograph (GHRS) and Space Telescope Imaging Spectrograph (STIS) instruments onboard the Hubble Space Telescope (HST) (e.g., Meyer et al. 1998; Cartledge et al. 2001, 2004; André et al. 2003) to the Far Ultraviolet Spectroscopic Explorer (FUSE) (e.g., Moos et al. 2002; Jensen et al. 2005). The difficulty in measuring any atomic species in the ISM comes from the necessity to measure absorption lines that are either very weak or very strong. If lines are weak enough, then they can be assumed to be free of saturation, and the equivalent width (a measure of total absorption relative to the continuum) of the absorption line is linearly proportional to the column density of absorbing material. Unfortunately, it is difficult to find lines that are both weak enough to meet this requirement yet strong enough to be detected and accurately measured, especially in lines of sight with heavy reddening (due to dust) where the signal-tonoise ratio in the UV range is low. Additionally, some species are expected to be "depleted" (that is, in forms other than the gas phase), and subsequently the gas-phase column density that creates these absorption lines is even smaller. Weak lines are still useful if they have slight saturation, if the velocity structure 2 of another atomic species can be assumed and an appropriate saturation correction is applied. Conversely, when very strong absorption lines exhibit "damping wings" as a result of the Heisenberg Uncertainty Principle applied to the transition lifetimes, the shape of the wings is fairly sensitive to the total column density, and the total equivalent width of the absorption feature is proportional to the square root of the column density. In these cases, as long as the spectral resolving power is sufficient to clearly delineate the damping wings, column densities can be accurately measured either through fitting the shape of the wings or by measuring the equivalent width of the absorption feature. However, most elements of astrophysical interest do not have transitions strong enough to exhibit such wings at typical column densities, especially when potential depletions are factored in. Instead, in many cases astronomers are left with lines of intermediate strength, which are relatively insensitive to column density, making it difficult to derive accurate abundances from them, even in some cases where the velocity structure is accurately known. If the only available lines are of moderate strength, the alternative is to construct a curve of growth. A curve-ofgrowth method fits the equivalent widths of multiple absorption lines to a curve that represents the expected values of the equivalent width of a line for a range of combinations of oscillator strength and column density. The free parameter in this fitting process is the column density. The shape of the curve also depends on the line-of-sight velocity structure. Either a certain velocity structure must be specified, or a single-component velocity dispersion can be assumed. In the latter case, the ¿-value of the dispersion is a second free parameter in the fitting process, with the various ¿-values creating a family of curves. The accuracy of the curve-of-growth method relies on having a wide range of absorption line strengths that are measured and the validity of the assumed velocity structure (see Spitzer 1978 for a review). Once oxygen column densities have been determined, abundances relative to hydrogen can be determined in diffuse (or denser) clouds by measuring hydrogen through profile fitting of the Lyman-a line (for atomic hydrogen) and the many low-/ (rotational) states of H 2 . The major observational barriers are the spectral type of the background star (where cooler stars may introduce significant contamination of the Lyman-a line for the atomic hydrogen measurement) and the difficulty of obtaining good data quality in the far-UV, where most H 2 transitions lie. 1

"Velocity structure" refers to the distribution in velocity space of the absorbing material that is reflected in the profile of an absorption line. This structure includes the number of cloud components and their relative column densities; the velocity offsets of the clouds relative to Earth; and the velocity dispersion of each cloud (a "¿-value" quantifies a dispersion that is assumed to be Gaussian).

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The early Copernicus studies of York et al. (1983) and Keenan et al. (1985) relied primarily on the weak 1356 A line of oxygen, using a curve-of-growth of neutral nitrogen to infer any possible saturation corrections. Neither study found any evidence of correlations in the oxygen abundance with the total hydrogen column density. Meyer et al. (1998), André et al. (2003), and Cartledge et al. (2001, 2004) studied O I abundances with the GHRS and STIS instruments onboard HST. The Meyer et al. (1998) study examined 13 lines of sight and did not find any statistically significant variation of O/H with respect to iV(Htot), the total line-of-sight column density of atomic and molecular hydrogen, measured in particles per unit area; n H , the average volume density of hydrogen in a line of sight, equal to iV(Htot)/r, where r is the line-of-sight pathlength; or the fraction of hydrogen atoms in molecular form,/(H2)=2iV(H2)/[2iV(H2) + iV(H I)], where iV(H2) and iV(H I) are the molecular and atomic hydrogen column densities, respectively. The more extensive study by Cartledge et al. (2004), however, found that O I depletion from the gas phase did increase with the average volume density of hydrogen, n H . Cartledge et al. found that O/H is approximately 100 parts per million (~25% of the total observed gas-phase O/H ratio) smaller in lines of sight with larger n H . Cartledge et al. interpreted this as representative of two phases in the ISM (warm and cold), with a transition at n H ~ 1 cm -3 . Jensen et al. (2005) used the FUSE satellite to probe reddened lines of sight with large fractions of molecular hydrogen. FUSE does not provide information on the weak 1356 A line or the potentially damped 1302 A line, but does provide information on a wide range of absorption lines of intermediate strength that are beyond the spectral range of HST, and at greater sensitivity and/or resolution than past instruments such as Copernicus. The Jensen et al. (2005) study did not find statistically significant correlations of O/H with any important line of sight parameters, but could not rule out possible correlations within the large errors. However, a small subset of lines of sight with both HST and FUSE data available (and subsequently smaller systematic and statistical errors) did show the hint of a correlation of decreasing O/H with an increasing ratio of total to selective extinction, R v = Av/EB.v. Typical interstellar grains will preferentially cause extinction at shorter wavelengths, and EB_V is often correlated with the total column density of typical interstellar dust. Larger grains, however, will cause extinction in visible light in a non-preferential manner (i.e., grey extinction). Therefore, increasing values of R v , a measure of significant extinction in the visual band (Ay), but relatively little preferential extinction at shorter wavelengths (EB_V), are thought to correlate with increasing grain size. Other ionization states of atomic oxygen. The ionization potential of O I is nearly identical to that o f H I ( 1 3 . 6 1 8 e V and 13.598 eV, respectively). This is convenient, because combined with the fact that the ionization of O I relative to H I is also governed by a charge-exchange reaction (Field and Stiegman 1971), we can draw the simple conclusion that in regions dominated by H I and H 2 (from diffuse atomic clouds to dense molecular, star-forming clouds) the vast majority of oxygen will be either neutral or found in a molecular or solid-state form. More specifically, a gas cloud in collisional equilibrium at temperatures up to IO4 K should have over 99% of its oxygen in the form of O I (Sutherland and Dopita 1993). Observing other ionization states of oxygen poses significant observational difficulties in addition to their relatively small fractions. O II, O VII, and O VIII do not have ground-state transitions with wavelengths in the UV, and the few such transitions of O III, O IV, and O V have such small/-values that they are unlikely to be observed in interstellar absorption (Morton 2003). These elements are, however, observed in emission. For example, O III is commonly observed in planetary nebulae. Abundances relative to certain other elements can be accurately determined in these cases, but direct measurements of the O/H ratio are not possible in such regions, where the hydrogen is largely ionized. Again, however, O III does not play a significant role in diffuse clouds.

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While it does not compose a significant fraction of the total oxygen in the Galactic ISM, and is never the dominant form of oxygen (Sutherland and Dopita 1993), the lithium-like ion O VI is an important diagnostic of the hot intercloud medium, as well as gas in the Galactic halo and the intergalactic medium. This is due in part to its detectability, and also because its large ionization potential, which is above the ionization edge of He II, means that its creation is dominated by collisional excitation rather than photoionization. O VI is observed through a pair of relatively strong transitions at 1032 A and 1038 A, in a spectral range detectable with FUSE (and past instruments such as Copernicus) but not HST. The broad thermal distributions of O VI that are typically seen imply that O VI absorption is not significantly saturated and O VI column densities can be determined directly and accurately. The study of O VI is important for studying gas at the interface between hot and cold gas; at collisional equilibrium temperatures of ~3xl0 5 K, O VI reaches its peak ionization fraction of ~20% (Sutherland and Dopita 1993). However, most O VI is observed in non-equilibrium regions, such as shock fronts, and traces cooling and heating, itself being an important coolant. Oxygen in gas-phase molecules. The dominant form of oxygen in molecules is carbon monoxide. All isotopologues 3 of CO have been detected in the interstellar medium through radio emission (from rotational transitions) except for 1 3 C 1 7 0 (Morton and Noreau 1994), though detection through ultraviolet absorption (from electronic transitions) is also common. Sheffer et al. (2002) and Federman et al. (2003) have presented some of the most recent observations of many different forms of CO in UV absorption with the STIS and GHRS instruments onboard HST. These results show just how small the abundance of molecules is in the diffuse interstellar medium. For example, the 12 C le O column density for the line of sight toward X Persei found by Sheffer et al. (2002) is a factor of ~50 smaller than the column density of gas-phase atomic oxygen found by Jensen et al. (2005). Other isotopologues of CO are even less abundant. The study of CO is particularly important for the cooling of dense molecular clouds; CO is collisionally excited and then radiates at wavelengths where the cloud is optically thin. In these cases, the CO abundance jumps up dramatically; as much as one-sixth of the total carbon abundance may be in CO. Other molecules of potential importance include H 2 0 , OH, 0 2 , and H 2 CO (formaldehyde). Goldsmith et al. (2002) presented a tentative detection of 0 2 in a high-velocity outflow seen in the line of sight toward p Ophiuchi, later confirmed by Liseau et al. (2006). The abundance relative to H 2 is approximately 10~5. In general, though, limits have been placed on the 0 2 abundance to be less than 10~7 relative to hydrogen. (See Goldsmith et al. (2002) for further discussion of a scenario explaining the large abundance in this line of sight.) Aggregate measurements of oxygen in gas and dust. Measuring oxygen in all of its forms is possible in some cases. The inner K-shell absorption edge of oxygen can be detected through X-ray spectroscopy, measuring oxygen along the line of sight in all of its forms— atomic, molecular, and in dust. Fine structure can further reveal whether the oxygen is in the gas form or tied up in solids. The results in the case of X Persei (Cunningham et al. 2004) are consistent with the results of gas-phase oxygen abundance studies and the requirements of dust models based on UV extinction. To summarize, the gas-phase atomic oxygen abundance is relatively constant over a wide range of conditions. However, the evidence from Cartledge et al. (2004) suggests that oxygen depletion from the gas phase increases somewhat with increased average gas density, and the 3

Isotopologues are chemicals with different isotopes of the same atomic constituents. An isotopomer refers to chemicals with the same isotopes of atomic constituents but different atomic arrangements (isotopomer is a contraction of isotopic isomer). Though isotopomer is commonly applied to the various forms of CO in astrophysical literature, isotopologue is the correct term.

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evidence from Jensen et al. (2005) hints that grain size is related to an increase in oxygen depletion. However, astronomers still do not have a perfectly formed idea of how oxygen varies in the most extremely reddened environments, or even in translucent clouds (a phase of the ISM that has long been postulated but for which direct evidence is sparse; see further discussion in Rachford et al. 2002). The Cosmic Origins Spectrograph (COS), which will be installed on HST as a part of Servicing Mission 4 (currently projected for August 2008), will have several times the sensitivity of STIS, but at somewhat lower resolution. COS will be able to probe oxygen column densities through detection of the weak 1356 A line of oxygen even at Ay > 5 mag, much more reddened than has ever been observed at high S/N in the ultraviolet. This will allow probing of the gas-phase oxygen abundance in the much denser environments that are not yet fully understood. Isotope measurements from gas-phase oxygen and carbon monoxide Atomic isotope shifts. Attempting to measure relative isotopic atomic abundances of elements in the interstellar medium is very difficult. In the case of single-electron (hydrogenlike) atoms, the difference in the reduced mass of the electron and the nucleus, |i = m e m nllcklls / (me+mnllcklls), is what provides the change in the energy and subsequent shift in wavelength of a given transition between two different isotopes. While this can be significant for the shift between hydrogen and deuterium ( ~ 3 x l 0 ^ of the transition energy, corresponding to the absolute value of the fractional shift in the wavelength), it is increasingly small with increasing atomic mass. For example, the difference in the transition energy due to the different reduced masses of l e O and l s O is ~4xl0 - 6 (assuming hydrogen-like O VIII). For species with more than one electron, the shifts depend on more than just the reduced mass, as the total momentum of all electrons in the system must be considered. The resulting energy shifts are difficult to calculate (see Clark 1984, who calculates shifts for many atomic species but not O I); however, they are unlikely to be substantially larger than the above approximation (which is equivalent to a shift of ~1 km/s for far-UV lines), which puts it at the threshold of the resolution of even a precision instrument such as STIS. No study of interstellar atomic oxygen has turned up clear evidence of other isotopes of atomic oxygen, due to a combination of the following limitations: (1) The inherently small shift between isotopes in the energy for a given transition, especially compared to the resolution of available instruments; (2) The broadening of interstellar profiles (~3-20 km/s for oxygen; see Jensen et al. 2005), which is a combination of thermal broadening, multiple clouds (with a range of velocity offsets) in the line of sight, and in the case of strong, damped lines, "natural" broadening; and (3) The relatively small abundances of isotopes other than le O. Isotopologues of carbon monoxide and other oxygen-bearing molecules. The situation with molecules is more promising. The difference in isotope mass of the atoms in a molecule influences the energy of its rotational and vibrational states, in some cases splitting the energies of electronic transitions by resolvable amounts. Further helping matters is the fact that molecules tend to form in cold environments with small thermal widths, resulting in absorption features that are more easily resolved. Radio observations of rotational transitions can be used to detect various isotopologues; in fact, as mentioned above, almost all variations of carbon monoxide have been measured in the ISM through radio observations. The ratio of 12CO to 13CO does seem to show a spatial variation that can be correlated with stellar evolution that turns 12C into 13C (Langer and Penzias 1990). However, the relative abundances of the various isotopologues of CO are unlikely to be good proxies for the relative isotopic abundances of oxygen. Both Sheffer et al. (2002) and Federman et al. (2003) discuss

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the fractionation of CO isotopologues. In all combinations, the fractionation is more severe than the relevant isotopic ratios; e.g., 12 C 16 0/ 12 C 18 0 is found to be ~3xl0 3 , compared to the Solar System 1 6 0/ 1 8 0 ratio of ~500 (Lodders 2003). This is because the less abundant isotopologues remain optically thin and therefore susceptible to photodissociation, while the more abundant isotopologues may begin to become optically thick and self-shielded from photodissociating radiation. The impact of slightly different RMS velocities—due to different isotope masses—on molecular formation rates may also play a small role in the difference between the isotopologue ratios of CO and the isotopic ratios of C and O. The isotope 17O is less abundant than both l e O and l s O. Penzias (1981) attempted to constrain the C 1 8 0/C 1 7 0 ratio, and found a value of ~3.5, consistent within the errors at all measured galactocentric radii. Wouterloot et al. (2005) found a ratio of -4.1, while Ladd (2004) found ratios of ~4.0 or ~2.8, depending on the model used to interpret the line intensities. All of these measurements, however, are somewhat smaller than the Solar System 1 8 0/ 1 7 0 ratio, -5.5 (Anders and Grevesse 1989; Lodders 2003). These low ratios cannot be due to differences in self-shielding effects, because C l s O is more likely to be self-shielded than C 1 7 0, which would increase the observed C 1 8 0/C 1 7 0 ratio (Ladd 2004); deep within molecular clouds, both are likely to be well-shielded. Vastly different 1 8 0/ 1 7 0 ratios are also observed in the Large Magellanic Cloud (LMC) and two nuclear starburst galaxies, implying that metallicity4 might play a role in determining this ratio (Wouterloot et al. 2005 and references therein). Studies of isotopologues of OH have also been used to explore the isotopic abundances of oxygen. Bujarrabal et al. (1983) found ratios of 16OH/18OH to be -370 and -440 in two different cloud components toward Sgr B2, with a lower limit of -200 in a third cloud component. Within the errors, these derived ratios are in reasonable agreement with the Solar System ratio, and Bujarrabal et al. suggest that previous interstellar results with lower ratios are likely the result of excitation of the more abundant 16OH. The 18OH/17OH ratio is found to be 3.6+0.5, however, again lower than the Solar System value of -5.5 and in agreement with the results from CO isotopologues. Polehampton et al. (2003, 2005) also used isotopologues of OH to explore the interstellar isotopic abundances of oxygen, finding 16OH/18OH ratios also roughly consistent with the Solar System 1 6 0/ 1 8 0 ratio (17OH was also detected, but 18OH/ 17 OH was not independently derived). They also found no evidence of a gradient in 1 6 0/ 1 8 0 with galactocentric distance, but their results also do not conclusively rule one out. The lack of a gradient in the isotope ratio would be in contrast with several other studies that show a low value in the galactic center (e.g., the compilation of 1 6 0/ 1 8 0 ratios determined from H 2 CO isotopologues in Wilson and Rood 1994). For further discussion, including information about the uncertainties of the various studies, see Polehampton et al. (2005) and references therein. Inferring gas-phase depletions of oxygen While there is some consensus on the gas-phase abundances of oxygen in environments that are not too extreme, what is less clear is the interpretation of the gas-phase oxygen measurements as it relates to forms other than the gas phase. Typically, derived abundances of gas-phase oxygen are compared to a certain standard (usually the solar oxygen abundance) and the "missing" oxygen is inferred to be the total amount of oxygen in molecules and dust. The problem with this method is two-fold. First, the appropriate standards are not well known. For example, measurements of the solar oxygen abundance vary by nearly a factor of two, even within the last decade. Secondly, and perhaps more importantly, there is reason to doubt that the solar oxygen abundance is the correct standard, or that any such standard exists at all. What is the solar oxygen abundance? Large ranges are found when considering photospheric models. Anders and Grevesse (1989) derived a value for the solar O/H of 853 4

In astrophysics, the term metallicity is often used to mean the mass fraction contained in "metals," i.e. elements heavier than helium, in an astrophysical environment.

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parts per million (ppm), while the more recent 3-D photospheric models of Holweger (2001) and Asplund et al. (2004) put solar O/H at 545 and 460 ppm, respectively. More recently, Ayres et al. (2006) have argued for a solar O/H ratio of 700 ppm based on chromospheric CO as a tracer of the total oxygen abundance. For a time, it was thought that many elemental abundances (e.g., carbon, oxygen, and krypton) in the ISM were very nearly 2/3 of the solar value, but this is no longer a widely held opinion, based largely upon the revisions to, or at least the uncertainty in, the solar values of the carbon and oxygen abundances (the disparity with krypton may still exist, however; see Sofia and Meyer 2001). Instead of the solar photospheric elemental abundances, perhaps protosolar abundances— i.e., "Solar System abundances"—are better estimates of the cosmic abundances. Estimates for abundances in the protosolar disk are most commonly derived from meteorites, but this is problematic for oxygen because it readily forms gaseous compounds, and it is therefore almost certain that there was much oxygen in the protosolar disk that did not condense into rocks. Protosolar abundances, such as that for oxygen, are derived by applying a systematic factor to the photospheric abundances; this systematic factor is based on understanding of condensation in the protosolar disk, including the meteoritic abundances of some of the heavier elements. Two examples of this method, Cameron (1973) and Lodders (2003), derive a range of protosolar O/H ~ 575-700 ppm, due in part to the different photospheric abundances used. For further discussion on the protosolar oxygen abundance and related issues, see the chapter in this volume (Davis et al. 2008) Perhaps our Solar System and Sun are unique and a better oxygen abundance standard is the oxygen abundance in other stars. Studies such as Snow and Witt (1996) and Sofia and Meyer (2001) have examined this issue. Specifically, Snow and Witt found O/H = 380 ppm for field B stars, and O/H = 490 ppm for cluster B stars and field F and G stars. For B stars, Sofia and Meyer found a weighted average of O/H = 350 ppm, and an average O/H of 445 ppm for F and G stars (not weighted, as it was based on a very small sample). The observed gas-phase ISM abundances of oxygen rule out the field B star abundances of Snow and Witt (1996) and the B star abundances of Sofia and Meyer (2001) as potential standards. Various theoretical models of interstellar dust, such as those cited in Snow and Witt 1996, attempt to explain the nature of interstellar extinction, including features such as the 10- and 18-|im bands believed to be from silicates; these models generally require >120 ppm of oxygen relative to hydrogen. Combined with the ISM averages of André et al. (2003) and Cartledge et al. (2004) of gas-phase 0/H=400 ppm (the "warm ISM" in the Cartledge et al. case), even the F and G star abundances fall a little short of explaining the total gas- and dust-phase oxygen in the ISM. What might cause the oxygen abundance in the ISM to differ from the oxygen abundance of the Sun and other stars? It is important to note that the spread in stellar abundances is greater than the spread in ISM abundances between various lines of sight, leaving the question of how such varied stars could form from a well-mixed ISM. A likely answer is that processes occurring during stellar formation heavily influence stellar abundances. Snow (2000) discusses two possibilities that may leave stars metal-poor: ambipolar diffusion and sedimentation. Ambipolar diffusion would magnetically exclude refractory elements, possibly resulting in stellar abundances that are smaller by a factor of 2-3. Sedimentation is the process by which large grains form in a protostellar disk but do not accrete onto the star, decoupling refractory elements from the gas, and resulting in metal-poor stars. Another process that could potentially factor into producing stars that are metal-poor relative to the ISM is the photoablation of protostellar disks. If protostellar disks are destroyed soon after stellar formation, then the late infall of potentially metal-rich material is disrupted. Additionally, if there is some late infall of gas (either metal-rich or metal-poor) after stellar convection has begun, this gas may remain in the photosphere, distorting observed versus true stellar abundances.

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O X Y G E N IN I N T E R S T E L L A R DUST On first sight, with an overall gas-to-dust ratio in the interstellar medium of 100-200, dust may not appear to be an important contributor to the total mass budget of oxygen. However, almost all of the gas in the ISM is actually in the form of hydrogen or helium, while these two species rarely compose a significant fraction of the dust mass. Thus, for the remaining elements, including oxygen, the solid phase appears to be important. In fact, at a metallicity of Z=0.02, and with a gas-to-dust ratio of 100, about half of the metals—including oxygen—are contained in the solid phase. Most of the oxygen that is found in the solid state in the diffuse interstellar medium is found in silicates, while some may also be present in metal oxides (Whittet 1984). In the dense interstellar medium, the remaining gas phase oxygen might condense out in the form of ices, in particular H 2 0 and C 0 2 , as evidenced by infrared spectroscopy of protostars (Gibb et al. 2004). Ices are much more volatile than silicates and are therefore less likely to contain a record of their formation and processing history. In addition, the phase in which ices can be observed astronomically spans only a short period of time in the total evolution cycle of dust. We will therefore limit the discussion on interstellar oxygen in the solid state to the silicates and will discuss such properties as composition and lattice structure, in the context of the evolution of silicate dust. Solar System silicates The most recent phase in the evolution of silicate dust, Solar System silicates, is discussed extensively in other chapters in this book. The most primitive silicates found in the Solar System today are believed to be the GEMS (Glasses with Embedded Metals and Sulfides; Bradley 1994), which are present in some interplanetary dust particles (IDPs), and the silicates that have been locked up in comets. Evidence for crystallinity is found in the 8-13 |im spectroscopy of both short and long period comets (Hanner et al. 1994; Honda et al. 2004). Ground-based observations of the Deep Impact encounter with comet 9P/Tempel 1 show that the silicates observed before and after the impact were mostly amorphous, and predominantly in the form of large grains, while the dust plume caused by the impact contains an abundance of smaller and crystalline silicates (Harker et al. 2005). This is interpreted as evidence that cometary silicates initially are significantly crystalline, and that these silicates gradually become more amorphous upon passing the Sun several times. A full 2-45 |im spectrum, obtained for Oort Cloud comet Hale-Bopp (Crovisier et al. 1997) supports this view, by showing strong forsterite features in the 20-40 |im range. Estimates for the degree of crystallinity of the silicates in Hale-Bopp typically fall in the 20-30% range (Brucato et al. 1999; Galdemard et al. 1999; Wooden et al. 1999; Hay ward et al. 2000; Harker et al. 2002; Bouwman et al. 2003), but more recently Min et al. (2005) have convincingly shown that the degree of crystallinity is probably closer to 7.5%. Because very little processing occurs inside comets in the Oort cloud, the crystalline fraction in long-period comets observed today essentially reflects the composition of the silicates in the comet-formation zone in the planet-forming disk around the young Sun. The fact that cometary silicates are more crystalline than GEMS and than interstellar silicates (Kemper et al. 2004), indicates that crystallization occurs in the planet-forming disk itself, or perhaps in the dense molecular cloud that pre-dates the Solar System. The process of crystallization of the amorphous interstellar silicates, as they get incorporated in the Solar System, is probably best studied in other planetary systems in formation. Silicates in circumstellar environments of young stars One can distinguish two types of circumstellar disks around young stars. Pre-main sequence stars often have relatively stable and massive primordial planet-forming disks around them, while young main-sequence stars can exhibit debris disks in which the sources of dust are planetesimal collisions. When the era of planetesimal collisions ends, the disk rapidly clears

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out due to Poynting-Robertson drag and radiation pressure. Debris disks are seen around stars up to an age of about 400 million years (Habing et al. 1999). The spectral properties in such disks are consistent with the planetesimal origin of the grains. Recent Spitzer observations have shown that typical debris disks are populated by large (Jura et al. 2004) and crystalline (Beichman et al. 2005) silicate grains. The planetesimals from which these grains originated may bear resemblance to either asteroidal or cometary bodies (Beichman et al. 2005), both of which, in our Solar System, are known to contain crystalline silicates. While crystallization in dense planet-like bodies occurs due to the heat of formation, the presence of crystalline silicates in loosely packed cometary bodies is not entirely understood. Formation heat never reaches sufficiently high temperatures to anneal the silicates, so the crystalline fraction observed in comets reflects the composition of the circumstellar environment where the comets are formed. Indeed, crystalline silicates are observed in the planet-forming disks around pre-main-sequence stars (e.g., Waelkens et al. 1996). A range of crystalline fractions is observed—always the Mgrich end-members forsterite and enstatite—however, it rarely exceeds a few percent of the total silicate mass. The crystallinity appears to be more enhanced for larger overall grain sizes in the circumstellar disk (Bouwman et al. 2001). Although the evolutionary status of the disk may correlate with the particle size of the silicate grains, a clear correlation between stellar age and crystallinity remains to be confirmed (Bouwman et al. 2001). It is clear that energetic processing is required to build up the observed degree of crystallinity (a few percent), but it remains unknown how this level of crystallinity can be achieved in the comet-forming region, where the radiative heating from the young star is not sufficient to heat the grains to crystallization temperatures. At least two different mechanisms have been proposed. First, radial mixing of grains from the warm inner radius of the disk toward the outer regions has been suggested as a mechanism to transport crystalline material to the comet-forming region (Bockelee-Morvan et al. 2002). This theory is supported by the recent results from the Stardust mission, where the presence of refractory Ca-, Al-rich grains among the cometary particles, as well as the low abundance of grains with presolar isotopic compositions, in particular the 18Q/16Q r a j j 0 ; suggests large-scale radial mixing (Brownlee et al. 2006; McKeegan et al. 2006; Zolensky et al. 2006). Alternatively, it might be possible that shocks propagating through the pre-solar nebula sufficiently heated the grains in situ to cause crystallization, even at distances of 10 AU (Harker and Desch 2002). Recent interferometric observations of three pre-main sequence stars clearly show a gradient in crystallinity in the circumstellar disk, with the highest crystalline fractions near the inner radius of the disk (van Boekel et al. 2004), thus providing supporting evidence for the radial mixing theory. Dust properties in the interstellar medium In all three phases of the interstellar medium (ISM), silicates are the most important depletion sink for oxygen. Studies of silicate properties have concentrated mostly on the diffuse interstellar medium however, which contains only 1% of the dust mass in the ISM, although it occupies about 90% of its volume. The properties of the dust composition in the denser phases of the interstellar medium have received less attention. In particular, silicate properties in young stellar objects (YSOs)—the solar nebula analogs—are relatively unexplored. Kessler-Silacci et al. (2005) present a sample containing nine approximately solar mass (M @ ) and three high mass (M > 8 M®), deeply embedded YSOs, showing silicate absorption features in the 8-13 micron wavelength range. They find that the silicates toward these protostars reveal an amorphous lattice structure, and only in more evolved objects, where the silicate feature appears in emission, can crystalline forsterite be observed. Demyk et al. (1999) have set an upper limit of 1-2% on the degree of crystallinity in the silicates towards two massive protostars, and find that the shape of the silicate absorption feature suggests an amorphous pyroxene composition. Early observations of the silicate profile towards the stars in the Trapezium cluster suggested an amorphous lattice structure for the silicates in the diffuse interstellar medium (Forrest et al.

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1975). The shape of this silicate feature has been used to construct the widely used optical constants for interstellar silicates (Draine and Lee 1984), under the assumption that silicates in the ISM are completely amorphous. In 2001, Li and Draine investigated the composition of interstellar dust from infrared emission, and were able to set an upper limit of 5% to the degree of crystallinity. More recently, the silicates in the 8.5 kpc-long line of sight towards the galactic center were shown to be 1 will condense carbon and carbides and one with C/0 Si0 2 ( s ) (Thiemens 1988, 1999, 2006; Marcus 2004). Thus, such models would predict that the Sun has essentially the same isotopic composition as the bulk Earth. It should be noted that the condensation reactions such as the latter one have not been shown in the laboratory to produce mass-independent fractionation of oxygen isotopes. Ozima et al. (2007) have considered planetary objects according to size and observed that the scatter in observed non-mass-dependent fractionation, expressed as A 1 7 0, decreases with increasing size. They performed a bootstrap statistical computation to derive a relationship between the standard deviation of A 1 7 0 and the diameter of planetary objects that is linear on a log-log plot. They argue that this is to be expected for hierarchical growth of planetary bodies and that the average oxygen isotopic composition of the planets should be the same as that of the solar nebula, and by extension, the Sun. Spectroscopic constraints on the oxygen isotopic composition of the Sun The photospheric abundance of oxygen is dominated by the contribution from le O. The abundance of the neutron-rich isotope l s O is obtainable from the array of CO fundamental and first-overtone vibration-rotation lines of the ground electronic state. The most recent analysis is that by Scott et al. (2006), who used a hydrodynamic model and examples of classical model photospheres. The carbon abundance obtained from the hydrodynamic model with an oxygen abundance of 8.66 is log e (C) = 8.39 ± 0.05, a value in perfect agreement with the carbon abundance from the 872.7 nm [C I] line (Allende Prieto et al. 2002) and with other atomic

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and molecular lines (Asplund et al. 2005). Thus, the hydrodynamic model gives a consistent carbon abundance from multiple indicators, as it does for oxygen from the suite of oxygenindicators. The relative isotopic abundances obtained were 12C/13C = 86.8^3' and 1 6 0/ 1 8 0 = 479^8 and are in good agreement with the terrestrial values, 89.45 ± 0.23 and 498.71 ± 0.12, respectively (Coplen et al. 2002). The corresponding 8 13 C (813C = [(13C/12C)samplc/(13C/12C)PDB - 1] x 1000, where PDB is Peedee Belemnite, a type of calcium carbonate shell from the Peedee Formation in South Carolina) and 8 l s O values of the Sun are 31^®%o and 4ltg3%o, respectively. The latter value may be difficult to reconcile with solar 8 1 7 0, 8 l s O of -50%o or lower, but it is not clear how robust the error estimate is for the spectroscopic 8 l s O value. The abundance of 1 7 0 is too low to result in detectable lines of 12 C 17 0, even in the strongest fundamental lines. Ayres et al. (2006) derived lower 12C/13C, 1 6 0/ 1 8 0, and 1 6 0/ 1 7 0 ratios and higher carbon and oxygen abundances on the basis of a semi-empirical ID model, but Scott et al. (2006) have criticized their methods. Identification of the solar isotopic composition trapped in lunar samples Hydrogen, carbon and nitrogen in lunar soils. Numerous efforts have been made during the past four decades to decipher the isotopic compositions of solar ions implanted in lunar soil samples. We highlight here recent efforts regarding the light elements carbon, nitrogen and oxygen, which are of particular interest, since they exhibit wide differences in their isotopic compositions among various planetary materials. The solar isotopic compositions of these elements are keys to the understanding of the origin and evolution of solids in the Solar System. One of the greatest difficulties in the search for solar components trapped in lunar samples is that the surfaces of mineral grains may show isotopic records not only of the solar ions, but also of various kinds of extraselenial materials having different origins. Different lunar soil grains have different exposure histories; they may have been exposed at quite different periods and for different durations, thus possibly registering, for example, variations in the accretion rates of extraselenial materials to the surface of the Moon. As lunar soil grains are occasionally excavated by the so-called gardening effect, components of different origins are often mixed and/or overlaid on the surfaces of the grains. As described later for the case of nitrogen, it was not easy to untangle this complex situation until techniques were developed to measure isotopic compositions on the micrometer scale. Isotopic analyses of 100-|im-sized single grains are now possible either by a microscopic laser ablation technique combined with low-blank, high-sensitivity mass spectrometry (Wieler et al. 1999; Hashizume et al. 2002; Hashizume and Marty 2004), or by isotopic depth profiling with a depth resolution of a few nm using secondary ion mass spectrometry (SIMS) (Hashizume et al. 2000, 2004; Hashizume and Chaussidon 2005; Ireland et al. 2006). These techniques were first applied to the challenging and controversial lunar record of solar wind nitrogen isotopic composition, a long-standing puzzle that arose fairly soon after the first recovery of lunar samples (Kerridge 1975), and continued to be a debated issue until quite recently. The nitrogen isotopic composition trapped in lunar soil samples is quite variable. The 8 15 N values (815N = [ ( " N ^ N ^ ^ N / 1 4 ] ^ - 1] x 1000) observed among bulk lunar soil samples typically range between -200%o and +100%o. These differences were found to be especially prominent between samples recently exposed at the surface of the Moon and samples with very long exposures (~Ga) (Kerridge 1993). The apparent correlation between concentrations of solar noble gases and nitrogen among different regolith samples, regardless of the nitrogen isotopic compositions, led Kerridge (1975, 1993) to conclude that the nitrogen isotopic composition of the solar wind has changed with time. Wieler et al. (1999) criticized this argument, based on their grain-by-grain nitrogen and 36Ar abundance measurements. In fact, although nitrogen appears to be roughly correlated with solar 36Ar when bulk abundances of gases extracted from millions of grains are compared, nitrogen and 36Ar are actually not correlated grain by grain, suggesting that the presence of these two elements in the grains came

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about in different ways. Wieler et al. (1999) concluded that the predominant portion of nitrogen in the recently exposed lunar sample they studied is of nonsolar origin. Hashizume et al. (2000) have further demonstrated, by measuring isotopic compositions of hydrogen and nitrogen in surface layers of grains from several lunar samples, that not all lunar samples are suitable for detecting the solar wind component. They have shown that certain grains, especially the young lunar sample that Wieler et al (1999) studied, are enriched in deuterium, where a nitrogen component with a characteristic ( 15 N-rich) isotopic composition was also observed. This nonsolar component apparently was mistakenly regarded as a solar wind component in earlier studies, but, since the solar wind is strongly depleted in deuterium because the Sun destroyed most of its deuterium by nuclear reactions in its earliest stages of evolution, this component cannot have come from the Sun. The exact nature of the 15N-rich nitrogen component commonly observed among lunar soil samples, however, is still debated. Hashizume et al. (2001, 2002) argued that interplanetary dust particles (IDPs) accreting to the lunar surface with a rate similar the terrestrial IDP accretion rate (Love and Brownlee 1993) could be the source of the nonsolar lunar nitrogen. The highly positive 8 1 5 N values measured in many IDPs (Floss et al. 2006) are consistent with such an argument. Ozima et al. (2005), however, proposed a different possible source for the nonsolar nitrogen: terrestrial nitrogen might be efficiently delivered to the Moon, particularly when the Earth had no geomagnetic field. Although this cannot account for the bulk of the nonsolar lunar nitrogen, which is characterized by 8 1 5 N values much higher (up to 100%o) than the terrestrial value, it strengthens the possibility that several different sources deliver nonsolar nitrogen to the Moon's surface. Hashizume et al. (2000) studied grains from an ancient lunar soil, which is estimated to have been exposed at the surface of the Moon 1-2 billion years ago, and detected a hydrogen component with an extremely low deuterium content (8D < -950%o) that is very clear evidence for a solar component. They found that a 15N-depleted component (8 15 N < - 2 4 0 ± 25%o) was associated with the deuterium-poor hydrogen, leading to the conclusion that nitrogen in the Sun is at least this isotopically light and that planetary solid materials are systematically enriched in 15 N relative to solar composition. This conclusion was later confirmed by spacecraft data, e.g., by the isotopic composition of the Jovian atmosphere measured by the Galileo probe (8 15 N = - 3 7 0 ± 80% c ; Wong et al. 2004) and the Cassini probe (8 15 N = - 3 9 0 ± 140%«; Fouchet et al. 2004), and by the reevaluated result of the solar wind directly measured by the SOHO probe (8 15 N = +40 ± 490% c ; Kallenbach 2003). Meibom et al. (2007) have recently reported the nitrogen isotopic composition of osbornite (TiN) enclosed in a refractory inclusion in the Isheyevo carbonaceous chondrite. They argue that this mineral condensed from the solar nebula at high temperature and that it sampled the isotopic composition of the Solar System, and, by extension, the Sun. They found that its 8 15 N value is - 3 5 9 ± 5%o, quite 15N-poor compared to nitrogen in other planetary solid materials, and in good agreement with lunar and Jovian values. On the other hand, preliminary results from the Genesis mission suggest that the contemporary solar wind may have a nitrogen isotopic composition similar to that of the Earth's atmosphere (Marty et al. 2007). Carbon isotope ratios show significant variations among lunar samples, although the range is much less than that of nitrogen. The 8 1 3 C values observed among bulk samples range from -20%o to +25%o (Becker 1980). Using the same approach as that used for nitrogen, carbon isotopes were measured by SIMS in several silicate grains containing pure and abundant solar hydrogen from an ancient lunar soil, an upper limit of - 1 0 5 ± 20%o was obtained for the solar wind 8 1 3 C by Hashizume et al. (2004). This estimate barely agrees with the Jovian atmospheric value of - 3 9 ± 45%o (Niemann et al. 1996), although further tests might be necessary to obtain a firm result. A noteworthy observation was that among the lunar grains rich in solar gas, there were two surface-correlated components with different isotopic compositions, situated at slightly different depths. Because the concentrations of both components were correlated with those of the solar hydrogen, it was proposed that these two components derived from the Sun

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(Hashizume et al. 2004). The 13 C-poor component was assigned to the solar wind (SW), the solar ions with a kinetic energy of ~1 keV/nucleon that dominate the contemporary solar wind flux. This component is observed within a depth interval from 30 to 70 nm, which agrees with the expected stopping depth when ions with a kinetic energy equivalent to that of the SW are implanted into silicate (Ziegler et al. 1985). The second 13C-rich component (813C up to +10%o), observed mostly at depths > 100 nm, was ascribed to the solar energetic particles (SEP), solar ions with higher kinetic energies than the SW. Such solar ions are very rare in the contemporary solar ion flux. For example, the mean fluences of solar helium, oxygen or iron ions with 10 keV/nucleon, compiled from data obtained by several spacecraft, are more than four orders of magnitude smaller than those of the slow solar ions with 1 keV/nucleon (Mewaldt et al. 2001). Nevertheless, the SEP component, especially of the noble gases, is abundantly observed among lunar soil samples and certain solar gas-rich meteorites (Wieler 1998). The abundances of SEP noble gases are roughly comparable to those of the SW gases. The isotopic compositions of the SEP noble gases are mass-dependently fractionated relative to the SW component (Wieler 1998): the SEP component is enriched in the heavier isotopes, which is qualitatively consistent with the above assignment of the two solar carbon components. At present, the exact reason for the prominence of the SEP component compared to the SW component among the gas-rich planetary materials, and the fractionation mechanism of the SEP isotopic compositions are both poorly understood (Wieler 1998). An interesting result was recently reported regarding the neon isotopic composition from metallic glass samples flown on the Genesis solar wind sample return spacecraft. The SW neon extracted using the closed-system stepwise etching method revealed substantial variation in the 20 Ne/ 22 Ne ratio, monotonically decreasing from 15.0 near the surface to 11.8 at depth, covering almost the entire range commonly observed for the SW and SEP compositions (Grimberg et al. 2006). The decrease of the 20 Ne/ 22 Ne ratio could be explained by a fractionation due to a slight difference in penetration depths of the two isotopes; thus no separate SEP component would be needed to explain the isotopic variations observed among the solar particles implanted in solid materials. Future studies of the carbon, nitrogen, oxygen and noble gas isotopic compositions of solar ions implanted in lunar soil samples might contribute to the understanding of these issues, shedding light on the exact origin in the Sun of SW and SEP components, their acceleration mechanisms, implantation mechanisms and the secular variations in their fluxes. Meibom et al. (2007) found that osbornite in a refractory inclusion in the Isheyevo meteorite had a 12 C/13C ratio of 88.8 ± 0.6, in agreement with the SIMS measurement of the carbon in ancient lunar soils assigned to SEP, indicating that the Sun and planets have the same carbon isotopic composition. Oxygen isotopes in lunar metal grains—two studies and two conclusions. Attempts at measuring the oxygen isotopic composition of the solar component trapped in lunar soil samples have been done by SIMS following the previous studies of solar nitrogen and carbon. There are two studies, which reported very different isotopic compositions (Hashizume and Chaussidon 2005; Ireland et al. 2006). In each, metallic iron grains extracted from lunar soil samples were used as long-term collectors of solar composition because they originally contained no oxygen. Both studies concluded that the solar oxygen isotopic composition is several percent away from the terrestrial fractionation line (8 1 7 0 = 0.52 x 8 l s O), but in opposite directions! Hashizume and Chaussidon (2005), hereafter HC05, inferred a solar composition enriched in l e O (8 1 7 0 ss 8 l s O < - 4 0 ± 8%c, or A 1 7 0 < - 2 0 ± 4%c), based on analyses of a component implanted at depths of 0.1-1 |im in metal grains from the lunar regolith breccia 79035. This ancient breccia experienced lithification some 1-2 b.y. before present, and was previously studied for determination of solar carbon and nitrogen isotopic compositions. HC05 prepared mounts containing the magnetic fraction of the lunar soil, which was pressed directly into indium foil. They found that the metal grains from 79035 had thick (order of 1 )im) oxide coatings, which

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generally gave oxygen isotopic compositions close to the terrestrial mass fractionation line. In a few instances, the A 1 7 0 of the analysis became lower with time, down to -20%o, indicating the presence of an le O-rich component at depth, corresponding to solar energetic particles. It is this component that is inferred to represent solar composition. Several examples of oxygen isotope depth profiles of lunar metallic grains obtained by HC05 are shown in Figure 2. Meanwhile Ireland et al. (2006), hereafter IHNC06, inferred a solar composition depleted in l e O (8 l s O ~ +100%o; A 1 7 0 = +26 ± 3%o), based on analyses of a component residing at a shallow depth (20-100 nm, corresponding to the SW implantation) in metal grains from a recent lunar soil, Apollo 11 sample 10084. IHNC06 separated individual metal spherules from the soil and mounted the spherules in gold foil. Two of the grains had very thin oxide coatings (300 nm depth part of grain 0 4 C mostly represent those of a lunar silicate grain sitting close to the metallic grain. The solar signature, the negative A 1 7 0 values, are observed in rather deep parts of the grains (200 nm to >1 depth), where other surface-correlated components such as the oxide layer (likely acquired after the recovery of the sample) were mostly removed.

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83

for solar wind implantation was a key indicator of the solar wind, and the corresponding carbon and nitrogen compositions at those depths were inferred to be solar. IHNC06 chose a recent soil because it offered the best prospect for recent exposure to the solar wind, and would therefore be most analogous to the exposure of synthetic materials on the Genesis mission (see below). The high proportion of grains with high track densities and the 36 Ar concentration attest to this exposure. HC05 argue that the bulk soil contains many components of diverse origins and therefore the young soils are compromised. However, the issue is one of ascertaining contributions on a grain-by-grain basis, such as was carried out by HC05 for carbon and nitrogen. In the case of carbon and nitrogen, the isotopic compositions are indeed at the appropriate depth and associated with low D/H. For oxygen isotope measurements, the oxide rims potentially compromise the analyses, with that oxygen coming from an unknown source. For this reason, IHNC06 chose not to analyze this material. HC05 and IHNC06 attempted to perform D/H isotopic measurements on the metal grains, but the samples have D/H at terrestrial levels and this is interpreted as a result of diffusion from the terrestrial atmosphere after the samples were returned to Earth. The depths of the solar signatures in the two studies are quite different and require quite different sources. SRIM 3 calculations show that for solar wind energies, the depth of penetration is less than 50 nm for oxygen implantation into iron metal. Even for extreme solar wind energies (1000 km s -1 ), oxygen ions are stopped by 150 nm. These results are broadly consistent with the depth of oxygen measured by IHNC06, given some degree of uncertainty in the sample geometry and sputtering calibration, although the abundance of the oxygen at deeper depths is higher than expected based on the relative proportions of modern normal solar wind velocities (400 km s - 1 , corresponding to an energy of a few keV/nucleon) to extreme solar wind velocities (1000 km s -1 ). Implantation of the oxygen into the 79035 iron metal grains is attributed to solar energetic particles (HC05). SRIM calculations show that for energies on the order of 10 MeV/nucleon, penetration depths for oxygen into Fe metal are indeed on the order of 3 |im. There appears to be an issue, however, regarding the relative contributions of SW and SEP in lunar samples. The solar wind contribution should be much larger in lunar materials compared to the rare events responsible for SER It appears that SW is preferentially lost relative to the SEP contribution, however; for example, the SEP contribution of noble gases composes 10-20% of neon, argon, and krypton in lunar soil 68501 (Becker and Pepin 1994). Possible mechanisms for this loss include ablation and diffusion. The SW contribution in the 79035 metal grains is masked because of the presence of the oxide layers. Only when the oxide contribution declines with depth does the SEP component become apparent. Its presence in only very few of the grains suggests that not all grains experience the same exposure conditions. The oxygen isotopic compositions of two metal grains from 10084 reveal high 1 7 0 and O at depths consistent with solar wind implantation. Two other grains had oxide layers and these analyses were discontinued after isotopically normal A 1 7 0 values were obtained. Ireland et al. (2007) report additional analyses of metal grains from the Apollo 16 and 17 sites. In lunar soil 61141, several grains with negligible oxide layers and several with thicker oxide layers were analyzed. None of these grains has nonterrestrial A 17 0. In lunar soil 78481, oxide layers were apparent and little deviation from terrestrial A 1 7 0 was noted. These data suggest that the presence of implanted oxygen may be specific to individual grains within individual soils. ls

To elucidate the discrepancy currently observed in solar oxygen studies of lunar samples, it is imperative to make clear the inventory and time variations of extraselenial oxygen fluxes supplied to the lunar surface. 3

SRIM, an acronym for Stopping and Range of Ions in Matter, is a widely used computer program and is available at http://www.srim.org.

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The implications of the two different reports of the solar oxygen isotopic composition, which are at present difficult to reconcile, are treated separately. Implications of an uO-rich solar composition. In this section, the value of HC05 is adopted for discussion, whereas the discussion based on the result of IHNC06 is done in the following section. In Table 1, the solar isotopic compositions are compiled, and are compared with the isotopic compositions of the terrestrial upper mantle, as an example of a large and well-studied planetary reservoir possibly representing, or at least not far from, the average of solid materials in the inner Solar System. The 8 salal values, the deviation of the average planetary isotopic composition relative to the solar composition, are calculated from these values. The true 8 salar values may be even larger, since the solar isotopic ratios obtained from the lunar samples are upper limits. For example, if we adopt the Jovian atmospheric isotopic composition of nitrogen (S^Njjj = - 3 8 0 ± 120%o), as the solar value instead of the upper limit value obtained from lunar samples ( S ^ N ^ < - 2 4 0 ± 25%o), the 8 15 N solar value will be 2 times larger than the one given in Table 1. Here, to enable comparison of the 8 salar values among carbon, nitrogen and oxygen, we adopt the lunar-based values. The 8D solar value, however, is calculated based on the Jovian atmospheric value (D/H = (2.0 ± 0.35) x 10~5; Geiss and Gloeckler 2003), representing the initial solar value before the Sun destroyed its deuterium. This is because the protosolar D/H ratio is not preserved in the present (main sequence) Sun or in lunar soil samples, whereas the solar photosphere is representative of the whole Sun for carbon, nitrogen and oxygen isotopic composition (Kallenbach et al. 2003). In Table 1, there are large differences between the planetary and solar isotopic compositions. It was argued for a long while that the observed range of oxygen isotopic composition among meteorites is primarily due to contributions to the planetary solid materials of presolar components with strikingly different isotopic compositions (e.g., Clayton 1993). However, the available data on oxide and silicate presolar grains do not favor this hypothesis. The majority of presolar oxide and silicate grains are enriched in 1 7 0 but depleted in l s O compared to solar composition (reviewed by Zinner 2007, and Meyer et al. 2008). The major population of presolar SiC grains is enriched in 14 N (Zinner 1998). Nanodiamonds separated from carbonaceous chondrites may be of presolar or of partial Solar System origin (Dai et al. 2002), and also are 14 N-rich (Zinner 1998). The observed differences in the oxygen and nitrogen bulk

Table 1. Comparisons of elemental and isotopic abundance ratios between planetary and solar compositions. Isotopic ratios (Conventional ô values, in %c)2

Elemental abundances 1 (log10 values) Meteorite - Photosphere 3

R

(%o) 4

Solar

Earth mantle

is ' 'stillic

H

-3.75

D/H

- 8 7 0 ± 205

-80

+6100

N

-1.53

15N/14N

< - 2 4 0 ± 256

-5

+3105

-0.99

12

6

-6

+1105

-0.27

17

0 /

16

0

+3

+455

18

0 /

16

0

+6

+475

C

o

C/

13

C

< - 1 0 5 ± 20 0.047 predicted at equilibrium below 1200 K must occur by Fe-Mg interdiffusion in the olivine. Although the condensation time-scale in the reconnection region is only ~1 day, the diffusion coefficients are much higher in this temperature interval than at the lower temperatures considered above. As a result, the mean XFll of 0.1 |im radius grains reaches ~0.09 before diffusion stops, reasonably close to the mean XFa of chondrule precursors. Alternatively, if rapid cooling in such a region caused olivine condensation to be delayed by ~235 K due to supercooling, grains with the required mean XFa could have condensed directly out of the gas, and diffusion of additional Fe2+ would be unnecessary. Caution should be exercised in concluding that this is the solution to the problem, however, as this prediction is based on application of equilibrium thermodynamics to condensation in an environment subjected to repeated ionization by solar flares, where the equilibrium assumption would not be expected to hold.

•

Photolysis of CO(gj. Clayton (2005) suggested another idea, also based on interaction of the protosun with matter in the innermost part of the nebular disk. He argued that mass-independent variations of oxygen isotopic compositions observed in refractory inclusions and chondrules may be the result of isotopic self-shielding during photodissociation of CO(g) by UV radiation from the young Sun. According to this model, at some distance from the Sun, a nebular region would exist where destruction of CO(g) would produce monatomic oxygen greatly and approximately equally enriched in 1 7 0 and l s O relative to le O, and that this could be the source of the oxygen which exchanged with nebular condensates and chondrules to produce isotopic compositions at the high- ls O end of the CCAM line. Clayton (2005) suggested that the observed correlation between FeO/(FeO + MgO) ratio and enrichment in 17 O and ls O within individual chondrites and among the different chondrite chemical classes is a consequence of a process that affected both parameters. In this view, photolysis of CO(g) molecules is not only responsible for the mass-independent oxygen isotopic variations but also for creating a much more oxidizing environment than a normal solar gas with equilibrium speciation, not by increasing the PU2C/PH2 r a t i ° significantly but by increasing the partial pressure of monatomic oxygen by a large factor. Lyons (2006) calculated the photolytically induced variation of the relative abundances of gas species in the system H-O-He as a function of height above the nebular midplane

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Grossman et al. that would be expected at a heliocentric distance of 3 AU at 800 K and 1.6xl0~ 6 bar for an enhancement of the protosolar UV flux by a factor of 103 relative to the modern Sun. While virtually all of the oxygen is present as H 2 0 at the midplane, its photolysis products increase in abundance with height, until virtually all of the oxygen is monatomic above a height of 0.6 AU.

Evaluating these ideas is difficult. Ebel and Grossman (2001) investigated the relative thermodynamic stabilities of condensates in a system of solar composition in which no polyatomic gas molecules were allowed to form. Entirely different condensates form at high temperature, including more Si0 2 -rich and more FeO-rich minerals, than those seen in equilibrium calculations. Although this is a poor analogy to the situation envisioned by Clayton (2005) in that thermodynamic equilibrium between gas and minerals was assumed in the absence of equilibrium gas-phase speciation and no ionized species were considered, it does demonstrate that completely different mineral assemblages could prevail in systems far from equilibrium, including those that contain high FeO at high temperature. Change of FeO/(FeO + MgO) during chondrule melting Reduction of FeO-bearing, non-equilibrium condensates. For the case of enrichment in oxygen by evaporation of water ice from radial migrators, what would the continued evolution of the system be like beyond the point where the magnesium silicates fail to equilibrate with the gas because of slow diffusion? One could imagine a nebular region containing a fine-grained mixture of forsterite, enstatite, feldspars [(Ca,Na)Al(Al,Si)3Os], feldspathoids and metallic nickel-iron. As the temperature continues to fall, troilite (FeS) condenses, coating some of the metal grain surfaces, and the magnesium silicates become more unstable. Eventually, the water will condense as water ice and/or by hydrating magnesium silicates, and may oxidize exposed metal grain surfaces, forming considerable amounts of FeO. If clumps of such relatively oxidized material were then heated, further mineralogical evolution of this material would depend on the temperature to which it was heated, the time spent at high temperature and the composition of the surrounding gas. If it were melted in vacuum for an hour or less, for example, fayalitic olivine would be one of the resulting phases, as seen in some of the Wang et al. (2001) evaporation experiments. If, instead, it were melted in the gas from which it condensed, it might eventually reach an equilibrium state characterized by the relatively low/o 2 and XFa corresponding to the temperature at which it was heated on the curves labeled "ice" in Figures 5 and 6, respectively. If, however, the time spent above the solidus were short compared to the reduction time, olivine with a higher XFa than the equilibrium value would be a likely product. Everything else being equal, the higher the fQl of the ambient gas, the slower the rate of reduction would be, and the longer the heating time that would allow persistence of relatively fayalitic olivine. Unmelted, or relict, olivine grains discovered in some chondrules in primitive chondrites by Nagahara (1981) and Rambaldi (1981) have higher XFa than olivine that apparently crystallized from their host chondrule melts. Rambaldi suggested that the tiny metallic Fe inclusions found inside these so-called "dusty" olivine grains are a product of partial reduction of the FeO in the relict olivine, showing that the FeO/(FeO + MgO) ratios of the precursors of at least some chondrules in primitive chondrites were modified during chondrule formation, in this case to lower values. Jones and Danielson (1997) measured the amount of metallic Fe in the dusty regions of relict olivine grains. For a final XFa of 0.01 to 0.06, they estimated that the precursors of those grains had values of XFa ranging from 0.06 to 0.11 based on 2 vol % metal, the minimum metal content they found, and from 0.26 to 0.29 based on 9%, the maximum. Leroux et al. (2003) synthesized dusty olivine by heating olivine with XFa = 0.16 at 1610 °C and log/o 2 = -15.2, 0.5 log units below that of a system of solar composition at the same temperature. Due to reduction of FeO, which formed )im-sized blebs of metallic Fe, XFa of the associated olivine fell to 0.058 in five minutes and 0.001 in 100 minutes, time-scales comparable to chondrule cooling times.

Redox Conditions

in the Solar

Nebula

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What fraction of the precursor material of each chondrule and of the chondrule population as a whole underwent this process is unknown. Johnson (1986) assumed, as did Wood (1967), that most iron in chondrule precursors was oxidized, that chondrule formation occurred by melting in a relatively reducing nebular gas, that individual chondrules cooled before their FeO could be totally reduced, and that the very wide range of XFa exhibited by the chondrules in individual UOCs and carbonaceous chondrites is a record of the range in the degree of reduction experienced by those objects. Johnson (1986) performed precise electron microprobe analyses of coexisting olivine and low-Ca pyroxene in the chondrules of five different chondrites, assumed that the lowest-FeO grains she found had reached equilibrium with the surrounding gas, and calculated its f0l from Equation (13). At 1500 K, the lowest and highest log f0l's are -17.3 in Chainpur and -15.4 in Krymka, respectively, corresponding to 0.9 and 2.8 log units above that of a solar gas at the same temperature. Relaxation of the assumption that the FeO/(FeO + MgO) ratio remains invariant during chondrule formation shifts the question away from how to make systems rich enough in oxygen to produce fayalitic olivine precursors. The question now becomes how to make the gas compositions with which chondrule melts approached equilibration. Plotting Johnson's (1986) results on Figure 5 shows that the required /o 2 for the Chainpur chondrules is readily achievable in both the water ice-enriched system and the system enriched in OC dust by a factor of 120 at 10~3 and 10~5 atm, but the Krymka data can be accounted for only in the latter system at 10~5 atm. The facts that silicate liquid is unstable in all three systems, metal is unstable in the dust-enriched system at 10~5 atm and that olivine of any composition is unstable in the ice-enriched system at 1500 K serve to underscore the importance of the kinetics of mineral transformations and silicate evaporation to theories of this kind. Nevertheless, the theory of Wood (1967) and Johnson (1986) seems not only capable of accounting for the wide range of XFll in chondrules but also requires systems enriched in oxygen by amounts that are within the range of those produced in nebular transport models. Whether the rate of redox equilibration of chondrule melts with their ambient gas relative to the rates of metal and silicate evaporation during typical chondrule thermal histories is permissive of this kind of theory remains to be determined (Fedkin et al. 2006). In a TEM study of the matrix of ALHA77307, the least equilibrated C 0 3 chondrite, Brearley (1993) found regions of siliceous amorphous material containing abundant, 0.1 to 0.3 |im olivine grains whose XFll varies from 0.09 to 0.69, along with similarly-sized grains of kamacite, pyrrhotite and magnetite. If nebular formation of fayalitic olivine can only be accomplished by incomplete reduction of FeO in chondrule melts, all of this matrix olivine would have to have been derived from chondrules as well, a phenomenon that is hard to imagine. Although indigenous hydrated silicates are absent from Brearley's assemblage, it may be related to the non-equilibrium, low-temperature condensates postulated above, and might be a good candidate for an oxidized precursor for chondrules. Brearley (1993) suggested that the olivine grains in this material formed by incipient annealing of the amorphous phase prior to accretion. This again points to the need for a very oxidizing nebular gas. Oxidation of reduced chondrule precursors. Jones (1990) argued that the partially resorbed forsterite cores of some olivine crystals in Type II chondrules in Semarkona are relict grains. If true, this indicates, in contrast to the above examples, that at least one precursor component of these chondrules may have formed at lower f0l than that which prevailed during crystallization of the bulk of the material in their host chondrules. This leads to the alternative hypothesis that, in these chondrules, the olivine crystals with the highest, rather than the lowest, fayalite contents are the ones that approached high-temperature equilibrium with the surrounding gas most closely. As discussed above, this would require enrichments in oxygenrich (CI) dust relative to gas by factors >500 compared to solar abundances, much greater than can be produced in radial and vertical transport models. In the context of a shock-wave model for chondrule melting, Cuzzi and Alexander (2006) calculated that chondrule-sized particles would have had to be enriched by factors of 300-500 in order to suppress isotopic fractionation

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of relatively volatile elements. It is interesting that the magnitudes of these enrichments approach those needed for nebular equilibration of assemblages consisting of silicate liquid + olivine with XFa >0.15, but these workers found that only a small proportion of chondrules can be so concentrated by their favored mechanism, turbulent concentration. Zanda et al. (1994) found that the concentrations of Si and Cr in metal correlate directly with one another and inversely with the FeO content of olivine in the chondrules of primitive chondrites, and suggested, as did Scott and Taylor (1983), that the distributions of these elements between metal and silicate were established during chondrule formation. Using the dependences of these distributions onfQl, Zanda et al. (1994) found that values of log/o 2 derived from each of the three elemental distributions were in reasonable agreement with one another, and varied from -8.4 to -12.2 from one chondrule to another at a reference temperature of 1873 K. This range of values is from 6.3 to 2.5 log units more oxidizing than a system of solar composition at the same temperature. Similarly, a value of -12.5 was calculated at the same temperature from the Si content of a metal grain in Bishunpur, which was inferred by Lauretta et al. (2001) to have come from a Type I chondrule. Based on experimental simulations of chondrule melting and crystallization, Connolly et al. (1994) suggested that the reducing agent responsible for both dusty olivines and the observations of Zanda et al. (1994) was reduced carbon that was present in the chondrule precursor material. If so, the f 0 l ' s estimated from these studies may represent internal/o 2 's generated during chondrule formation, and may thus have little or no relation to the/o 2 of the surrounding gas, such as might be expected if the timescale for chondrule formation were too rapid for significant gas-droplet interaction. The degree to which reducing agents in chondrule precursors affected the compositions of chondrule minerals and estimates of/o 2 's derived from them depends on the amounts of reducing agents initially present, and these are unknown. Lauretta et al. (2001) sought to determine the gas composition responsible for formation of fayalite + troilite assemblages on the edges of kamacite grains in the Bishunpur chondrite, using thermodynamic calculations and kinetic constraints. At the derived temperature of ~1200 K, P"" of ~10 -5 bar and elemental composition of the gas, the equilibrium P ^ c J P ^ ratio is ~0.27. This value, very far off the scale of Figure 4, is well beyond the values achievable in any of the nebular transport models considered here.

REDOX CONDITIONS INFERRED FROM OTHER IRON-BEARING NEBULAR MATERIALS Amoeboid olivine aggregates Amoeboid olivine aggregates are fine-grained mixtures of Ca- and Al-rich phases, such as spinel and aluminous diopside, with olivine and metallic nickel-iron. These objects have been interpreted as aggregates of nebular condensates (e.g. Grossman and Steele 1976; Krot et al. 2004a,b) but, since Ti3+ has not been detected in their clinopyroxene, it is not certain that they condensed in a system of solar composition. According to Krot et al. (2004b), most olivine in AOAs in the least metamorphosed C 0 3 chondrites, ALHA77307, Acfer 094 and Adelaide, has XFa < 0.02 but values up to 0.15 are known from ALHA77307. The lowest XFa reported in AOAs in Acfer 094 is -0.002 (Krot et al. 2004a). An XFa as low as 0.002 is only twice as large as the maximum that can be expected for olivine that formed in a system of solar composition (Fig. 7a), but this value is based upon 106-yr cooling times and grain radii of only 0.1 |im. Petrographic study of AOAs in these objects reveals patches of olivine up to 5 |im in diameter (Chizmadia et al. 2002). Unless these patches underwent post-condensation grain coarsening or FeO metasomatism, an origin by condensation in a system of solar composition can only be entertained if they are composed of crystallites « 0 . 1 |im in radius.

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Metal grains in CH chondrites Using a combined condensation, grain growth and diffusion model, Petaev et al. (2003) and Petaev (2006) calculated fits to concentration profiles of major, minor and trace elements in zoned grains of metallic nickel-iron in CH chondrites. In the latest calculations (Petaev 2006), the best fits to the profiles in a metal grain in PAT91456 were obtained under the assumption that it is a direct condensate from a system enriched in dust relative to gas by only a factor of 3 compared to solar composition. From the dust composition used by Petaev et al. (2003), a plausible case can be made from this calculation that the grain condensed from a system whose /o 2 was one full log unit greater than that of a system of solar composition. This conclusion is non-unique, however, as it depends on a number of assumptions, on physico-chemical data such as diffusion coefficients and activity coefficients which are uncertain, as well as on a number of adjustable parameters such as P"", degree of isolation of various condensate phases, cooling rate and size of grain nucleus which are unknown. If, on the other hand, the grain could be shown to be part of an equilibrium assemblage containing an element in two valence states whose activity ratio had been experimentally calibrated as a function of fQl, the/o 2 would be known with such certainty that its value could be used to constrain the adjustable parameters. FORMATION CONDITIONS OF ENSTATITE CHONDRITES Mineralogy of EH3 enstatite chondrites The EH3s are the least equilibrated enstatite chondrites. According to Nehru et al. (1984), Parsa, a typical member of this group, contains ~57 wt% enstatite, 17% metallic nickel-iron, 14% troilite, 6% plagioclase, 4% forsterite, 1.4% silica (Si0 2 ), 0.6% schreibersite (Fe3P), 0.6% daubreelite (FeCr2S4), 0.4% oldhamite (CaS) and 0.1% niningerite (MgS). Using mineral composition and modal abundance data for several EH3 chondrites, Wasson, Kallemeyn and Rubin (unpublished data) calculated that 64% of the Ca is in oldhamite, and 4% of the Mg in niningerite. Lusby et al. (1987) showed that approximately half the pyroxene grains in these meteorites contain 0.90 (Fig. 9). It is possible that the thermodynamic data for some of these relatively exotic phases are uncertain by sufficient amounts that stability fields for CaS and MgS may exist at a C/O ratio where enstatite formation occurs at a higher temperature. Formation conditions of E H 3 enstatite chondrites If the constituents of EH3 chondrites condensed in a system with a C/O ratio of 0.83, log /q2 was IW-8.9 before removal of high-temperature condensates, but varied from ~IW10.1 at 1200 K to ~IW-13 at 900 K after grain removal increased the C/O ratio of the residual gas (Fig. 10). Formation of such a system requires enrichment of a nebular region in carbon relative to oxygen by 66% compared to solar composition. It is interesting to note that each of the mechanisms described above for enriching nebular regions in oxygen relative to carbon also creates complementary regions depleted in oxygen relative to carbon. While it is tempting to associate the condensation site of EH3 chondrites with these complementary regions, the latter cannot be responsible. The mechanism for oxygen depletion is condensation of oxygen into silicates at relatively low temperatures in the coagulation and settling models, and into water ice at even lower temperatures in the radial migrators case, followed by transport of the oxygen-containing condensates to other regions in both cases. The compositions of the oxygen-depleted regions would have been even more strongly depleted in all other condensable elements than they would have been in oxygen, and thus could not have supported condensation of EH3 chondrites. The most frequently encountered ferrosilite content in pyroxene from EH3 chondrites is ~1 mol%, with values ranging as high as an amazing 34% (Weisberg et al. 1994). Ferrosilite contents in excess of trace levels are completely incompatible with the formation conditions derived here based on the Si content of the metal and the presence of CaS and MgS, raising the possibility that different constituents of these meteorites formed under different redox conditions, possibly in different nebular regions. Lusby et al. (1987) and Weisberg et al. (1994) presented petrographic evidence for partial reduction of the FeO-rich pyroxene into metallic iron and pyroxene with lower FeO content, implying that the relatively FeO-rich pyroxene formed

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under oxidizing conditions but, at a later stage, was exposed to more reducing conditions. Ebel and Alexander (2005) performed condensation calculations in a system enriched in "C-IDP" dust, an analog to anhydrous cluster IDPs, identical to the OC composition used in this work but with sufficient carbon added to yield an atomic C/Si ratio of 0.7561. In a system enriched by a factor of 1000 in C-IDP dust, the C/O ratio is quite low, ~0.20, and partial pressures of condensables are quite high, causing FeO-rich olivine and pyroxene to condense with silicate liquid at high temperature. As the temperature falls, however, a large fraction of the oxygen condenses in silicates while carbon remains in the gas, and the fQl falls sharply, leading to reduction of previously condensed FeO and condensation of oldhamite and niningerite above 1000 K. At high temperature, the curve of log/o 2 vs. temperature lies well above the curve for C/O = 0.83 shown in Figure 10 but approaches the latter with falling temperature and is nearly coincident with it between 1400 and 1150 K. Thus, Ebel and Alexander (2005) found a novel way of accounting for the two-stage redox history recorded by pyroxene in EH3 chondrites without appealing to mixing of materials formed in distinct nebular regions with different /o 2 's but the magnitude of the dust enrichment required to do so is much higher than has been produced in vertical transport models, and it is not known whether removal of FeO-bearing, high-temperature condensates could yield a residue with the combination of bulk Mg/Si and Fe/Si ratios observed in EH3 chondrites.

CONCLUSIONS Crystallization experiments on liquids with compositions similar to those of compact Type A, Type B1 and Type B2 refractory inclusions were conducted under controlled temperature and/o 2 conditions. Application of the results to the compositions of coexisting Ti3+ -bearing fassaitic clinopyroxene + melilite pairs in natural examples of these inclusion types shows that, if they crystallized at ~1509 K, they did so at log/o 2 = -19.8 ± 0.9. This is only slightly below the equilibrium log/o 2 calculated for a partially condensed system of solar composition at the same temperature, -18.1^3, or IW-6.8. Fassaite is the only f 0 l indicator that shows that anything in chondrites formed in a system that was close to solar in composition. Solar composition is so reducing that metallic nickel-iron is the stable condensed form of iron over a very wide temperature range. Equilibrium calculations predict vanishingly small amounts of condensed FeO until temperatures fall below 800 K, where the FeO/(FeO + MgO) ratio of the condensate begins to increase very gradually with falling temperature by oxidation of metallic iron and formation of fayalite in solid solution in previously condensed forsterite. The mechanism for the latter process is diffusion of Fe 2+ through the crystal structure of forsterite, but the diffusion rate is nearly zero at these temperatures. By comparison to what is achievable in a system of solar composition, the mean FeO/(FeO + MgO) ratio of the olivine in chondrules in unequilibrated ordinary chondrites is very high, ~0.15. Making such ratios in chondrule precursors by solar nebular processes requires sufficiently high/o 2 for iron to become oxidized above temperatures where diffusion of Fe 2+ becomes very low. In an attempt to do so, dynamic models have been proposed for enrichment of oxygen relative to carbon and hydrogen in specific nebular regions. Two such models were investigated quantitatively in the present work: radial transport of water ice-rich migrators across the snow line into the inner part of the solar nebula where the ice evaporates; and coagulation, vertical settling and evaporation of anhydrous dust in the median plane of the inner nebula. In both cases, the maximum achievable fQv ~IW-4.5, produces a maximum XFa before diffusion ceases that is a factor of >7 less than would be required for UOC chondrule precursors, even for grains only 0.1 |im in radius and nebular cooling times as high as 106 yr. The same dynamic models are also incapable of creating environments sufficiently oxidizing to produce olivine with XFa = 0.15 during formation of chondrules by melting of FeO-poor precursors. If, instead, chondrule precursors were made of very FeO-rich, non-equilibrium condensates, reduction of chondrule melts by nebular gas

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may have been arrested before the meanX F a of chondrule olivine could fall below 0.15 because chondrules were hot for such a short time. In contrast to the chondrules in unequilibrated ordinary chondrites, a nebular origin for the mineral assemblage of unequilibrated enstatite chondrites requires/o 2 significantly below that of a system of solar composition. In particular, after fractionation of specific amounts of predicted high-temperature condensates, equilibrium condensation in a system whose P"" = 1 0 ^ atm and whose initial composition is solar except for a C/O ratio of 0.83 yields an assemblage characterized by a very large enstatite/forsterite ratio, the presence of oldhamite and niningerite, metallic nickel-iron containing several wt% Si, and small amounts of pure silica and albitic plagioclase, very similar to the mineral assemblage of EH3 chondrites. L o g / o 2 in this system varies from IW-8.9 at 1500 K to IW-13 at 900 K. The mechanisms proposed above for enriching nebular regions in oxygen relative to carbon produce other regions with complementary fractionations, but the latter are unsuitable sites for high-temperature condensation of EH3 chondrites because the very mechanism that separated oxygen from these sites left them vastly depleted in all condensable elements due to lowtemperature condensation. It is seen that the mechanisms proposed to date for fractionation of C, O and H from one another are quantitatively insufficient to produce the magnitude of nebular/Q 2 variations needed to account for primitive features of unequilibrated ordinary and enstatite chondrites. Perhaps it is time to consider the idea that the solar nebula inherited spatial heterogeneities in the relative proportions of these elements from its parent interstellar gas cloud, just as it did in the case of isotopic compositions.

ACKNOWLEDGMENTS The experimental part of this paper is part of John Beckett's 1986 PhD thesis at the University of Chicago. The authors are grateful to A. Rubin for permission to use unpublished mineralogical data on EH3 chondrites, and to R. N. Clayton, J. N. Cuzzi, S. Desch, A. El Goresy, J. N. Grossman, I. Hutcheon, K. Keil, H. Palme, and Y. Lin for helpful discussions. The paper benefited from thorough reviews by H. C. Connolly, M. Petaev, and E. R. D. Scott. This research was supported by funds from the National Aeronautics and Space Administration through grants NNG05GG00G (to Lawrence Grossman) and NNG04GG14G (to Edward Stolper), and from the Carnegie Institution of Washington (to Fred Ciesla).

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Palme H, Hutcheon ID, Spettel B (1994) Composition and origin of refractory-metal-rich assemblages in a Ca, Al-rich Allende inclusion. Geochim Cosmochim Acta 58:495-513 Perry RH, Chilton CH (1973) Chemical Engineers Handbook, fifth edition. McGraw-Hill, New York Petaev MI (2006) Modeling major and trace element chemistry of zoned metal grains from the CH and CB chondrites. Lunar Planet Sci XXXVII:1681 Petaev MI, Wood JA (1998) The condensation with partial isolation (CWPI) model of condensation in the solar nebula. Meteoritics Planet Sci 33:1123-1137 Petaev MI, Wood JA, Meibom A, Krot AN, Keil K (2003) The ZONMET thermodynamic and kinetic model of metal condensation. Geochim Cosmochim Acta 67:1737-1751 Prinn RG, Fegley B Jr. (1989) Solar nebula chemistry: Origin of planetary, satellite, and cometary volatiles. In: Origin and Evolution of Planetary and Satellite Atmospheres. Atreya SK, Pollack JB, Matthews MS (eds) University Arizona Press, Tucson, p 78-136 Rambaldi ER (1981) Relict grains in chondrules. Nature 293:558-561 Reed SJB, Ware NG (1973) Quantitative electron microprobe analysis using a lithium drifted silicon detector. X-Ray Spectrom 2:69-74 Rietmeijer FJM (1998) Interplanetary dust particles. Rev Mineral 36:2-1—2-95 RuzickaA (1997) Mineral layers around coarse-grained, Ca-Al-rich inclusions in CV3 carbonaceous chondrites: Formation by high-temperature metasomatism. J Geophys Res 102:13387-13402 Sakao H, Elliott JF (1975) Thermodynamics of dilute bcc Fe-Si alloys. Metall Trans 6A:1849-1851 Scott ERD, Taylor GJ (1983) Chondrules and other components in C, O, and E chondrites: Similarities in their properties and origins. Proc 14th Lunar Planet Sci Conf, Part 1, J Geophys Res 88 Supp:B275-B286 Shu FH, Shang H, Gounelle M, Glassgold AE, Lee T (2001) The origin of chondrules and refractory inclusions in chondritic meteorites. Astrophys J 548:1029-1050 Simon JI, Young ED, Russell SS, Tonui EK, Dyl KA, Manning CE (2005) A short timescale for changing oxygen fugacity in the solar nebula revealed by high-resolution 26Al-26Mg dating of CAI rims. Earth Planet Sci Lett 238:272-283 Simon SB, Davis AM, Grossman L (1999) Origin of compact type A refractory inclusions from CV3 carbonaceous chondrites. Geochim Cosmochim Acta 63:1233-1248 Simon SB, Grossman L (2004) A preferred method for the determination of bulk compositions of coarsegrained refractory inclusions and some implications of the results. Geochim Cosmochim Acta 68:42374248 Simon SB, Grossman L (2006) A comparative study of melilite and fassaite in Types B1 and B2 refractory inclusions. Geochim Cosmochim Acta 70:780-798 Simon SB, Grossman L, Davis AM (1991) Fassaite composition trends during crystallization of Allende Type B refractory inclusion melts. Geochim Cosmochim Acta 55:2635-2655 Simon SB, Sutton SR, Grossman L (2005) Valence of Ti and V in fassaite: A recorder of oxygen fugacity during crystallization of coarse-grained refractory inclusions. Workshop on Oxygen in the Earliest Solar System, LPI Contrib 1278:35 Simon SB, Sutton SR, Grossman L (2007) Valence of titanium and vanadium in pyroxene in refractory inclusion interiors and rims. Geochim Cosmochim Acta 71:3098-3118 Stolper E (1982) Crystallization sequences of Ca-Al-rich inclusions from Allende: An experimental study. Geochim Cosmochim Acta 46:2159-2180 Stolper E, Paque JM (1986) Crystallization sequences of Ca-Al-rich inclusions from Allende: The effects of cooling rate and maximum temperature. Geochim Cosmochim Acta 50:1785-1806 Sutton SR, Karner J, Papike J, Delaney JS, Shearer C, Newville M, Eng P, Rivers M, Dyar MD (2005) Vanadium K edge XANES of synthetic and natural basaltic glasses and application to microscale oxygen barometry. Geochim Cosmochim Acta 69:2333-2348 Vogel IA, Palme H (2004) Activity coefficients of silicon in iron-nickel alloys: Experimental determination and relevance for planetary differentiation. Lunar Planet Sci XXXV:1592 Wang J, Davis AM, Clayton RN, Mayeda TK, Hashimoto A (2001) Chemical and isotopic fractionation during the evaporation of the Fe0-Mg0-Si02-Ca0-Al 2 0 3 -Ti02 rare earth element melt system. Geochim Cosmochim Acta 65:479-494 Wark DA, Boynton WV (2001) The formation of rims on calcium-aluminum-rich inclusions: Step I-Flash heating. Meteoritics Planet Sci 36:1135-1166 Wark DA, Lovering JF (1977) Marker events in the early evolution of the solar system: Evidence from rims on Ca-Al-rich inclusions in carbonaceous chondrites. Proc Lunar Sci Conf 8th, Pergamon, New York, p 95-112 Wasson JT (1985) Meteorites: Their Record of Early Solar-System History. WH Freeman and Co., New York Weisberg MK, Prinz M, Fogel RA (1994) The evolution of enstatite and chondrules in unequilibrated enstatite chondrites: Evidence from iron-rich pyroxene. Meteoritics 29:362-373 Williams RJ, Mullins O (1976) A system using solid ceramic oxygen electrolyte cells to measure oxygen fugacities in gas-mixing systems. NASA Tech. Mem. X-58167

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 141-186,2008 Copyright © Mineralogical Society of America

Oxygen Isotopes of Chondritic Components Hisayoshi Yurimoto Department of Natural History Sciences, Isotope Imaging Laboratory Hokkaido University Sapporo 060-0810, Japan yuri@

ofCRIS

ep.sci.hokudai.ac.jp

Alexander N. Krot Hawai'i Institute of Geophysics and Planetology School of Ocean and Earth Science and Technology University of Hawai'i at Manoa Honolulu, Hawai'i 96822, U.S.A.

Byeon-Gak Choi Seoul National University, Earth Science Education Gwanak-Gu, Silliing-Dong Seoul 151-748, South Korea

Jerome Aleon Centre de Spectrometrie Nucleaire et de Spectrometrie de Masse Bat 104, 91405 Orsay Campus, France

Takuya Kunihiro Institute for Study of the Earth's Interior Okayama University Yamada 827, Misasa, Tottori 682-0193, Japan

Adrian J. Brearley Institute of Meteoritics, Department of Earth and Planetary Sciences University of New Mexico Albuquerque, New Mexico 87131, U.S.A. ABSTRACT We review the oxygen isotopic compositions of chondrite components (refractory inclusions, chondrules, and matrix) and their inter- and intra- crystalline oxygen isotopic distributions. Primary oxygen isotopic compositions, acquired before planetesimal accretion, are easily disturbed by parent-body processes such as aqueous alternation and thermal metamorphism. Primary or original oxygen isotopic compositions of refractory inclusions (Ca-, Al-rich inclusions and amoeboid olivine aggregates) distribute along a slope-1 line on the three-oxygen isotope diagram over the range of -60%c < SnO « 5 l s O < +10%c. The variations suggest that oxygen isotopic compositions of the solar nebular gas temporally and spatially varied between l s O-rich and 1 7 0-, l s O-rich during refractory inclusion formation. On the other hand, primary minerals of most chondrules have small isotopic variations and are enriched in 17 0 and l s O relative to refractory inclusions, suggesting that chondrule formation occurred in " O - , l s O-rich nebular gas. However, rare l s O-rich chondrules have been found, suggesting some 1529-6466/08/0068-0008$05.00

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Yurimoto et al. overlap in the timing of formation of chondrules and refractory inclusions in the solar nebula. Oxygen isotopic compositions of matrix grains distribute along the slope-1 line over the same range as refractory inclusions and chondrules. The similarity in oxygen isotopic distributions suggests that matrix was originally a mechanical mixture of nebular dusts co-generated with chondrules and refractory inclusions. Presolar grains with oxygen isotopic compositions that are clearly distinct from those of solar nebular materials are a trace component of chondrite matrices. Based on these oxygen isotopic characteristics, more than 99.5% of the solid materials in the nebula formed locally in the solar nebula, and the remainder formed in interstellar space. The astrophysical setting of chondrite component formation in the early Solar System is also discussed. Refractory inclusions and ls O-rich matrix dusts formed around the inner edge of the solar nebula. On the other hand, most chondrules and 1 7 0-, ls O-rich matrix dusts seem to have formed elsewhere in the solar nebula. Efficient, large-scale radial mixing of the solar nebular materials may have been an essential process in the formation of chondritic planetesimals.

INTRODUCTION Oxygen is the third most abundant element in the Solar System and the most abundant element of the terrestrial planets. The presence of oxygen in both gaseous and solid phases makes oxygen isotopes (terrestrial relative abundances: l e O = 99.757%, 17 O = 0.038%, and l s O = 0.205%) important tracers of various fractionation processes in the solar nebula, which are essential for understanding the evolution of gaseous and solid phases in the early Solar System. Primitive meteorites, i.e., chondrites, are composed of aggregations of solar nebular materials. The solar nebula contained presolar materials as well as materials formed, metamorphosed or altered in the nebula. These nebular materials were also metamorphosed/altered after chondrite parent body formation. Chondrite parent bodies are believed to represent the first-stages of planet formation. Because they did not melt internally, and metamorphism and alteration were incomplete, chondrites are recorders of early Solar System history and evolution. Oxygen isotopic compositions are normally expressed in 8 units, which are deviations in parts per thousand (per mil, %o) in the 1 7 0 / 1 6 0 and 1 8 0 / 1 6 0 ratios from Standard Mean Ocean Water (SMOW) with 1 7 0 / 1 6 0 = 0.0003829 and 1 8 0 / 1 6 0 = 0.0020052 (McKeegan and Leshin 2001): 8 17 ' 18 O SMO W= [( I 7 4 8 0/ I 6 0) s a m p l c /( 1 7 ' 1 8 0/ 1 6 0) S M O W-I]xl000. Onathree-isotopeplotof 8 l s O vs. 8 1 7 0, compositions of nearly all terrestrial samples plot along a single line of slope 0.52 that is called the terrestrial fractionation (TF) line. This line reflects mass-dependent fractionation from a single homogeneous source during chemical and physical processes that results from differences in the masses of the oxygen isotopes (Fig. 1). The slope 0.52 results from changes in 1 7 0 / 1 6 0 that are nearly half those in 1 8 0 / 1 6 0 because of isotopic mass differences; the precise value of the slope depends on the nature of the isotopic species or isotopologues (e.g., Thiemens 2006). In contrast, O-isotopic compositions of the vast majority of extraterrestrial samples, including primitive (chondrites) and differentiated (achondrites) meteorites, deviate from the terrestrial fractionation line, reflecting mass-independent fractionation processes that preceded accretion of these bodies in the protoplanetary disk. Samples from bodies that were largely molten and homogenized, such as Earth, Mars and Vesta, lie on lines that are parallel to the terrestrial fractionation line. Lunar samples show no detectable deviations from the terrestrial fractionation line. The deviation from the terrestrial fractionation line is commonly expressed as A17OSMOW = 817OSMOW - 0.52 x 818OSMOW- For additional discussion of isotopic abundances and notation, see Criss and Farquhar (2008). Mass-independent O-isotopic variations are also observed among components of any unequilibrated chondrite (Fig. 2). The O-isotopic variation range among components within a chondrite is typically larger than the O-isotopic variation range among bulk chondrites. The variations are due to mass-independent fractionation and plot along a line of slope ~1 called the carbonaceous chondrite anhydrous mineral (CCAM) line (Clayton 1993). The slope of the

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6180 (%. rei. SMOW) Figure 2. Oxygen isotopie compositions of CAI and chondrules from C, O and E chondrites measured by conventional gas mass spectrometry. TF: terrestrial fractionation line. CAI: corresponding to carbonaceous chondrite anhydrous mineral mixing (CCAM) line. [Reprinted, with permission, from Clayton (1993) Annii Rev Earth Planet Sci, Vol. 21, Fig. 2, p. 123, by Annual Reviews www.annualreviews.org.]

CCAM line is slightly variable and usually less than 1.00. A line from unaltered minerals represents a slope of 1.00 and oxygen isotopic compositions enriched in 1 7 0 and l s O relative to this line can be explained by mass dependent fractionation or isotope exchange by aqueous alteration probably on the parent body (Young and Russell 1998; Young et al. 2002). The largest variation of O isotopes is observed in Ca-, Al-rich inclusions (CAIs), a component of chondrites. The mass-independent O-isotopic variations of chondrite components lie along a slope-1 line on the three-isotope diagram. These variations indicate that Solar System solids were not derived entirely from a chemically homogeneous, well-mixed reservoir, but from mixing of le O-rich and 17 0-, ls O-rich reservoirs. The origin of the mass-independent

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O-isotopic variations, or anomalies, in Solar System materials has been a major puzzle for planetary scientists since they were discovered over 30 years ago (Clayton et al. 1973), and is still one of the most important outstanding problems in cosmochemistry. The O-isotopic compositions of chondritic components reflect integral mixing of isotopic anomalies from various stages of Solar System history prior to planet formation. In order to understand each stage of Solar System history from the chondrite records, a stratigraphy of oxygen isotopic compositions based on mineralogy and petrology for individual chondrite components would be valuable. Oxygen isotopic measurements at micron-scale resolution are essential for this task and can be obtained using secondary ion mass spectrometry (SIMS). In this report we review the O-isotopic distributions in chondritic components as seen using |im-resolution O-isotopic measurements by SIMS that have become available in the last 10 years, and discuss implications for the astrophysical settings for the formation of these components.

CHONDRITES AND THEIR COMPONENTS Chondritic meteorites consist of three major components that may have formed at separate locations and/or times in the solar nebula: refractory inclusions [Ca-, Al-rich inclusions (CAIs) and amoeboid olivine aggregates (AOAs)]; chondrules; and fine-grained matrix. CAIs are ~1 |im to ~1 cm-sized, irregularly-shaped or spheroidal objects composed mostly of oxides and silicates of calcium, aluminum, titanium, and magnesium. CAIs are sub-classified into fine-grained (FGIs) and coarse-grained (CGIs) inclusions, according to the grain sizes of the minerals that compose them. The sizes of the smallest CAIs are similar to those of minerals in the fine-grained matrix. FGIs are composed of minerals that are too fine-grained to identify in thin section with an optical microscope, whereas CGIs are mainly composed of coarse-grained minerals that can be identified optically. The critical grain size is on the order of the thickness of a standard thin section, typically about 30 |im. FGIs and CGIs are normally distinctive in the Allende CV3 chondrite, apparently due to severe aqueous alteration. In contrast, no distinct boundary is observed in less aqueously-altered chondrites, such as reduced CVs and COs. However, we use this classification because removing effects of the alteration/metamorphism is essential for micro-scale oxygen-isotopic characterizations, and the effects are mostly distinguishable between FGIs and CGIs. AOAs are physical aggregates of individual grains of forsterite (Mg 2 Si0 4 ), Fe, Ni-metal, and FGI minerals. The grain sizes of each mineral are typically less than 50 |im. Coarsegrained AOAs have not been observed, but there is a continuum in texture and mineralogy between AOAs and FGIs (Itoh et al. 2002). Evaporation and condensation appear to have been the dominant processes during formation of refractory inclusions; subsequently some CAIs experienced extensive melting and partial evaporation (MacPherson 2003) and are considered to be igneous. Chondrules are igneous, rounded objects, 0.01-10 mm in size, composed largely of crystals of ferromagnesian olivine (Mg 2 _ x Fe x Si0 4 ) and pyroxene (Mg!_ x Fe x Si0 3 , where 0 < x < 1) and Fe,Ni-metal with interstitial glassy or microcrystalline materials, or metastasis. Most chondrules have textures that are consistent with crystal growth from rapidly cooled (100-1000 K hr - 1 ) silicate melts (e.g., Hewins et al. 2005). Some chondrules contain relict fragments of refractory inclusions and earlier generations of chondrules, and are surrounded by igneous rims, suggesting chondrule formation was repetitive (e.g., Jones et al. 2005; Russell et al. 2005). Additionally, chondrules in primitive chondrites are commonly surrounded by finegrained rims. Based on these observations, it is generally inferred that chondrules formed by

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varying degrees of melting of dense aggregates of ferromagnesian silicate, metal and sulfide grains during multiple flash-heating events, possibly by shock waves, in the dusty, inner (50 )im (and typically >100 |im across) and altered parts are limited to less than tens of |im in width along grain boundaries and cracks, many fresh and clear crystals survived aqueous/thermal alteration of CV3 chondrites. Complex thermal histories (MacPherson et al. 1988) are suggested by some elemental zoning patterns in crystals, e.g., vanadium in pyroxene (Fig. 27). The CGIs are petrographically classified into Type A, composed mostly of melilite; Type B, composed of abundant pyroxene and melilite; and Type C, containing abundant anorthite (Grossman 1980; MacPherson et al. 1988). Most CGIs have a thin (typically ten to a few tens of |im thick) sequence of monomineralic layers referred to collectively as the Wark-Lovering rim (Wark and Lovering 1977). The Wark-Lovering rim is overlain by an accretionary rim mainly composed of olivine. Wark-Lovering and accretionary rims are believed to be direct condensates from the nebular gas.

Figure 26. Greyscale mineral-image overlapped on a transmitted-light photomicrograph of a Type B1 CGI, HN3 from the Allende CV3 chondrite. A melilite-rich mantle encloses a fassaite-, anorthite-, spinel-rich core. An: anorthite; Mel: melilite; Fas: fassaite; sp: spinel; Alt: alteration products.

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Figure 27. Map of vanadium concentration in a fassaite crystal in Type B1 CGI HN3 from the Allende CV3 chondrite. The fassaite crystal is surrounded by anorthite. Many spinel grains are enclosed in the fassaite and anorthite. As the fassaite crystal grew, its vanadium concentration gradually decreased and abruptly increased, then gradually decreased again. The wavy boundary just inside of the abrupt increase suggests that melting occurred before the second growth. No oxygen isotopic heterogeneity is observed in this crystal (Yurimoto etal. 1994).

Oxygen isotopic heterogeneity is observed among constituent mineral phases in almost all CGIs. In many cases, spinel, fassaite and Wark-Lovering rim minerals are enriched in l e O, whereas melilite and anorthite are enriched in 1 7 0 and l s O (Fig. 28). The oxygen isotopic distribution, however, is not readily explained by crystallization sequences from either liquid or gas (Fig. 29). If CGIs formed in a cooling gas or liquid, the oxygen isotopic composition of the gas or liquid must have changed more than two times during cooling. However, the observed elemental zoning would not be consistent with such a simple cooling model. The time interval between CAI melting and solidification may be determined by 26 A1-Mg chronology, using the short-lived radionuclide 26A1, which decays to 26 Mg with a half-life of 0.73 m.y. (MacPherson et al. 1995). The initial 26A1/27A1 of the Solar System is estimated to be ~5 x 10~5; this is referred to as the canonical value. If CAIs did have complex thermal histories, the oxygen isotopic distribution in each CGI mineral should be more complex than that shown in Figure 28. Recent in situ microanalysis of oxygen isotopes has revealed such complex oxygen isotopic distributions in several CGIs. Oxygen isotopic heterogeneity of CGIs is not only observed between different phases in a CAI, but also between different grains (inter-crystal) of the same phases, as well as within single grains (intra-crystal).

Figure 28. Typical oxygen isotopic distribution among minerals of CGIs. TF: terrestrial fractionation line; CCAM: carbonaceous chondrite anhydrous mineral mixing line. After Clayton ( 1993); Krot et al. (2002a).

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after Wood and Hashimoto (1993) and Stolper (1982)

7R-19-1, a compact Type A CGI 7R-19-1 is a compact Type A inclusion from the Allende CV3 chondrite (Yurimoto et al. 1998; Ito et al. 2004). In it, spinel is enriched in l e O, with 8 l s O = -40%o. Melilite and fassaite show various degrees of 1 6 0-enrichment along the C C A M line, with -40%o < 8 l s O < 15%o and -40%o < 8 l s O < 0%c, respectively. The oxygen isotopic distribution in a single melilite crystal analyzed by Yurimoto et al. (1998) is bimodal (Fig. 30); its oxygen isotopic composition changes abruptly over a distance of less than 1 |im. The profile suggests that the 17 O-, l s O-rich melilite grew on the l e O-rich melilite after incomplete melting of the inclusion precursor. The crystallization sequence of this CGI inferred from the petrography is: l e O-rich spinel; l e O-rich fassaite; l e O-rich melilite; intermediate l e O-rich fassaite; 1 7 0-, l s O-rich melilite; and 1 7 0-, l s O rich fassaite, as a result of multiple heating events (Yurimoto et al. 1998). This crystallization sequence is clearly controlled by a disequilibrium kinetic melting process during multiple heating events, but the mechanism is not fully understood. The oxygen isotopic compositions of partial melts of the inclusion precursor were controlled by exchange with the surrounding nebular gas. The oxygen isotopic composition of the nebular gas changed f r o m 1 6 0-rich to 1 7 0-, l s O-rich. The time interval of oxygen isotope change of the nebular gas is estimated by the crystallization age difference of l e O-rich melilite and 1 7 0-, l s Orich melilite, which is less than 0.4 m.y., based on 26 Al- 26 Mg chronometry (Ito et al. 2007).

E49, a compact Type A CGI The distribution of oxygen isotopes in E49, a compound compact Type A inclusion ~4 m m across, f r o m the Efremovka reduced C V 3 chondrite (El Goresy and Zinner 1994), suggests that this particular inclusion underwent isotopic exchange during crystallization (Aleon et al. 2007). O isotopes in melilite are correlated with both the crystal chemistry of the melilite and location within the inclusion: melilite near the Wark-Lovering rim is nearly pure gehlenite and is l e O-rich. The l e O-excess decreases smoothly toward the interior of the CAI as the M g content increases (Fig. 31). About 300 |im inside the CAI, O isotopes in melilite reach terrestrial values (A 1 7 0~0), though relatively depleted in l e O, while the M g content reaches the average interior range of Ak20-30- By contrast, a spinel-, pyroxene-rich xenolith in the inclusion is systematically enriched in l e O but the pyroxene is locally depleted in l e O, and melilite islands show a broader range of isotopic composition. A compound Type B2 inclusion from Allende, containing isotopically distinct xenoliths, has also been reported (Kim et al. 2002).

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Figure 30. Oxygen isotopie distribution in two melilite crystals of 7R-19-1 CGI. From a linear traverse from point M 2 l to M l 6 in the backscattered electron (BSE) image (top panel), the âkermanite content profile (middle) and oxygen isotopie profile (bottom) are shown. Dashed line indicates the location of a grain boundary between two melilite crystals. The inset in the BSE image is a SCAPS (stacked CMOS active pixel sensor) isotopography (Yurimoto et al. 2003) indicating two-dimensional oxygen isotopie distribution on the position. Mel: melilite. After Yurimoto et al. ( 1998) and Ito et al. (2004).

Âk(mol%) Figure 31. Oxygen isotopic composition of melilite from C A I E 4 9 as a function of location and akermanite content. Black spots: rim melilite (outer 300 |jm); Grey spots: interior melilite. After Aleon et al. (2005a).

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Since the increase in akermanite proportions in melilite starting from the exterior is usually interpreted as the result of fractional crystallization during inward growth of melilite (e.g., MacPherson et al. 1989), and chondrite-normalized rare earth element patterns in E49 are typical of Type A inclusions and are also suggestive of fractional crystallization (Simon et al. 1999), it is likely that the O-isotopic exchange occurred during fractional crystallization. This implies that reservoirs of l e O-rich and 1 7 0-, l s O-rich gases co-existed in the nebula, with transport between the two regions possible. Both E49 and the xenolithic inclusion have approximately canonical initial 26A1/27A1, 4.1 ± 0.6 x 10~5, which indicates that the two Oisotope reservoirs coexisted extremely early in the history of the Solar System.

SS-02, a Type B2 CGI SS-02 is a spherical Type B2 inclusion, 10 m m in diameter, f r o m Allende (Kim et al. 2002). It is a compound inclusion consisting of at least three components: palisade body #1 (PB#1); palisade body #2 (PB#2); and the host. The palisade bodies are areas enclosed in shells of spinel grains. The oxygen isotopic distributions among individual minerals differ among the three components (Fig. 32).

518Osmow(%°) Figure 32. Oxygen isotopie compositions of each component in the SS-02 type B2 CGI from Allende. TF: terrestrial fractionation line; CCAM: carbonaceous chondrite anhydrous mineral mixing line. After Kim et al. (2002).

PB#1 is melilite-rich. Spinel is enriched in l e O, with 8 l s O = - 5 0 to -30%o. Melilite is enriched in 1 7 0 and l s O, with 8 l s O = - 5 to 10%o. Fassaite has variable oxygen isotopic compositions along the C C A M line with 8 l s O = - 4 0 to +10%o. PB#2 consists of spinel, melilite, fassaite and anorthite. Spinel and fassaite are enriched in l e O, with 8 l s O = -30%o, whereas melilite and anorthite are enriched in 1 7 0 and l s O with 8 l s O = -5%o. The host also consists of spinel, fassaite, melilite and anorthite. The oxygen isotopic compositions of the minerals are enriched in l e O but vary along the C C A M line, with spinel: 8 l s O = - 4 0 to - 2 0 % c ; fassaite: 8 l s O = - 3 0 to - 2 0 % c ; melilite: 8 l s O = - 2 5 to -15% c ; and anorthite: 8 l s O = - 2 5 to -15% c . The distinct O-isotopic distributions in the three components suggest that they formed separately as small inclusions in different nebular environments and then accumulated together to form a large CAI. The host formed in the last melting event. The existence of compound CGIs such as SS-02 and E49 suggests that accretion of CAIs was a basic inclusion-forming process. The accretion may have started with FGIs.

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TTV1-01, a Type B2 CGI TTVl-01 is a Type B2 inclusion from the Vigarano CV3 chondrite (Fig. 33) (Yoshitake et al. 2005). Oxygen isotopic compositions of individual minerals plot along the CCAM line. Spinel and fassaite are enriched in le O, with 5 l s O = - 4 0 to -30%o. Melilite is enriched in 1 7 0 and ls O, with 5 l s O = 5 to !5%o. Minerals of the WarkLovering rim and accretionary olivine were originally 16 0-rich, with 5 l s O = —40%,o, but disturbance of the oxygen isotopic composition is observed in parts of melilite grains of the Wark-Lovering rim. Anorthite has a continuous oxygen isotopic distribution along the CCAM line, based on point analyses (Fig. 7 of Yoshitake et al. 2005), but isotopic imaging reveals that anorthite has a bimodal oxygen isotopic distribution, with 5lsO = and 5 l s O = ~0%c, and a sharp boundary less than 1 |im in width (Fig. 33). The continuous distribution of oxygen isotopes in the point analyses is due to beam overlap onto the boundary. We note that continuous oxygen isotopic distribution in a mineral phase must be treated carefully because beam overlap cannot be fully evaluated by point analysis; linear traverses or two-dimensional analysis is essential for evaluation of the true distribution. Yoshitake and Yurimoto (2005) measured Figure 33. Backscattered electron image of a part of the TTV-01 CAI (top) and the oxygen initial 26A1/27A1 ratios for each mineral phase isotopic distribution of the area (bottom). Small in this inclusion. le O-rich spinel and 17 0-, ls Opits in the top image are from point analyses by rich melilite have close to canonical values of SIMS. Line in the bottom image corresponds to inferred initial 26A1/27A1, 5.1±0.7 x 10"5 and the anorthite grain boundary. After Nagashima 5.2±1.0 x 10~5, respectively, and that of le O-rich et al. (2004b). fassaite is4.2±0.5 x 10" 5 . 16 0-rich and 17 0-, ls Orich anorthite show no clear excess, with inferred initial 26A1/27A1 2 m.y. before its host chondrule. For the three other chondrules, either the Al-Mg systematics were reset by the chondrule melting event, or the relict CAIs and host chondrules formed >2 m.y. after CAIs that had the canonical 26A1/27A1. Summary of oxygen isotopic characteristics of chondrules Most ferromagnesian chondrules from primitive (unmetamorphosed) chondrites have small isotopic variations (within 3-4%e in A 17 0) and are enriched in 1 7 0 and l s O (A 17 0 >

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OU U9tl

A hitH-nn'l

Figure 45. Oxygen isotopie compositions of minerals in CAI-bearing chondrites. hib: hibonite; mei: melilite; ol: olivine; sp: spinel,; plag: plagioclase; px: pyroxene; mes: mesostasis; TFL: terrestrial fractionation line; CCAM: carbonaceous chondrite anhydrous mineral mixing line. Data from Krot et al. (2006).

-50

-40

-30

-20 -10 s | 8 o , %

0

10

20

—l%o) relative to amoeboid olivine aggregates and most CAIs (A 17 0 < -20%o). Chondrules enriched in l e O (A 17 0 < —10%c) are rare. The distribution of oxygen isotopic compositions of chondrules suggests that they mainly formed in an 17 0-, ls O-rich nebular gas. The existence of le O-rich chondrules and refractory inclusion-bearing chondrules suggests, however, that chondrule formation was not completely unrelated to refractory inclusion formation.

OXYGEN ISOTOPIC COMPOSITIONS OF MATRIX The origin of chondrite matrix is still controversial, with two main hypotheses. One is that because presolar grains that have extreme isotope compositions are found in primitive chondrites (e.g., Zinner et al. 2003), presolar grains might be the main component of matrix. The other hypothesis is that matrix consists of grains newly formed from gas or from presolar materials in the Solar System. There are two possible places in the Solar System where dust formation or the reprocessing of grains took place: the solar nebula (Larimer and Anders 1967); or the chondrite parent bodies (Brearley 2003). It is worth emphasizing that, with a few rare exceptions, no chondrite is entirely unaltered or completely primitive. For example, the matrix of CM chondrites was reprocessed by aqueous alteration, and phyllosilicates were generated on the CM parent body (Brearley 2003). Recent TEM studies reveal that some meteorites have distinctly primitive matrices, however. The matrix of Acfer 094, which is considered to be one of the most pristine, least altered meteorites, contains 40% amorphous silicates, crystalline silicates including 30% olivine, 20% pyroxene and 10% other (Greshake 1997). The matrices of pristine chondrites such as Acfer 094, ALHA 77307 and Adelaide lack phyllosilicates and consist largely of fine, crystalline silicates that are embedded in amorphous silicates (Brearley 1993). There are two possible origins of amorphous silicates (Jensen et al. 2008): nebular condensation; and formation from the crystalline state by irradiation in the interstellar medium. It is unlikely that they formed by shock-related quenching on a parent body (Greshake 1997). Amorphous silicates are fairly susceptible to alteration, and the preservation of them implies that primary material survives in the matrices of some meteorites.

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et al.

Scott and Krot (2003) summarized the matrix mineralogy of various chondrite groups (Table 2). The matrices of CI, CM, CR, CV3 oxidized, and most CO chondrites contain phyllosilicates that formed by aqueous alteration. The matrices of the reduced CV3s contain minor phyllosilicates (Lee et al. 1996). The matrices of highly pristine chondrites, such as Acfer 094, ALHA77307 and Adelaide lack phyllosilicates, and those meteorites are thought to have been unaffected by aqueous alteration and/or metamorphism. The grain sizes of the matrices of the latter meteorites range from 5 |im down to 10 nm. Below, we mainly focus on matrices of meteorites that lack phyllosilicates. Table 2. Summary of phases in matrices of chondrites (Scott and Krot 2003). The underlined components are thought to have formed by aqueous alteration.

Chondrite

Matrix phases

CI

serpentine, saponite. ferrihvdrite. magnetite. Ca-Ma carbonate, pvrrhotite

CM

serpentine, tochilinite. pvrrhotite. amorphous phase, calcite

CR2

olivine, serpentine, saponite. FeS. pentlandite. pvrrhotite. calcite

CV3 oxidized

favalitic olivine. Ca-Fe pvroxene. nepheline. pentlandite

CO >3.1

favalitic olivine, phvllosilicates. ferric oxide

CV3 reduced

fayalitic olivine, low-Ca pyroxene, low-Ni metal, FeS

C03.0

amorphous silicate, Fo30-98, low-Ca pyroxene, Fe,Ni metal, magnetite, sulfides

Acfer 094

amorphous silicate, forsterite, enstatie, pyrrhotite, ferrihydrite

Adelaide

amorphous silicate, fayalitic olivine, enstatite, pentlandite, magnetite

Existence of submicron silicate grains with extreme non-solar oxygen isotopic compositions Because of its fine grain size, matrix is much more susceptible to alteration by aqueous fluids and heating in asteroids than the other chondritic components. It is well known, however, that submicron presolar grains are commonly found in primitive chondrites (Clayton and Nittler 2004). Submicron (down to 100 nm) presolar silicates with distinct oxygen isotopic compositions are not only embedded in the matrices of extremely primitive chondrites (Acfer 094, ALHA77307, Adelaide and Semarkona), but also in those of normal, petrologic types 2 and 3 primitive chondrites (C, O, and E) (Fig. 46) (Mostefaoui and Hoppe 2004; Mostefaoui et al. 2004; Nagashima et al. 2004a; Nguyen and Zinner 2004; Bland et al. 2005; Kobayashi et al. 2005; Nguyen et al. 2005; Stadermann et al. 2005; Ebata et al. 2006; Tonotani et al. 2006; Nagashima and Yurimoto, in prep). The presolar silicates are olivine, pyroxene and amorphous materials, which are the most abundant species in chondrite matrix. Submicron amorphous silica with extreme oxygen isotopic anomalies, reaching 200 and 50 times the solar 1 7 0/ 1 6 0 and 1 8 0/ 1 6 0 ratios, respectively, have been reported from acidinsoluble organic matter recovered from the Murchison CM2 chondrite (Aleon et al. 2005). This material seems to have been produced by irradiation of solar gas by stellar energetic particles in the early Solar System. Its oxygen isotope compositions are distinct from the CCAM line, and have 1 8 0/ 1 7 0 ratios between 0.6 and 1 (solar 18 0/ 17 0~5.25). The existence of submicron silicate grains in primitive chondrite matrices suggests that oxygen isotopes in the matrices were not completely modified by aqueous alteration. The fact that sub-micron presolar silicates with extreme oxygen isotope compositions are preserved in the fine-grained matrix implies that aqueous alteration could not equilibrate oxygen isotopes of sub-micron grains (at least of 100 nm size) in the matrices of primitive chondrites.

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Isotopes of Chondritic

Components

177

Figure 46. Backscattered electron (BSE) image and the oxygen isotopogram of presolar grain-bearing matrix areas from Vigarano and Acfer 094 (Nagashima et al. 2004a; Nagashima and Yurimoto 2007). Arrows point to presolar grains.

Oxygen isotopic heterogeneity of matrix Although grains with extreme oxygen isotopic anomalies are sparsely scattered in matrix (Fig. 46; their maximum abundance is estimated at 5500 ppm, from interplanetary dust particles (Messenger et al. 2003)), we need to measure oxygen isotopic variations of individual matrix grains that are not presolar. The first attempt has been performed using isotopography, a highprecision oxygen isotopic technique with micron to submicron spatial resolution (Kunihiro et al. 2005a). In a matrix area with no detected presolar grains, there is ~50%e heterogeneity in oxygen isotopic analyses for 5 1 7 0 and 5 l s O at < ~1 )im resolution for the primitive chondrites Vigarano (CV3) and Acfer 094 (Kunihiro et al. 2005a; Nagashima and Yurimoto, in prep) (Fig. 47). The oxygen isotopic variations distribute along the CCAM line (Nagashima and Yurimoto, in prep) (Fig. 48), with a range similar to that observed in refractory inclusions and chondrules. Most relatively refractory matrix minerals, such as olivine and pyroxene, are enriched in 0 and ls O, but a small fraction of relatively refractory minerals is enriched in le O. Most of the refractory minerals, such as spinel and melilite, are enriched in le O. Although there are no clear differences in chemical compositions between 17 0-, ls O-rich and le O-rich samples of the same mineral species, the le O-rich phases are considered to be related to refractory inclusions while the 17 0-, ls O-rich phases may be related to chondrules. The volume fractions of le O-rich phases (~1% of Vigarano matrix, ~5% of Acfer 094 matrix) support such a relationship because they are similar to the abundances of refractory inclusions observed in the chondrites. 17

Let's focus on the 17 0-, ls O-rich submicron grains, which typically compose ~95 vol% of the matrix. The oxygen isotopic heterogeneity of this component is about ±5%c (a) for a 1.2 x 1.2 jinr area of a Vigarano section, which corresponds the measurement uncertainty between pixels (1.2 |im length) of the isotopogram (Kunihiro et al. 2005a). Therefore, the real heterogeneity could be smaller than 5%c. The same heterogeneity of oxygen isotopic composition is observed in the Acfer 094 matrix (Fig. 47). Because the spatial resolution of the isotopogram

178

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et al.

%o

rr 20 -

0

Figure 47. Backscattered electron images (left) and oxygen isotopograms (right) of matrix from Acfer 094 and Vigarano CV3 chondrites, in both sets of images, arrows point to 1 6 0-rich micro-grains of pyroxene and olivine, px: pyroxene; ol: olivine. After Kunihiro et al. (2005a).

is limited to 1 )im and the grain size in the region is from 0 3

(2)

After 0 3 * formation in Equation (1), an energy redistribution ("intramolecular energy randomization") occurs among the vibrations and rotations of the 0 3 * molecule. This redistribution is typically assumed to provide, on the average, an equipartitioning of the excess energy among all the coordinates of the vibrationally hot 0 3 , subject to the constraint of the fixed total energy E and total angular momentum J of this 0 3 * that exists prior to the next collision. The energy redistribution among the coordinates is due to anharmonic couplings of the molecular vibrations of 0 3 * and to coriolis and other couplings of the vibrations and rotations. The standard theory used to treat bimolecular recombination and unimolecular dissociation reactions in the literature is a statistical theory, known as "RRKM" theory (Rice, Ramsperger, Kassel, Marcus) (Marcus 1952; Wardlaw and Marcus 1988; Gilbert and Smith 1990; Holbrook et al. 1995; Forst 2003). Symmetrical isotopomers such as 1 6 0 1 6 0 1 6 0 , 1 6 0 1 7 0 1 6 0 and 1 6 0 1 8 0 1 6 0 have fewer intramolecular dynamical couplings for an energy redistribution (fewer "quantum mechanical coupling matrix elements"), because of symmetry restrictions, as compared with asymmetric 0 3 isotopomers, such as 1 6 0 1 6 0 1 7 0 and 1 6 0 1 6 0 1 8 0. Because of the reduced number of coupling elements in the symmetric isotopomers, we have assumed that the symmetric isotopomers have less redistribution of the energy of the newly formed chemical bond among the other coordinates than do the asymmetric isotopomers. Thereby, the symmetric 0 3 * occupy less "phase space" and, consequently, in terms of unimolecular reaction theory, have shorter

194

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lifetimes than the asymmetric isotopomers. The shorter lifetime means that there is less chance of the symmetric 0 3 * being stabilized by loss of energy in a collision, so the rate of formation of the stabilized molecule is less. This assumption of less energy redistribution in the symmetric isotopomers remains to be tested by ab initio quantum mechanical calculations and by a direct experiment noted below. This property of reduced number of coupling elements for symmetric systems is the same for all symmetrical isotopomers regardless of isotopic masses, since all have the same common symmetry property. The simultaneous formation of symmetric heavy atom isotopomers, 1 6 0 1 7 0 1 6 0 and 1 6 0 1 8 0 1 6 0, dilutes the magnitude of the MIF by about 1/3 but does not eliminate it. In summary, because of a dynamical consequence of symmetry, the vibrationally excited asymmetric isotopomers QOO*, where Q is 1 7 0 or l s O, have approximately equal lifetimes that are longer than that of the symmetric isotopomers, OOO* and OQO*. At low pressures they thereby have an improved chance of being deactivated by a collision and of forming a stable ozone molecule, leading to an equal (mass-independent) excess of the heavy isotopes in the ozone. This symmetry/asymmetry behavior is not restricted to 0 3 but would apply to all triatomic or larger molecules that have the possibility of forming symmetric and asymmetric isotopomers, although its extent will depend on the molecule and the temperature, e.g., it can differ in magnitude for 0 3 *, C0 2 *, Si02*, and O^*. Experimental studies permit tests of these ideas. For example, the universality of the effect among all types of isotopomers is seen in oxygen mixtures heavily enriched in 1 7 0 and l s O (Mauersberger et al. 1999). The effect of pressure on the MIF (Morton et al. 1990; Thiemens and Jackson 1990) has been measured and the theory tested by comparison with the data (Gao and Marcus 2001, 2002). The effect of temperature has been studied (Morton et al. 1990) and at present is perhaps qualitatively understood: the higher the temperature, the higher the energy of the 0 3 *, the shorter its lifetime, the less time there is for energy redistribution in the excited ozone, and so the greater the non-statistical effect and thereby the MIF. Missing, however, is a direct experiment: the dissociative lifetime behavior of the vibrationally excited 0 3 * has not been studied under well-defined, collision-free conditions. Under suitable conditions, 0 3 * could be prepared with a known vibrational energy and its time-evolution could be studied using a two laser "pump-dump" method in a molecular beam. A single-exponential decay of the 0 3 * would indicate full statistical intramolecular mixing of the energy, while a more complex time decay would indicate incomplete mixing ("non-RRKM" behavior) (e.g., Marcus et al. 1984). Different isotopomers of 0 3 * could be similarly studied, together with the effect of increased energy on the distribution of lifetimes. An increased energy is expected to increase the difference in the lifetimes of the symmetric and asymmetric isotopomers, based on an interpretation of the observed effect of temperature on the MIF. There is a large body of experimental data on ozone formation obtained under very special experimental conditions (Mauersberger et al. 1999) in which the isotopic exchange ("isotopic scrambling") is minor. In these experiments, the ratios of recombination rate constants, such as the ratio of k's of [ l e O + 1 8 0 1 8 0]/[ 1 6 0 + 16 O le O], are measured directly. These results show strikingly large isotope-specific quantum effects, very different from MIF but well understood in terms of differences of zero-point energies of the two competing modes of dissociation of a an ozone molecule, aObOcO* O + b O c O and a O b O + cO, e.g., Marcus and Gao (2001, 2002). However, as interesting as these special "exit-channel" effects are in their own right, it has been shown that because of a cancellation, contrary to some reports in the literature, they have no bearing on the MIF phenomenon (Hathorn and Marcus 1999). We have stressed this point, since occasionally the two very different effects, the mass-independent effect of "scrambled" systems and the anomalously large mass-dependent effect for reactions of the type Q + OO —> QOO* —> QOO and QO + O, where Q denotes a heavy isotope, appear to have been confused

Oxygen

Isotope Variation in the Solar

Nebula

195

in the literature. This cancellation arises since both sides of the reaction, Q + OO —> QOO* —> QO + O, contribute to the observable effect, not just one side. Isotopic effects on the formation of ozone in the laboratory have been studied at temperatures ranging from ~100 K to 373 K (Morton et al. 1990) and at pressures from 10~2 to 102 bar (Morton et al. 1990; Thiemens and Jackson 1990). Such conditions are very different from the temperatures and pressures at which CAIs are presumed to have formed. With these remarks as background, we turn next to the MIF in CAIs. Conditions for a chemical MIF in the formation of CAIs Detailed calculations have shown that at 1500-2000 K and 10~7 bar, storage of the heavy O isotopes by purely gas phase recombination reactions in molecules such as Si0 2 and A102 is not possible (Marcus 2004; Chen and Marcus, to be published); gaseous Si0 2 and A102 are too unstable at the high temperatures to serve as storage reservoirs. When formed, they also immediately redissociate at these low pressures, since collisions occur too infrequently to stabilize them before redissociation. Thus, any purely chemical explanation for MIF in the early solar nebula should invoke, instead, reactions on surfaces, such as on existing dust grains. Reaction on a surface instead of in the gas phase changes by many orders of magnitude the adverse entropy effect that occurs in a gas phase bimolecular recombination at low pressures, because of the concentrating effect of the surface (e.g., Marcus 2004). At the same time, by not requiring the existence of a long-lived, unstable triatomic molecule for a storage reservoir in an H2-rich atmosphere, the surface reaction mechanism avoids a second major drawback of an MIF due to purely gas phase reactions (Marcus 2004). A recent study of dissociation of 0 3 on silica surfaces (Chakraborty and Bhattacharya 2003a,b; Janssen 2003) suggests that MIF effects can occur by surface-induced dissociation, lending experimental support for the concept of MIF with surface-mediated reactions. A possible chemical mechanism for MIF in CAIs A purely chemical mechanism for the MIF in the CAIs is sketched schematically in Figure 6, using SiO as an example of a reacting diatomic molecule. In the case of minerals rich in Al, the deposition onto the surface would involve Al atoms, since AlO is expected to be too unstable in the gas phase to be a significant contributor. The Al atom deposition could be followed by reactions such as Alads + O ads A10 ads and A10ads + Oads A102,ads*, where the Oads has arisen from the reaction in Figure 6 or, under a more oxidizing atmosphere, from oxygen atoms, Ogas + surface —> Oads.

Competing Processes on Surface

SiO 160

rich

H,0 V

1

Si0 a d s + 0 a d s

Ca

0

Al

SiO + O H,

170

& 1 8 0 rich

SiO 2 ads

Si

Figure 6. Schematic diagram of competing processes on the surface of a growing CAI mineral.

In the mechanism in Figure 6, the principal step for forming adsorbed SiO on the surface is the reaction of H 2 0 with the surface, the reverse of a possible reaction for accelerating the vaporization of forsterite by reaction with H2. The H2 reaction with forsterite has been studied in the laboratory (Nagahara and Ozawa 1996; Tsuchiyama et al. 1998, 1999). The reason for choosing H 2 0 instead of O as the reactant in the Figure 6 case is based on estimated rate constants and concentrations (Marcus 2004; Chen and Marcus, to be published) using the data on the H 2 -catalyzed vaporization.

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In Figure 6, a vibrationally excited Si0 2 , Si0 2 ads*, is formed on the surface by recombination of Oads and SiOads (or A10 2 ads* from A10ads + Oads). These Oads and SiOads undergo two competitive reactions: (a) recombination on the surface to form the vibrationally excited Si0 2ads *; and (b) entrance into the crystal lattice, together with other species such as Ca, Mg and Al, to form a Ca-, Al-rich mineral. The Si0 2ads * itself undergoes two competing reactions in Figure 6: (a) redissociation into SiOads + Oads; and (b) evaporation into SiO and O in the gas phase. We next adapt the argument used for interpreting the MIF in 0 3 * formation in the gas phase to the system in Figure 6. There are conditions, however, for a viable MIF by a mechanism of this type to produce SiO (gas) + O(gas) that is mass-independently enriched in 1 7 0 and l s O and SiOads + Oads that is mass-independently enriched in l e O upon entering the lattice of a protoCAI. These conditions are expressed with reference to the following reactions: SiOgas H> SiOads

H 2 O gas H> H 2 + Oads

(2a)

SiOads + O ads ** Si0 2ads * (rate constants k3, k_3)

(3)

SiOads + O ads + (Ca,Al)

(4)

CAI (rate constant k4)

Si0 2ads * —> SiOgas + Ogas (rate constant k5)

(5)

Si0 2 ads* + lattice —> Si0 2 ads + lattice (rate constant k6)

(6)

where k3 and k_3 are the rate constants of the forward and reverse, respectively, of Equation (3). The conditions are pre-equilibrium of Si0 2ads *, SiOads, and Oads, corresponding to k3 » ¿4, and k_3 » k5, k6, where the Ca and Al surface concentrations have been absorbed into the definition of Laboratory experiments can test these conditions. Analogs of Eqs. (3) through (5) for an Al-bearing mineral such as MgAl 2 0 4 are A10ads + Oads H> A10 2 ads*

(3')

A10ads + Oads + Mg ads ^ CAI

(4')

A10 2 ads* —> A!Ogas + Ogas

and

Algas + 20 g a s

(5')

The available data on deactivation of vibrationally hot molecules on surfaces are sparse. A study using complex molecules suggests that the efficiency of deactivating vibrationally excited molecules on surfaces decreases with increasing temperature (Flowers et al. 1981). In addition, the intrinsic lifetime of the Si0 2 ads* or A10 2 ads* with respect to dissociation decreases with increasing energy and hence with increasing temperature, according to unimolecular reaction rate theory. Both factors favor the possibility that the desired surface pre-equilibrium Reaction (3) and (3') may occur. There is also the question of time scale. If the overall processes are very slow, some equilibration of the isotopic separation might occur in some form before the species are stabilized by permanently entering the crystal lattice. There is presently no information on this time scale, since a study of the formation of solids from SiO, H 2 0, Ca (atoms) and Al (atoms) at 1500 to 2000 K remains to be made. Consequences of chemical mechanism for MIF in the early water We consider here the consequence of the chemical reaction scheme in Equations (2a)-(5') and Figure 6 for the early water formed in the Solar System. In this scheme there would be an equal enrichment of 1 7 0 and l s O (MIF) in water formed as a byproduct of CAI formation. To see this result we add to the Equations (2a)-(5') the fast Reactions (6') and (7') that would inevitably accompany that scheme: Ogas + H2gas —» OHgas + Hgas

(6')

OHgas + H2gas —> H 2 O gas + Hgas

(7')

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Isotope Variation in the Solar

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197

the O atoms coming from the Oads in Reactions (5) and (5') and Figure 6, and also ultimately from recycling of the SiOgas or A10gas emitted in Reactions (5) and (5'). All are enriched in 11 0 and l s O in an MIF manner. As noted in Appendices B and E of Marcus (2004), Reactions (6') and (7') are very rapid, Reaction (6') occurring at every collision at ~ 2000 K and Reaction (7') occurring at 1 of every 10 collisions. In this way, the heavy O atom enrichment in the early H 2 0 formed in this scheme balances their MIF deficiency in the formation of the CAIs in the scheme. Testing the hypothesis: experiment to test gas phase MIF at high temperature Prior to testing any postulated surface chemical reaction scheme for the MIF effect in solids, it is desirable to know, in what appears to be a simpler experiment, if an MIF is even chemically possible in the gas phase at the high temperatures (1500 to 2000 K) relevant for CAI formation. In laboratory experiments on ozone formation at much lower temperatures (100 to 373 K), the magnitude of the MIF was observed to increase with increasing temperature (Morton et al. 1990). Thus, it is possible that an MIF will exist at high temperatures. If an MIF does not occur in the gas phase at these high temperatures, it is unlikely to occur on a surface at those temperatures. For this reason, an initial exploration of a possible chemical MIF at 1500-2000 K in a gas phase experiment is particularly desirable, since the corresponding surface experiment and its interpretation will be more difficult. In such a gas phase experiment, the newly-formed, vibrationally-excited molecules that have subsequently been collisionally stabilized have to be extracted from the reaction system quickly, before these product molecules redissociate. One example of a possible reaction is CO + O —> C0 2 * in an H 2 -free atmosphere (Bhattacharya and Thiemens 1989), with pressures high enough for there to be some collisional stabilization of the vibrationally hot C0 2 * together with reaction times short enough for the back-reaction C 0 2 + M —> C0 2 * + M to be negligible. Further, the source of the O atoms should be such that any complicating transient species (such as 0 3 ) are absent. One potential source of the O atoms is the photodissociation of N 2 0 to yield N 2 + O. If, as an example, CO is used as the second reactant, the subsequent isotopic scrambling of the O atoms, via O + CQ *» OCQ* *» OC + Q, would eliminate any isotopic fractionation resulting from the photodissociation itself. The recombination reaction CO + O —> C 0 2 has been studied isotopically (Bhattacharya and Thiemens 1989), though not yet under high-temperature conditions. At first glance, a potential objection (that is easily dismissed) to observing an MIF in a reaction such as CO + O —> C0 2 *, is that the desired reaction is electronically spin-forbidden (singlet CO + triplet O —> singlet C0 2 ). The presence of an odd-numbered nucleus, 1 7 0, in a reactant reduces the spin-forbidden impediment for this reaction, due to electron spin-nuclear spin coupling, and so catalyzes the recombination and destroys any mass-independence. However, this effect would occur both in the formation, CO + O —> C0 2 *, and in the redissociation, C0 2 * —> CO + O, of the C0 2 *. Since the redissociation of the C0 2 * dominates over collisional stabilization of the C0 2 * at low pressures, this spin-spin effect favoring formation of C0 2 * containing 1 7 0 also favors this reverse process, relevant here, however, there would be little redissociation of the C0 2 * because of collisional deactivation, and then the spin-spin coupling would indeed favor 1 7 0 enrichment of the C0 2 . To see if any MIF exists in gas-phase reactions at 1500 to 2000 K, it would be useful to explore the low-pressure behavior of a gas-phase reaction such as the above. Further, undertaking such studies at sufficiently low pressures, in conjunction with unimolecular reaction rate theory, provides information in the deactivational effect of the surface at these temperatures. It will be interesting to see from such experiments whether a chemically-based MIF can occur in the gas phase at 1500-2000 K, and also to explore the surface experiments.

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At room temperature, a surface causes the MIF associated with 0 3 formation to disappear at pressures < 7xl0~ 3 bar (Morton et al. 1990). There is thus a contrast with the scheme in Figure 6, since now there is no competition with oxygen entering into a lattice. When in a surface study of ozone formation at high temperatures most of the 0 3 * formed on the surface is deactivated on the surface, then a pre-equilibrium O + 0 2 = 0 3 * needed for an MIF cannot exist. Only when most of the 0 3 * is not deactivated, either on the surface at low pressures or in the gas phase at high pressures, can a pre-equilibrium needed for an MIF exist.

PHOTOCHEMICAL MASS-INDEPENDENT OXYGEN ISOTOPE FRACTIONATION: CO SELF-SHIELDING Isotope-selective photodissociation of molecules (especially CO and 0 2 ) combined with self-shielding has long been regarded as a possible explanation for the anomalous oxygen isotope distribution in Solar System rocks (Kitamura and Shimizu 1983; Thiemens and Heidenreich 1983; Navon and Wasserburg 1985), but the idea languished until recently, primarily because of the likelihood that back reactions among products at high temperatures (e.g., Navon and Wasserburg 1985) would erase the isotopic effects. A revival of the concept as it pertains to CO and N 2 is in full swing, however, as a result of a suggestion by Clayton (2002b) that self-shielding by CO at the inner annulus of the proto-solar circumstellar disk might be the cause of the slope-1 trend of CAI data on plots of 8 1 7 0 vs. 8 l s O. Doubt has been raised about the efficacy of CO self-shielding at the disk inner annulus (Lyons and Young 2003), though more experimental and computational work is required to fully assess the model. Nonetheless, the idea has spawned two additional models. One model supposes that the oxygen isotope effects of CO self-shielding are inherited from the placental molecular cloud (Yurimoto and Kuramoto 2004) while the other considers CO self-shielding at the surfaces of the disk (Lyons and Young 2005a). In either case, the l s O and 1 7 0 liberated by photodissociation of CO is most likely sequestered in water ice (Young and Lyons 2003; Yurimoto and Kuramoto 2004; Lyons and Young 2005a). These models predict that one should see an excess of C l e O relative to C l s O and C 1 7 0 in protoplanetary disks as a consequence of self-shielding by C l e O. The prediction offers the prospect of a direct test of the CO self-shielding hypotheses in their various forms. In this section we will consider all three proposed astrophysical settings for mass-independent oxygen isotope fractionation by CO photodissociation and self-shielding in the early Solar System: 1) the inner annulus of the protostellar disk; 2) the surfaces of the protostellar disk; and 3) the placental molecular cloud of the disk. Although not yet verified by laboratory experiments, CO self-shielding, especially at low temperatures, is an attractive explanation for A 1 7 0 variability in the Solar System because it is a process observed in the interstellar medium and, perhaps, in disks as well (see below). It also has the advantage that a principle by-product is le O-depleted H 2 0 . The latter appears to be an important feature of the distribution of oxygen in the early solar nebula (Clayton and Mayeda 1984; Young 2001, 2007a; Sakamoto et al. 2007).

CO photodissociation and self-shielding Carbon monoxide experiences destruction by photodissociation induced by far-UV (FUV) stellar radiation at wavelengths of 91 to 110 nm. Photodestruction takes place primarily through predissociation. Predissociation involves passage to bound excited states prior to dissociation. Since the excited molecules have well-defined rovibrational states, photolysis by predissociation is isotope-specific. In particular, rovibrational lines in the UV absorption bands of the 1 7 0 and l s O CO oxygen isotopologues are red-shifted by ~25 cm - 1 per amu relative to the lines for C l e O (e.g., Ubachs et al. 2000). The observed shifts are consistent with the prediction based on an assumption of a relatively strong C-O bond in the excited state, where the shifts in frequency are v C o _

Oxygen

Isotope Variation in the Solar

199

Nebula

Vc'o = [1 - ()^cc/)^c'o)1/2]rae> where coe is the vibrational constant and |i; is the reduced mass for isotopologue i. Line spacings for a given isotopologue are on the order of 3.6 cm - 1 as prescribed by the dependence of rotational constants on reduced mass. The result of these isotope-specific shifts is a tight intermingling of lines (e.g., C 1 7 0 R and C l e O P transitions) with separations on the order of 1 cm -1 . This separation can be compared with line widths. If linewidths are less than the separation of intercalated lines, then absorption will be isotopespecific. Excited lifetimes (x) for CO are > 3xl0 - 1 1 s (Eikema et al. 1994) but may be as long as 2xlO~10 s (Ubachs et al. 2000), resulting in natural line widths T (FWHM) of < 0.16 cnr 1 , and perhaps as narrow as 0.03 cm - 1 (T = l/(x 2 n c)). A purely Lorentzian line shape suggests that the wings of a 0.16 cm - 1 FWHF peak extend 7 cm - 1 from the peak center out to intensities of \%o relative to the maximum intensity. Wings for a 0.03 c n r 1 line extend 2.2 c n r 1 from the peak center at the \%o threshold. These figures serve to demonstrate that the FUV absorption spectra of the CO isotopologues are sufficiently distinct that CO dissociation is profoundly wavelength-dependent. As a result of the distinct UV absorption spectra of C l e O, C 1 7 0 and C l s O, CO will absorb FUV wavelengths in proportion to the column densities of the constituent oxygen isotopologues (Fig. 7). Because solar and interstellar l e O/ l s O and 1 6 0/ 1 7 0 ratios are ~500 and ~2600, respectively, spectral lines for C l e O will be more optically thick than C 1 7 0 and C l s O lines (van Dishoeck and Black 1988). This process of saturation of the photodissociating lines of the abundant isotopologue is referred to as self-shielding and is a process known to be important in the interstellar medium (ISM) (Bally and Langer 1982; Sheffer et al. 2002). However, as the separation between intercalated lines from the different isotopologues are separated by distances smaller than the full baseline widths of the peaks themselves, there is mutual shielding of one isotopologue on another; the isotope specificity is not simple (van Dishoeck and Black 1988; Lyons and Young 2005a). More work on the specifics of the mutual

CO photodissociation

self shielding

x c16o = 1

X c18o = 1

x c17o = 1

uv UV source

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Optical depth: x1 = Np^ = J^ ",-CKifc Rate constant: kcl6Q = k0 e x p ( - ^ „ / ^ „ J e x p ( - o . > c „ 0 N c „ 0 ) e x p ( ~ a ^ > o N c » o ) e x p ( - am, ^h, ) exp(

yAv)

Figure 7. Schematic illustration of the process of CO self-shielding. Optical depths for the relevant wavelengths of far ultraviolet light (zig-zag arrows), x, depend upon the wavelength-specific absorption cross sections a>. and column depths of Nj of CO isotopologues /. The column unit optical depths depend largely on the number densities », of the isotopologues. The rate constant for photodissociation of each of the isotopologues, A-„ is a product of shielding functions of form exp(-a>.j Nj) and the visual extinction of dust. A,, modified for the FUV (y). Atomic oxygen liberated by photodissociation of CO between the unit x for C 1 6 0 and C 1 7 0 is enriched in l s O and 1 7 0 relative to the reactant CO.

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shielding effect is required as the slope of the self-shielding effect in oxygen three-isotope space depends on these relationships. By assuming a solar-like ratio of CO to H2 in the gas phase (e.g., gas in an accretion disk or in a molecular cloud in the interstellar medium), one can use the FUV shielding functions for CO (Lyons and Young 2005a; van Dishoeck and Black 1988) to derive a general rule relating hydrogen column density (NH, see Fig. 7) to oxygen isotope-specific photodissociation of CO. The result is that oxygen isotope fractionation by CO photodissociation is expected in regions where NH ~ 1019 to 1022 cm -2 . One can also use typical relationships between NH and extinction of visible light, or Ay, expressed in magnitudes (Am = mag = 2.5 log(I2/Ii) where I2 and Ij are two different intensities of light), by ~ 0.1 micron-sized dust grains (Lee et al. 1996) to show that mass-independent oxygen isotope fractionation by CO photodestruction is expected where A,, ranges from ~ 5xl0~ 3 to 5. Larger-sized dust grains, such as might be found in a disk, reduce Ay in CO self-shielding regions by 10 fold. These ranges in NH and Av serve as general guides for assessing the likely locations for CO isotope-selective photodissociation. Astronomical observations of oxygen isotope fractionation by CO self-shielding Molecular clouds. Carbon monoxide is the second most abundant molecular gas in molecular clouds and also the most abundant oxygen-bearing species. Anomalously high CleO/ C l s O in molecular clouds was first attributed to isotope-selective CO photodissociation and self-shielding by Bally and Langer (1982). Since then, large mass-independent oxygen isotope fractionation effects attributed to CO self-shielding in interstellar clouds have been observed. Indeed, evidence for isotopic fractionation between the relatively abundant isotopologues 13CO and C l s O (composing ~1 and ~0.3% of total CO, respectively, if solar isotopic abundances are used) in molecular clouds has been accumulated for more than two decades (e.g., Frerking et al. 1982; Lada et al. 1994). Mass-independent fractionation of CO in molecular clouds based on observations of C 1 7 0 as well as C l s O has been reported (e.g., White and Sandell 1995; Bergin et al. 2001; Wouterloot et al. 2005). Sheffer et al. (2002), for example, using UV spectroscopy, report 12 C 16 0/ 12 C 17 0 and 12 C 16 0/ 12 C 18 0 approximately five times (i.e., 8 ls O ~ 8 1 7 0 ~ 4000%o) the ambient values for the local interstellar medium on a line of sight towards X Persei. Enrichments in both heavy oxygen isotopologues are consistent with a slope of 1 on a plot of 8 1 7 0 vs. 8 ls O, within large uncertainties. Many such observations are conducted by cross-correlating visual extinction (Ay) with column density of CO isotopologues along the lines of sight. As noted above, Ay reflects the column density of dust grains that are the primary absorber of visual light in a molecular cloud. Hence the visual extinction also provides a measure of the attenuation of interstellar UV in clouds. The column densities of CO isotopologues can be determined from the intensity of mm-wave emission due to the transition of rotational energy state. Each isotopologue has its own rotational energy levels that can be used to obtain column densities for each of the isotopologues. In Figure 8, CO isotopologue ratios observed in the dark cloud IC5146 (Lada et al. 1994; Bergin et al. 2001) are shown as a function of visual extinction. The 13CO/ClsO ratio depends on visual extinction, decreasing with increasing visual extinction with a peak at several mag. This represents the isotopic fractionation across the molecular cloud. On the other hand, the C 18 0/C 17 0 ratio is nearly constant independent of visual extinction, meaning that the isotopic fractionation in CO is "mass independent." Note that the column density of C l e O is difficult to determine directly because the intensity of line emission from this abundant isotopologue is saturated and no longer proportional to its column abundance. Note also that the fractionation between 13CO and 12CO is expected to be significantly diluted because of the exothermic 12C+ + 13CO (van Dishoeck and Black 1988; Warin et al. 1996). The reaction 13C+ + 12CO mass-independent isotopic fractionation of CO in Figure 8 is interpreted to be the result of self-shielding by CO isotopologues (e.g., van Dishoeck and Black 1988; Viala et al. 1988; Warin et al. 1996).

Oxygen

t .

Isotope Variation in the Solar

'*

Nebula

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•

•

Figure 8. The isotopomeric ratios of gaseous CO as a function of visual extinction in the dark cloud IC5146. 1 3 CO/C l s O is after Lada et al. (1994) and C 1 8 0 / C 1 7 0 is after Bergin et al. (2001).

0

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Visual Extinction (magnitude)

Circumstellar disks. Observations of at least one circumstellar gaseous disk provide a hint of C l e O overabundance. HL Tau is a low-mass, pre-main sequence star with a disk that is regarded as a prototype for the solar nebula. The surrounding disk extends out as far as 1400 AU from the star. Accretion is occurring at a rate of 10~5 M® yr -1 (Brittain et al. 2005). Brittain et al. (2005) presented high-resolution infrared spectra of the embedded star showing broad CO emission lines and narrow CO absorption lines. The broad emission lines represent the hot (~1500 K) inner portion of the circumstellar disk while the absorption lines originate from the disk gas at an effective temperature of about 100 K. In general, one expects CO in the distal regions of a circumstellar disk (R > 100's of AU) to represent the canonical oxygen isotope ratios, e.g., C l e O/C l s O = 560 ± 25 (Wilson and Rood 1994). Brittain et al. found instead that the ratio of column densities for C l e O and C l s O, N c i6o/N c ls o, is 800 ± 200. These authors suggest that the overabundance of C l e O (albeit at only the 2 a level) could be the result of isotope-selective photodissociation resulting from the HL Tau UV field. Detection of high C l e O/C l s O in the HL Tau disk (C l e O/C l s O = 800 ± 200 vs. 560 ± 25) in gas in the outer disk is tantalizing evidence that CO self-shielding of stellar or interstellar FUV may be a feature of the chemical evolution of disks. Enrichments in C l e O/C l s O (and C 1 6 0/C 1 7 0) in other protoplanetary disks around young stellar objects have not been identified, but neither have the observations been made with this explicit goal in mind. Sensitive measurements of several lines of all three oxygen isotopologues in several objects of various masses would permit a systematic investigation of oxygen isotope ratios in the disks. One would require high signal-to-noise ratio in order to use line shape information to control for optical depth effects in the main isotope, as well as to make sensitive measures of the two rare isotopes. The pivotal role of H 2 0 The importance of H 2 0 in the isotopic evolution of the solar nebula is suggested by numerous studies showing that A 1 7 0 of water during chondrite formation was substantially greater than that of the rock component (Clayton and Mayeda 1984; Choi et al. 1998; Young et al. 1999; Lyons and Young 2005b; Sakamoto et al. 2007). This is an important observation because water plays a pivotal role as a carrier of le O-poor oxygen in CO self-shielding models. Here we review briefly predictions and evidence for high and variable A 1 7 0 of water in the inner (R < 5 AU) solar protostellar disk.

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Solar abundances of the elements and the kinetics of gas phase reactions indicate that approximately half of the total oxygen in a protoplanetary disk resides in CO. Another onethird is contained in H 2 0 , with the remainder as lithophile oxides (Anders and Grevesse 1989; Fegley 2000). Water plays a central role as the carrier of the le O-depleted oxygen in models for CO self-shielding described below. Clayton and Mayeda (1984) recognized the importance of H 2 0 as a reservoir for le O-poor oxygen in the solar nebula based on the isotopic compositions of secondary minerals in chondrites. Yurimoto and Kuramoto (2002, 2004) pointed out that there should have been a reciprocity in A 1 7 0 between H 2 0 and CO in the protosolar molecular cloud. Lyons and Young later showed that the same isotopic reciprocity likely existed between H 2 0 and CO in cold portions of the protosolar accretion disk high above the midplane (Young and Lyons 2003; Lyons and Young 2005a). Water is the most likely driver for the oxygen isotopic evolution of the Solar System because H 2 0 forms readily from O liberated by CO photodissociation, and because H 2 0 vapor exchanges oxygen isotopes with silicate rapidly while CO does not. Yu et al. (1995) showed that 50% oxygen isotope exchange occurs between molten silicate and H 2 + H 2 0 gas at 1773 K and 1 bar after 5 minutes. These authors argued that at nebular pressures (~10 - 5 to 10~8 bar), 50% exchange would take 10 hours, demonstrating that exchange of oxygen between molten silicate and H 2 0 gas is efficient. Conversely, analogous experiments involving CO and silicate show no evidence of isotopic exchange under similar conditions (Bosenberg et al. 2005). A protoplanetary disk is born as an accretionary gas disk, where gas is transported toward the central star, and subsequently evolves to a quiescent disk. During disk evolution, chemical heterogeneity is caused by gas-dust fractionation within (e.g., Morfill et al. 1985). Fractionation of gas and dust includes the sedimentation of dust grains toward the disk midplane and the preferential inward migration of dust grains. In particular, the latter process may have caused significant change in the mean isotopic composition of the inner disk (Yurimoto and Kuramoto 2004). Radial migration of dust grains occurs in response to gas drag. Gas composing the circumstellar disk rotates around the central star slightly slower than the Keplerian rate at each radial distance because of the outward pressure gradient. Solid grains immersed in the gas tend to follow the Keplerian rotation, resulting in gas drag and loss of angular momentum. Loss of angular momentum causes inward migration. Such inward migration of dust grains occurs during active accretion (Weidenschilling and Cuzzi 1993). Assuming the profile of density and temperature of disk gas follows those of the minimum-mass solar nebula, mm- to m-sized grains migrate most rapidly inward. The scenario is robust and is not sensitive to disk parameters such as surface density profile (Cuzzi and Zahnle 2004; Kuramoto and Yurimoto 2005). Because the disk temperature increases inward, a migrating dust grain releases volatiles upon encountering the appropriate sublimation temperature. Under the realistic range of disk gas density, the sublimation temperature of H 2 0 is about 150 K (Lewis 1972). The annulus region where H 2 0 sublimes is called the snow line. The snow line migrates inward with time as the disk cools. It may have been as close as 3 AU from the proto-sun at some interval based on a minimum-mass solar nebula model (Hayashi et al. 1985). After the passage of the snow line, dust grains leave behind H 2 0 vapor in the surrounding gas which may accrete inward more slowly. This causes enrichment of H 2 0 vapor within the inner disk. Numerical analyses show that enrichments of H 2 0 of more than 2x relative to the solar proportion defined by solar C/O/H are possible inside the snow line, using realistic physical parameters (Cuzzi and Zahnle 2004; Kuramoto and Yurimoto 2005). Numerous studies of meteorites suggest that the inner solar nebula became enriched in 0 , l s O with time (e.g., Choi et al. 1998; Wasson et al. 2004), and there is plentiful evidence for oxygen isotopic exchange between surrounding nebular gas and meteoritic constituents, including CAIs (Yurimoto et al. 1998; Faganet al. 2004; Yoshitake et al. 2005) and chondrules 17

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(Maruyama et al. 1999; Wasson et al. 2004). These observations appear consistent with the scenario of gradual H 2 0 enrichment in the inner disk during disk evolution. Previous studies attempting to use the oxygen isotopie compositions of minerals produced by aqueous alteration of carbonaceous chondrites to infer the original isotopie composition of H 2 0 in the early Solai' System ( Clayton and Mayeda 1984; Choi et al. 1998; Clayton and Mayeda 1999; Young 2001) were hindered by the fact that waters from which these minerals grew were likely to have already exchanged oxygen with rock (Young et al. 1999). The recent findings by Sakamoto et al. (2007) of incipient aqueous alteration in a phase in a carbonaceous chondrite with 5 l s O ~ 5 1 7 0 ~ 180 %o appears to circumvent this hindrance. The oxygen isotope data for aqueously altered chondrites strongly supports the hypothesis that H 2 0 in the inner regions of the solar protoplanetary disk was severely depleted in l e O and enriched in 1 7 0 and l s O. CO self-shielding at the inner annulus of the solar circumstellar disk R. N. Clayton (2002b) argued that self-shielding by CO at the inner annulus of the solar circumstellar disk might be the origin of mass-independent oxygen isotope fractionation in the Solar System (Fig. 9). This hypothesis is based on the assumption that solar 5 1 7 0 and 5 l s O are coincident with the low end of the data trend shown in Figure 1 (i.e., le O-rich) and that the signal incurred by self-shielding by CO could be exported back to distal regions of the disk as material entrained in the X-wind described by Shu et al. (2001). Self-shielding by CO in the region of the X-wind of the protostellar disk is attractive as an hypothesis to the extent that the X-wind itself was responsible for forming and/or modifying primitive solids like CAIs and chondrules in the early Solar System. There are three potential difficulties associated with the scenario in Figure 9. The first is that at the high temperatures that prevail in the i?x region ( T > 1500 K), one might expect line broadening to cause substantial overlapping of the predissociation spectra of the CO isotopologues, and thus eliminate the isotope selectivity. Recent calculations, however, suggest that line broadening may be insufficient to preclude the effect, even at the high temperatures that obtain in the inner annulus of the disk (Lyons et al. 2007). The second possible difficulty is that at high temperatures and number densities, back reaction between liberated O and other

Self shielding by CO near Rx (high T)

Figure 9. Schematic illustration of the Clayton hypothesis for CO self-shielding at the inner annulus of the solar circumstellar disk. Circumstellar radius Rx marks the inner edge of the disk. Here material flowing inward (dark grey arrows) in the disk is periodically ejected from the Rx region back out into the disk in the magnetically-driven "X wind" ( Shu et al. 1996) ( dotted lines). The region of CO self-shielding between the x = 1 surfaces for C 1 6 0 and C 1 7 0 and C l s O, together denoted CQ, (heavy black lines) is calculated to be on the order of km in scale based on typical disk number densities (Lyons and Young 2003).

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species may well eliminate the possibility of preserving the isotopic consequences of selfshielding in a reservoir of oxygen. The preservation of le O-poor O as H 2 0 at high temperatures requires further investigation. The third potential difficulty is that where FUV fluxes are high, CO is destroyed so rapidly that there is no substantial isotope fractionation (see the next section for a fuller explanation of this last issue). CO self-shielding at the surfaces of the solar circumstellar disk Possible difficulties with preserving the isotopic signal of CO self-shielding at high temperatures are overcome where CO photodissociation occurs at low temperatures (e.g., < 100 K), as in the interstellar medium clouds. Within the circumstellar disk that was the solar nebula, the surfaces of the disk at H2 number densities of ~105 to 106 cm - 3 (pressure ~ 10~14 bar) are prime locations for preserving the isotopic effects of CO self-shielding (Fig. 10). There are two sources of FUV at the surfaces of a circumstellar disk. One is ambient UV from the surrounding environs entering the disk at a high incident angle. The other is the central star itself where light travels on a line of sight from the star through the disk. A flared geometry, as shown in Figure 10, is though to be typical of many disks (Kenyon and Hartmann 1987). A disk such as that in Figure 10, with concave upper surfaces relative to the midplane, intercepts more light along a sight line to the central star than a non-flaring disk. This effect increases with distance from the central star (Kenyon and Hartmann 1987) and maximizes the potential for photochemistry within the optically thin surfaces, especially in the outer regions of the disk.

100 -

80 -

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o CO

1, the principal chemical reactions add hydrogen to other atomic species on the ice surface, resulting in simple hydrides like H 2 0, CH 3 OH, NH 3 , and CH4. This results in highly polar ices dominated by H 2 0. If, however, the local gas has H/H2 < 1, gas-grain reactions yield ices dominated by less-polar molecules like CO, C0 2 , 0 2 , and N2. Infrared spectra of the ices in dense clouds support this basic concept (Sandford et al. 1988; Tielens et al. 1991). Provided they were not greatly warmed during their incorporation into comets during formation of the Solar System, it is possible that these icy materials could be directly preserved in comets. As with the products of gas phase ion-molecule reactions, however, these icy compounds would not have efficiently survived in meteoritic materials; further chemical processing was required to "set" the oxygen in these ices into the more refractory forms that survived in meteoritic materials. One process that may play a role in generating more refractory materials is irradiation chemistry (Fig. 2). The ice mantles surrounding dust grains in dense clouds and the protosolar

Oxygen in Comets and Interplanetary

ULTRAVIOLET

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PHOTONS

QUI NONE

ORGANI C REFRACTORY MANTLE

Figure 2. A large fraction of the volatile materials in dense interstellar clouds condense onto grain mantles. These ices are generally rich in oxygen in the form of molecules like H 2 0 , CH 3 OH, CO, and C 0 2 . Cosmic rays and UV photons can break down these molecules into ions and radicals that can recombine to produce considerably more complex, and more refractory, O-bearing organic species.

nebula can be further processed by cosmic rays, UV radiation from the attenuated diffuse ISM field, or UV produced by nearby stars and cosmic ray interactions (Norman and Silk 1980; Prasad and Tarafdar 1983). These forms of radiation can break bonds within molecules in the ice and generate ions and radicals that can subsequently react. This results in additional chemistry that forms a wide variety of more complex, largely C-rich, species. Laboratory studies show that some of the species produced by charged particle and UV irradiation contain oxygen and resemble those found in meteoritic and cometary materials. These include molecular species like amphiphiles, amino acids, and aromatic ketones and alcohols (cf. Dworkin et al. 2001; Bernstein et al. 1999, 2002a, 2003). Of particular interest is the irradiation of PAHs in mixed molecular ices, since this results in the addition of excess H and a variety of chemical side groups (=0, -OH, -NH 2 , -CN, -CH 3 , -OCH 3 , etc.) derived from other species in the ice (Bernstein et al. 2002b, 2003). As shown in Figure 3, in H 2 0-rich ices, i.e., the ices seen to dominate most dense clouds, the principal additions to PAHs are = 0 and -OH groups linked to peripheral C atoms on the PAH, and bridging oxygen spanning "bay" regions (aromatic ethers) (Bernstein et al. 1999). Many of these species resemble those found in meteorites. For additional, more detailed discussion of oxygen in the interstellar medium, see the chapter by Jensen et al. (2008). The oxygen in organics in meteorites, cosmic dust, and comet samples. Some of the oxygen in C-rich carriers in primitive meteorites resides in soluble species like amino acids, amphiphiles, carboxylic acids, etc. (cf. Cronin et al. 1988; Rrishnamurthy et al. 1992; Huang et al. 2005). However, a major C-rich carrier of oxygen in meteorites is an insoluble macromolecular material ("kerogen") of uncertain origin (Cronin et al. 1988; Cody et al. 2002). In the Murchison carbonaceous chondrite, this material has relative abundances of C, O, and N of 100:18.3:3.8, respectively (Cody et al. 2002). The oxygen in this macromolecular material is found in both aromatic and aliphatic moieties (Gardinier et al. 2000; Cody et al. 2002). The meteoritic macromolecular material appears to be much more O-rich than the Crich solid-state material seen in the diffuse ISM, but considerably less O-rich than the average of organics in collected IDPs (Flynn et al. 2006) and the organics seen in cometary grains returned from comet 8IP/Wild 2 by the Stardust mission (Sandford et al. 2006).

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DiSanti,

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1 Bay Region

1 1

( ^ Y y

PAH

> etc.

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J-H H

UV radiation H 2 0 ice at 10 K

Figure 3. Polycyclic aromatic hydrocarbons (PAHs) are abundant in space. Their exterior carbon rings are easily functionalized with oxygen when they are irradiated in O-bearing ice mantles in dense interstellar clouds. The mix of functional groups depends somewhat on the composition of the ice, but =0, -OH, -O(ether), and -O-CH3 groups are commonly seen in laboratory simulations.

Because of their small sizes, considerably less detail is available about the nature of the organics in IDPs. They contain abundant aromatic species (see, for example, Allamandola et al. 1987; Wopenka 1988; Quirico et al. 2005), and a variety of specific PAHs have been identified in them (Clemett et al. 1993; 1998). The total oxygen content of the organics in IDPs is generally much higher than that seen in the macromolecular material in meteorites (Flynn et al. 2006; Sandford et al. 2006), but the nature of the molecular carriers is still unclear. Given the abundance of aromatic species in these particles, aromatic ketones and alcohols may be possible carriers. Amino acids have been identified in larger Antarctic micrometeorites (Matrajt et al. 2004), but these are unlikely to be a major reservoir of the O found in IDPs since they are minor components in meteorites and micrometeorites. Furthermore, these amino acids may be products of aqueous processing (Strecker synthesis), and the parent bodies of anhydrous IDPs have not experienced hydrous alteration. The organics returned from comet 8IP/Wild 2 by the Stardust spacecraft are generally considerably richer in O than meteoritic organics, but are comparable, on average, to those of measured stratospheric IDPs (Sandford et al. 2006). X-ray Absorption Near Edge Spectroscopy (XANES) studies of these cometary samples demonstrate that the O is present in a wide variety of bonding states, including aldehydes, alcohols and esters. While the 81P/Wild 2 samples have O/C ratios somewhat similar to those seen in IDPs, it is clear that the two types of samples are not identical. The Wild 2 samples appear to contain a labile organic component that is missing from anhydrous IDPs (Sandford et al. 2006). If anhydrous IDPs have a cometary origin, this suggests that they may have lost a more labile fraction of their original organics during atmospheric entry or during their transit through interplanetary space from their parent body to Earth. Also, as noted above, the organics in IDPs are generally dominated by aromatic materials. In contrast, many Wild 2 samples contain little or no aromatic materials. The organics in these aromatic-poor grains contain large amounts of O, indicating that the O in Comet Wild 2 organics cannot be solely associated with aromatic species. In these particles, the carrier may be dominated by a "polymeric" material akin to polyoxymethylene or related "irregular" molecular structures similar to those made when mixed molecular ices are processed by high energy radiation (Schutte et al. 1993; Bernstein et al. 1995; Sandford et al. 2006). The generally increasing O content of organics from the diffuse ISM, to meteorites, to IDPs and cometary samples suggests that meteoritic, IDP, and cometary organics do not simply consist of unaltered organics from the diffuse ISM, but must contain contributions from either dense cloud, protosolar nebula, and/or parent body processes. The fact that material in

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IDPs and from comet 8IP/Wild 2 generally shows higher O contents than meteoritic organics further suggests that dense cloud and/or protosolar nebular processes (radiation processing, gas and gas-grain chemistry, etc.) must generally act to drive up the overall O content of organics more than do parent body processes (aqueous alteration, heating, etc.) In summary, oxygen is seen in a variety of forms in stellar outflows, the diffuse ISM, and dense molecular clouds. Silicates and other oxides are the main carriers of solid-state oxygen in the diffuse ISM and are major carriers in dense clouds. Most of the gas phase O in the diffuse ISM is in the form of atomic O; very little oxygen seems to be associated with the C-rich materials seen in the diffuse ISM. Most of the non-mineral oxygen seen in the dense ISM is in relatively volatile, "non-meteoritic" forms, and much of it is in the form of ices. "Fixing" interstellar oxygen into C-rich forms that can survive incorporation into asteroidal and cometary parent bodies requires significant processing in dense cloud and protostellar nebular environments. These chemical processes result in organics that contain oxygen in a diverse set of molecules ranging from small, soluble species to macromolecular "kerogens." In the specific case of comets, the organics (at least for comet 8IP/Wild 2) are extremely O-rich and the oxygen is present in a wide variety of bonding states. Direct detection of oxygen-bearing volatiles in comets Whipple (1951) was the first to recognize the importance of comets for the history of our Solar System. Late in the last century it was accepted that comets represent the best-preserved material from the early Solar System. Because they retain their volatiles, comets also appear to have retained their full complement of solar nebular oxygen (Fig. 4). Furthermore, since some molecules found in comets, e.g., C 4 H (Geiss et al. 1999), can be traced back to the dark molecular cloud from which our Solar System formed, we can use cometary volatiles and their abundances to study the processes that led from the molecular cloud through accretion into the solar nebula, to the present constituents of cometary nuclei (e.g., Altwegg et al. 1999). Until about 1980, it was assumed that comet nuclei primarily contained frozen H 2 0, NH 3 , and CH 4 . However, models based on this assumption, even when supplemented with a few other minor species, failed to explain the radicals and ions identified in spectra of cometary

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Hydrogen / Silicon Figure 4. Elemental abundances of carbon (C), oxygen (O) and nitrogen (N) in different bodies of our Solar System as a function of hydrogen/silicon ratio (after Geiss 1988).

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comae. One of the first models to come close to explaining the observations was based on the assumption that cometary nuclei consisted of frozen interstellar molecules and grains (Biermann et al. 1982; Greenberg 1982). Since that time, considerable progress has been made in finding new molecular species in comets and in modeling and understanding their compositions in this context (e.g., Altwegg et al. 1999). Figure 5 shows the abundances of cometary molecules compared with those of molecules in molecular clouds having C/O ratios near solar. Such data are becoming increasingly available (e.g., Ehrenfreund and Charnley 2000). The relative abundances of most cometary molecules measured so far are similar to those seen in interstellar dense clouds; the ranges of observed molecular abundances, relative to H 2 0, generally overlap in these objects. This suggests that much of the volatile component in comets has been preserved from the preceding cold molecular cloud stage. However, the degree of processing experienced by their ices is a fundamental question in cometary science, and indeed this is the primary driver for building a new taxonomy of comets based on composition. One of the defining characteristics of comets is their high abundance of volatiles (ices). As mixtures of nebular and interstellar ices, their compositions may be sensitive indicators of spatial and temporal variations in the nebular thermal environment. The dominant cometary ice phases (polar and apolar) are major reservoirs of primordial volatile oxygen. Measuring the compositions of cometary native ices (i.e., those contained in the nucleus) can therefore provide unique constraints on the origin and history of O-bearing icy materials in the Solar System. These data reflect the degree to which the composition of organic pre-cometary ices varied with distance from the young Sun (and with time) in the early solar nebula. Their structure and composition depend on local conditions (chemistry, temperature, degree of radiation processing) prevalent when and where they formed (Mumma et al. 1993; Irvine et al. 2000; Bockelee-Morvan et al. 2004). Measured cometary ice compositions can be compared with interstellar ices, with laboratory-processed analogs, and with formation models of comets and the proto-solar nebula.

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Comparison of molecular abundances Comets-Interstellar Ice- Hot cores A Interstellar Ices • Hot cores • Comet coma abundance p] Range of measured " data in clouds

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Figure 5. A comparison between the relative abundances of various molecular species in ices found in dense molecular clouds (triangles), in gases in "hot cores" within dense molecular clouds (squares), in the gas phase in cometary comae (circles), and the range of abundances found in gas phase molecules in dense molecular clouds (cross-hatched bars). All abundances are normalized to the abundance of H 2 0 , which has been assigned a value of 100. Arrows indicate upper limits. After Crovisier (1998).

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Oxygen plays a prominent role in the volatile composition of comets. Most of the volatile oxygen is in the form of H 2 0, and since this is the most abundant ice in comets, it is generally used as the reference against which the abundances of other ices are measured. In terms of oxidized carbon, abundances of CO (relative to H 2 0) are highly variable among comets, ranging from less than 1% to nearly 20%. Abundances of H2CO and CH3OH as high as ~3% and ~7%, respectively, have been reported (see, e.g., Fig. 12 of Bockelee-Morvan et al 2004). C0 2 , although observed in only a few comets, appears to be present at levels of several percent. By contrast, hydrocarbon molecules (CH4, C2H2, C2H6) provide an overall lower contribution to the total budget of volatile carbon. Abundances of the simplest fully reduced one-carbon molecule, CH4, although variable among comets, are at most CH4/H20 ~ 2% (see Gibb et al. 2003), and C2H2 and C2H6 are even less abundant (e.g., C2H6/H20 ~ 0.6% and C 2 H 2 /H 2 0 ~ 0.2% in most Oort cloud comets observed to date). This is surprising in view of the very hydrogenrich composition of the solar nebula, in which comets formed. It is consistent, however, with material formed in an interstellar cloud or in a kinetically controlled solar nebula, where CO can either condense directly in polar (H 2 0-rich) or apolar (H20-poor) ices (see below). Cometary nuclei warm when approaching the Sun, causing their ices to sublime, releasing volatiles into their comae (i.e., atmospheres), where they can be sensed spectroscopically at infrared wavelengths (principally between 2.8 and 5.0 jim) and sub-millimeter wavelengths. A fundamental challenge for cometary observations is distinguishing direct release by the nucleus (parent volatiles) from sources of extended release in the coma (e.g., by thermal degradation of grains or by chemistry). The spectral signature of volatiles consists of emission lines arising from fluorescent vibrational excitation by incident solar radiation. Although the presence of such lines in cometary spectra was predicted previously (Mumma 1982; Crovisier and Encrenaz 1983; Weaver and Mumma 1984; Bockelee-Morvan and Crovisier 1987), their detection required the development of astronomical spectrometers having sufficiently high sensitivity and spectral resolving power (e.g., X/AX ~ 2 x 104 or higher). Modern IR spectrometers have small (sub-arc-second) pixels, and so are well-suited for measuring molecular abundances of volatiles in comets. They also provide spatial coverage of the sky, thereby permitting an accurate measure of the spatial distribution of emission in the coma. The emission intensity for direct release is highly peaked at the nucleus and decreases approximately as inverse projected distance from the nucleus, whereas distributed release gives rise to a flatter spatial profile of emission. This is particularly relevant to CO and H2CO, for which significant, or even dominant, distributed source contributions have been observed in some comets (Eberhardt et al. 1987; Meier et al. 1993; Wink et al. 1997; Eberhardt 1999; DiSanti et al. 1999, 2003). Advances in instrumentation over the past decade now enable the routine detection of multiple parent volatiles in comets, primarily through resonant (i.e., fundamental-band) rovibrational transitions occurring between excited and ground vibrational states. Detected molecules include CO (Fig. 6A), monomelic formaldehyde (H2CO, Fig. 6B), and methyl alcohol (CH3OH, Fig. 6C). Together with C0 2 , these represent the principal reservoirs of volatile oxidized carbon in comets. Ground-based detection of H 2 0 requires sensing non-resonant ("hot band") transitions (Figs. 6A, 6D) that occur between two excited vibrational states that are not significantly populated in the terrestrial atmosphere (Dello Russo et al. 2000). Prompt emission from OH was also proposed for sounding H 2 0 in comets (Mumma 1982; Bockelee-Morvan and Crovisier 1989; Mumma et al. 2001). In contrast to the relatively flat spatial distribution of emission from fluorescent OH, the spatial profile of OH prompt emission faithfully traces that of parent (H 2 0). Through simultaneous measurement with H 2 0, a method has been developed for using OH prompt emission lines to quantify H 2 0 production in comets (Bonev et al. 2004; Bonev 2005). These OH lines can therefore be used to establish both the spatial distribution and production rate of H 2 0 when emissions from H 2 0 itself are not available for observation within the band pass used.

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Figure 6. Spectra of oxygen-bearing molecules in comet C/2004 Q2 (Machholz) obtained with the Near IR Spectrometer (NIRSPEC) at the Keck Observatory. In each panel, the upper trace shows the extracted spectrum (solid) including dust continuum emission, and the superimposed modeled atmospheric transmittance (heavy dashed trace). The bottom trace shows residual cometary emission in excess of the continuum. A. Simultaneous measure of CO fundamental and H 2 0 hot band emission near k = 4.7 |jm. CO lines are labeled by their rotational designations, and are blue-shifted relative to their terrestrial counterparts due to the geocentric velocity of the comet ( 20 km s _1 ). B. Region of H 2 CO emission near 3.6 |.im, showing the Vj Q-branch plus four lines of OH prompt emission, used as a proxy for H 2 0 . C. CH 3 OH v 3 band emission, centered near 3.52 |jm, with P-, Q-, and R-branches indicated. D. H 2 0 hot band emissions near 2.9 |jm. Detailed treatment of these and other comet Q2 spectra can be found in Bonev (2005).

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The abundances (i.e., the production rates) of parent volatiles are determined by applying a quantum mechanical fluorescence model to the observable ro-vibrational lines. In most cases the analysis requires development of new (or extension of existing) fluorescence models appropriate to the low temperatures (20-150 K) typical of cometary comae. Such fluorescence models need to be formulated from high-resolution laboratory spectra, synthesized at the desired rotational temperature, and convolved to the instrumental resolution. Interpretation of infrared emission from H 2 CO (Reuter et al. 1989; DiSanti et al. 2006), CH 3 OH (Reuter 1992), and H z O (Dello Russo et al. 2000, 2005) has been approached in this manner. Because CO is a linear molecule having a relatively simple spectral signature, its rotational temperature can be measured directly by comparing the flux (W m~2) contained in lines spanning a range of rotational energies (Herzberg 1950; DiSanti et al. 2001). More recently, a general method has been developed for measuring rotational temperatures of any molecular species for which a fluorescence model exists over a range of temperatures (Dello Russo et al. 2005; DiSanti et al. 2006). Besides H 2 0, CO, CH 3 OH, and H 2 CO, the other main O-containing gas phase species in cometary comae is C 0 2 (e.g., Feldman et al. 1986; Crovisier et al. 1997), which has typical abundances relative to H 2 0 of 1-10% (Bockelee-Morvan et al. 2004). A host of other O-containing species have also been identified in comets, including HCOOH, HCOOCH 3 , CH 3 CHO, NH 2 CHO, HNCO, OCS, and S0 2 , but these species are generally present at abundances well below 1% that of H 2 0 (see Bockelee-Morvan et al. 2004 for a review). The discovery of cometary CO during spacecraft ultraviolet observations of Comet West (1976 VI) (Feldman and Brune 1976; Feldman 1978) established this molecule as an important component in these primitive Solar System objects. Since then, CO has been observed extensively at UV (Festou et al. 1982; Feldman et al. 1997), IR (Mumma et al. 2003, and references therein), and radio (Biver et al. 2002) wavelengths, and it has become the cornerstone for studies of oxidized carbon in comets. Pure CO ice has the lowest sublimation temperature (~25 K) of any parent molecule, so fractionation of CO-rich, pre-cometary ices should depend strongly on local temperatures. Alternatively, if trapped in H 2 0-rich ices, CO can be captured at higher temperatures (~50 K, or perhaps even to 150 K; Sandford and Allamandola 1988, 1990; Crovisier and Encrenaz 2000). In the IR, CO emission has been detected in all Oort cloud comets observed since 1996, and the large variation in its native abundance relative to H 2 0 (1-20%) suggests that pre-cometary ices experienced a range of temperatures (indeed, Oort cloud comets formed in the giant planets' region, ~5-30 AU from the Sun, over which local temperatures would have varied greatly). However, the abundances of CO and CH 4 (the next most volatile parent molecule in comets, after CO) are not correlated among the comets studied (Gibb et al. 2003), suggesting that thermal effects alone cannot explain volatile abundances in cometary nuclei. A plausible means of converting CO to H 2 CO and H 2 CO to CH 3 OH involves hydrogen atom addition reactions on the surfaces of pre-cometary grains. This process is analogous to the H-atom addition (e.g., to C 2 H 2 ) proposed to explain the high abundance ratio of C 2 H 6 /CH 4 observed in C/1996 B2 (Hyakutake) (Mumma et al. 1996) and subsequent comets. The relative abundances of native CO, H 2 CO, and CH 3 OH to each other and to H 2 0 in comets provide a measure of the efficiency of hydrogenation of CO on grain surfaces. UV and proton irradiation of laboratory analogs of pre-cometary ice (mixed H 2 0 and CO) demonstrated the production of HCO, H 2 CO, CH3OH, and formic acid (Bernstein et al. 1995; Hudson and Moore 1999). Hydrogen-atom irradiation, both of polar (mixed H 2 0, CO) ice and of apolar (pure CO) ice, showed the conversion efficiency to be highly dependent on temperature in the 10-25 K range and on the density of H atoms (Hiraoka et al. 2002; Watanabe and Kouchi 2002; Watanabe et al. 2004)., Hydrogen-atom addition to CO proceeds through the highly-reactive formyl radical (HCO). This subsequently converts to formaldehyde polymers such as polyoxymethylene (POM) and related derivatives (Huebner et al. 1987; Meier et al. 1993; Schutte et al. 1993), and also to H 2 CO and then to CH3OH. The relative abundances of CO, H 2 CO and CH 3 OH test their inter-relationship in comets, and the laboratory yields provide a basis for comparison in assessing natal conditions.

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The presence of formaldehyde and related compounds in comets may also have been important for the origin of life. For example, H 2 CO has been proposed as the principal one-carbon molecule capable of generating complex organics of biological importance (Weber 2000; 2002). H 2 CO may also have played a central role in the formation of amino acids in primitive meteorite parent bodies during aqueous alteration via Strecker-cyanohydrin synthesis. Alternative mechanisms for producing amino acids in pre-cometary ices without liquid water have also been identified (Bernstein et al. 2002a; Muñoz Caro et al. 2002). These studies either employ HCN, NH 3 , and H 2 CO directly, or generate them in situ (Moore and Hudson 2003). Cometary organic molecules may also have contributed to the formation of nucleobases (Oró 1960), sugars (Weber 2002), possible pre-RNA backbones (Nelson et al. 2000), and a host of biochemical intermediates (Oró et al. 1990) either on the icy body or after accretion to the early Earth. The oxygen-bearing minerals in "cometary" IDPs and samples from comet 811'/Wild 2 The remote detection of oxygen-bearing minerals in comets. The presence of O-bearing minerals in comets can be detected at infrared wavelengths, primarily through the use of spectroscopy in the 8-13 |im region where the characteristic Si-O stretching vibrations of silicate minerals fall, but silicate features can sometimes be detected out to 40 |im. When small cometary grains are ejected from a cometary nucleus, they are heated by solar radiation. This energy is subsequently re-radiated in the infrared, resulting in a quasi-blackbody continuum upon which a superimposed silicate emission feature is sometimes seen. To produce a strong feature in emission, the emitting silicate particles must be smaller than 1 |im in radius; particles larger than this will, for the most part, only contribute to continuum emission, and their mineralogy is difficult to assess in this manner. The strength and profile of the emission features are therefore dependent on grain composition, size and albedo, and emission models must be used to interpret these features (see Hanner and Bradley 2004 for a review). The infrared emission features of dust in cometary comae are generally interpreted to be largely dominated by olivine, pyroxene, and "glassy" silicates (e.g., Hanner et al. 1994; Wooden et al. 1999; Crovisier et al. 2000; Hanner and Bradley 2004). Models of the silicate emission feature suggest that only a minor fraction of the silicates (15-30%) are in crystalline form; the remainder of the silicates may be amorphous. A unique opportunity to study cometary dust remotely occurred when the Deep Impact mission sent a 364 kg impactor into the nucleus of comet 9P/Tempel 1 at 10.2 km/sec. The impact produced large amounts of ejecta from the surface layers of the comet, and infrared spectra of the ejecta were collected by the Spitzer Space Telescope (Lisse et al. 2006). Emission signatures due to amorphous and crystalline silicates, amorphous carbon, carbonates, phyllosilicates, polycyclic aromatic hydrocarbons, H 2 0 gas and ice, and sulfides were reported. Many of these materials are seen in chondritic porous IDPs and samples from comet 8IP/Wild 2 (see below), but others, for example phyllosilicates and carbonates, have yet to be confirmed in Wild 2 samples. The reason for these apparent differences has yet to be fully resolved. Overall, the spectral features observed in cometary infrared emission spectra imply a fairly complex mineralogy that includes both amorphous and crystalline grains, with olivine and pyroxene being the dominant crystalline phases. The mineralogical mixture is largely consistent with the composition of chondritic porous IDPs and samples returned from comet 8IP/Wild 2, although it is difficult to assess the abundance of glassy silicates in the Wild 2 samples (see the following two sections). Oxygen-bearing minerals in "cometary" IDPs. Interplanetary dust particles are collected in the stratosphere at 20-25 km altitude and are typically 5-15 |im in diameter. There are two major types of IDPs: chondritic porous (CP) IDPs that are typically anhydrous; and chondritic smooth (CS) IDPs that have undergone aqueous alteration that formed clay minerals and carbonates. The CP IDPs have been linked to cometary sources from their inferred orbital characteristics (Brownlee et al. 1995), fine-grained mineralogy (Bradley and Brownlee 1986), infrared spectral properties (Sandford 1991), and high abundances of presolar materials

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(Messenger et al. 2003). These particles have escaped the thermal metamorphism and aqueous alteration that affected even the most primitive meteorites, and are characterized by high carbon and nitrogen abundances (Thomas et al. 1993; Keller et al. 2004; Flynn et al. 2006), unequilibrated mineralogy (Keller and Messenger 2005), and the presence of non-solar hydrogen and nitrogen isotopic signatures and abundant presolar silicates (Messenger et al. 2003). Typical CP IDPs are highly porous particles that consist of fine-grained crystalline silicates, GEMS (glass with embedded metal and sulfides, Bradley 1994) grains, and Fe-Ni sulfides, all bound together by an organic-rich carbonaceous matrix. The constituent grains in IDPs are much smaller ( 90 km. Moderate-albedo S-types prevailed in the inner belt (a < 2.4 AU) while low-albedo objects dominated the outer belt (a > 2.4 AU). Zellner (1979) found that, for the whole main belt population, ~75% are C-types, ~15% S-types, and ~10% are other types.

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Gradie and Tedesco (1982) showed that the main belt is highly structured, with at least six major compositionally distinct regions, and a particular class or set of classes dominating each region. They proposed that the asteroids formed at or near their present locations, because this distinct zoning of the distribution could not be explained by random transport of objects over large distances of the Solar System. In particular, the S-types peak at ~2.3 AU and make up 40% of the local population, while the C-types have their maximum at ~3.0 AU, accounting for 80% of the local population. Assuming the usual linkages between taxonomic classes and particular asteroids, the distribution of objects found by Gradie and Tedesco (1982) mimics part of the equilibrium condensation sequence (Table 1), with enstatite-rich bodies being found in the inner belt and asteroids becoming more oxidized with larger heliocentric distances. High-albedo E-types (linked with the enstatite-rich aubrites) are found in the innermost part of the belt. S-types, which tend to contain ferrous olivines and pyroxenes, are most abundant in the inner part of the belt. The low-albedo C-complex objects (usually linked with some type of carbonaceous chondritic material) peak in the middle of the belt. Lower-albedo P- and D-types, believed to be organic-rich, peak in the outermost part of the belt. In the Chapman et al. (1975) analysis, the same general characteristic was seen: moderatealbedo objects dominate the inner belt and the low-albedo objects were preferably found in the outer belt. The main difference between the results of Chapman et al. (1975) and the work of Gradie and Tedesco (1982) concerns the distribution of C-class objects in the main belt. Gradie and Tedesco (1982) see a peak in the distribution at around 3.0 AU, which is not as apparent in the Chapman et al. (1975) distribution. Since the last decade, high throughput long-slit spectrographs employing charge-coupled devices (CCDs) have become widely used in measuring the visible spectra of asteroids. Initiated in 1991, data from SMASS II (Bus and Binzel 2002a) became available in 2001, and produced an internally consistent dataset of spectra for over 1,400 asteroids in the 0.44-0.92 |im spectral range. The SMASS II survey focused on regions surrounding the asteroid 4 Vesta, and between 2.69 and 2.82 AU, with detailed investigations of dynamical families in these regions. Initiated some years later, another survey, S30S2 (Lazzaro et al. 2004), was completed and obtained visible spectra for over 700 asteroids from 0.5-0.92 |im. Like SMASS II, S30S2 is also biased, because part of the survey focused on some large asteroid families or clans like Flora, Eunomia and Themis, as well as the "Magnya region" around 3.14 AU, and the regions around some resonances. Another focus of the S30S2 was on asteroids located at high eccentricities and inclinations. Asteroid spectra in both surveys were classified according to Bus' (1999) taxonomy, which is based on the Tholen's (1984) taxonomic system. Therefore, most of the objects belonging to one of the classes of Tholen are also members of one of the Bus (1999) complexes. Bus (1999) used ~1,200 objects located between 2.10 and 3.278 AU, with most objects having D < 20 km, to determine the heliocentric distribution of the Bus (1999) classes. The diameters of the objects were estimated using the mean IRAS albedos for each spectral class. He divided the belt into three zones, with limits of 2.100 < a < 2.501 AU, 2.501 < a 1

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the percentages for the HMB and LMB, respectively. The outlying classes Ld, D, and T, which are not presented by Mothe-Diniz et al. (2003), have been added to Figure 31 for a better comparison with the distributions produced by Bus (1999). The results of Bus (1999) confirm the main results of Gradie and Tedesco (1982). Considering only objects larger than 20 km, they found that the percentage of objects of the C-complex increases from ~25% in the inner to almost 65% in the outer main belt. On the other hand, the percentage of asteroids in the S-complex decreases from about 55% to 5% in the same interval of semi-major axis, while the distribution of asteroids in the X-complex is quite flat, at a level of 20%. The results of Mothe-Diniz et al. (2003) are also similar to the previous works of Gradie and Tedesco (1982) and of Bus (1999), but only if they used a "restricted sample:" either the objects larger than 30 km diameter, or those with high inclinations or eccentricities (HMB) (Fig. 32a). Using the entire database, the distribution of taxonomic classes is significantly different from previous work because it does not show a steep decrease of S-types with semimajor axis. The S-types are rather evenly distributed along the main belt, even when families and small asteroids are not included in the analysis (smaller than 13 km, which represents the size above which the population of asteroids in the three complexes of Bus (1999) is complete). By studying the spatial distribution (HMB vs. LMB) of the asteroidal population, Mothe-Diniz et al. (2003) noticed that the distribution in the LMB (Fig. 32b) is very different from that in

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the HMB (Fig. 32a), with the latter distribution being similar to the results of previous authors. These results indicate that the LMB seems more "mixed" than the HMB. In addition to the new spectroscopic data, an enormous database of asteroid spectrophotometric data from SDSS has been released (Ivezic et al. 2001). Five-color CCD photometry was determined almost simultaneously in five bands: u, g, r, I, and z, centered respectively at 0.3557 |im, 0.4825 |im, 0.6261 |im, 0.7672 |im, and 0.9097 |im, that have 0.10.3 jim band widths (Fukugita et al. 1996). This catalogue provides important information on asteroids because it has a detection limit much fainter (V ~ 21.5) than other asteroid surveys ( V ~ 18) (Ivezic et al. 2002) (Fig. 33). Using SDSS observations, Ivezic et al. (2002) estimated that 90% of asteroids in the main belt are parts of families. They found that objects that did not belong to the most populous families and are usually thought of as background objects did show color clustering. One flaw in this argument is that just because objects have similar colors does not mean that they are genetically related. The main limitation of any study of the distribution of taxonomic classes is the availability of spectral data. The observations that allow taxonomic classifications cover only a small fraction of the numbered asteroids. This leads to large bias factors in some regions that have not been well-sampled. Another limitation is the incompleteness of the numbered population, presently starting at about absolute magnitude 11.75 in the external part of the main belt (3.03.5 AU), 12.25 in the intermediate main belt (2.6-3.0 AU), and 12.75 in the inner main belt (2.0-2.6 AU) (Jedicke and Metcalfe 1998). One could also expect a limitation caused by the magnitude cutoff, i.e. the diameter cutoff, which is a function of heliocentric distance, and is different for each class. In fact, this causes the bias analysis to be limited to diameters larger than the cutoff diameter of the most distant zone.

Absolute Magnitude (H) Figure 33. Histograms showing the number of objects observed in SMASSII and S30S2, compared to those observed in SDSS, per half-magnitude bin.

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This point does not represent a real limitation, because most of the small objects come from the breakup of larger bodies, and in all analyses, the asteroid families were considered as one single body. Therefore, most of the small objects in the sample were not taken into account. The meaning of some taxonomic classes has changed significantly from the first classification scheme of Chapman et al. (1975) and the first work on the distribution of classes of Zellner (1979). However, the cores of some classes have suffered little change. This is the case of the S and C classes, which were recognized very early as distinct groups. In the Bus taxonomy, all the S-types, and their subtypes according to other taxonomies, are included in the S-complex. The same is true for the C-complex. Thus, the conclusions from research on the C- and S-complexes should be directly comparable to previous studies that involved the C- and S-types. The same is not true for X-types. The X-complex of Bus (1999) is composed of, among others, objects from the earliest E-, M-, and P-classes, which are subtypes of Tholen's (1984) X-class. As discussed earlier, objects from the E-, M- and P-classes have similar, featureless visible spectra, but have been separated according to their albedo. However, even if an old Xtype object belongs to the X-complex, it is difficult to compare the distribution of this complex from 2.064 to 3.278 AU to previous studies. These earlier works included the heliocentric distributions of types E, M and P individually, and because albedo data are not used in the Bus (1999) taxonomic scheme, we cannot say whether or not the distribution of the X-complex or of one of its subtypes is similar to that of earlier works. There are definite trends in the distribution of X-types in the far inner and outer main belt regions that appear consistent with mineralogical differences. Asteroids in the Hungaria group (1.78-2 AU) are known to contain a considerable number of asteroids classified as E (Clark et al. 2004a), X, and Xe (Carvano et al. 2001). The distribution of different classes in the Cybele region was found to be a function of size, with the spectrally red D-type asteroids generally more numerous at small diameters, while larger asteroids tended to be more spectrally neutral (flatter spectral slopes), such as X-types. The number of P-types decreases with increasing distance from the Sun. Dahlgren et al. (1997) found that objects in the Hilda group (3.70-4.20 AU) are typically D types, with very few C- and P-types, and with relatively low albedos (0.03-0.11). Still farther out, Fornasier et al. (2004) found that Trojan asteroids (~5.2 AU) are predominantly D-types with only a few P-types. Taxonomic classifications are based on spectral features which, among ground-based spectroscopic measurements, are the best indicators of an asteroid's underlying composition. Therefore, they should provide some indication of an asteroid's mineralogy, as well as an interpretation of their distribution along the main belt. Based on the results of Gradie and Tedesco (1982), Bell et al. (1989) attempted to relate the distribution of classes with mineralogy, by proposing a scenario in which the "igneous" asteroids (objects formed from a melt and represented by the classes V, R, S, A, M and E) were in the inner part of the main belt; the "metamorphosed" (those that have been sufficiently heated to exhibit spectral changes, represented by the classes B, F, G, and T) in the middle; and the "primitive" (those that have undergone little or no heating, with representatives in the classes D, P, C, K, and Q) in the outer part of the belt. These three large groups were called "superclasses" and agreed with the intuitive picture of a heliocentric heating gradient. This interpretation is dubbed the "Big Picture." The main complaint with the "Big Picture" is the mineralogical interpretation of the igneous classes. Bell et al. (1989) considered the S-types as igneous since, at that time, they were commonly associated with achondrites and stony-iron meteorites, even if an alternative association with ordinary chondrites was already postulated through dynamical studies and meteorite fall rates. With the increase of the quantity and quality of the observations, as well as a better understanding of the space weathering phenomenon, however, this paradigm has changed. There is now stronger evidence linking some S-types to ordinary chondrites and a

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better understanding of space weathering processes (e.g., Chapman 1996; Binzel et al. 1996; Pieters et al. 2000; Chapman 2004). There is evidence for partially differentiated S-complex asteroids (e.g., Gaffey et al. 1993a; Hardersen et al. 2006; Gaffey 2006), but it is still unclear whether chondritic or partially differentiated objects predominate among the S-types. Sunshine et al. (2004) suggest that some members of the Merxia and Agnia families might be differentiated. However, the location of these two "differentiated" families in the main belt (around 2.74 AU and 2.78 AU, respectively) does not seem to support the kind of heating gradient in the asteroid belt that Bell et al. (1989) suggested. Another question is the absence of differentiated families associated with M-types. Such families were to be expected if M-types are metallic cores of differentiated parent bodies that formed locally and were shattered by catastrophic collisions. Moreover, the discovery of hydration bands in some Mtypes (Rivkin et al 2002) suggests that many M-types are not metallic. Finally, a basaltic Vtype (1459 Magnya) was found in the outer belt, without any apparent dynamical relation with the Vesta family (Lazzaro et al. 2000; Michtchenko et al. 2002). This discovery is inconsistent with the strict zoning of the belt proposed by Bell et al. (1989). Mothe-Diniz et al. (2003) showed that S-types are almost as abundant in the outer part of the belt as in the inner belt. In summary, the scenario of the "Big Picture" is no longer valid but, on the other hand, at present no clear and/or intuitive geological scenario for the distribution of the taxonomic classes can be proposed. In fact, a picture emerges of a main belt where both the dynamical and thermal evolution is much more complex than what is depicted in the Bell et al. (1989) "Big Picture." There seems to be much more mixing than was previously proposed, specifically among small asteroids. This mixing is consistent with the knowledge that Yarkovsky effects can lead to considerable semi-major axis mobility for small asteroids (Bottke et al. 2002, 2006b). The Yarkovsky effect is a non-gravitational thrust produced when small bodies absorb sunlight, heat up, and then reradiate the energy after a short delay due to the thermal inertia of the surface. In this context, we can suggest a possible explanation for why the taxonomic distribution in the upper main belt is more similar to what is seen when only bigger asteroids are considered: in general, higheccentricity asteroids can move only through a small range in semi-major axis before hitting a resonance and being removed. The abundance of high-albedo objects with featureless spectra consistent with aubrites in the interior Hungaria region of the belt appears to indicate that objects in this region were heated to temperatures high enough for them to melt. The abundance of C-complex asteroids in the outer part of the belt and D- and P-objects still farther out indicates that asteroids in these regions were not significantly heated. But without knowing how ordinary chondrites, partiallydifferentiated, and fully-differentiated asteroids are distributed in the S-complex population, it is very difficult to understand how the heating gradient varied throughout most of the belt. Was there a sharp dropoff in heating with heliocentric distance, or was the dropoff more gradual?

DISTRIBUTION OF HYDRATED ASTEROIDS IN THE MAIN BELT Looking at the belt as a whole, it is clear that there is a correlation between semi-major axis and various measures of hydrated mineral concentration. Figure 34 shows the fraction of objects testing positive for hydrated minerals using the 0.7-|im and 3-|im regions. An understanding of the distribution of water in the asteroid belt is inevitably shaped by an understanding of how meteorite samples are delivered to Earth. The fraction of hydrated meteorites found on Earth is skewed by many factors, including the relatively low strength of hydrated meteorites relative to other meteorites, and the observation that impacts into carbonaceous chondrites seem to generate fewer large pieces of ejecta and a greater mass of dust (Tomeoka et al. 2003). When combined with the greater contribution made to the meteorite population by objects in

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Semi-Major Axis (AU) Figure 34. Hydrated C-types in the main belt. The histogram shows the number of C-complex (shaded) and Ch/Cgh-subclass (solid) objects in the belt vs. semi-major axis (AU). In areas with good statistics, the Ch/Cgh objects represent roughly half of the C-complex. The open symbols represent this fraction, using the axis on the right. It is uncertain whether the increase from 2.1-2.3 is an artifact due to small number statistics. Given that some C-types can be hydrated without having a 0.7-|.im band (which would make them Ch/Cgh), the true fraction of hydrated C asteroids is higher than 0.5.

the inner asteroid belt and near resonances, it is unlikely that the meteorite collection is a true measure of the relative proportions of materials present in the asteroid belt, and furthermore some unsampled mineralogies may easily be present. As discussed by Burbine et al. (2002b), there is a huge difference between the number of distinct parent bodies (~100-150) that we have evidence for in our meteorite collections and the number of asteroids in the main belt. In terms of the carbonaceous chondrites, roughly 2/3 of falls are hydrated. This ratio is in good agreement with observations of main-belt C-types, of which roughly 2/3 show a 3-|im band (Rivkin et al. 2002). Nearly 400 different objects in the C-complex have been identified in the SMASS and S20S3 surveys (Bus and Binzel 2002b; Lazzaro et al. 2004). Figure 34 shows a histogram of this population and its Ch/Cgh fraction as a function of semi-major axis. In addition, the fraction of the C-complex sample that appears hydrated (Ch- and Cgh-subclasses) is shown. The fraction of hydrated C-types, by this measure, is 0.51 ± 0.07 between 2.3 and 3.2 AU, where the majority of this class is found. The majority of hydrated C-types have spectra with band shapes similar to those of the CM chondrites. However, roughly 1/3 of the objects have band shapes like Ceres, which differs significantly from the typical CM-like shape (Rivkin et al. 2004b). It appears that within a given asteroid family, all members tend to share the same hydration state (Bus 1999). However, several objects have been seen to have varying 0.7- or 3-|im bands with rotation (Howell et al. 2001a), which is not obviously consistent with the family observations. In the outer belt, the number (and fraction) of objects which show evidence of hydrated minerals drops dramatically. Observations at 3 jimby Jones et al. (1990), Lebofsky et al. (1991), Howell (1995), and Emery and Brown (2003) found no clear-cut evidence of any Trojan with a

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3-|im band. Given the position of the Trojans, between hydrated C-types and the icy satellites of Jupiter, this is particularly surprising. Jones et al. (1990) interpreted this to mean that the Trojans, which presumably formed from water ice and anhydrous silicates, never got hot enough to create hydrated minerals, and that any ice present would be found beneath the surface. Cruikshank et al. (2001) obtained data on 624 Hektor and analyzed it using a Hapke mixing model. They placed a relatively high upper limit of 40% for phyllosilicates mixed with low-albedo constituents, and noted that the lack of a 3-|im band did not rule out surficial hydrated minerals per se. They also reinterpreted the trend of decreasing frequency of hydrated minerals with increasing solar distance as due to an increase in macromolecular carbon compounds, which could mask the spectral signature of hydrated minerals. The Trojan asteroids are an interesting contrast to the moons of Jupiter. The Galilean satellites (other than Io) appear to not only have copious amounts of water ice, but also hydrated non-ice material (McCord et al. 1997; McCord et al. 1999). The small, irregular moons of Jupiter also show evidence of hydrated minerals: a 3-|im band is present on Amalthea (Takato et al. 2004) and a 0.7-|im band is seen on Himalia, along with a possible 3-|im band (Jarvis et al. 2000; Chamberlain and Brown 2004). Recent work (Morbidelli et al. 2005) suggests that the Trojans may have been captured from the Kuiper Belt during the period of planetary migration. In that case, the connection between the Trojans and main-belt asteroids becomes more confusing, and the lack of hydrated minerals even more puzzling. The inner asteroid belt is dominated by objects that are usually interpreted as the parent bodies of anhydrous meteorites, such as the ordinary chondrites and achondrites. There is some evidence of aqueous alteration in ordinary chondrites, with the identification of phyllosilicates (Hutchison et al. 1987) and fluid inclusions in halite (NaCl) (Zolensky et al. 1999) in some ordinary chondrites. We may expect some inner belt S-types to show 3-|im bands. Indeed, 6 Hebe, proposed as the H chondrite parent body by Gaffey and Gilbert (1998), has been found to have a weak 3-|im band (Rivkin et al. 2001). About 1/3 of the M-types reported by Rivkin et al. (2000) have a 3-|im band, as do 4 of the 6 E-types surveyed (Rivkin et al. 1995; Rivkin 1997). The mineralogies of hydrated M-types, labeled as the W-class by Rivkin et al. (1995), and of hydrated E-types have not been fully understood, perhaps representing objects with mineralogies not present in our meteorite collections, such as having large concentrations of high-albedo salts. Hardersen et al. (2005) included bencubbinites as a possible interpretation of a number of M- and W-types using 0.8-2.5 |im data, which may be a preferred interpretation because some of these meteorites contain hydrated minerals (Krot et al. 2002). The hydrated E-types have no obvious analogs among meteorites, although hydrated enstatite chondrite clasts exist in the Kaidun meteorite (e.g., Zolensky and Ivanov 2003). The hydrated E-types are all in the main part of the asteroid belt rather than the Hungaria region, which dominates the inner edge of the belt and is where the majority of E-types are found. This is consistent with the Hung aria-region objects being different mineralogically from the E-type asteroids further out in the belt.

NEAR-EARTH ASTEROIDS Near-Earth asteroids, or NEAs, are asteroids with perihelia less than 1.3 AU. Over 5,000 NEAs are currently known. Because of their short dynamical lifetimes, resupply from the main belt is necessary to provide the number of NEAs currently seen. Mars-crossers (MCs) are objects whose orbits cross that of Mars and are thought to be in the process of being dynamically pushed out of the main belt into the NEA population. Binzel et al. (2004) looked at the classifications and spectral properties of ~400 NEAs. The NEAs tend to be dominated by objects that are part of the S- (S, Sa, Sk, SI, Sr, K, L, and Ld) and Q- (Q and Sq) complexes, as defined by Binzel et al. (2004), with approximately 2/3 of all objects belonging to one of those

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two complexes. More recent surveys (Lazzarin et al. 2004) have results that are consistent with Binzel et al. (2004). Marchi et al. (2005b), de León et al. (2006), Duffard et al. (2006), and Davies et al. (2007) have also identified a number of NEAs with V-type spectra. Marchi et al. (2006a) found a correlation between the visible spectral slope and perihelion distance. They found that the average spectral slope decreases with decreasing perihelion distance, which means that Q- and Sq-types are more abundant at smaller perihelion distances. NEAs with small perihelion distances are more dynamically evolved, meaning that they have had more close encounters with planets. Nesvorny et al. (2005) proposed that close encounters with planets should change the optical properties of a surface through tidal effects that can remove a significant proportion of space-weathered material or can bury the space-weathered material under fresher subsurface material. There appears to be fewer hydrated NEAs than one would expect from looking at either the main-belt or meteorite data. Because of the faintness of typical NEAs, 3-|im observations are difficult, and those which are observable typically require much larger thermal corrections than are necessary for main-belt objects. Therefore, these objects are obvious candidates for use of the 0.7-|im proxy band. Binzel et al. (2004) found 18 NEAs belonging to the C complex, including objects with and without the 0.7-|im band. Because of their comparable albedos, observing biases between hydrated and anhydrous objects should be minimal or nonexistent. In contrast to the fraction of hydrated C-types found in the main belt and the fraction of hydrated carbonaceous chondrites, only one Ch-type has been found among the NEAs, and the Cgh-class was the only Bus (1999) class absent from the NEO population in this sample. The C-complex objects on Mars-crossing orbits are more evenly split (3 Ch-types out of 7). While this result is based on a relatively small sample, the numbers are surprising, because 9 Ch/Cgh NEAs would be expected if there were the same correlation between the 0.7- and 3-|im bands as seen in the main belt. Even assigning a 50% correlation, much lower than that seen among the main-belt objects, 6 objects would be expected. It is unclear what the cause is for this difference. The temperature (~400 °C) needed for the 0.7-|im band to disappear (Hiroi et al. 1996b) is much higher than the sub-solar temperatures for even very low-albedo objects at 1.0 AU, which may reach 150 °C at the most. However, because the Hiroi et al. (1996b) result is an upper bound on this band's disappearance, heating due to solar radiation still may play a role. Heating due to micrometeorite impacts could also dehydrate carbonaceous chondritic material. In addition, the C-complex population of NEAs may contain a contribution from extinct comets (e.g., Binzel et al. 2004). Binzel and Lupishko (2005) estimate that 15 ± 5% of NEAs may be extinct or dormant comets, although most of these are on short-lived, Jupiter-crossing orbits (Bottke et al. 2002). The extinct comet population might not contain hydrated minerals, or at least may not show evidence for them based on the available spectra (e.g., Abell et al. 2005). Near-IR spectra of 1373 Cincinnati and 2906 Caltech, which have cometary orbits, have spectral properties (Ziffer et al. 2005) roughly consistent with cometary nuclei and primitive asteroids.

SPACECRAFT MISSIONS The exploration of NEA 433 Eros by the Near Earth Asteroid Rendezvous (NEAR)Shoemaker mission opened another chapter in linking asteroids and meteorites. Like previous ground-based and fly-by measurements, NEAR-Shoemaker relied heavily on the multi-spectral imager (MSI) and near-IR spectrometer (NIS), which provided reflectance spectra from 0.452.6 |im, for mineralogical interpretation and asteroid-meteorite linkages. Ground-based observations (Murchie and Pieters 1996) were interpreted to indicate that Eros was heterogeneous in composition, with an olivine-rich side and a pyroxene-rich side. Despite high spatial resolution, however, little color variation was observed by NEAR-Shoemaker, although significant variations in albedo are present on Eros (McFadden et al. 2001). As had also been observed from

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numerous telescopic studies of other S-types, NEAR-Shoemaker reflectance spectra of Eros are a relatively poor fit to the ordinary chondrites, although the band centers and Band Area Ratios are similar to those of ordinary chondrites. McFadden et al. (2001) calculated an ol/(ol+pyx) ratio of 0.58 ± 4% for Eros. This mismatch between spectral matching and derived band parameters has plagued much of asteroid spectroscopy and, for the first time, NEAR-Shoemaker provided the tools to resolve this issue in the form of X-ray and gamma-ray spectrometers. Eros was the first asteroid where in situ geochemical measurements (Trombka et al. 2000) have been made. These measurements provided not only a seemingly solid link between Stypes and ordinary chondrites, but also provided some unexpected insights into the processes occurring on asteroid surfaces. Evans et al. (2001) used gamma-ray data and Nittler et al. (2001) used x-ray data to determine the geochemistry of the surface of Eros. Calculated elemental weight ratios (Mg/Si, Al/Si, Ca/Si, Fe/Si), determined from x-ray data, and Si/O and the K abundance are all within the range of ordinary chondrites. However, the calculated S/Si, Fe/Si (determined from gamma-ray data), and Fe/O elemental weight ratios are low compared to values for ordinary chondrites. The two most plausible explanations for these differences are small degrees of partial melting and space weathering. Nittler et al. (2001) rejected the former hypothesis, because, while removal of metal and sulfide could explain the apparent fractionation of these components relative to silicates, studies of partially melted meteorites universally find non-chondritic Al/Si ratios in meteorites depleted in metal and troilite. Instead, it appears that a combination of space weathering and mechanical metal-silicate separation are responsible for the aberrant S/Si, Fe/Si (determined from gamma-ray data), and Fe/O ratios. Depletion of S likely results from dissociation of FeS due to solar wind exposure and micrometeorite bombardment, with loss of the volatile S. In contrast, separation of metal from silicates to explain the low Fe/Si and Fe/O ratios likely results from the physical processes operating in the regolith to separate the dense metal from the lighter silicates. These interpretations have been bolstered by recent determinations of chondritic Cr/Fe, Mn/Fe, and Ni/Fe ratios from X-ray spectra (Foley et al. 2006). While these interpretations paint a dynamic picture of the regolith of an asteroid, they were only possible because of the coupling of mineralogical information from reflectance spectra and chemical information from X-ray and gamma-ray spectroscopy. Even with this powerful combination, uncertainties in the data prevented the NEAR-Shoemaker team from definitively linking 433 Eros to one of the chemical subgroups of ordinary chondrites (H, L, LL) (McCoy et al. 2001,2002). It appears theoretically possible to use remotely-sensed Si/O ratios to determine the olivine/pyroxene ratio with a level of uncertainty that, coupled with precise mineralogical information, would allow one to remotely determine the/o 2 of formation of an asteroid. Such information and a definitive meteorite-asteroid link at the subgroup level, while theoretically possible, may well have to await sample return. NEA 25143 Itokawa was the target of the Hayabusa mission, which orbited Itokawa and attempted to retrieve a sample. A reflectance spectrum of Itokawa by Binzel et al. (2001) appears similar to a reddened LL chondrite (Fig. 35), with MGM modeling of the spectrum also consistent with an LL chondrite composition. Burbine et al. (2003) interpreted the resulting Band Area Ratio of Itokawa to imply an ol/(ol+pyx) ratio of ~0.70. Folco et al. (2005) argued that Itokawa could possibly be an impact-melted H or L chondrite where olivine has been enriched in the surface after impact melting. A mineralogical analysis by Abell et al. (2006) of their spectrum of Itokawa indicates a pyroxene composition of Wo16±5Fs44±5, which is much more FeO-rich than the average composition of pyroxene in ordinary chondrites, and an ol/(ol+pyx) ratio of ~0.70. Abell et al. (2006) believe that Itokawa is a partially differentiated object. Results from Hayabusa's near-IR spectrometer confirm the presence of olivine and pyroxene on Itokawa's surface (Abe et al. 2006). Abe et al. (2006) find that Itokawa's surface is olivinerich and that Itokawa's l-|im band is most similar in shape to those of LL5/LL6 chondrites.

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Wavelength (|jm) Figure 35. Reflectance spectra of S-type NEA 25143 Itokawa (Binzel et al. 2001) and LL6 chondrite Bandong (Burbine et al. 2003). Spectra are normalized to unity at 0.55

Results from Hayabusa's X-ray spectrometer are consistent with L and LL chondrites (Okada et al. 2006). Using near-IR spectra data, Hiroi et al. (2006) found that one dark region on Itokawa was significantly more space-weathered than a nearby bright region. They also found that the spectra of the two regions were consistent with LL5/LL6 chondrites with the spectral continuum removed, and that the spectral difference between the dark and bright region was consistent with a higher abundance of nanophase iron particles for the dark region. Sample collection was attempted, with Hayabusa twice landing on Itokawa's surface. It is unclear at this writing if a metal projectile was successfully fired to cause Itokawa fragments to be ejected from the surface and into the sample return canister. Due to thruster problems, the plan is now for the Hayabusa spacecraft to return to Earth in 2010.

COLLISIONAL AND DYNAMICAL EVOLUTION OF ASTEROIDS The collisional and dynamical history of the main belt is strongly linked with the growth and evolution of the planets, with the events occurring during this primeval era recorded in the orbits, sizes, and compositional distributions of the asteroids and in the meteorites reaching Earth. By studying asteroids and meteorites and placing them into the appropriate geologic context, we can use these objects as probes into the processes by which planetesimals and planets formed from the solar nebula over 4.5 Ga ago. Hence, to understand the constraints provided by asteroids, we need to first understand how the asteroid belt was affected by planet formation. The classical view of planet formation in the inner Solar System, which involves the gradual coalescence of many tiny bodies into rocky planets, can be divided into four stages: (i) the accumulation of dust in the solar nebula into km-sized planetesimals; (ii) runaway growth of the largest planetesimals via gravitational accretion into numerous protoplanets isolated in their feeding zones; (iii) oligarchic growth of protoplanets fed by planetesimals residing between their feeding zones; and (iv) mutual perturbations between Moon-to-Marssized planetary embryos and Jupiter, causing collisions, mergers, and the dynamical excitation of small body populations not yet accreted by the embryos (e.g., Greenberg et al. 1978). It is

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believed that runaway growth occurs over a timescale of 0.01-1.0 m.y. while the latter stages required a few tens of m.y. During this time, the main belt probably contained perhaps as much as an Earth-mass of material, enough to allow the planetesimals in the main belt region to accrete on short timescales. Eventually, most of the mass was agglomerated into planetary embryos on the scale of Moon- to Mars-size bodies (e.g., Wetherill 1992; Chambers and Wetherill 1998). In turn, these protoplanets dynamically excited the smaller bodies in the main belt region enough to cause some degree of radial mixing (Petit et al. 2001). This likely explains the semi-major axis distribution of large S- and C-types observed by Chapman et al. (1975) and Gradie and Tedesco (1982). At the same time, the dynamical excitation of these bodies increased their collision velocities enough to terminate accretion among most planetesimals and initiate fragmentation (Bottke et al. 2005a,b). The elimination of planetary embryos and numerous planetesimals from the primordial main belt was triggered by the formation of Jupiter several m.y. after the birth of the Solar System (e.g., Petit et al. 2002). Numerical simulations indicate that combined perturbations of Jupiter and the embryos dynamically ejected more than 90% of the bodies out of the main belt zone. The main belt may have experienced a second dynamical depletion event associated with the so-called Late Heavy Bombardment that took place roughly 3.9 Ga ago (Gomes et al. 2005; Strom et al. 2005). Together, both events depleted the main belt not only of planetary embryos but also of over 99% of its population. At the same time, collisions during this phase produced numerous disruption events. Bottke et al. (2005a,b) argued that the wave-like shape of the main belt size distribution observed today is mostly a "fossil" left over from this early comminution phase. After the Late Heavy Bombardment, the main belt population took on its current characteristics. Over the last 3-4 Ga, small bodies have been slowly lost from the main belt through a combination of collisions, dynamical evolution via Yarkovsky thermal forces, and resonances (e.g., Bottke et al. 2006a). The Yarkovsky effect compels asteroids smaller than 30-40 km to drift in semi-major axis over long time scales. In some cases, bodies can drift far enough that they enter a powerful mean motion or secular resonance. From here, their eccentricities can be pumped up to planet-crossing (and Earth-crossing orbits), where a small fraction (1-2%) are delivered to Earth. This process tends to keep the NEA and meteoroid (asteroids too small to observe) populations in a quasi-steady-state, though they can be affected by large breakup events or smaller stochastic breakups occurring near resonances. This quasisteady-state explains why the lunar and terrestrial cratering rate appears to have been relatively constant (within a factor of 2) over 0.5-0.8 to 3 Ga ago (e.g., McEwen et al. 1997). Dynamical and collisional simulations based on this scenario have allowed us to glean insights into the nature of the asteroids and meteorites discussed above.

DELIVERY OF METEOROIDS TO EARTH One intriguing question is how well calculated fall percentages, the percent of the falls that belong to each meteorite group, correlate with abundances of asteroids in the main belt or NEO population. Meteorite falls are heavily biased toward ordinary chondrites, with approximately 80% of the roughly 1,000 meteorites that have been observed to fall being ordinary chondrites. While the Yarkovsky effect is a "democratic" process in that it allows small objects across the main belt to drift into main belt resonances, this does not necessarily mean that the meteorite record itself represents a uniform sample of material from all main-belt regions. This idea was tested by Bottke et al. (2006a), who tracked the delivery efficiency of test bodies started from various main-belt resonances to strike the Earth (Fig. 36). Their results indicate that meteoroids

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Figure 36. The delivery efficiency of test bodies from various main belt resonances striking the Earth. To create this plot, we updated the work of Gladman et al. (1997) and tracked the dynamical evolution of thousands of test bodies started in all major main-belt resonances. For reference, we have also plotted the proper semi-major axis (a) (AU) and inclination (/) (degrees) of 71,323 numbered asteroids with absolute magnitude (H) < 14. The stars represent values taken from test bodies started in the v 6 secular resonance. In order of increasing a. we gave them / = 2.5. 5.0. 7.5, 10.0. 12.5. and 15.0 (Bottke et al. 2002). The filled square represents test bodies placed in the intermediate-source Mars-crossing region located adjacent to the main belt between a = 2.0-2.5 AU (Bottke et al. 2002). The filled circles are values from test bodies placed in numerous mean motion resonances with Jupiter. Objects escaping the main belt with a < 2.3 AU are more than 2 orders of magnitude more likely to strike Earth than those with a > 2.8 AU. This implies that our meteorite collection is significantly biased toward the innermost regions of the main belt.

escaping the main belt from resonances located at (semi-major axis) a < 2.3 AU have a 1-4% chance of striking the Earth. This fraction drops by two orders of magnitude, however, as we move to resonances with a >2.5 AU. Because the fluxes of material escaping out of various main-belt resonances are thought to be comparable to one another (Bottke et al. 2002), the known meteorite collection is likely to be biased in favor of samples from the inner main belt. With the dominant taxonomic type found in the inner main belt being S-type asteroids, this may imply that S-types are primarily ordinary chondritic material. There may also be additional biases in the fall statistics. For example, Gaffey (2006) argues that ordinary chondrites (H, L, LL) are derived from just three parent bodies (one being Hebe) that are located favorably near meteorite-delivering resonances (though this may be an oversimplification of a more complicated problem in terms of meteorite delivery; see Bottke et al. 2005c). Gaffey (2006) argues that partially-differentiated asteroids are more abundant in the main belt. The argument of Gaffey (2006) is a combination of theoretical calculations that predict that most asteroids were not raised to high enough temperatures to fully melt, and mineralogical analyses of a number of main-belt S-types (e.g., Gaffey et al. 1993a; Hardersen et al. 2006). They show that most of these objects do not have interpreted mineralogies consistent with ordinary chondrites, but rather with partially differentiated assemblages. Partially differentiated meteorites (primitive achondrites), however, are less than 1% of known falls. If the Gaffey (2006) idea is true, then the spectral trends found by Binzel et al. (2004) and Marchi et al. (2006b) would not necessarily imply space weathering. The S, Sq, and Q NEAs and the main-belt S-types would tend to have different mineralogies, with NEAs tending to have ordinary chondrite mineralogies and main-belt S-types tending to be partially differentiated. The

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assumed dominance of ordinary chondritic material in the NEA population may also be due to the location of the "three" ordinary chondrite parent bodies favorably near meteorite-delivering resonances. Also, CI and CM chondrites are extremely friable (Sears 1998) and are significantly less likely to make it through the atmosphere than the much stronger ordinary chondrites.

THE EFFECTS OF PLANETARY EMBRYOS AND RADIAL MIXING IN THE MAIN BELT In the main belt evolution scenario above, we described how planetesimals undergoing gravitational interactions with planetary embryos are the likely explanation for the radial mixing seen among the large S- and C-complex asteroids. The implications of this idea, however, have yet to be fully explored. For example, recall that planetary embryos were also present in the terrestrial planet region, such that they, too, could have scattered planetesimals. To test this idea, Bottke et al. (2006a) used numerical simulations (Fig. 37) to track thousands of test bodies evolving amid a swarm of Moon/Mars-sized planetary embryos spread between 0.5-3.0 AU. They found that planetary embryo perturbations produce the same kinds of radial displacement of bodies observed in the main belt (Chapman et al. 1975; Gradie and Tedesco 1982). The more interesting part of the Bottke et al. (2006a) simulation, however, was concerning the objects scattered into the main belt zone through a combination of resonant interactions and close encounters with planetary embryos. They found that 0.01-0.1%, 1%, and 10% of

1 2 3 Semi-Major Axis (AU)

Figure 37. A snapshot of inner Solar System planetesimals and planetary embryos after 10 m.y. of dynamical evolution (Bottke et al. 2006a). The starting conditions and methods used were the same as those used by Levison and Agnor (2003). Here we tracked a set of 100 embryos (grey dots) distributed between 0.5-3.0 AU. Interspersed among the embryos, we placed 1,000 test bodies with uniform semi-major axes between 0.5-2.0 AU and low initial eccentricities and inclinations. The squares, crosses, and triangles show what happens to 1,000 planetesimals started with 0.5-1.0,1.0-1.5, and 1.52.0 AU, respectively. The black line represents the location of the main asteroid belt. Bottke et al. (2006a) found that numerous planetesimals (one square, several crosses and triangles) were driven into the main belt by gravitational interactions with embryos, with the highest concentration in the inner main belt region, the same region that is more likely to deliver meteorites to Earth.

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the particles that started with a = 0.5-1.0, 1.0-1.5, and 1.5-2.0 AU, respectively, achieved main belt orbits. Once there, Bottke et al. (2006a) found that these objects were dynamically indistinguishable from the rest of the main belt population. Even though many of these "interlopers" should be ejected over time via interactions with planet embryos, resonances, etc., the proportion of "interlopers" to indigenous material in the main belt should stay the same. Finally, Bottke et al. (2006a) found that most of the "interloper" material was emplaced into the inner main belt, the same region that is most likely to deliver meteorites to Earth. This implies that "interloper" material should be an important component in the meteorite collection.

COULD IRON METEORITES HAVE COME FROM THE TERRESTRIAL PLANET REGION? If planetesimal material from the terrestrial planet region can actually be found in the main belt, what would it look like? Observations show a broad-scale taxonomic stratification among large main belt asteroids, with S-complex asteroids dominating the inner main belt and C-complex asteroids dominating the outer main belt. This trend, if followed inward toward the Sun, implies that inner Solar System planetesimals experienced significantly more heating than main belt asteroids (with the most plausible planetesimal heat source being short-lived 26 Al; halfiife of 0.73 m.y.) (Bizzarro et al. 2005). Hence, bodies that accrete quickly stand the best chance of undergoing differentiation. While precise accretion timescales across the inner Solar System are unknown, modeling work suggests they vary with swarm density and distance from the Sun, such that accretion timescales increase with increasing heliocentric distance. Accordingly, if main belt interlopers are derived from regions closer to the Sun (Bottke et al. 2002), their shorter accretion times would lead to more internal heating, and thus they would probably look like heavily-metamorphosed or differentiated asteroids. At this point, a reasonable connection can be made between the putative interlopers and iron meteorites. Cooling rate and textural data from irons indicate that most come from the cores of small (D ~ 20-200 km) differentiated asteroids (Mittlefehldt et al. 1998). Isotopic chronometers also indicate that core formation among iron meteorite parent bodies occurred 1-2 m.y. before the formation of the ordinary chondrite parent bodies (e.g., Kleine et al. 2005; Baker et al. 2005). The paradox is that if small asteroids differentiated in the main belt at such early times, it would be reasonable to expect larger bodies forming near the same locations to have differentiated as well (e.g., Grimm and McSween 1992). Hence, if iron meteorites are indigenous to the main belt, large numbers of differentiated bodies and their fragments should reside there today. This is not observed. Instead, Bottke et al. (2006a) argued that a more probable formation location for many iron meteorite parent bodies was the terrestrial planet region, where accretion occurred quickly, and thus differentiation was more likely to occur among small bodies. The protoplanets emerging from this population not only induced collisional evolution among the remaining planetesimals but also scattered some of the survivors into the main belt, where they resided for billions of years until escaping via a combination of collisions, Yarkovsky thermal forces, and resonances. If true, this means that some asteroids are main-belt interlopers, with a select few possibly being remnants of the long-lost precursor material that formed the Earth. High-temperature condensates (Table 1) would also be expected to be scattered into the inner main-belt. Some of these "interlopers" may be enstatite chondritic and aubritic material, which are believed to have initially condensed very near the Sun (e.g., Rubin and Wasson 1995). The inner-belt Hungaria region has a high proportion of X-types (Clark et al. 2004a), which includes the high-albedo E-types, the more moderate albedo M-types, and the Xe-types. X-types have spectral properties similar to aubrites, enstatite chondrites, and iron meteorites, which are all bodies that would be expected to have formed in the inner Solar System and scattered into the inner main belt in the Bottke et al. (2006a) model.

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SUMMARY R e m o t e sensing can d e t e r m i n e the p r e s e n c e of F e 2 + and Fe 3 + and w a t e r ( H 2 0 or O H ) on asteroid surfaces. Distributions a m o n g d i f f e r e n t asteroid classes can b e seen in the asteroid belt; however, w e currently d o not k n o w h o w w e l l e a c h asteroid class g r o u p s o b j e c t s of similar s u r f a c e c o m p o s i t i o n s . It is also difficult to d e t e r m i n e w h a t these distributions actually m e a n d u e to the e f f e c t s of space w e a t h e r i n g , w h i c h appears to a f f e c t the reflectance spectra of the s u r f a c e s of asteroids, and d y n a m i c a l m i x i n g , w h i c h affects the orbital distribution of objects. S a m p l e return m i s s i o n s should allow u s to better d e t e r m i n e h o w well w e c a n derive the surface c o m p o s i t i o n s of asteroids f r o m r e m o t e sensing.

ACKNOWLEDGMENTS A l l meteorite reflectance spectra are f r o m the B r o w n University K e c k / N A S A R e l a b S p e c t r a Catalog. A s t e r o i d spectra w e r e primarily taken f r o m the S M A S S w e b s i t e (smass.mit. edu). T H B a c k n o w l e d g e s s u p p o r t f r o m N A S A C o s m o c h e m i s t r y grant N A G 5 - 1 2 8 4 8 and a F i v e - C o l l e g e A s t r o n o m y D e p a r t m e n t F e l l o w s h i p . T h e w o r k of T M D w a s s u p p o r t e d by the C o n c e l h o N a c i o n a l d e D e s e n v o l v i m e n t o Científico e T e c n o l o g i c o - C N P q / B r a s i l . W e w o u l d like to t h a n k Takahiro Hiroi and D u c k M i t t l e f e h l d t f o r very h e l p f u l reviews.

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Trombka JI, Squyres SW, Brückner J, Boynton WV, Reedy RC, McCoy TJ, Gorenstein P, Evans LG, Arnold J R , Starr RD, Nittler LR, Murphy M E , Mikheeva I, McNutt R L Jr, McClanahan TP, McCartney E, Goldsten JO, Gold RE, Floyd S R , Clark PE, Burbine TH, Bhangoo J S , Bailey SH, Petaev M (2000) The elemental composition of asteroid 433 Eros: Results of the NEAR- Shoemaker X - R a y spectrometer. Science 289:2101-2105 UedaY, Hiroi T, Pieters CM, Miyamoto M (2002a) Changes of Band I Center and Band Ii/Band I Area Ratio in reflectance spectra of olivine-pyroxene mixtures due to the space weathering and grain size effects. Lunar Planet Sei X X X I I I : 2 0 2 3 UedaY, Miyamoto M, Hiroi T (2002b) Expanding the modified Gaussian model: Estimating the composition of olivine-rich asteroid. Meteor Planet Sei 37:A141 Vernazza P, Birlan M, Rossi A, Dotto E, Nesvorny D, Brunetto R, Fornasier S, Fulchignoni M, Renner S (2006a) Physical characterization of the Karin family. Astron Astrophys 460:945-951 Vernazza P, Brunetto R, StrazzullaG, Fulchignoni M, Rochette P, Meyer-Vernet N, Zouganelis I (2006b) Asteroid colors: a novel tool for magnetic field detection? The case of Vesta. Astron Astrophys 4 5 1 : L 4 3 - L 4 6 Vernazza P, Mothe-Diniz T, Barucci MA, Birlan M, Carvano JM, Strazulla G, Fulchignoni M, Miglorini A (2005) Analysis of near-IR spectra of 1 Ceres and 4 Vesta, targets of the Dawn mission. Astron Astrophys 436:1113-1121 Viateau B (2000) Mass and density of asteroids (16) Psyche and (121) Hermione. Astron Astrophys 354:725731 Vilas F (1994) A cheaper, faster, better way to detect water of hydration on solar system bodies. Icarus 1 1 1 : 4 5 6 467 Vilas F, Cochran AL, Jarvis K S (2000) Vesta and the Vestoids: A new rock group? Icarus 147:119-128 Vilas F, McFadden L A (1992) CCD reflectance spectra of selected asteroids: I. Presentation and data analysis considerations. Icarus 1 0 0 : 8 5 - 9 4 Weisberg M K , McCoy TJ, Krot AN (2006) Systematics and evaluation of meteorite classification. Lauretta DS, McSween H Y Jr (eds), University of Arizona Press, Tucson, p 19-52 Weissman PR, A'Hearn MF, McFadden LA, Rickman H (1989) Evolution of comets into asteroids. In: Asteroids II. Binzel RP, Gehreis T, Matthews M S (eds) Univ Arizona Press, Tucson, p 880-920 Wentworth SJ, Keller LP, McKay DS, Morris RV (1999) Space weathering on the Moon: Patina on Apollo 17 samples 75075 and 76015. Meteor Planet Sei 34:593-603 Wetherill G W (1992) An alternative model for the formation of asteroids. Icarus 1 0 0 : 3 0 7 - 3 2 5 Wood B J (1974) Crystal field spectrum of Ni 2 + in olivine. Am Mineral 59:244-248 X u S, Binzel RP, Burbine TH, Bus S J (1995) Small main-belt asteroid spectroscopic survey: Initial results. Icarus 115:1-35 Yamada M, Sasaki S, Nagahara H, Fujiwara A, Hasegawa S, Yano H, Hiroi T, Otake H (1999) Simulation of space weathering of planet-forming materials: Nanosecond pulse laser irradiation and proton implantation on olivine and pyroxene samples. Earth Planets Space 51:1255-1265 Zappalä V, Cellino A, Dell'Oro A, Paolicchi P (2002) Physical and dynamical properties of asteroid families. In: Asteroids III. Bottke WF, Cellino A, Paolicchi P, Binzel RP (eds), University of Arizona Press, Tucson, p 619-631 Zappalä V, Cellino A, Farinella P, Milani A (1994) Asteroid families: II. Extension to unnumbered multiopposition asteroids. Astronom J 107:772-801 Zellner B (1979) Asteroid taxonomy and the distribution of compositional types. In: Asteroids. Gehreis T (eds), University of Arizona Press, Tucson, p 7 8 3 - 8 0 6 Zellner B, Leake M, Morrison D, Williams J G (1977) The E asteroids and the origin of the enstatite achondrites. Geochim Cosmochim Acta 41:1759-1767 Zellner B, Thirunagari A, Bender D (1985a) The large-scale structure of the asteroid belt. Icarus 62:505- 511 Zellner B, Tholen DJ, Tedesco E F (1985b) The eight-color asteroid survey: Results for 589 minor planets. Icarus 61:355-416 Zellner B, Wisniewski WZ, Anderson L, Bowell E (1975) Minor planets and related objects. X V I I I . U B V photometry and surface composition. Astron J 80:986-995 Ziffer J, Campins H, Licandro J, Pinilla-Alonso N, Fernandez Y, Bus S J (2005) Near infrared spectra of two asteroids with low Tisserand invariant. Earth Moon Planets 97:203-212 Zolensky ME, Bodnar RJ, Gibson E K , Nyquist LE, Reese Y, Shih C-Y, Wiesmann H (1999) Asteroidal water within fluid inclusion-bearing halite in an H5 chondrite, Monahans (1998). Science 2 8 5 : 1 3 7 7 - 1 3 7 9 Zolensky M, Ivanov A (2003) The Kaidun microbreccia meteorite: A harvest from the inner and outer asteroid belt. Chem Erde 63:185-246

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 345-397, 2008 Copyright © Mineralogical Society of America

Oxygen Isotopes in Asteroidal Materials Ian A. Franchi Planetary and Space Sciences Research Institute Open University Milton Keynes, MK7 6AA, United Kingdom i.a.franchi@ open.ac. uk

ABSTRACT Measurement of the variation of the 1 7 0 / 1 6 0 and 1 8 0 / 1 6 0 ratios of extraterrestrial materials has proven to be a key tool in understanding the formation and evolution of the Solar System. This chapter attempts to collate the huge data set of oxygen isotopic measurements of asteroidal material, as sampled by meteorites, in order to understand some of the processes involved in the formation of these primitive bodies and how such events have affected the oxygen isotopic ratios, ultimately offering a window back to the very origin of Solar System and its primordial oxygen isotopic heterogeneity. Oxygen is generally a major element in the gas, liquid and solid phases that have interacted throughout the evolution of the Solar System, its chemical properties ensuring that there is abundant opportunity for reactions and exchange, the details of which can be recorded in the isotopic ratios of the products. Such processes can be crudely subdivided into the following six categories: •

Early solar nebular isotopic heterogeneity

•

High-temperature modification of nebular components

•

Metamorphism and melting in asteroids

•

Aqueous alteration in asteroids

•

Mixing of components/brecciation

•

Terrestrial weathering

It is clear from the isotopic variation in the earliest-formed nebular components, such as refractory inclusions and chondrules, that there was considerable oxygen isotopic heterogeneity in the early solar nebula, spanning over 50%c, and perhaps much greater, in both 5 1 7 0 and 5 l s O . The origin of this variation is discussed elsewhere in this volume. However, isotopic exchange between the l s O-rich solids and 1 7 0- and l s O-rich gas phase(s) resulted in a wide range of isotopic variation in these early-formed components, offering insight into the conditions of these processes and the evolution of the reservoirs with time and/or space. Once accreted onto the early planetisemals, much of the primitive nebular components experienced considerable modification. The metamorphism and melting experienced by some of these bodies, most probably the result of heating from the decay of short-lived radionuclides (e.g., 26A1), resulted in progressive homogenization, within individual bodies, of any original isotopic heterogeneity. Complete homogenization appears to have been achieved in the HED parent body, indicating very high levels of melting and the development of a large magma ocean. The carbonaceous and ordinary chondrite parent bodies experienced more modest heating, sufficient to mobilize water ice that accreted along with the more refractory materials. The subsequent reactions between the liquid and solid phases at low temperature imparted large isotopic fractionations which record some of the conditions of this alteration, such as watenrock ratio, temperature, and isotopic signature, and thus potentially the origin of the water ice. 1529-6466/08/0068-0013$ 10.00

DOI: 10.2138/rmg.2008.68.13

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Franchi The evolution of the asteroidal bodies has been far from simple, and an important use of oxygen isotope measurements has been in the identification of the components present in some meteorites, the result of breakup, reassembly and mixing of different bodies during impact processing. The final oxygen isotopic signatures imparted into meteorite samples are those associated with terrestrial weathering. The isotopic compositions of the weathering products depend on location (primarily latitude) while the magnitude of the weathering is dependent on many factors, with some meteorites (e.g., highly reduced enstatite chondrites and the matrix-rich C 0 3 s ) particularly susceptible to such effects.

INTRODUCTION Determination of the relative ratios of the three stable isotopes of oxygen, l e O, 1 7 0 and in asteroidal materials via the analysis of meteorites has developed over the past 30 years into a major cornerstone for the study of early Solar System science. Pioneered by Robert Clayton and co-workers of the Enrico Fermi Institute at the University of Chicago, it first came to prominence with the discovery of l e O excesses, relative to 1 7 0 and l s O, compared to normal terrestrial oxygen in refractory components from the Allende meteorite (Clayton et al. 1973). Virtually all types of extraterrestrial samples were analyzed by the Chicago group in the following years, developing a picture of the complex oxygen isotopic heterogeneity that exists in the sampled parts of the Solar System (e.g., Clayton 1993). In the late 1980s and early 1990s a proliferation of new laboratories and analytical techniques also began generating oxygen 3-isotope data. These new techniques, including laser heating fluorination, laser ablation fluorination and secondary ionization mass spectrometry (SIMS), extended the range of applications and types of material that could be studied, continually reducing the required sample size and improving analytical precision or spatial resolution. l s O,

More or less concurrent with the expansion of the analytical base was the rapid increase in numbers of new meteorites collected through search programs, initially in Antarctica and later in hot deserts (Nullarbor, Sahara, Atacama, Oman, etc.). These collections have yielded exciting new samples of previously unknown or very rare meteorite types. This offered fresh insight into the early Solar System, while the sheer numbers of new meteorites provided the opportunity to investigate systematic patterns within meteorite groups in far greater detail than had previously been possible. The objective of this chapter is to synthesize the available information on oxygen isotopic variation in asteroidal materials sampled by our extensive meteorite collections. Data have been drawn from a wide variety of sources, including reaction bomb and laser heating fluorination, laser ablation and SIMS studies. The intention is to attempt to understand the variation in oxygen isotopic composition between different meteorite types, the variations between meteorites with a common origin, and the variation that exists at the microscopic scale within individual meteorites. These data are then used to constrain the processes that were responsible for establishing the observed isotopic patterns and to discuss the implications for the asteroidal parent bodies. Some of the variation reflects processes in the very early solar nebula and potentially prior to solar nebula formation—these aspects, although discussed briefly where appropriate, are largely beyond the scope of this chapter and are covered elsewhere (Young et al. 2008; Yurimoto et al. 2008). Oxygen isotope data within this chapter are reported as conventional 8 1 7 0 and 8 l s O values (for details, see Criss and Farquhar 2008). Much reference is made to the A 1 7 0, the 1 7 0 excess or depletion relative to the terrestrial fractionation line, defined as A 1 7 0 = S 1 7 0 - (0.52 x S l s O)

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However, the latest generation of laser fluorination instruments provide levels of analytical precision which make the effects of the linear approximation of the delta notation and the exact magnitude of the slope of the terrestrial fractionation line significant parameters. In a number of sections within this chapter the definition of the 1 7 0 excess is revised after Miller (2002) , SO , SO A 17 0* = 1000 In 1 + - A, 1000 In 1 + 1000 1000 where X = slope of the terrestrial fractionation line. The slope of the terrestrial fractionation line has been reported differently from various labs over the years, although there may be convergence on a value around 0.525 for most silicate phases (Rumble et al. 2007). Use of A 17 0* in this chapter is limited to discussion of meteorite groups where the original reports used such a notation—primarily the basaltic achondrites and associated groups (i.e., HEDs, angrites, mesosiderites and pallasites). Within this chapter, oxygen isotopic data are presented on traditional (8 1 7 0 vs. 8 l s O) three-isotope plots and on A 1 7 0 versus 8 l s O plots. In the former (for example, Fig. 1) the terrestrial fractionation line, and all other mass fractionation lines, have a slope of ~0.52. Any group of samples which originated from a single homogeneous reservoir and whose oxygen was only affected by mass-dependent fraction processes would fall along a line of slope ~0.52 on a three-isotope plot of 8 1 7 0 versus 8 l s O. Any group of samples defining an array with a slope distinct from ~0.52 are the result of mixing between two or more oxygen reservoirs with distinct oxygen isotopic compositions. Mass-independent fractionation can generate fractionation lines with steeper slopes (e.g., ozone formation fractionates 1 7 0 and l s O equally relative to le O, generating a slope = 1 fractionation line on a three-isotope plot). Direct evidence of such processes affecting asteroidal materials is difficult to establish, however, and is covered in more detail elsewhere (Young et al. 2008; Yurimoto et al. 2008). Plots of A 1 7 0 versus 8 l s O essentially orient the graph so that the terrestrial fractionation line is horizontal (for example, see Fig. 6). Such plots are useful for displaying small variations in A 17 0, particularly where there is a large spread in 8 l s O. Throughout this chapter, three reference lines are shown in most plots: the terrestrial fractionation line (A 17 0 = 0%c by definition); and two interpretations of the primitive nebular mixing line initially identified by Clayton et al. (1973). The CCAM line (slope =0.94) is defined by all anhydrous components from refractory phases in CV chondrites (Clayton 1993) while the Y&R=1 line (slope =1) is defined by a much smaller number of analyses of alteration-free components in the Allende meteorite (Young and Russell 1998). These are discussed further in the CV Chondrite section. The range in whole-rock isotopic compositions of meteorites is quite limited, spanning little more than 20%o in 8 1 7 0 and 8 l s O. Analytical precision is such that even with this relatively limited range the primitive chondritic meteorites (Fig. 1) clearly display many of the isotopic patterns already discussed. Groups such as the equilibrated ordinary chondrites (H, L and LL) and R chondrites show trends parallel to the terrestrial fractionation line, indicating that the dominant processes in their ultimate formation involved mass fractionation of largely homogeneous reservoirs. Other groups, such as the CRs and CMs, fall along lines of slope ~0.7, indicative of mixing between two reservoirs, while the CV and CK meteorites fall along a line of slope ~1, a different mixing line which was the result of different processes or starting compositions. The achondrites, having experienced high levels of thermal metamorphism or melting, display much less oxygen isotopic variation (Fig. 2) than the chondritic meteorites. Many of the groups fall within l%c of the terrestrial fractionation line, which has necessitated the need for high analytical precision and appropriate use of A 17 0*, and generally data are better displayed on A 1 7 0 vs. 8 l s O plots.

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Ô180(°/oo) Figure 1. Schematic plot of oxygen isotopic variation of whole-rock samples of chondritic meteorites. TFL is the terrestrial fractionation line; CCAM is a line defined by anhydrous components from hightemperature chondrules and refractory inclusions in CV meteorites; and Y&R=1 is a line from alterationfree components in one refractory inclusion from the CV3 Allende. Data from many sources. See text throughout chapter for references.

CCAM

Ô180(°/oo) Figure 2. Schematic plot of ranges of oxygen isotopic compositions of whole-rock samples of asteroidal achondritic meteorites. IAB/Win = IAB complex; HE = HE iron meteorites; Acap/Lod = acapulcoites and lodranites; Ang = angrites; Aub = aubrites; Brach = brachinites; HED/Mes = howardites, eucrites, diogenites and mesosiderites; Pall = pallasites; Ure = ureilites. Data from many sources. See text throughout chapter for references.

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Materials

ORDINARY CHONDRITES Introduction In this section, the oxygen isotopic variations of the ordinary chondrites are considered. The H, L and LL chondrite groups are very closely related and appear to be constructed from largely identical components, but in different proportions—primarily in the relative proportions of Fe-Ni metal and sulfides to silicate-rich chondrules. Differences in their oxygen isotopic signatures, however, indicate that oxygen-containing components from the different groups sampled different reservoirs or that there were variable inputs from more than one reservoir. In terms of attempting to understand these differences it is useful to consider the metamorphosed, or equilibrated ordinary chondrites (EOCs), separately from the unequilibrated ordinary chondrites (UOCs). The EOCs permit characterization of the differences between the groups on a broad scale while the limited thermal metamorphism of the UOCs allows investigation of the primary components which make up the ordinary chondrites. Ordinary chondrites - whole-rock Both the UOCs and EOCs display limited ranges of whole-rock oxygen isotopic compositions, and therefore, when determining the extent of such variation, the best information is obtained from observed falls. In contrast to virtually all other meteorite types, the ordinary chondrites are extremely abundant, and therefore such a limitation does not significantly affect our ability to identify patterns or ranges of isotopic variation. Clayton et al. (1991) reported a large number of results from ordinary chondrite falls, and the results are summarized in Figure 3 together with a small number of other results. The equilibrated H, L and LL chondrites, with mean A 17 0 values of 0.73 ± 0.09%c, 1.07 ± 0.09%c and 1.26 ± 0.12%c, respectively (Table 1), can be resolved into three distinct groups on an oxygen 3-isotope plot, albeit with some overlap, particularly between the L and LL chondrites. The scatter within each of the ordinary chondrite (OC) groups is more or less consistent with mass fractionation of a common reservoir. Indeed, a set of H chondrites analyzed using a laser fluorination technique, which offers higher analytical

O

H Chondrite (Folco et al. 2004) H Chondrite

n •

L Chondrite L L Chondrite

S• •

"»O e

v i

^

1

1

M z

o -

1—

a Burnwell"

r

TFL^—

ô

i

1 8

i

0

i

i

(%o)

Figure 3. Oxygen isotopic compositions of equilibrated ordinary chondrites observed to fall. The bulk of the data are from Clayton et al. (1991) with additional data, not shown separately, from Jabeen et al. ( 1998), Stepniewski et al. ( 1998), Grossman (2000), Paliwal et al. (2001), Reimold et al. (2004), Russell et al. (2004), and Simon et al. (2004). The anomalous H chondrite Burnwell is shown (Russell et al. 1998), as are a set of high-precision analyses of a suite of H chondrites (Folco et al. 2004).

350

Franchi Table 1. Mean oxygen isotopic compositions of ordinary chondrite (falls) groups.

Class

5170

5180

A170

H3 H4 H5 H6 Mean H4,5,6 Mean H4,5,6 (Folco et al. 2004)

2.77 2.88 2.87 2.79 2.85 2.70

±0.10 ±0.13 ±0.18 ±0.12 ± 0.15 ±0.16

4.04 4.13 4.11 3.97 4.08 3.72

±0.12 ±0.24 ±0.25 ±0.16 ± 0.22 ±0.29

0.68 ± 0.04 0.74 ± 0.09 0.73 ±0.11 0.72 ± 0.05 0.73 ± 0.09 0.77 ±0.04

L3 L4 L5 L6 Mean L4,5,6

3.40 3.55 3.57 3.46 3.52

± 0.08 ± 0.20 ± 0.09 ±0.11 ± 0.14

4.71 4.79 4.75 4.60 4.70

±0.16 ± 0.34 ±0.18 ±0.19 ±0.24

0.97 ±0.15 1.06 ± 0.07 1.10 ±0.07 1.06 ±0.11 1.07 ± 0.09

LL3 LL4 LL5 LL6 Mean LL4,5,6

4.01 4.01 3.90 3.79 3.88

±0.17 ±0.11 ± 0.22 ±0.11 ±0.16

5.60 5.25 5.02 4.92 5.04

± 0.33 ± 0.25 ± 0.21 ±0.17 ± 0.24

1.10 ±0.10 1.28 ± 0.09 1.29 ±0.12 1.23 ±0.13 1.26 ± 0.12

Data from Clayton et al. (1991) and Folco et al. (2004).

precision than most of the other results, shows good evidence that the samples fall along a mass-fractionation line (Fig. 3) with a A 1 7 0 value of 0.77 ± 0.04% c (Folco et al. 2004). The scatter is still greater than that observed among samples from larger, melted bodies (e.g., Earth, Mars, Vesta)—although this appears to be largely due to inclusion of the potentially anomalous, highly shocked (S6) and veined meteorite Rose City. Whether similar, more clearly defined mass fractionation lines exist for the L and LL groups has yet to be established. Table 1 shows that within the EOCs there is a hint of a trend of decreasing 8 l s O with increasing metamorphic temperature from type 4 to type 6 (Clayton et al. 1991), although the amount of variation within each petrological type considerably limits any confidence in this observation. The oxygen isotopic variation of the EOCs is consistent with essentially closed-system metamorphism of common precursors for each of the three OC groups—H, L and LL (Clayton et al. 1991). There are numerous examples of OC meteorites which are difficult to unambiguously assign to one of the groups on the basis of petrographic information, but these meteorites have oxygen isotopic compositions within the range of the main H, L and LL meteorites. The Burnwell meteorite is an anomalous equilibrated chondrite, described as a low-FeO chondrite. It displays a number of mineralogical and chemical features which indicate that it formed under more reducing conditions than the H chondrites (Russell et al. 1998). It also has an oxygen isotopic composition distinct from the H, L and LL chondrites (Fig. 3), with a A 1 7 0 value of 0.51%o, and therefore Russell et al. (1998) suggested it is evidence for an extension of the H, L and LL pattern of reducing/oxidizing conditions and oxygen isotopic signatures. However, the 8 l s O value of this meteorite is more than 2%c heavier than the expected oxygen isotopic trend of the OCs, perhaps indicating that another process, in addition to those responsible for the variation observed in other OCs, was in some part responsible for the anomalous signatures observed in Burnwell. Observed UOC falls are far less common than those of EOCs, but a sufficient number have been analyzed, primarily by Clayton et al. (1991), to establish their general oxygen isotopic signatures. Two observed H3 falls have a mean A 1 7 0 value of 0.68 ± 0.04%o, essentially indistinguishable from the H4 to H6 meteorites analyzed in the same study (Fig. 4; Table 1),

Oxygen

Isotopes in Asteroidal

351

Materials

1.4

1.0

O 0.6

0.2 3

4

5 6

1 8

0

6

(%O)

Figure 4. Oxygen isotopic compositions of unequilibrated ordinary chondrites. Shaded areas show the ranges for H, L and LL equilibrated ordinary chondrite falls (see Fig. 3). Falls (large symbols) and finds (small symbols) are shown separately. Data from Clayton et al. (1991) and Li et al. (2000).

although possibly distinct from the 0.77 ± 0.04%e reported by Folco et al. (2004). Isotopic compositions of a number of H3 finds are also plotted in Figure 4, and they reveal a similar pattern, but with somewhat greater scatter, most probably the result of terrestrial weathering. The L3 meteorites have oxygen isotopic signatures close to that of the equilibrated L chondrites, although once again the L3 mean A 17 0 (0.97 ± 0.15%o) hints at a lower value than the equilibrated L chondrites. L3 finds display a similar pattern to the L3 falls (Fig. 4), but with greater scatter, generally to lower A 17 0 values, most likely as a result of terrestrial weathering. Tieschitz and Bremervörde, both often classified as H/L3 meteorites, have oxygen isotopic compositions indistinguishable from the L3 meteorites. The mean oxygen isotopic composition of the LL3s reported by Clayton et al. (1991) is different from that of the more equilibrated LL chondrites. The mean A 17 0 value is 1.10 ± 0.10%e, compared to the LL4 to LL6 meteorite mean of 1.26 ± 0.12%e. Only two meteorites, Bo Xian (Li et al. 2000) and St. Mary's County, have A 17 0 values that approach the mean value of the equilibrated LL chondrites. Also of note is that the 5 l s O of some of the LL3s is shifted to heavier values, by up to l%c compared to the mean equilibrated meteorites, with a clear trend of increasing 5 l s O away from the mean equilibrated LL3 value, with decreasing petrologic type. Ordinary chondrites - components The limited degree of mineral equilibration in the UOCs allows investigation of the oxygen isotopic composition of individual components and how these isotopic heterogeneities evolve as thermal metamorphism progresses. Chondrules. A considerable number of individual chondrules from ordinary chondrites have been analyzed by conventional or laser fiuorination techniques (e.g., Clayton et al. 1991; Bridges et al. 1998; Li et al. 2000). In general, the oxygen isotopic compositions of the chondrules, while overlapping with the whole-rock EOC and UOC measurements, also extend to higher 5 l s O values. A small number of chondrules have 5 1 7 0 and 5 l s O values several %o beyond the range of the whole-rock measurements (Fig. 5). A striking feature of these measurements is that chondrules from any one group, or indeed any one meteorite, display a range of oxygen isotopic compositions independent of the oxygen isotopic composition of the

Franchi

352

6

-2

-2

0

2

4 s

1 8

0

6

8

10

(%o)

Figure 5. Oxygen isotopic compositions of individual chondrules from ordinary chondrites (Clayton et al. 1991; Bridges etal. 1998). The fields of H, L and LL EOC finds (Clayton et al. 1991) are superimposed.

host meteorite. There is no significant indication of any clustering of chondrule compositions f r o m different meteorites or types. Indeed, for those meteorites where a sizeable number of chondrules have been analyzed, there are very similar ranges in isotopic composition. Bridges et al. (1998) identified interdependent correlations between chondrule mineralogy, chemical composition, size and oxygen isotopic composition f r o m a suite of 40 individual chondrules f r o m the L L 3 chondrites Chainpur and Pamallee. The main controlling factor for the oxygen isotopic composition of the chondrules was f o u n d to b e the nature of the mesostasis, with smaller chondrules containing m o r e glassy, S i 0 2 - r i c h mesostasis being m o r e enriched in 17 0 and l s O . A n earlier study f o u n d no chondrule features correlated with oxygen isotopic composition (Clayton et al. 1991). T h e analysis of individual c h o n d r u l e s is limited in its ability to accurately describe all the o x y g e n isotopic systematics of ordinary chondrites, as the sample size r e q u i r e m e n t s limit the sample selection to the largest chondrules. This is best exemplified by the H chondrites, w h e r e individual c h o n d r u l e s display a r a n g e of o x y g e n isotopic c o m p o s i t i o n s almost entirely outside the range of the w h o l e - r o c k values for H chondrites (Fig. 6)—indicating the presence of other, significant o x y g e n c o m p o n e n t s (Clayton et al. 1991). C o m p o s i t e s a m p l e s of small chondrules f r o m the H 4 chondrites b e c o m e increasingly l e O - r i c h as the size fraction decreases below ~300 |im. M a t r i x material separated f r o m the D h a j a l a meteorite has an o x y g e n isotopic composition even m o r e enriched in l e O than the smallest chondrule or mineral f r a g m e n t size fraction and lies on an extension of the composite sample trend (Fig. 6). M a t r i x f r o m A L H A 77299 (H3.7) has even m o r e a n o m a l o u s o x y g e n , with a A 1 7 0 value similar to that of D h a j a l a but with a 5 l s O value of almost -2%c (Brearley et al. 1989). It is possible that isotopically light Antarctic ice water h a s affected the o x y g e n isotopic composition of this fine-grained material (Clayton et al. 1991). Very little data exist on the o x y g e n isotopic composition of O C matrix. W h i l e the H chondrites typically have the most l e O - r i c h matrix compositions, a matrix sample f r o m Weston, an H4, has an o x y g e n isotopic composition indistinguishable f r o m bulk H chondrites. Similarly, while most L and L L chondrite matrix samples are indistinguishable f r o m their w h o l e - r o c k compositions, M e z o - M a d a r a s (L3.7) contains matrix with a A 1 7 0 value similar to that of D h a j a l a , but has a slightly different 5 l s O value of ~3%e (Nehru et al. 1991).

Oxygen

Isotopes in Asteroidal —I

1

Individual chondrules

^

Materials

353

r

Matrix

H 3-6 WR

Chondrule Composites Fragment Composites

6 1 8 0 (%O) Figure 6. Oxygen isotopic compositions of chondrules and other components in the H chondrites Dhajala and Weston. Data are from Clayton et al. (1991). The chondrule composite and fragment composite samples are various size fractions ranging from CM > H > EH. Refractory siderophile elements follow the same trend of decreasing abundance with chondrite group. There are relatively few refractory siderophile elements, however, and scatter in the data make it difficult to determine whether they are at uniform abundance within a given chondrite group. They appear to be in uniform abundance for H and EH chondrites, but may not be for CV and CM chondrites. To first order, there is no clear correlation with condensation temperature or cosmochemical behavior. The most plausible cause for the decreasing abundance sequence given above is variation in the amount of high-temperature condensates (solids that formed before the major silicate and metal phases) that accreted to the different chondrite parent bodies (Palme and Jones 2003). Moderately volatile elements show a different pattern. Figure lb is a plot of moderately volatile siderophile element/Si ratios vs. 50% condensation temperature for the same chondrite groups. There is a general trend of falling abundance with decreasing temperature. The CV and H chondrites show the greatest depletions in the most volatile elements plotted, while CM chondrites show the least. The EH chondrites follow this basic pattern, but with a slight twist. Their moderately volatile elements first show an increase in abundance above that of the main component, and then decrease with decreasing condensation temperature. The EH chondrites have abundances of moderately volatile siderophile elements greater than CI for those with 50% condensation temperatures >1000 K (Fig. lb). Similar patterns are shown by the moderately volatile lithophile elements. The pattern of decreasing abundance with decreasing condensation temperature is generally considered to reflect separation of dust from gas as temperature was falling in the nebula (Wai and Wasson 1977; Humayun and Cassen 2000; Palme and Jones 2003). Thus, the bulk of the dust that went into forming CM chondrites maintained equilibrium with the nebular gas to lower temperatures than did that of the H chondrites. Evidence for metal-silicate fractionation is shown on a plot of Cl-normalized Ir/Si vs. Al/Si (Fig. 2a). Aluminum and Ir both have high condensation temperatures, 1653 and 1603 K, respectively (Lodders 2003), but the former is lithophile and the latter siderophile. In roughly half of the chondrite groups, chondrite-normalized Ir/Si = chondrite-normalized Al/Si. The

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5 0 % condensation temperature (°K) Figure 1. Plots of Cl-normalized element/Si ratios vs. 50% condensation temperature for CM, CV, H and EH chondrites illustrating element fractionations in chondrites, (a) Refractory lithophile elements show uniform abundances uncorrelated with temperature, but with different mean ratios in different chondrite groups (dashed lines). Similar patterns hold for refractory siderophile elements. The three elements at the lowest temperatures are main component lithophile elements, (b) Moderately volatile siderophile elements show decreasing abundances with decreasing temperature. Similar patterns hold for moderately volatile lithophile elements. The four elements at the highest temperatures are main component siderophile elements. Chondrite data are from Lodders and Fegley (1998); condensation temperatures are from Lodders (2003).

CV and K chondrites have lower mean refractory siderophile element/Si ratios than mean refractory lithophile element/Si ratios. Clearly anomalous are the three ordinary chondrite groups. They form a vertical trend on Figure 2a, with identical Al/Si but with Ir/Si varying by over a factor of 2 from H chondrites to LL chondrites (arrow). Decreases in siderophile element/Si ratios in the sequence H-L-LL are seen for all but the most volatile siderophile elements. There is a general trend of decreasing (siderophile element/Si) H /(siderophile element/ Si)LL with decreasing condensation temperature (Fig. 2b). These observations are most simply explained as being caused by separation of the metal phase from the silicates (Wasson 1972), and that this process was more efficacious at higher temperature. The last chemical fractionation we will discuss is variation in oxygen content, manifested in variation in oxidation state. At high temperatures, grains formed in the nebula are reduced, and all Fe is metallic (Grossman 1972; Grossman and Larimer 1974). With falling temperature, Fe metal begins to be oxidized by gaseous H 2 0, and can react with magnesian silicates to form ferroan solid solutions of olivine and pyroxene (Grossman 1972; Grossman and Larimer 1974).

Oxygen

in Meteorites

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Figure 2. Graphs illustrating metal-silicate fractionation in ordinary chondrites, (a) Cl-normalized Ir/Si vs. Al/Si for average compositions of major chondrite groups. Most chondrites follow a 1:1 line indicating little or no fractionation of Ir (siderophile) f r o m A1 (lithophile). Ordinary chondrites have identical Al/Si ratios, but very different Ir/Si as a result of metal silicate fractionation. (b) Siderophile element/Si ratios for H chondrites normalized to those of L L chondrites show a decreasing trend with decreasing condensation temperature for the element, suggesting metal-silicate fractionation decreased as the nebula cooled. Tungsten and M o have low ratios, possibly due to partial oxidation of these elements. Silver is anomalous. Data sources as in Figure 1.

800

50% condensation T (K)

Chondritic meteorites have recorded this process at various stages of completion. Figure 3 is a schematic Urey-Craig plot, a graph of reduced Fe/Si vs. oxidized Fe/Si for bulk chondrites. The enstatite chondrite groups are highly reduced, and all Fe is in metal and sulfide. The CI and CM chondrites are highly oxidized, and almost all Fe is FeO or Fe 2 0 3 . Note, however, that the highly oxidized character of these chondrites is in part due to parent body processes, not nebular. On this diagram, lines of slope - 1 are lines of constant EFe/Si. The three ordinary chondrite groups show that metal-silicate fractionation, with EFe/Si decreasing in the sequence H-L-LL, correlates with increasing oxidation (arrow). Oxygen isotope anomalies Although the basic observations of anomalous isotopic behavior in meteoritic oxygen have been well-known for over 30 years (Clayton et al. 1973), consensus on the underlying mechanism has not been achieved. The simplest illustration of the anomalous behavior is seen in a comparison of the isotopic compositions of silicon and oxygen in a suite of calcium-, aluminum-rich inclusions (CAIs) from the Allende CV3 meteorite. These are the two most abundant elements in the CAIs, and each has three stable isotopes. Oxygen and silicon were extracted simultaneously from 15 CAIs (Clayton et al. 1988). Figure 4 shows the behavior of silicon on a plot of 5 29 Si vs. 530Si, where: 8 Si =

( 2 9 Si/ 2 8 Si)sample 29

( Si/

2S

Si), , ,

i

1000

(units of per mil)

with analogous definitions for 530Si, 5 1 7 0, and 5 l s O (the latter two based on ratios to le O). The measured line has a slope of 0.50, which is determined by the ratio of mass-difference

404

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0.6

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oxidized Fe/Si (molar) Figure 3. Schematic Urey-Craig plot of reduced Fe/Si vs. oxidized Fe/Si for chondrite groups demonstrating the range in oxidation state. Reduced Fe is Fe as metal and sulfide. Lines of slope - 1 on this diagram have constant EFe/Si. Ordinary chondrites demonstrate decreasing EFe/Si (metal-silicate fractionation) with increasing oxidation. Diagram modified after Krot et al. (2003).

: O Allende CAIs |

CM CO

m

CO -2

-2

0

Ô J0 Si

Figure 4. Three-isotope graph for silicon from 15 Allende whole CAIs. The line of slope 1/2 is a mass-dependent fractionation line, with variations resulting from processes of evaporation (positive 5 values) and condensation (negative 5 values). Points near zero are near chondritic in isotopic composition. NBS-28 is a terrestrial quartz standard. Data from Clayton etal. (1988).

(%o rel. N B S - 2 8 )

on the y-axis (29-28) to mass-difference on the x-axis (30-28). These variations in isotopic compositions are the result of high-temperature evaporation of Si from partially molten CAIs into space. Oxygen isotopic compositions of the same CAIs are shown in Figure 5, and a comparison of the two elements is given in Figure 6. It is obvious that there is no correlation between the mass-dependent fractionation in silicon and the non-mass-dependent variations in oxygen. The slope of the oxygen array is very near unity (0.98 for the data in Fig. 5), which corresponds to a near-constant ratio of 1 7 0 / l s 0 , with variable le O. There are two general classes of processes that could account for the slope 1 array: nuclear processes, based on the fact that l e O is synthesized in different nuclear foundries than 1 7 0 and l s O; and chemical processes, based on the rarity of 1 7 0 and l s O relative to l e O (1/2600 and 1/500, respectively). A nuclear origin would be expected to lead to correlations with nuclear abundance variations in other elements, such as magnesium and silicon. Figure 6 shows that this is not observed. A chemical origin might result in correlation of oxygen isotope abundances with other chemical properties in primitive materials, such as the compositions of

Oxygen

in Meteorites

and Terrestial

405

Planets

O Allende CAIs | Figure 5. Three-isotope plot for oxygen from exactly the same samples (and same chemical preparations) as those shown for silicon in Figure 4. The data fall on a mixing line of slope 0.98, probably resulting from interaction of two isotopic reservoirs, one 16 0-rich (lower end), one 16 0-poor (upper end). The line marked "terrestrial fractionation" is the locus of points from terrestrial rocks, minerals and waters with variations due to mass-dependent fractionation. The isotopic standard is SMOW (Standard Mean Ocean Water). In the photochemical shielding model, discussed in the text, the 16 0-rich end-member is the expected solar composition. Data from Clayton et al. (1988).

terrestrial fractionation

O 5

-10

CO

-20

-30 -30

-20

-10

6 1 8 0 (%o rel. SMOW)

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i

1

0

m

-2

o ,

-30

o

o,

-20

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O Allende CAIs") -10

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chondmles and/or matrix, and the abundance of refractory element-rich CAIs. The search for such correlations is the subject of this chapter. Mechanisms of non-mass-dependent isotope fractionation Within the category of chemical processes, two quite different reaction types have been proposed: (1) a symmetry-dependent, non-mass-dependent effect, first discovered in the laboratory by Thiemens and Heidenreich (1983), which has been especially thoroughly studied in the synthesis of ozone from molecular oxygen (for recent reviews, see Mauersberger et al. 2005; Thiemens 2006); and (2) a photochemical self-shielding effect, extensively discussed in the context of molecular clouds (Bally and Langer 1982; van Dishoeck and Black 1988), and proposed for the solar nebula by Clayton (2002) and Thiemens and Heidenreich (1983). Both of these processes involve non-equilibrium chemistry, and thus may impose recognizable signatures on other chemical properties of primitive materials, such as oxidation states of transition metals (Clayton 2005). These processes are discussed in detail in the chapter by Young et al. (2008), and are briefly reviewed here. In ozone synthesis, oxygen atoms, generated non-thermally, attack oxygen molecules in a process in which the intermediate species with one of the heavy rare isotopes ( 1 7 0 or ls O) occu-

406

Mittlefehldt,

Clayton,

Drake,

Righter

pying a terminal position has a longer lifetime than the symmetrical 1 6 0 - 1 6 0 - 1 6 0 , and hence has a greater probability of deactivation to form a stable ozone molecule (Gao and Marcus 2001). The result is enrichment of the product ozone in 1 7 0 and l s O in equal proportions, giving a slope 1 array on the three-isotope plot that closely resembles the meteoritic array (Fig. 7). This process occurs naturally in the Earth's stratosphere, producing ozone enriched in heavy oxygen isotopes by 80-100%o relative to 0 2 (Mauersberger et al. 2005; Thiemens 2006). The solar nebula was very reducing because of the high abundance of H (Grossman 1972), and neither 02 nor 0 3 was likely to have been in significant abundance. An attempt has been made by Marcus (2004) to explain 16 0-enrichment in CAIs by invoking reactions analogous to the ozone synthesis, with a metal atom in the central position. The symmetry-dependent non-mass-dependent model has not yet led to specific predictions of the solar oxygen isotope abundances or to expected correlations with other meteoritic properties. However, reactions of several gaseous metal oxides, including AlO, SiO, CaO and FeO with O or OH can potentially result in non-mass-dependent O anomalies being transferred to the metal oxides in the nebula (Thiemens 2006). Thiemens and Heidenreich (1983) suggested that photodissociation of nebular 02 may have led to the oxygen isotope anomalies observed in meteorites, but calculations by Navon and Wasserburg (1985) indicated that this was unlikely. The current photochemical self-shielding hypothesis is based on the ultraviolet photodissociation of gaseous carbon monoxide. This process involves absorption of narrow lines from the incident starlight continuum, with the wavelength of the lines at characteristic values for each isotopic species of CO. Because of the much greater abundance of l e O relative to 1 7 0 and l s O, the lines leading to dissociation of C l e O become saturated while those of C 1 7 0 and C l s O are unsaturated. Thus, the abundances of atomic 1 7 0 and l s O are enhanced in the interior of the gas cloud or solar nebula. These chemically reactive atomic species can then react with H 2 to form water, or with metallic atoms to form solid condensates. The constant 1 7 0 / l s 0 ratio results from these reactions, and the residual CO in the cloud interior is therefore depleted in C 1 7 0 and C l s O relative to C l e O. This depletion, by a factor of five, has been directly observed in molecular clouds (Sheffer et al. 2002). Because the photochemical process acts only in one direction—to enrich the reaction products in the heavy rare isotopes—it follows that the initial CO in the solar nebula had an isotopic composition near the le O-rich end of the mixing line in Figure 5, and that all inner Solar System solids with greater 1 7 0/ 1 6 0 and l s O/ l e O values have gone through the photochemical enrichment process. Thus, this model predicts that solar oxygen is about 5%

Figure 7. Oxygen three-isotope plot for ozone (filled symbols) and residual oxygen (open symbols) from the laboratory synthesis of O , from 0 2 by Thiemens and Heidenreich (1983). The slope and range of the compositions of the residual 0 2 are very similar to those seen in the CAI data in Figure 5, leading to the suggestion that such a chemical process may have been the cause of the meteoritic variations. -80

-80

-60

-40

-20

0

6 1 s O (%O rel. initial 0 2 )

20

40

Oxygen

in Meteorites

and Terrestial

407

Planets

enriched in l e O with respect to terrestrial oxygen. Photochemical processes, operating in a gas of solar composition, lead to decidedly non-equilibrium concentrations of oxygen species, notably to a high concentration of atomic oxygen (Sternberg and Dalgarno 1995). One might, therefore, expect a positive correlation in primitive meteoritic materials between heavy-isotope enrichment and degrees of oxidation of iron and other transition metals. CHONDRITIC METEORITES Micro- and meso-scale correlations In this section, we consider whether there are correlations between the compositions of components in chondrites and oxygen isotopic compositions that may constrain the origin of A 17 0 variations in the solar nebula. We are not concerned with asteroidal processes, for example aqueous alteration, that may have changed the isotopic compositions of accreted materials. Nor will we present a detailed overview of isotopic compositions of chondritic components; that is part of the subjects of the chapters by Franchi (2008) and Yurimoto et al. (2008). If the photochemical self-shielding picture is the appropriate framework for understanding the non-mass-dependent effects in meteoritic oxygen, then concentrations of oxidizing agents such as H 2 0 or atomic oxygen in the nebular gas may greatly exceed equilibrium values. This could be expected to yield a positive correlation between 17 0- and ls O-enrichment and the oxidation state of transition-metal ions, particularly iron. The scale of such a correlation is expected to be at the level of individual chondrules, their constituent minerals, and their rims because the FeO contents of these vary (e.g. Clayton et al 1983; McSween 1985; Pack et al. 2004). Several studies have sought correlations of FeO of mafic silicates with A 17 0, with mixed results (Fig. 8). Isolated olivine grains in the CI chondrite Orgueil show a positive correlation of fayalite content with A 17 0 (Leshin et al. 1997). In contrast, isolated olivine grains in the CV chondrite Allende show no obvious correlation, although the most ferroan olivine does have the highest A 17 0 (Fig. 8). Olivine in chondrules from Allende shows

10

20

olivine F a (mole %) Allende

•

.

°

• l ^ t S c P o B j e « « D 0 o O aé>

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5

10

• chondrule Ochondrule rim O grain

15

20

25

30

olivine F a (mole %) C O chondrites

O

• •

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cu %

OY-81020 • ALHA77307

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20

40

60

olivine F a (mole %)

Figure 8. A 1 7 0 vs. olivine fayalite content for separated ferromagnesian components in carbonaceous chondrites, (top) Isolated olivine grains from CI chondrite Orgueil show a positive correlation (data from Leshin et al. 1997). (middle) Bulk chondrules, chondrule rims and isolated olivine grains in CV chondrite Allende show a positive correlation (data from Clayton et al. 1983; McSween 1985; Rubin and Wasson 1987; Rubin et al. 1990; Choi et al. 1997; Pack et al. 2004). (bottom) Isolated grains, relict grains in chondrules and grains crystallized in chondrules in the CO chondrites ALHA77307 and Y-81020 show at best a weak positive correlation (data from Jones et al. 2000; Kunihiro et al. 2004). Note change in .v-axis scale in the bottom panel, and variable v-axis scales.

408

Mittlefehldt,

Clayton,

a positive correlation of A 17 0 with fayalite content, but olivine grains in chondrule rims show no such correlation. Olivine grains from a variety of textural settings in two CO chondrites, ALHA77307 and Y-81020, show at best a weak positive correlation of fayalite content with A 17 0 (Fig. 8). Figure 9 shows similar data for unequilibrated ordinary chondrites. Analyses of individual, essentially FeO-free olivine grains from ordinary chondrites show a wide range in A 17 0. In contrast, ordinary chondrite chondrules have a much smaller range in A 17 0, and no correlation with fayalite content.

Drake,

Righter

o O

-io -12

o

p j i i i i i i i i i i i i i i i 10

20

Ograin • chondrule

30

40

50

olivine Fa (mole %)

Rubin et al. (1990) measured oxFigure 9. A n O vs. olivine fayalite content for individual ygen isotopic abundances in Allende olivine grains and bulk chondrules f r o m ordinary ferromagnesian chondrules and their chondrites. Data are f r o m Bridges et al. (1998), Gooding coarse-grained rims, and found that et al. (1983) and Pack et al. (2004). all data fell on a line with a slope of 0.9 on the three-isotope plot, and that the rims are all iron-rich and le O-depleted relative to their host chondrules. This presents a stratigraphic sequence, with later-formed material being relatively richer in both ferrous iron and the heavier oxygen isotopes. Wasson et al. (2000) showed that bulk chondrules from the ungrouped carbonaceous chondrite LEW 85332 show a positive correlation of A 17 0 with bulk chondrule molar FeO/(FeO+MgO), which they suggested possibly indicates that A 17 0 of the nebular gas increased with time. At the scale of chondrite classes, there is a clear progression of both ferrous iron content and abundance of 1 7 0 and l s O from H- to L- to LL-chondrites (Clayton et al. 1991). However, the correlation may not be due directly to nebular chemical processes, since individual chondrules in the meteorites do not show such a correlation (Fig. 9). It should also be noted that the highly reduced enstatite chondrites have lower A 17 0 than ordinary chondrites, and highly oxidized R chondrites have the highest A 17 0 of all chondrites. The data on mafic silicate minerals and components in chondrites thus present a mixed signal regarding correlation of oxidation state with oxygen isotopic composition. There is some evidence suggesting that more ferroan components are enriched in 1 7 0 and l s O among the carbonaceous chondrites studied (Fig. 8). This is possibly evidence for oxidation of Fe by reactive O produced either by photodissociation or molecular reactions. But with conflicting evidence from ordinary chondrite chondrules and olivine grains (Fig. 9), no conclusions can be reached. Correlations among chondrite groups As summarized in the Introduction, nebular processes caused chemical fractionations that are observed in chondrite bulk compositions (Palme and Jones 2003). If nebular processes also produced non-mass-dependent O-isotopic variations, then these may correlate with chondrite bulk chemistry. We have just seen that there is some, but no compelling, evidence for correlated chemical-oxygen isotopic properties of some components in chondrites. Here we will examine whether correlations exist for bulk chondrites.

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The three general models that have been invoked to explain O isotopic heterogeneity in Solar System materials are summarized in Table 1. The photodissociation of CO model has several variants, with different inferred locations for the occurrence of photodissociation. Clayton (2002) has suggested that it occurred in the inner regions of the solar nebula, near the T-Tauri-stage proto-Sun. The Sun's X-wind (Shu et al. 1996) then distributed material carrying 0 isotopic anomalies further out, to where planetesimals formed. Lyons and Young (2005) suggested that photodissociation occurred near the surface of the solar nebular disk through irradiation by nearby O or B stars. Yurimoto and Kuramoto (2004), on the other hand, envisioned that O isotopic anomalies originated in interstellar molecular clouds via photodissociation, and were transported to the solar nebula in ice-coated silicate grains. The predictions given in Table 1 are for the model favored by Clayton (2002, 2005), which concludes that the Sun is enriched in le O. This latter may or may not be the case. Hashizume and Chaussidon (2005) performed O-isotopic analyses of metallic grains from the lunar regolith containing implanted solar wind. They found a component with A 17 0 of ' ' : ; »' , •150 km) asteroid ~8.3 ± 0.5 m.y. ago (Nesvorny et al. 2006). This would have been the largest recent disruption of an asteroid. There is also evidence for this disruption on Earth. A pronounced peak in the concentration of 3 He from 8 m.y.-old, slowly-accumulating pelagic terrestrial clays indicates that Earth experienced a pulse of Veritas-derived dust "immediately" following the disruption (Farley et al. 2006). Thus, immediately following the impact event that created and dispersed the Veritas family asteroids, dust from this source dominated the flux to Earth. Calculations (Dermott et al. 2002; Nesvorny et al. 2006) suggest that dust from the Veritas family still provides from 25-50% of current interplanetary dust particles (IDPs) to Earth. The Veritas family asteroids are predominantly classified as type Ch (Bus and Binzel 2002), though a few are relegated to the X type, which indicates very uncertain mineralogy (Clark et al. 2004). Ch-type asteroids have a broad 0.7 |im absorption band, interpreted as resulting from Fe in hydrous minerals. Therefore the dust originating from the Veritas family asteroids can be found among the hydrous IDPs. Approximately 50% of chondritic IDPs contain hydrous phases; these are called hydrous chondritic IDPs (Zolensky and Lindstrom 1992; Zolensky and Barrett 1994). Therefore, it would appear that the majority of the hydrous chondritic IDPs derive from the Veritas

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~i—•—I—1—I—1—I—*—|MAC88107 fayalitd O , nKaidun carbonates CV fayalites

01 e • S I

C M carbonates —A 1

•o 00

è

F i g u r e 17. Relative 5 3 Mn- 5 3 Cr and 129 I129 Xe ages of secondary minerals formed by aqueous activity on the CI, CM, CV, M A C 8 8 1 0 7 and Kaidun chondrite parent bodies. Upper panel: M n - C r ages of the secondary carbonates and fayalite in carbonaceous chondrites relative to the L E W 8 6 0 1 0 angrite (data f r o m Endress et al. 1996; Hutcheon and Phinney 1996; Hutcheon et al. 1997, 1998, 1999; Brearley and Hutcheon 2000; Krot et al. 2000a; Brearley et al. 2001; H u a et al. 2005). Lower panel: I-Xe ages of the CV chondritic components (CAIs, chondrules, matrix, dark inclusions) and mineral fractions (magnetite, phyllosilicates) relative to the Shallowater aubrite (data f r o m Swindle et al. 1983, 1998; Krot et al. 1999; Hohenberg et al. 2001; Pravdivtseva et al. 2003b,c).

Id J

C I carbonates

-14

-12

-10

-6 -4 age, Ma Kuba, pti>l,Al E39

-2

0

Bali mei mi—w_ i|ih}l ti] ¿ti——I

D a r k i n c l u s i o n s iti

the •reduced CVs LV1

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• Allende CAIs Allcruk' chottdrules Atlende matrix . I . I i L.

-6

-4

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0

2 4 age, Ma

6

8

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family asteroids. We therefore know exactly the source of this one IDP type. The hydrous chondritic IDPs consist of olivine (Fa^s), and orthopyroxene (Fs0.25) and clinopyroxene (diopside, augite) and the hydrated silicates saponite and serpentine (Fig. 18) (Zolensky and Lindstrom 1992; Zolensky and Barrett 1994). The range of olivine and low-Ca pyroxene compositions is considerably narrower, lacking the more Fe-rich compositions observed in the anhydrous chondritic IDPs. It is likely that these particular ferromagnesian minerals have been preferentially lost during aqueous alteration. These particles are typically volatile element-rich, with bulk carbon at ~3-6xCI. They contain abundant pyrrhotite, pentlandite and sulfides with intermediate compositions, the so-called "intermediate sulfides or "monosulfides." The Ni-rich sulfides apparently formed during aqueous alteration, since they are absent from the anhydrous chondritic IDPs (Zolensky and Thomas 1995). Ca-Fe-Mg-carbonates are commonly present within the hydrous chondritic IDPs, both as separate euhedral crystals and as aggregates. The alteration mineralogy of the hydrous chondritic IDPs is, in many essential ways, similar to that for the CI and CR chondrites. Unlike the C chondrites, in the IDPs, carbonates are more typically associated with Fe-Ni sulfides rather than with magnetite or tochilinite. An exception to this is the single-known tochilinite-bearing hydrous IDP that appears identical to CM2

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Figure 18. Transmission electron images of one hydrous chondritic interplanetary dust particle, showing typical mineralogy and textures, (a) Microtomed slice of the entire particle; view measures 15 |jm across, (b) Crystals of siderite (Sid, F e C 0 3 ) contained within an aggregate of ankerite (ank, F e C a ( C 0 3 ) 2 ) - most of the black grains are sulfides, (c) Typical microtexture of the particle, predominantly phyllosilicates (principally serpentine and saponite) with embedded Fe-Ni sulfides; view measures approximately 1 |jm across, (d) Grains of typical sulfides, pyrrhotite (pyrr) and Ni-rich Fe-Ni monosulfide, the latter having formed during aqueous alteration, (e) High-resolution image of a saponite flake, showing 1 nm basal lattice fringes, (f) High-resolution image of a serpentine flake, showing 0.7 nm basal lattice fringes.

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matrix material (Bradley and Brownlee 1991). Given all that we know, the hydrous chondritic IDPs have not received nearly the same attention as the anhydrous IDPs; we need more thorough mineralogical work on hydrous IDPs. McKeegan reported the O-isotopic composition of one hydrous chondritic IDP (named "Skywalker"), and found no measurable l e O excess (McKeegan et al. 1985; McKeegan 1987). It has an O-isotopic composition that plots on the terrestrial fractionation line. Clearly more than one measurement is necessary to fully characterize these IDPs. S ASTEROIDS ORDINARY AND R CHONDRITES Aqueous alteration of ordinary chondrites and its oxygen isotope record Few ordinary chondrites show any evidence of low-temperature aqueous alteration, and these are primarily the LL3 chondrites (Hutchison et al. 1987 1998; Alexander et al. 1989a,b; Bridges et al. 1997; Krot et al. 1997b). Even in these meteorites the evidence is generally sparse. The effects of aqueous alteration, resulting in the formation of phyllosilicates, magnetite, maghemite, Fe,Ni-carbides, calcite, Ni-bearing sulfides, ferrous olivine, and alkali-rich secondary phases, are best documented in chondrules and matrices of the type 3 ordinary chondrites Semarkona (LL3.0), Bishunpur (LL3.1), Krymka (LL3.1), Parnallee (LL3.4), Chainpur (LL3.4) and Tieschitz (H/L3.6) (Hutchison et al. 1987, 1998; Alexander et al. 1989a,b; Bridges et al. 1997; Krot et al. 1997b; Keller 1998; Choi et al. 1998; Grossman et al. 2000, 2002). There are only a few measurements of O-isotopic compositions of secondary minerals in ordinary chondrites, however. Choi et al. (1998) reported a large range in mass-dependent fractionation of oxygen (À ls O ~ 13%o) of magnetite grains, indicative of Rayleigh fractionation as a result of growth in the presence of a limited water reservoir (Fig. 19; Choi et al. 1998). Because of their limited indication of aqueous alteration, it is ironic that the ordinary chondrites are host to the best-studied fluid inclusions in meteorites. Two H5 chondrite regolith breccias (Monahans 1998 and Zag) contain fluid inclusion-bearing halite and sylvite (Zolensky et al. 1999, 2000; Rubin et al. 2002). Since these halides are present only within the noble

15

Figure 19. Oxygen-isotope compositions of magnetite in the unequilibrated ordinary chondrites Semarkona (LL3.0) and Ngawi (LL3.1). The large (~13%o) range in A l s O observed suggests that a limited supply of oxidant was largely consumed during magnetite formation, resulting in Rayleigh-type fractionation. This favors an asteroidal setting for the origin of magnetite, because formation of magnetite from the solar nebula (a nearly infinite H 2 0 reservoir) is expected to yield relatively constant A l s O (from Choi et al. 1998). -5

-5

0

5

5 l s O, %0

10

15

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gas-rich matrix component of these meteorites, it is likely that they originated on other bodies. The presence of a CI clast in Zag suggests one possible source (Zolensky et al. 2004); fluid inclusions have been reported within CI chondrite carbonates. Unfortunately, the nanomole quantities of fluid in individual meteoritic fluid inclusions have precluded detailed analysis.

Aqueous alteration of R-chondrites and its oxygen isotope record The asteroid type that produced the R chondrites is very uncertain. We place them with the S asteroids, but who knows? Most known R (Rumuriti-like) chondrites are finds (Rumuruti is the only exception), which experienced thermal metamorphism and brecciation to various degrees (e.g., Krot et al. 2003 and references therein). Greenwood et al. (2000) reported O-isotopic compositions of magnetite grains in an unequilibrated clast from R-chondrite PCA91002. On a three-oxygen isotope diagram, these compositions define a linear array with A 17 0 values of 34%c and A ls O ranging from -15%o to -10%o, which are different from those of magnetite grains from Ngawi and Semarkona (Fig. 19). Greenwood et al. (2000) concluded that the PCA91002 magnetites formed by oxidation of Fe,Ni-metal during aqueous alteration in an asteroidal setting, and that the original A 17 0 value of H 2 0 in the Rumuruti chondrite parent body was similar to that incorporated into the ordinary chondrites estimated by Choi et al. (1998).

M AND E ASTEROIDS - INCLUDING ENSTATITE CHONDRITES Although the visible and near-IR spectra of the M and E asteroids are nearly featureless, approximately half of the best-characterized ones are now known to have hydration features at 3 |im that are absent from the spectra of 15 others (Rivkin et al. 1995, 2000). The enstatite chondrites may derive from among the M asteroids, as do the metal cores of differentiated asteroids. The E asteroids are probably the source of the enstatite achondrites, and while we have no samples of such meteorites displaying the effects of aqueous alteration, most of the largest E asteroids show spectroscopic evidence of hydrous phases. The M class asteroids are believed to include the parent bodies of iron meteorites as well as the enstatite chondrites. Both materials have a paucity of iron in silicate minerals, and formed under reducing conditions. Since these mineralogical matches to the M asteroids are waterpoor, most workers conclude that M asteroids, as a class, are very water-poor (Keil 1989). Accordingly, the highly reduced nature of typical enstatite-rich meteorites (enstatite chondrites and enstatite achondrites) suggests that aqueous alteration was very limited on any M asteroid parent body. There are, however, spectroscopic observations of several M-class asteroids that suggest the presence of hydrated phases (Rivkin et al. 1995, 2000). Examination of the Kaidun meteorite reveals additional evidence for the presence of hydrated phases on M asteroids. The unique Kaidun meteorite completely consists of a disparate assemblage of mm-sized chips of asteroidal materials ranging from carbonaceous chondrites, to ordinary chondrites, to basaltic achondrites, to enstatite chondrites (Zolensky and Ivanov 2003). Many of these materials exhibit the complete range of alteration state, from fully anhydrous lithologies through lithologies that are completely altered on a very fine scale. There are even clasts that are half-altered and half-unaltered (Fig. 20). This observation requires aqueous alteration to have occurred at a different location than the place of final assembly of Kaidun. Indeed, much, if not most, of the materials within Kaidun must have formed on many different asteroids and possibly other bodies (e.g., comets, moons, etc.) as well. Half of the lithologies in Kaidun appear to be materials we have not yet seen as full-scale meteorites, i.e. they represent otherwise unsampled asteroid types. In most regards, the EH lithologies in Kaidun are like the typical EH meteorites (Zolensky and Ivanov 2003): Enstatite, plagioclase, silica, Fe,Ni metal, schreibersite, troilite, niningerite and unusual Fe,Cr-sulfides are present; carbon and perryite are present but less abundant;

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Figure 20. BSE images of a unique partially-aqueously altered EH lithology in Kaidun. (a) The entire clast. Above the black line the sample is pristine, below the line all phases are completely aqueously altered, (b) Close-up of an altered chondrule [see arrow in (a)]. Metal and silicates are completely altered to iron-oxyhydroxides and silica gel (grey). Troilite has been transformed to pyrrhotite (white).

oldhamite, sphalerite, djerfisherite, schollhornite, daubreelite, roedderite and Ca,Fe-phosphate are rare. The critical feature is that hydrated phases are abundant in practically all EH lithologies in Kaidun, in contrast to all other examples of EH chondrites, which are entirely anhydrous. There is evidence that E, C, and D asteroids, at the very least, provided the various materials now in Kaidun (Zolensky and Ivanov 2003). The mechanism for transport of these diverse materials to a single site on one parent asteroid must have involved numerous impacts. A record of some of these impact events remains in Kaidun in the form of melt clasts and shock melt veins, some even in carbonaceous chondrite lithologies. The different Kaidun lithologic units are micrometeorite-sized, and dynamical studies reveal that sampling of the Solar System is much more representative for such small objects than for much larger meteorites (Gounelle et al. 2005). Some EH chondrite lithologies in Kaidun show evidence of incomplete, late-stage aqueous alteration otherwise unknown from these classes of meteorites. This observation buttresses the spectroscopic data indicating the presence of water of hydration on some Easteroids (Keil 1989).

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The enstatite achondrites are believed to derive from the E asteroids. The common perception that the E class asteroids are all completely reduced and anhydrous is driven by the biased sampling we have from these meteorites. However, even the most reduced objects in the Solar System could not escape periods of aqueous alteration, due to the influx of hydrated bodies from heliocentric distances beyond the "snowline," the point where ice was stable. Thus it should be no surprise that the majority of the largest E asteroids show spectroscopic evidence of hydrous phases at their surfaces (Rivkin et al. 1995). A strong water absorption band at 3 |im on several E asteroids is an unambiguous indicator of hydrated minerals (Jones et al. 1990; Rivkin et al. 1995). A second feature, at ~0.7 |im, is seen in spectra from several E asteroids. Vilas (1994) noted a correlation for many asteroids between weak bands due to Fe 2+ and Fe 3+ transitions at 0.65-0.75 |im, which she has interpreted to indicate the presence of Fe-bearing secondary minerals. Until we have documented samples of these asteroids, however, we can only guess at what is actually present.

OXYGEN ISOTOPIC COMPOSITION OF ASTEROIDAL WATER AND EVOLUTION OF OXYGEN ISOTOPIC COMPOSITION OF THE INNER PROTOPLANETARY DISK The oxygen isotopic composition of water responsible for alteration of the chondritic meteorites described above can be inferred from isotopic compositions of secondary minerals which either formed by replacement of O-free minerals (e.g., magnetite replacing Fe,Ni metal or sulfides) or directly precipitated from an aqueous solution (e.g., fayalite, andradite, Ca-, Fe-rich pyroxenes, carbonates) if the temperature of alteration and water-mineral fractionation factors are known. Although the temperature of alteration and isotopic fractionation factors are poorly known, A 17 0 values of the secondary minerals listed above will be dependent to some degree upon the A 1 7 0 value of the fluid from which they precipitated. During progressive aqueous alteration of anhydrous silicates, which appear to have variable O-isotopic compositions, all of which are also different than that of the initial asteroid fluid (e.g., Clayton and Mayeda 1999), the A 1 7 0 value of the fluid should evolve toward those of anhydrous silicates (e.g., Benedix et al. 2003). Another way of looking at it is to say that the bulk composition of the meteorite is altered toward the oxygen isotopic composition of the fluid. Of course, the magnitude of the change in oxygen isotopic composition directly depends upon the water:rock ratio (Clayton and Mayeda 1999). Another problem with using secondary minerals produced by water-rock reactions to infer the original isotopic composition of Solar System water (fluid) is the sluggish growth rates of the secondary phases relative to the exchange rate of fluid with rock—to some degree, the secondary minerals formed from Solar System fluids that had already exchanged O isotopes with the precursor rocks, erasing some of the original isotopic differences. Therefore, the O-isotopic compositions of secondary minerals probably never record the exact initial Solar System fluid composition. It is generally believed that asteroidal water accreted together with anhydrous silicates in the form of water ice, which subsequently experienced melting due to the heat produced largely by decay of the short-lived radionuclide 26Al (e.g., Zolensky and McSween 1988; Brearley 2005 and references therein). Oxygen isotopic compositions of secondary minerals reviewed above indicate that asteroidal water had le O-poor compositions which show small variations between different chondrite groups; the inferred A 1 7 0 values range from -3%o for CV chondrites to 4-7%o for ordinary chondrites (Figs. 3, 16, 19). These compositions are very different from the inferred compositions of the Sun based on O-isotope measurements of the solar wind implanted into lunar soil metal grains, A 1 7 0 < -25%o and A 1 7 0 ~ +30%o, by Hashizume and Chaussidon (2005) and Ireland et al. (2006), respectively. They are, however, are similar to the composition of the Sun inferred by Ozima et al. (2006).

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Based on the observations that bulk O-isotopic compositions of chondrites, achondrites and terrestrial planets plot near the terrestrial fractionation line (Clayton and Mayeda 1984, 1999), and that these bodies formed from progressive random accretion of planetesimals, Ozima et al. (2006) concluded that the average A 1 7 0 value of an entire planetesimal population should represent the solar nebula value. We note that this model does not consider O-isotope exchange between nebular gas and solids during high-temperature processing that resulted in formation of chondritic components (CAIs, AOAs, chondrules, and fine-grained matrix) which accreted into chondrite and, probably, achondrite parent bodies. It also does not explain the observed mass-independent fractionation of oxygen isotopes recorded by meteorites and their components. At the same time, the value reported by Hashizume and Chaussidon (2005) is in agreement with O-isotopic compositions of the oldest solids formed in the solar nebula—CAIs and AOAs, and can be interpreted in terms of the CO-self-shielding models of Clayton (2002), Yurimoto and Kuramoto (2004) and Lyons and Young (2005). According to these models, UV irradiation of CO gas in the initially le O-rich molecular cloud (Yurimoto and Kuramoto 2004) or protoplanetary disk (Clayton 2002; Lyons and Young 2005) results in preferential photodissociation of C 1 7 0 and C l s O. The residual CO becomes 16 0-enriched. The released atomic 1 7 0 and l s O are incorporated into water that freezes out to form le O-depleted water ice. Subsequent enrichment of the inner protoplanetary disk in isotopically heavy water relative to isotopically light CO by a mechanism discussed by Yurimoto and Kuramoto (2004), Cuzzi and Zahnle (2004), and Ciesla and Cuzzi (2005) could explain evolution of O-isotopic composition of the nebular gas in the inner Solar System toward the terrestrial fractionation line. The le Odepleted water that accreted into the meteorite bodies 1-3 m.y. after formation of CAIs is generally consistent with the CO self-shielding model (Lyons and Young 2005).

SUMMARY AND FUTURE WORK Mineralogical, petrographic and isotopic observations indicate that most groups of chondritic materials experienced parent-body alteration to various degrees, resulting in formation of secondary minerals such as phyllosilicates, magnetite, carbonates, ferrous olivine (Fa40.100), salite-hedenbergite pyroxenes (Fs10_50Wo45_50), wollastonite, andradite, and nepheline. The alteration occurred in the presence of aqueous solutions under variable conditions (temperature, water:rock ratio, fQl, and fluid compositions), and in many cases was multistage. Fluids were present over a wide range of temperatures, ranging from very low temperatures comparable to terrestrial weathering on Earth, through true hydrothermal alteration at temperatures exceeding 100-150 °C, to fluid-assisted metamorphism at 300 °C and above. Although water:rock ratios, /o 2 , etc. certainly played an important role, the range of temperatures is probably the most important factor in influencing the range of mineral assemblages observed in chondrites. The Al-Mg, Mn-Cr, and I-Xe dating of secondary minerals suggests that alteration may have started within 1-2 m.y. after formation of the CV CAIs, that have an absolute Pb-Pb age of 4567.2±0.6 Ma (Amelin et al. 2002), and lasted up to 15 m.y. These observations probably indicate that chondrite parent bodies accreted various proportions of anhydrous silicates and water ice, and were subsequently heated by radioactive decay of short-lived radionuclides, resulting in melting of ice and water-rock interaction. Oxygen isotopic compositions of secondary minerals, which either replaced primary high-temperature minerals (e.g., magnetite, phyllosilicates, nepheline, sodalite, grossular, monticellite, wollastonite, forsterite, ferrous olivine) or precipitated directly from the fluid (e.g., carbonates, fayalite, Ca,Fe-rich pyroxenes, andradite, wollastonite) suggest that water ice that accreted into the chondrite parent bodies was le O-poor (A 17 0 ~ ±3%o). The inferred oxygen isotopic composition of meteoritic water is very different from that of the Sun [A 17 0 —25%o

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(Clayton 2002; Yurimoto and K u r a m o t o 2004; H a s h i z u m e and C h a u s s i d o n 2005; L y o n s and Young 2005)], suggesting extensive evolution of the O-isotopic c o m p o s i t i o n of the g a s e o u s reservoir in the inner protoplanetary disk over its lifetime. T h e r e are two c a r b o n a c e o u s chondrite g r o u p s not discussed in this chapter: the C H and the C B chondrites. T h e C H and C B chondrites show no clear evidence for in situ a q u e o u s alteration, but contain heavily aqueously-altered Cl-like clasts c o m p o s e d of phyllosilicates, f r a m b o i d a l and platelet magnetite, and c a r b o n a t e s ( G r e s h a k e et al. 2002). O x y g e n isotopic c o m p o s i t i o n s of secondary minerals in C H and C B chondrites have not b e e n reported. F u t u r e studies of o x y g e n isotopic c o m p o s i t i o n s of secondary mineralization of chondritic meteorites should b e f o c u s e d o n u n d e r s t a n d i n g their multistage alteration histories u s i n g c o m b i n a t i o n s of analytical tools, including S E M , E P M A , C L , T E M , S I M S , and I C P - M S .

ACKNOWLEDGMENTS W e t h a n k A d r i a n Brearley f o r a t h o r o u g h and h e l p f u l review of this paper. This w o r k w a s supported b y a N A S A C o s m o c h e m i s t r y P r o g r a m grant to M . Zolensky, N A S A Grants N A G 5 10610 and N A G 6 - 5 7 5 4 3 to A. K r o t and N A G 5 - 1 1 5 9 1 to K. Keil. This is H a w a i ' i Institute of G e o p h y s i c s and Planetology Publication N o . 1507 and School of O c e a n and Earth Science and Technology Publication No. 7207.

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 463-492, 2008 Copyright © Mineralogical Society of America

The Oxygen Cycle of the Terrestrial Planets: Insights into the Processing and History of Oxygen in Surface Environments James Farquhar* and David T. Johnston Department of Geology and ESSIC University of Maryland, College Park College Park, Maryland 20742, U.S.A. jfarquha @essic. umd. edu

ABSTRACT Physical and chemical processes that operate at low temperature in Earth's surface environments and in other planetary settings have left an indelible record of the evolution and interactions of planetary surface and near-surface reservoirs. This chapter reviews the chemistry of oxygen (elemental and iso topic) in lunar, planetary, andEarth surface reservoirs. The discussion begins with a brief review of the relative abundances of oxygen isotopes in planetary materials, including lunar samples, the SNC (Martian) meteorites, and components that are inferred to come from Martian surface pools. This is followed by a brief, largely historical account of the development of oxygen isotope geochemistry in Earth-surface materials. The discussion then transitions, using an Earth systems perspective, into an overview of developments in oxygen isotope geochemistry of atmospheric compounds. This treatment includes recent studies that (1) use high-precision measurements of 5 1 7 0 and 5 l s O to characterize atmospheric oxygen; (2) apply information about isotopically substituted species (isotopologs) for characterizing the position-dependent isotopic fractionations of some oxygen-bearing atmospheric species; and (3) investigate non mass-dependent isotopic fractionations in ozone and other atmospheric species. The discussion then examines the evidence and models that point to change in the oxygen cycle and oxidation chemistry in Earth surface environments over the course of geologic time. We use the subdivisions of the Earth systems - the hydrosphere, geosphere and biosphere, as a basis for examining these changes. This record is explored in the context of hypotheses and data that relate to the nature of the changes and transformations in the oxidation state of these different components of the Earth's surface environments. We conclude with a brief accounting of areas of research where scientific inquiry is presently very active.

INTRODUCTION B e f o r e c o n s i d e r a t i o n c a n b e g i v e n to t h e c o n s e q u e n c e s of o x y g e n c h e m i s t r y , it is v a l u a b l e to u n d e r s t a n d w h y t e r r e s t r i a l p l a n e t s h a v e t h e i r c o m p l e m e n t of this e l e m e n t a n d its t h r e e s t a b l e i s o t o p e s ( l e O , 1 7 0 a n d l s O ) . P r o d u c t i o n of l e O b e g a n in t h e first g e n e r a t i o n of stellar h e l i u m b u r n i n g t h a t o c c u r r e d a f t e r t h e B i g B a n g . T h e e s t a b l i s h m e n t of t h e c a r b o n - n i t r o g e n - o x y g e n ( C N O ) c y c l e of n u c l e o s y n t h e s i s in s u b s e q u e n t g e n e r a t i o n s of stars, h o w e v e r , h a s p r o v i d e d t h e p r e d o m i n a n t s o u r c e of o x y g e n ( l e O , 1 7 0 , a n d l s O ) in t h e t e r r e s t r i a l p l a n e t s . T h e m o s t a b u n d a n t i s o t o p e of o x y g e n , l e O , h a s a p a r t i c u l a r l y s t a b l e n u c l e a r c o n f i g u r a t i o n , w i t h 8 p r o t o n s , 8 n e u t r o n s , a n d filled n u c l e a r shells. O n E a r t h it m a k e s u p ~ 9 9 . 7 6 % of t h e total o x y g e n . T h e next most abundant isotope, l s O , accounts for approximately 0.2%, and has an even n u m b e r of n e u t r o n s (10), w h i l e t h e r a r e s t s t a b l e o x y g e n i s o t o p e , 1 7 0 , m a k e s u p t h e r e m a i n i n g ~ 0 . 0 4 % a n d h a s an o d d n u m b e r of n e u t r o n s (9). 1529-6466/08/0068-0016505.00

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Oxygen is one of the principal constituents of silicate minerals and, as a result, composes a significant fraction of the mantles and crusts of the terrestrial planets. It is also a stable component of many gases, liquids, and ices that may be present on the terrestrial planets, including molecular oxygen (0 2 ), carbon dioxide (C0 2 ), water (H 2 0) and ozone (0 3 ), and it is an important component of the atmospheres, cryospheres, and hydrospheres of these planets. The reactivity, or more specifically, the electronegativity of oxygen makes it a significant player in determining the oxidation state of the surface environments of the terrestrial planets. The proportion of molecular oxygen in the atmospheres of the terrestrial planets varies widely. Mercury has only a tenuous atmosphere and is not considered here, and while the atmospheres of Mars and Venus are substantial, they have only trace amounts of oxygen. On Mars, molecular oxygen is produced by photolysis of carbon dioxide, leading to an atmospheric abundance of approximately 0.13% (cf. Yung and Demore 1999; Lodders and Fegley 1998). On Venus, the oxygen in the atmosphere is consumed by reactions with reduced species, and it is present in only trace amounts (cf. Yung and Demore 1999; Lodders and Fegley 1998). On Earth, molecular oxygen gas (0 2 ) constitutes approximately 21% of dry air and is also dissolved into its oceans at concentrations that vary from a few tens to a few hundred ppm. This abundance of oxygen is a direct reflection of Earth's biosphere. A prebiotic Earth would have had oxygen levels controlled by reactions with reduced species, as they are on present-day Venus. In this chapter, we review the chemistry, evolution, and isotopic chemistry and composition of oxygen in the surface reservoirs (atmospheric, oceanic, biologic, and cryogenic). We begin with a brief historical account of the background of oxygen isotopic studies of atmospheric and oceanic oxygen with a focus on molecular oxygen, atmospheric species produced by oxygen photochemistry, and water. We then examine the evidence and models that point to change in the oxygen cycle and chemistry in Earth surface environments throughout geologic time, and we discuss the oxidation states of the atmospheres and surface environments of Mars and Venus. We use the principal subdivisions of the Earth systems- the hydrosphere, geosphere, biosphere, and atmosphere as a basis for discussion of the oxygen isotopic variations of other terrestrial planets (and Earth's Moon) because it provides a framework that can be used to compare and contrast the distinct behavior of oxygen isotopes on these different bodies. We conclude with a brief listing of features of the oxygen isotope system that remain to be documented and better understood.

ISOTOPIC VARIATIONS AMONG TERRESTRIAL MATERIALS The variations in the relative abundances of oxygen isotopes in planetary materials arise as a result of chemical, physical, and biological processes that occur in planetary environments and from variations in the compositions of the original sources of oxygen. Resolvable differences in the oxygen isotopic compositions of planetary materials have been revealed by measurements of terrestrial samples, lunar samples, and meteorites. These variations occur for 8 l s O and A17Ot, and a fully consistent reconciliation of both parameters has been devised that captures the genetic relationships between these different oxygen reservoirs (e.g., Lodders 2000), although the reader is referred to studies of Lodders and Fegley (1997) and Sanloup et al. (1999) for models that use oxygen isotopes as constraints on planetary accretion models for Mars. The Earth and Moon lie within error of a single mass-fractionation array, known as the terrestrial fractionation line (Clayton and Mayeda 1975; Wiechert et al. 2001), and are inferred to have

\0.5247

S180 = lOOOx

A O

= lOOOx

(Miller 2002)

Oxygen Cycle of Terrestrial Planets

465

sampled the same primordial oxygen reservoir. The isotopic heterogeneities for oxygen in lunar materials define a significantly smaller range than those for terrestrial materials, and this reflects the different types of physical and chemical processes that have acted on lunar oxygen compared to terrestrial oxygen. The active hydrosphere and biosphere on Earth produce a large and variable range of oxygen isotope fractionations compared to lunar materials (e.g., Onuma et al. 1970, Epstein and Taylor 1970, Clayton and Mayeda 1975; Wiechert et al. 2001). Analyses of SNC (shergottite-nakhlite-chassignite) meteorites, which are inferred to have come from Mars on the basis of their trapped atmospheric gases, young geologic ages compared to other meteorites, the unique A 1 7 0 of silicate minerals they possess, and dynamical arguments associated with the delivery of material from other evolved planets (Ashwal et al. 1982; Bogard and Johnson 1983; McSween 1984, 1994), yield information about the oxygen cycle of Mars. Significant differences exist between the oxygen isotopic composition of minerals from SNC meteorites and terrestrial minerals. Measurements of silicate minerals from SNC meteorites indicate that the silicate parts of Mars lie on a different mass-fractionation line than the one that is defined for terrestrial silicate minerals (Clayton and Mayeda 1983; Clayton and Mayeda 1996; Franchi et al. 1999). SNC silicate minerals have higher 1 7 0/ 1 6 0 and lower 1 8 0/ 1 6 0 ratios than average Earth values. Unlike low-temperature environments on Earth, analyses of water extracted from hydrous minerals (Karlsson et al. 1992; Romanek et al. 1998), oxygen in carbonate minerals (Farquhar et al. 1998; Farquhar and Thiemens 2000), and in sulfate (Farquhar and Thiemens 2000) indicate that low-temperature Martian minerals do not lie on the same mass-fractionation line as their silicate counterparts. These oxygen isotopic variations have been interpreted to reflect the lack of plate tectonics; the effects of liquid water; and a different (more active) role for atmospheric chemistry and atmosphere-surface exchange in Martian surface environments. Microscale variations in 8 l s O from carbonate minerals in the SNC meteorite ALH 84001 have been interpreted to reflect low-temperature conditions for carbonate formation. The oxygen isotopic compositions of Mercury and Venus remain to be determined. This information will provide important constraints that may help to reconcile and understand the 8 l s O and A 1 7 0 variations among the Earth, Moon, Mars, and differentiated meteorites. The most significant differences between the different planetary oxygen cycles are summarized in Figure 1. The isotopic consequences of these differences can be summarized as follows. First, materials from the Moon exhibit the smallest range of variation in 8 l s O values, and essentially no variation of A 1 7 0 values with the possible exception of implanted oxygen from the solar wind. These variations reflect high-temperature oxygen isotope partitioning. Second, materials from Earth preserve a significantly wider range of variation of 8 l s O values, reflecting low-temperature processing of oxygen, fractionation effects associated with the hydrologic cycle, and atmospheric processing of oxygen isotopes. The atmospheric processing also imparts a significant range of variation in A 1 7 0 values of atmospheric species and a small range of values for oxygen-bearing ions that readily exchange oxygen with water (Fig. 2). The largest water reservoir on Earth, the ocean, has an isotopic composition that is determined on long time-scales by exchange with the crust as a result of hydrothermal circulation at divergent plate boundaries (a link to tectonic processes), with a superimposed, short time-scale response resulting from evaporation and precipitation, and most notably to changes in ice volume. Third, materials from Mars also preserve a large range of oxygen isotopic variations (Fig. 3) because of low-temperature fractionation processes, but the absence of active tectonic processes and an active hydrologic cycle leads to the efficient transfer of nonzero A 1 7 0 values to products of weathering such as carbonate minerals and secondary iddingsite. Historical account of oxygen isotopic variations of terrestrial reservoirs Due to its abundance, low mass, and large isotopic fractionations in the geologic and meteoritic records, oxygen has inspired great interest, extending back to the 1930's and 1940's.

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aj fci u u"OÛ 1.2' 5-1 d 8. §>:

•s 'E _"5 aHI 1-8 f 1=1 ao S ii i3

ss al ss Sii S S t;o -a o fVIi 1" tu -C Si U ¡'T O t-C ¿u

H i ra t .t¡ fS i— a oi C >av O !< s«-;«

a p^ iS a m cen11 > ai F £ ai ¿ ä « ÍT= E — « > "cOS sìL — 1E aj ïr, .y113•2 TS " O S. Ç O S ai "iajS a— e E CT O"o tuÈ 511 H u o -s-5 i l l —-° ¡aS ïX2c° O l £ 1S 1&1 o tí3 5"He +16%o). Only small effects (< 2%o) can be attributed to pure fractional crystallization, as isotopic fractionations are small at high temperatures; this partly explains why felsic rocks are systematically higher in l s O by l-4%o than mafic and ultramafic rocks, but cannot explain the total ranges in terrestrial rocks. Earth's large l s O ranges require interaction or exchange of rocks and magmas with oxygen reservoirs located or formed at or near Earth's surface, where large enrichments or depletions in l s O are possible. Key processes that have been recognized include formation of minerals in equilibrium with the hydrosphere or atmosphere; subsolidus exchange of rocks with infiltrating fluids; and the incorporation into magmas of wallrocks having high or low 8 l s O values generated under low-temperature conditions. Subsolidus alteration effects are easily identified on 8-8 plots.

REFERENCES Bowen NL (1928) The Evolution of the Igneous Rocks. Princeton University Press, Princeton Chacko T, Cole DR, Horita J (2001) Equilibrium oxygen, hydrogen and carbon isotope fractionation factors applicable to geologic systems. Rev Mineral Geochem 43:1-81 Criss RE (1999) Principles of Stable Isotope Distribution. Oxford University Press, Oxford Criss RE, Farquhar J (2008) Abundance, notation, and fractionation of light stable isotopes. Rev Mineral Geochem 68:15-30 Criss RE, Taylor HP Jr (1986) Meteoric-hydrothermal systems. Rev Mineral 16:373-424 Eiler JM (2001) Oxygen isotope variations of basaltic lavas and upper mantle rocks. Rev Mineral Geochem 43:319-364

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Criss

Foulger GR, Natland JH, Presnall DC, Anderson DL (eds) (2005) Plates, plumes, and paradigms. Geol Soc Am, Special Paper 388. Geological Society of America, Denver, Colorado Gregory RT, Criss RE (1986) Isotopic exchange in open and closed systems. Rev Mineral 16:91-127 Gregory RT, Criss RE, Taylor HP Jr (1989) Oxygen isotope exchange kinetics of mineral pairs in closed and open systems: Applications to problems of hydrothermal alteration of igneous rocks and Precambrian iron formations: Chem Geol 75:1-42 Gregory RT, Taylor HP Jr (1981) An oxygen isotope profile in a section of Cretaceous oceanic crust, Samail ophiolite, Oman: Evidence for 8 O-buffering of the oceans by deep (>5 km) seawater-hydrothermal circulation at mid-ocean ridges. J Geophys Res 86:2737-2755 Henderson P (1982) Inorganic Geochemistry. Pergamon Press, New York Hoefs J (2004) Stable Isotope Geochemistry. Springer, New York Johnsen SJ, Dansgaard W, Clausen HB, Langway CC (1972) Oxygen isotope profiles through the Antarctic and Greenland ice sheets. Nature 235:429-434 Keith ML, Weber JN (1964) Carbon and oxygen isotopic composition of selected limestones and fossils. Geochim Cosmochim Acta 28:1787-1816 Knauth LP, Lowe DR (1978) Oxygen isotope geochemistry of cherts from the Onverwacht group (3.4 billion years), Transvaal, South Africa, with implications for secular variations in the isotopic composition of cherts. Earth Planet Sei Lett 41:209-222 KyserTK (1986) Stable isotopic variations in the mantle. Rev Mineral 16:141-164 Kyser TK (ed) (1987) Stable Isotope Geochemistry of Low Temperature Fluids. Mineralogical Society of Canada, Quebec Larson PB, Taylor HP Jr (1987) Solfataric alteration in the San Juan Mountains, Colorado: Isotopic variations in a boiling hydrothermal environment. Econ Geol 82:1019-1036 Lodders K, Fegley BJ (1998) The Planetary Scientist's Companion. Oxford University Press, Oxford MayedaTK, Shearer J, Clayton RN (1975) Oxygen isotope fractionation in Apollo 17 rocks. Proc Lunar Sei Conf 6th:1799-1802 Muehlenbachs K, Clayton RN (1976) Oxygen isotope composition of the oceanic crust and its bearing on seawater. J Geophys Res 81:4365-4369 Savin SM, Yeh HW (1981) Stable isotopes in ocean sediments. In: The Sea. Emiliani C (ed) John Wiley & Sons, New York 7:1521-1554 Sharp Z (2007) Principles of Stable Isotope Geochemistry. Pearson Prentice Hall, Upper Saddle River, New Jersey Singleton MJ, Criss RE (2004) Symmetry of flow in the Comstock Lode hydrothermal system: Evidence for longitudinal convective rolls in geologic systems. J Geophys Res 109:B03205 Spicuzza MJ, Day JMD, Taylor LA, Valley JW (2007) Oxygen isotope constraints on the origin and differentiation of the Moon. Earth Planet Sei Lett 253:254-265 Taylor HP Jr (1980) The effects of assimilation of country rocks by magmas on 1 8 0/ 1 6 0 and 87Sr/86Sr systematics in igneous rocks. Earth Planet Sei Lett 47:243-254 Taylor HP Jr (1986) Igneous rocks: II. Isotopic case studies of circumpacific magmatism. Rev Mineral 16:273317 Taylor HP Jr, Sheppard SMF (1986) Igneous rocks: I. Processes of isotopic fractionation and isotope systematics. Rev Mineral 16:227-271 Urey HC (1947) The thermodynamic properties of isotopic substances. J Chem Soc (London) 562-581 Valley JW (1986) Stable isotope geochemistry of metamorphic rocks. Rev Mineral 16:445-489 Valley JW, Cole DR (eds) (2001) Stable Isotope Geochemistry. Reviews in Mineralogy and Geochemistry, Vol. 43. Mineralogical Society of America, Chantilly, Virginia Valley JW, Lackey JS, Cavosie AJ, Clechenko CC, Spicuzza MJ, Basei MAS, Bindeman IN, Ferreira VP, Sial AN, King EM, Peck WH, Sinha AK, Wei CS (2005) 4.4 billion years of crustal maturation: Oxygen isotopes in magmatic zircon. Contrib Mineral Petrol 150:561-580 Valley JW, Taylor HP, O'Neil JR (eds) (1986) Stable Isotopes in High Temperature Geological Processes. Reviews in Mineralogy, Vol. 16. Mineralogical Society of America, Chantilly, Virginia Wiechert U, Halliday AN, Lee DC, Snyder GA, Taylor LA, Rumble D (2001) Oxygen isotopes and the Moonforming giant impact. Science 294:345-34

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 527-553, 2008 Copyright © Mineralogical Society of America

Basalts as Probes of Planetary Interior Redox State Christopher D. K. Herd Department of Earth and Atmospheric Sciences 1-26 Earth Sciences Building University of Alberta Edmonton, Alberta, T6G 2E3, Canada [email protected]

ABSTRACT Whether the redox state, quantified as oxygen fugacity, recorded in a planetary basalt is an accurate representation of the redox state of the planetary interior from which it was derived through partial melting, ascent, eruption and emplacement is a fundamental question in planetary geology. In the absence of mantle xenoliths in samples from the Moon, Mars and differentiated asteroids, the basalt-mantle source relationship must be extrapolated from what is known about the Earth in order to probe the redox state of these planetary interiors. A review of current knowledge regarding the basalt-mantle source relationship for the Earth provides insights into the advantages and pitfalls of determining mantle redox state. The range of currently available oxybarometers, including thermodynamic models based on ferrous-ferric mineral equilibria and multivalent cation analysis are surveyed and their limitations presented. The result is a basis for the informed interpretation of the oxygen fugacity of planetary basalts, and new insights into the role of C-H-O volatiles in the terrestrial planets.

INTRODUCTION For the purpose of elucidating the redox evolution of terrestrial planet interiors, the Earth represents a natural, if not ideal, laboratory in which to examine the relationship between the oxygen fugacity (fo2) of a basaltic sample and the redox state of its mantle source. Partial melting of terrestrial planet interiors to produce basaltic eruptives is a fundamental process that was initiated on the terrestrial planets shortly after their formation and has continued in some cases to the present day. The redox characteristics of a basalt are the direct results of the physical and chemical conditions of partial melting, ascent and eruption. The oxygen fugacity of a basalt from the Earth can be determined by the judicious selection of one or more pertinent oxybarometers based on ferrous-ferric mineral equilibria or multivalent trace element characteristics; comparison of its oxygen fugacity with that of a related mantle xenolith, complemented perhaps by insights from laboratory experiments, places constraints on the behavior of multivalent elements (and redox-sensitive volatiles) during partial melting. In this way, the influence of the physical and chemical conditions of partial melting of the mantle source on the oxygen fugacity of the basaltic product can be quantified, or at least characterized. The same cannot be done for samples of the Moon, Mars and differentiated asteroids (e.g., 4 Vesta), due to the apparent absence of mantle xenoliths from our samples of these suites. Instead, the basalt-mantle source relationship, as quantified from studies of terrestrial samples, must be adapted for and extrapolated to the suite of basaltic samples from these other planetary bodies. For this reason, summaries of the oxidation state of the Earth's mantle, the array of available oxybarometers applicable to basaltic samples, and a discussion of whether the 1529-6466/08/0068-0019$05.00

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oxygen fugacity of a basalt is reflective of the oxidation state of its mantle source are important precursors to the presentation of data for planetary basalts, and are essential to their proper interpretation. Oxygen fugacity is not a straightforward concept. A summary of the origin of the concept and a discussion of common misconceptions is provided by Frost (1991). The misconception most relevant to the following discussion is that oxygen is present as a fluid species in natural systems. Although we express oxygen fugacity in units of gas pressure (for example, 10~9,3 bars of 0 2 ), oxygen fugacity does not represent the partial pressure of a gas, but instead monitors the chemical potential, and can therefore be used to describe a condensed system in which no free oxygen is present (Frost 1991). Therefore, the " 0 2 " term in many of the equations presented in this chapter does not necessarily indicate involvement of free oxygen. Oxygen fugacity is most often expressed relative to an assemblage of pure phases that define a specific oxygen fugacity for a given temperature. These assemblages are known as solid oxygen buffers or petrologic buffers; they are not necessarily present in natural systems, but represent convenient references to describe the oxygen fugacity of a natural system. Common examples are iron-wiistite (IW), fayalite-magnetite-quartz (FMQ), nickel-nickel oxide (NNO), and magnetite-hematite (MH). A number of studies have been carried out over the past several decades to define the buffer equations, resulting in a number of formulations. A selection is given in Table 1, and other compilations can be found in Frost (1991) and Chou (1987). The oxygen fugacity defined by the buffers increases with increasing temperature, and the slopes of the buffers are nearly the same in all cases. An oxygen fugacity estimate calculated using an oxybarometer (see below) can be expressed relative to one of the buffers, obviating the need to state both the absolute oxygen fugacity and temperature. For example, an oxygen fugacity estimate of log/o 2 = - 9 . 3 at a temperature of 1200 °C and a pressure of 1 bar corresponds to 1 log unit below FMQ, or "FMQ - 1". The most recent data for solid oxygen buffers come from O'Neill and Pownceby (1993); regression of their data for the IW and NNO buffers is provided in Table 1. Also provided are regressions of data from O'Neill (1987a) for NNO, O'Neill (1988) for IW, and O'Neill (1987b) for FMQ. The latter equation should be used instead of the erroneous equation of Holloway et al. (1992) for FMQ based on the same data. Many of the formulations do not explicitly account for pressure effects, and often the buffers are not experimentally calibrated for high pressures; additional buffer equations that include pressure terms are provided by Ballhaus et al. (1991). The use of different buffer formulations can result in small discrepancies; for example, the use of the NNO formulation according to Schwab and Kiistner (1981) for log/o 2 = - 9 . 3 and T = 1200 °C yields NNO - 1.8, whereas use of the regression of the O'Neill and Pownceby (1993) data (Table 1) yields NNO - 1.6. Although the discrepancies tend to be small, typically less than the uncertainty on an oxybarometer estimate, an explicit statement of the formulation used to calculate and express oxygen fugacity provides the reader with the ability to compare oxygen fugacity estimates without introducing added uncertainties.

THE OXIDATION STATE OF THE EARTH'S MANTLE The differences between oxygen fugacity, oxidation state and oxygen content, as outlined above, are fundamentally important for understanding the Earth's mantle. Iron is the most abundant element that exists in more than one oxidation state in planetary interiors; for this reason the oxidation state of the Earth's mantle is expressed in terms of iron oxidation state, i.e., the relative proportions of the different valence states of iron (Fe°, Fe 2+ and Fe 3+ ; commonly given as Fe 3+ /EFe, where EFe = Fe 2+ + Fe 3+ ). Oxygen fugacity is related to iron oxidation state through equilibria between Fe-bearing minerals. The valence state of iron in minerals in a natural system, such as the Earth's mantle, however, is the result of the complex interplay of mineral crystal chemistry and the oxidation state of the system. The oxidation state of iron may thus change as a result of crystal chemical effects, without a correlative change in oxygen fugacity.

Basalts as Probes of Planetary Interior Redox State Table 1. E q u a t i o n s * Buffer IW

A -27589.7

B

529

for solid o x y g e n buffers.

C

6.790

Trange (°C) 769-1371

Fe - F e ^ O

Data Source(s) Regression o f data from O'Neill and Pownceby (1993)

-27654.0

6.849

769-1177

Regression o f data from

-26834.7

6.471

800-1260

Myers and Eugster ( 1 9 8 3 )

-27215

6.57

1050-1400

Eugster and Wones ( 1 9 6 2 )

O'Neill ( 1 9 8 8 )

0.056

using the data o f Darken and Gurry ( 1 9 4 5 ) FMQ

-24935.0

8.489

627-1147

Regression o f data from

-25096.3

8.735

573-1200

Frost ( 1 9 9 1 )

-24441.9

8.290

-25035

8.74

>914

-25738

9.00

600-800

-24525.4

8.944

Fe2Si04 - Fe304 - Si02

NNO

O'Neill ( 1 9 8 7 b ) 0.110

600-1140

427-1065

Ni - N i O

Myers and Eugster ( 1 9 8 3 ) Schwab and Kustner ( 1 9 8 1 ) Wones and Gilbert ( 1 9 6 9 ) Regression o f data from O'Neill and Pownceby (1993)

-24569.5

8.960

527-1147

Regression o f data from O'Neill ( 1 9 8 7 a )

-25025

9.46

-24930

9.36

0.046

771-1204

Schwab and Kustner ( 1 9 8 1 )

519-1319

Huebner and Sato ( 1 9 7 0 ) with " C " term from Chou (1987)

MH

-25700.6

14.55

-23847.6

13.48

-25632

14.62

-24912

14.41

0.019

682-1100

Frost ( 1 9 9 1 )

1040-1270

Myers and Eugster ( 1 9 8 3 )

Fe304 - Fe203

0.019

25-1227

Haas and Robie ( 1 9 7 3 )

800-1500

Eugster and Wones ( 1 9 6 2 ) using the data o f Norton (1955)

»where l o g / o 2 = A / r + B + C(P-1)/T, T is i n K a n d P i s in bars Note: T range is the range of temperatures used in data calibration

The relationship of oxygen fugacity to iron oxidation state through mineral equilibria can only be determined under the assumption of equilibrium. The quantification of this relationship is dependent on our knowledge of the stability of mineral assemblages at pressures and temperatures appropriate for the Earth's mantle. This knowledge is derived from studies of mantle xenoliths and complementary experimental data, as well as high-pressure experiments for deeper parts of the mantle, from which no direct samples exist. A brief overview of the oxidation state and oxygen fugacity of the lower and upper parts of the Earth's mantle is provided here; for a more complete review, see McCammon (2005).

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The lower mantle High-pressure experiments indicate that the lower mantle is dominatedby (Mg,Fe)(Si, A1)0 3 perovskite, with minor ferropericlase (Mg,Fe)0 and CaSi0 3 perovskite (e.g., Kesson et al. 1998). Recent experimental work on (Mg,Fe 2 + )Si0 3 perovskite has demonstrated that the substitution of Al in the structure has significant effects on the iron oxidation state in perovskite (McCammon 1997). The coupled substitution Mg 2 + + Si 4 + = Fe 3 + + Al 3+ is energetically favorable, and causes an increase in the Fe 3 + content of perovskite while maintaining charge balance in the structure. The Al content is correlated with Fe 3+ /EFe, even at low oxygen fugacity (Lauterbach et al. 2000). Thus, the Al-Fe 3+ substitution in Mg perovskite is an example of a crystal-chemical control on the iron oxidation state recorded by a mineral, independent of oxygen fugacity. The implication of these observations is that the lower mantle has a higher Fe 3 + content than previously assumed. Frost et al. (2004) calculate that the Fe 3+ /EFe of the lower mantle is 0.60, on the basis of their experimental results in conjunction with a bulk silicate earth composition in which the lower mantle contains 70 wt% perovskite containing approximately 5 wt% A1 2 0 3 . This implies that the lower mantle is enriched in Fe 3 + relative to the upper mantle, or that the formation of Al-substituted perovskite is accompanied by a complementary reaction to maintain the same overall oxidation state as the upper mantle. The latter could involve the reduction of volatile (C-H-O) species, or the disproportionation of iron through the reaction: 3 Fe 2 + = Fe° + 2 Fe 3 +

(1)

In experiments on Al-substituted perovskite, reaction (1) is manifested in the presence of discrete blebs of iron metal (Lauterbach et al. 2000; Frost et al. 2004). The volatile budget of the mantle (Wood et al. 1996) is inadequate to account for the amount of reduction required (Frost et al. 2004). Evidence for whole-mantle convection (e.g., van der Hilst et al. 1997) precludes a scenario in which the lower mantle has a higher oxidation state than the upper mantle. The disproportionation reaction has other implications for the redox evolution of the Earth's mantle. If the experimental results are representative, they imply that the lower mantle contains metal blebs, which are dispersed throughout the silicate assemblage; the recombination of iron metal and ferric iron would occur as material moves out of the perovskite stability field, resulting in no net change in bulk oxygen content (Frost et al. 2004). In the lower mantle of the early Earth, however, the metal formed by disproportionation may have been transported to the core during core formation, resulting in a net increase in O relative to Fe in the mantle (Wood and Halliday 2005). Assuming whole-mantle convection, this process may explain why the upper mantle is apparently out of equilibrium with an iron-rich core. Furthermore, the formation of iron blebs in the lower mantle would likely have implications for the abundances of siderophile elements in the Earth's mantle (Frost et al. 2004). Although ferric iron may be the dominant form of iron in the lower mantle, its bulk abundance does not constrain the oxygen fugacity of the lower mantle. Ultimately, the problem with determining the oxygen fugacity of the lower mantle is the lack of appropriate mineral equilibria. Unlike the upper mantle, the temperature dependency of the lower mantle mineral assemblage is poorly known. At relevant pressures and temperatures, experimental studies show that Ca-perovskite can contain Fe 3 + , although natural samples are chemically very pure (Harte et al. 1999; Stachel et al. 2000). Ferropericlase is the only phase in the mantle in which Fe 3+ /EFe reflects oxygen fugacity (McCammon et al. 1998; Frost et al. 2001). McCammon et al. (2004) used the composition and Fe 3+ /EFe of ferropericlase in diamond inclusions to estimate an oxygen fugacity for the lower mantle between the IW and Re-Re0 2 (e.g., Pownceby and O'Neill 1994) buffers. However, the estimate is largely qualitative and poorly constrained. Other, indirect approaches show promise, such as geophysical measurements of the lower mantle calibrated using experimental results (e.g., Wood and Nell 1991).

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The upper mantle The upper mantle is dominated by olivine, orthopyroxene, clinopyroxene, and spinel or garnet. All of these phases are iron-bearing; mineral equilibria involving them can be used to calculate oxygen fugacity in the spinel (lower pressure) and garnet (higher pressure) fades of the upper mantle. In the spinel fades, at depths less than ~60 km (~2 GPa), the dominant mineral equilibrium is '3

(2)

which is commonly referred to as the spinel peridotite reaction. Calculation of oxygen fugacity in spinel peridotites, as outlined in a later section, shows a range of over 4 orders of magnitude, and a relationship to tectonic environment, metasomatism, and partial melting (e.g., Mattioli et al. 1989; Ballhaus et al. 1990; Bryndzia and Wood 1990; Wood et al. 1990; Ballhaus 1993; Kadik 1997; Parkinson and Arculus 1999; McCammon et al. 2001). Generally, the more reduced samples have oxygen fugacities between 2 log units below the fayalite-magnetite-quartz buffer and the buffer (FMQ - 2 to FMQ), and are derived from suboceanic abyssal peridotites and undepleted, fertile subcontinental mantle xenoliths, whereas the more oxidized samples, up to ~ FMQ + 2, are peridotites that have been influenced by the effects of subduction and metasomatism. Summaries are provided by Wood et al. (1990), Wood (1991), Ballhaus et al. (1990), Ballhaus (1993), Ionov and Wood (1992), Woodland et al. (1992) and Amundsen and Neumann (1992). The effect of pressure on equilibrium (2) is to drive the reaction to the right, favoring smaller-volume phases and resulting in lower oxygen fugacity. The volume change for the solids in equilibrium (2) is about half that of the solids in the FMQ buffer (8.6 cm 3 vs. 17.95 cm 3 ; Wood et al. 1996). Therefore, all else being equal, the oxygen fugacity will decrease by about 0.25 log units per GPa pressure increase relative to the FMQ buffer (Ballhaus 1995; Wood et al. 1996). At greater depths, garnet becomes stable and the mineral equilibrium that dominates is

(3) Calibration of this equilibrium for calculation of oxygen fugacity is provided by Gudmundsson and Wood (1995). Woodland and Koch (2003) point out that an erroneous expression for skiagite (Fe 2+ 3 Fe 3+ 2Si30 12 ) activity was given by Gudmundsson and Wood (1995), and refer to Woodland and Peltonen (1999) for the correct expression. Application to garnet peridotite xenoliths from the Kaapvaal (Southern Africa) and Slave (Canada) Cratons yields oxygen fugacities of FMQ - 3 or lower, decreasing to below FMQ - 4 at about 6 GPa (Gudmundsson and Wood 1995; Woodland and Koch 2003; McCammon and Kopylova 2004). The decrease in oxygen fugacity with depth observed in garnet peridotite xenoliths is consistent with thermodynamic arguments, specifically volume effects. The volume change of reaction (3) is greater than that for reaction (2), and increasing pressure favors the incorporation of Fe 3 + into garnet. The expected decrease in oxygen fugacity is 0.9 log units/GPa (Wood et al. 1996). The volume effects on the oxygen fugacity-depth relationships within the garnet and spinel fades are examples of crystal chemical controls on iron oxidation state. The lower oxygen fugacity of the garnet fades relative to the spinel fades is attributable to the effect of the relative modal abundances of Fe 3+ -bearing mineral phases—whereas the Fe 3 + content (and therefore the O content) of the upper mantle is quite low, Fe 3+ /XFe ~ 0.023±0.010 (O'Neill et al. 1993), the relative oxygen fugacity of the spinel fades is quite high (near FMQ) because the Fe 3 + is concentrated into spinel, the least abundant phase, and virtually excluded from

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all other phases. In fact, the oxygen fugacity of the spinel facies would be 4 log units lower if Fe 3 + were concentrated equally in all mineral phases (O'Neill et al. 1993). Assuming the same bulk chemistry for the garnet, including Fe3+/EFe (O'Neill et al. 1993), the higher modal abundance of garnet at these pressures will dilute the concentration of Fe 3 + in the garnet, lowering the activity of the skiagite component and resulting in a lower relative oxygen fugacity. Furthermore, the increase in the modal abundance of garnet with depth contributes to a further decrease in relative oxygen fugacity with increasing pressure (Wood et al. 1990; Ballhaus 1995). Deeper into the upper mantle, the modal abundances of garnet and clinopyroxene increase at the expense of orthopyroxene. Oxygen fugacity may therefore be controlled by

(4) where F e S i 0 3 is the clinoferrosilite component in clinopyroxene (McCammon 2005). The breakdown of pyroxene to majorite garnet dilutes the Fe 3 + in garnet, further lowering the relative oxygen fugacity (Wood et al. 1996). This may be offset by the difference in volume change in equilibrium (4) relative to equilibrium (3). The decrease in oxygen fugacity relative to FMQ may therefore be muted at these depths (McCammon 2005). Uncertainties in the thermodynamic properties of the components currently prevent an accurate assessment. Oxybarometry of the lower part of the upper mantle is important for mapping the overall redox stratigraphy of the mantle, however, and for addressing specific redox-dependent questions, such as where the oxygen fugacity of the bottom of the upper mantle lies relative to the Ni precipitation curve (e.g., Ballhaus 1995). Carbon is considered to be another important element in the Earth's interior that is involved in redox-dependent equilibria. Fluids in the Earth's mantle are dominated by species involving C, H and O in oxidized (e.g., C 0 2 , H 2 0 ) and reduced (e.g., CH 4 , CO) forms, in equilibrium with C (as graphite in the upper mantle). The fluid species may also be dissolved in a melt. Two such equilibria are CH 4 + 0 2 = C + 2 H z O

(5)

c + o2 = c o 2

(6)

As written, reaction (5) involves the oxidation of methane and reaction (6) the oxidation of C; reaction (6) is often referred to as the CCO buffer. Because these equilibria involve fluids and/ or melt and graphite instead of mostly solid phases, the effect of pressure is different than that for equilibria (2), (3) and (4). This is simply due to volume effects—solid phases have smaller volumes than fluid phases. As a result, equilibria (5) and (6) have opposite slopes on plots of oxygen fugacity (relative to FMQ) as a function of increasing pressure. The implication is that mantle material that is buffered by Fe-bearing equilibria will have associated fluids that gradually shift with depth from H 2 0 - C 0 2 to H 2 0 - C H 4 , without any change in the activity of the ferric iron components. Opinions diverge as to the relative importance of Fe-bearing mineral equilibria and CH-O equilibria in controlling the oxygen fugacity of the upper mantle. Wood et al. (1996) argue that volatile speciation in the upper mantle depends on the oxygen fugacity determined by Fe-bearing mineral equilibria, because the concentration of C (80 ppm) is much less than that of iron (FeO ~ 8 wt%, F e 2 0 3 ~ 0.2 wt%). Ballhaus (1995) estimates the extent of relative oxidation with depth, taking into account all the factors that would influence this property, including those that would force a decrease, such as: volume effects ( - 0 . 3 to - 0 . 4 log units/ GPa, mainly for equilibrium (2)); solid solution ( - 0 . 1 log units/GPa); and the spinel-to-garnet transition ( - 1 . 5 to - 2 log units/GPa). However, he further argues that volatile equilibria (C-HO and S) are a moderating influence on the oxygen fugacity-depth gradient, providing enough

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oxygen for one-third to one-half of the F e 2 0 3 in the upper mantle. The average gradient would therefore be about - 0 . 6 log units/GPa. Thus, this perspective has volatiles playing a more important role, relative to ferrous-ferric equilibria, in influencing the oxygen fugacity of the upper mantle.

OXYBAROMETERS APPLICABLE TO BASALTIC ROCKS Basaltic volcanism is a fundamental process on differentiated planetary bodies; samples from the Moon, Mars, and differentiated asteroids (e.g., 4 Vesta) are dominated by basaltic samples. Like other physical and chemical factors involved in its petrogenesis, the oxygen fugacity of a basalt is the result of the complex history of partial melting, extraction, ascension, eruption and emplacement. Basalts can be used as probes of planetary interiors, provided that the effects of post-extraction processes can be adequately assessed. The effort is worthwhile— ultimately, insights into fundamental differences in the origin and evolution of the terrestrial planets can be gained, as exemplified by the comparative studies of the Basaltic Volcanism Study Project (BVSP 1981). The suite of planetary samples has expanded significantly since the completion of the BVSP, resulting in further insights (e.g., Wadhwa 2008). Igneous petrologists, it often seems, would be happy if magmas never crystallized, and instead quenched to glasses that are representative of the parent magmas from which the physical and chemical factors involved in their petrogeneses can be directly determined. The relationship between redox state and the concentrations of FeO and F e 2 0 3 in a multicomponent silicate liquid, i.e., 2 F e 2 + 0 + Yi 0 2 = F e 3 + 2 0 3

(7)

has been calibrated experimentally over a wide range of oxygen fugacity and bulk composition (Sack et al. 1980; Kilinc et al. 1983; Kress and Carmichael 1988, 1991) and is given by the empirical equation ln[X Fe2 0 3 /X F e 0 ] = a ln/o 2 + b/T + c + ZXA

(8)

where a, b, c and dt are constants determined by regression of experimental data; details are provided by Carmichael and Ghiorso (1990). Equation (8) allows calculation of oxygen fugacity from determinations of FeO and F e 2 0 3 in glassy lavas. More importantly, it has been demonstrated that the change in oxygen fugacity with temperature of a silicate liquid is parallel to the change in oxygen fugacity of a solid oxygen buffer such as NNO or FMQ. Otherwise said, the relative oxygen fugacity of a silicate liquid is independent of temperature. The implication is that in the absence of crystallization, determination of the redox state of a magma is straightforward. Furthermore, the redox states of different magmas can be easily compared by calculating the oxygen fugacity relative to a solid oxygen buffer at an arbitrary temperature. This has been done for glassy basic lavas from a range of tectonic environments on the Earth, demonstrating that oxygen fugacity varies more widely than any other variable in petrology, by over 7 orders of magnitude (Carmichael and Ghiorso 1990; Carmichael 1991). With a couple of rare exceptions, glassy lavas are absent from the planetary sample suite and we must rely on oxybarometers to see through the effects of crystallization. The ultimate goal of an oxybarometer is to employ a redox-sensitive element, or suite of elements, to provide a record of oxygen fugacity at a particular stage in a sample's petrogenesis. Given that the stage of interest is typically prior to crystallization, in order to assess the redox conditions of the parent magma and infer the nature of its source, a couple of caveats are worth noting. Oxybarometers are thermodynamic models, and as such, the inherent assumptions outlined by Ghiorso (1997) regarding the application of thermodynamic models to igneous systems are relevant: that the igneous system is in equilibrium everywhere along its evolutionary path;

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and that the processes being modeled are reversible. The example given by Ghiorso (1997) is one in which a magmatic system evolves from melting in the source region (state A) to final solidification in a shallow reservoir (state B); as noted, there is an infinite number of reversible paths that will take the system from state A to state B. We could equally imagine the evolution of the redox state of a magma where state X is pooling in a shallow magma body, and state Y is 20% crystallization of an assemblage of olivine, pyroxene and chromite phenocrysts. In the application of an oxybarometer, one would do well to remember the following: "Utilizing thermodynamic models of igneous processes, one cannot ever hope to attain a unique inversion to provide the unique history of magma evolution. That history is lost in the assumption of the applicability of the method'' (Ghiorso 1997) Instead of trying to uniquely determine the history of a magma, thermodynamic models ought to be used in forward modeling of magmatic evolution in order to assist in discrimination between competing hypotheses (Ghiorso 1997). Another potential pitfall worth noting is that the oxygen fugacity recorded by an oxybarometer depends on how readily the redox-sensitive elements can be reset by subsequent processes. Using ferrous-ferric iron equilibria as an example, it does not require much oxygen to affect the ferric iron in a system with initial Fe3+/XFe = 0.10. The effects of subsolidus reequilibration are a concern for any oxybarometer, as eloquently summarized by the following poem by Cin-Ty A. Lee (loosely in the haiku style): Fugacity has no memory It has no past Only what it sees last The memory of an oxybarometer has been compared to those of elephants and goldfish by John Delano: whereas elephants remember paths for yearly migration and the locations of burial grounds, goldfish cannot remember what happened in the previous few seconds such that every trip around the fishbowl is a new one. Whether a particular oxybarometer is an "elephant" or a "goldfish" will depend equally on the geochemical behavior of the redox-sensitive element involved and on the petrogenesis of the rock to which it is applied. Textural or compositional evidence for equilibrium among mineral phases will bolster application of the oxybarometer, and results need to be assessed in the context of petrologic studies of the sample. Forward modeling of changes in oxygen fugacity with crystallization are useful in assessing, for example, the effect of crystallization on the redox state of the melt. Ghiorso (1997) models the equilibrium and fractional crystallization of primitive MORB under closed system conditions using the MELTS program (Ghiorso and Sack 1995). In contrast to experiments in which oxygen fugacity is fixed relative to a buffer, MELTS allows modeling of a system in which the total oxygen content of the system (liquid + solids) is constant. The results show the expected trend of increasing ferric iron in the melt resulting in an increase in the relative oxygen fugacity of the melt due to crystallization of Fe 2+ -bearing olivine and pyroxene; in this example, the increase is a maximum (under equilibrium crystallization) of 0.7 log units. Subsequent crystallization of (Fe 3+ -bearing) spinel results in a decrease in relative oxygen fugacity of 1 log unit; the total excursion throughout the crystallization history is a maximum of 0.8 log units (Ghiorso 1997). This is consistent with Carmichael and Ghiorso (1990), who argue that in a closed system, the iron redox state of the liquid during crystallization will be regulated such that it may resemble a buffered path; specifically, that the increase in ferric iron in the melt as a result of crystallization of an early, Fe 3+ -poor phase will stabilize a later, Fe 3+ -rich phase such as spinel, and an immiscible (Fe-S-O) sulfide liquid, both of which will counteract the increase in relative oxygen fugacity of the melt. In a crystallizing liquid under open-system conditions, any addition or subtraction of oxygen should result (initially) in a

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change in the proportion of the solids; any change in oxygen fugacity should result in a change in the iron redox ratio in the solids and liquid (Carmichael and Ghiorso 1990). These insights represent a framework in which oxybarometry results can be interpreted; for example, although Fe-Ti oxides (titanomagnetite and ilmenite) typically appear later in the crystallization of a basalt, the oxygen fugacity that they record may be reflective (i.e., within a log unit) of magmatic redox conditions, if the system was closed with respect to oxygen. One approach to determination of the redox state of a basaltic sample is to apply an oxybarometer, or multiple oxybarometers, to a range of assemblages in the rock, to determine the fo2-T path of the rock as best as possible. The advantage of this approach is twofold: changes in oxygen fugacity during crystallization can be assessed, resulting in insights into the petrogenesis of the rock (e.g., open- vs. closed-system); and the results from the highesttemperature, presumably near-liquidus assemblages can be more confidently interpreted as representing the redox conditions of the magma. Several methods (oxybarometers) currently exist to determine oxygen fugacity in basalts, based on ferrous-ferric mineral equilibria, or multivalent trace elements (e.g., Eu, V). The user faces two challenges: choosing the appropriate method; and assessing the relevance of the results within the petrologic context. To assist the reader in determining the best method for his or her particular needs and applying it in an informed manner, a description of each method is provided below, and their respective strengths, assumptions and limitations are outlined. Oxygen fugacity from mineral equilibria Fe-Ti oxide. The Fe-Ti oxide oxybarometer owes its existence to A.F. Buddington, who suggested that the T i 0 2 content of magnetite was largely a function of temperature, on the basis of over 200 analyses of magnetite compiled from a range of igneous rock types (Buddington et al. 1955; Buddington 1956), and to J. Verhoogen, who demonstrated, on theoretical grounds, that the compositions of Fe-Ti oxides are significantly affected by the partial pressure of oxygen (Verhoogen 1962). Buddington and Lindsley (1964) overcame the lack of experimental data on the compositions of oxides as a function of T and oxygen fugacity, and extrapolated their data to experimentally inaccessible but petrologically important conditions, enabling the first practical application of the oxybarometer, albeit limited to the Fe-Ti binary system (i.e., oxide pairs containing exclusively the Fe and Ti cations). More information on the development of this oxybarometer can be found in Ghiorso and Sack (1991a), Lindsley and Frost (1992), and Lattard et al. (2005). The strength of the Fe-Ti oxide oxybarometer rests in the nature of the two oxides involved; the cubic oxide in the magnetite (Fe 2 + Fe 3 + 2 0 4 ) - ulvospinel (Fe 2 + 2 Ti0 4 ) series and the rhombohedral oxide in the hematite (Fe 3 + 2 0 3 ) - ilmenite (Fe 2 + Ti0 3 ) series are each solid solutions of end-members with different oxidation states. As such, the activities of magnetite in the cubic oxide and hematite in the rhombohedral oxide are used as the oxybarometer according to the reaction 4 Fe 2 + Fe 3 + 2 0 4 + 0 2 = 6 F e 3 + 2 0 3

(9)

Note that this is equivalent to the Magnetite-Hematite (MH) solid oxygen buffer. Exchange of Fe and Ti between the oxide pairs is strongly dependent on temperature and only weakly pressure-dependent. This exchange is expressed in the following equilibrium F e 3 + 2 0 3 + Fe 2 + 2 Ti0 4 = Fe 2 + Ti0 3 + Fe 2 + Fe 3 + 2 0 4

(10)

The main weakness of this oxybarometer is that Fe and Ti exchange between pairs continues during cooling. As such, the oxides do not retain their high-temperature compositions and are prone to resetting. Furthermore, in many basaltic rocks, they are among the last phases to crystallize. Thus, the oxides may record temperatures and oxygen fugacities that differ from

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those of the original magma. Fortunately, subsolidus equilibration between the oxides behaves in a somewhat predictable manner, and may be accounted for in many cases, as outlined by Lindsley and Frost (1992). Currently there exist two commonly-used formulations of the Fe-Ti oxide oxybarometer. Ghiorso and Sack (1991a) noted that the application of Buddington and Lindsley's formulation to multicomponent oxides required algorithms to project the compositions found in nature into the relevant binary systems. Citing a need for a more robust thermodynamic treatment of the Fe-Ti oxides that includes the main substituting cations, Ghiorso and Sack (1991a) presented a thermodynamic model that accounts simultaneously for all of the complex peculiarities of the Fe-Ti oxides (i.e., phase equilibrium constraints, cation order-disorder, and mixing and end-member properties), while minimizing projection schemes. This is the strength of the formulation, because proportions of some substituting cations, such as A1 and Cr in cubic oxides, can be significant. The formulation is available as a software package from the MSA website (Supplemental Material at http://www.minsocam.org), and is simple to use, requiring standard wt %-oxide compositional data for the two oxides. A new version of the oxybarometer, which is much improved for conditions o f / b 2 = NNO to NNO + 3 and T = 700 to 900 °C is expected in 2008 (Ghiorso, pers.comm.). The formulation is based on the quinary model for cubic oxides, using solution theory adopted from Sack and Ghiorso (1991a, 1991b) for cubic oxides in the system (Mg,Fe 2+ )(Al,Cr,Fe 3+ ) 2 0 4 - (Mg,Fe 2+ ) 2 Ti0 4 . Minor elements, specifically V, Mn, Ca, Zn and Ni, are included in the formulation; however, their inclusion is based on several assumptions. For example Mn, Zn and Ni are modeled as MnAl 2 0 4 , ZnAl 2 0 4 and NiAl 2 0 4 components, respectively, and are proxied by an equivalent amount of FeAl 2 0 4 . Non-zero concentrations of V 2 0 3 , MnO, CaO and ZnO are not rigorously accounted for in the calculation of equilibration temperatures; the authors note that the concentrations of these components in their sample dataset are less than 0.5 wt%, with the exception of MnO (up to 2.8 wt%). Therefore the user should be wary of using this oxybarometer when the concentrations of these components exceed the concentrations used in the oxybarometer formulation. The energetics of rhombohedral oxides are modeled in the quaternary system Fe 2 0 3 - FeTi0 3 - MgTi0 3 - MnTi0 3 . Minor elements include Al, V, Cr, Ca, Zn and Ni. Non-zero concentrations of A1 2 0 3 , V 2 0 3 , Cr 2 0 3 , CaO and ZnO are not rigorously accounted for, but A1 2 0 3 and V 2 0 3 are proxied as hematite, and the concentrations of each of these components in the authors' dataset is less than 0.2 wt%. Once again, the utility of the oxybarometer may be limited if the concentrations of these components exceed those used in the formulation. Since the calibration of the model does not account for the energetics of magnetic ordering in either oxide, it cannot be used for assemblages that equilibrated below 600 °C. Non-stoichiometry effects, which would be most significant above 1100 °C and at oxygen fugacity near the limits of cubic oxide stability, are excluded from the model; caution should be used in applying the formulation to assemblages suspected of equilibration under such conditions. The Buddington and Lindsley work enabled the estimation of oxygen fugacity in various oxide-bearing igneous rocks. Consequently, it was recognized that the oxygen fugacity was not only reflected in the compositions of Fe-Ti oxides, but also in the compositions of coexisting ferromagnesian silicates (Frost et al. 1988) and that much of the common subsolidus reequilibration of the Fe-Ti oxides could be circumvented through the use of equilibria such as Si0 2 + 2 Fe 2 Ti0 4 = 2 FeTi0 3 + Fe 2 Si0 4 quartz ulvdspinel ilmenite fayalite

(11)

commonly referred to as QUI1F (Frost et al. 1998). Lindsley and Frost (1992) and Andersen et al. (1993) presented an updated thermodynamic model for the Mg- and Ca-bearing system, referred to as Ca-QUIIF. This formulation includes

Basalts as Probes of Planetary Interior Redox State

53 7

equilibria between augite, pigeonite, orthopyroxene, olivine and quartz in addition to the Fe-Ti oxides. The main advantage of this formulation is that it can reduce the uncertainty in using the Fe-Ti oxide oxybarometer alone. For example, the four-component subsystem Fe0-Mg0-Fe 2 0 3 -Ti0 2 , with two phases (the Fe-Ti oxides) has a formal variance of four, but because the partitioning of Fe and Ti are coupled, and the effect of Mg on temperature and oxygen fugacity is minor, two intensive parameters (T and f0l) are tightly constrained. Furthermore, the equilibria can be used to assess equilibrium among phases. For example, if the temperature calculated from Fe-Ti oxides is consistent with a temperature calculated from the same oxides with co-existing olivine and pyroxene, then it supports equilibrium amongst all of these phases. By the same token, the model can be used to "see through" subsolidus re-equilibration of oxides, including oxyexsolution of the cubic oxide (i.e., titanomagnetite). For example, if pyroxene temperatures and oxide temperatures do not agree (assuming these phases were initially in equilibrium) then the original oxygen fugacity can be estimated, given the relative abundances of the oxides and assuming a closed system during cooling. In fact, the rhombohedral oxide is expected to gain FeTi0 3 and the cubic oxide is expected to gain Fe 3 0 4 upon equilibrium cooling in a closed system (Lindsley and Frost 1992). Thus, magmatic (or at least super-solidus) oxygen fugacity estimates can be made, even when the Fe-Ti oxides have undergone re-equilibration. Ca-QUIIF is available as a software package (QUI1F95) from the MSA website (Supplemental Material at http://www. minsocam.org). Mineral compositions must be expressed in terms of mole fractions of end-members (e.g., XHematite, XFaya]ite, XWollastonite) with the exception of the cubic oxide phase, for which the user must calculate the numbers of Ti, Mg and Mn cations per formula unit (NTi, NMg, NMn; cations per four oxygen). The program can be used to calculate temperature, pressure, oxygen fugacity, equilibrium mineral compositions, and activities of Si0 2 , Fe and Ti0 2 . As in the previous case, some limitations are useful to note. The calibration of the oxide models was done below FMQ + 2, and therefore may not be applicable to highly oxidized assemblages. With regard to cubic oxide compositions, Fe, Ti and Mg are the only cations accounted for, and the assumptions of the model provide for only two independent compositional parameters, iVxi and iVMg. Therefore the model is not appropriate for cubic oxides with significant A1 2 0 3 and Cr 2 0 3 contents. The formulation assumes that silicates have negligible Fe 2 0 3 and Ti0 2 contents and that CaO is unimportant in the oxides; therefore, caution should be used where, for example, pyroxene contains significant Ti0 2 . Lastly, the model cannot be used quantitatively for MgFe 3+ 2 0 4 -rich cubic oxides because of the limits of the solution models and because such oxides will tend to be nonstoichiometric; as with the Ghiorso-Sack model, nonstoichiometry is not included in the model. The reader is referred to Ghiorso and Sack (1991a,b), for comparisons of the GhiorsoSack formulation with Ca-QUIIF. It should be noted that the Ghiorso-Sack model uses thermodynamic data that are internally consistent with olivine and orthopyroxene solution theory from Sack and Ghiorso (1989) as well as the fayalite, ferrosilite, 0 2 (gas), and quartz data after Berman (1988); these data could be assembled in a separate formulation of QUI1F. The most obvious difference between the Ca-QUIIF and Ghiorso-Sack formulations of the FeTi oxide oxybarometer is the treatment of minor substituting cations, especially in the cubic oxides. In cases where the concentrations of minor cations are low, the two models agree well (e.g., Herd et al. 2001). Citing the lack of experimental calibration of the Fe-Ti oxide oxybarometer at high temperatures and over a wide range of oxygen fugacity, Lattard et al. (2005) carried out a series of experiments in the Fe-Ti-O system at temperatures of 1000 to 1300 °C and oxygen fugacity ranging from NNO - 5 to NNO + 5 (~IW to FMQ + 6). The results should lead to an improved thermodynamic model for rhombohedral oxide, thereby reducing the discrepancies

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between experimental results and those calculated using either the Ca-QUIIF or Ghiorso-Sack models, as well as enabling more accurate calculation of oxygen fugacity, especially under oxidizing conditions (> FMQ + 2). The data of Lattard et al. (2005) also include compositions of ilmenite and pseudobrookite ((Fe 3+ ,Fe 2+ ) 2 (Ti,Fe 3+ )0 5 ), which is an additional oxybarometer applicable to some terrestrial and lunar igneous rocks. In deciding which oxybarometer to apply, the user should exercise caution in applying a particular formulation to oxide compositions that deviate significantly from the Fe-Ti endmembers upon which the thermodynamic model is based. The best application of any formulation is informed by detailed petrography and an assessment of the degree of equilibrium. Olivine-pyroxene-spinel. The olivine-pyroxene-spinel oxybarometer, also referred to as the spinel peridotite oxybarometer, was developed for application to mantle xenoliths from the spinel facies. It is governed by equilibrium (2), repeated here: 6 Fe 2 Si0 4 + 0 2 = 2 Fe 2 + Fe 3 + 2 0 4 + 6 FeSi0 3 olivine

spinel

(2)

opx

This equilibrium has also been referred to as the fayalite-ferrosilite-magnetite (FFM) buffer (e.g., King et al. 2000). Oxygen fugacity is calculated using log (/o2) = - 6 log aolFe2sio4 + 2 log aspFe3o4+ 6 log a°pxFeSio3

(2a)

involving the activities (a) of the respective iron end-members. In principle, this equilibrium is applicable to low-pressure assemblages in basaltic samples, assuming equilibrium between olivine, spinel, and low-Ca pyroxene. Spinel in this case is typically chromian spinel. Because the mineral phases in Equation (2) are near the liquidi of many basaltic rocks, the results from this oxybarometer are presumably good indicators of magmatic temperatures and oxygen fugacities. Subsolidus re-equilibration is limited to Fe-Mg exchange. The involvement of three phases requires determination of whether all three are cogenetic and remained in equilibrium. Detailed petrography can assist in overcoming this obstacle. An overview of the development of the olivine-pyroxene-spinel oxybarometer is given by Wood (1991), where he presents the equation for oxygen fugacity relative to the FMQ buffer, as originally derived by Wood (1990): log (fo2)p,T = log/o 2 (FMQ)pj- + 220/T+ 0.35 - 0 . 0 3 6 9 P I T - 12 log XolFe - (2620/7XX ol Mg ) 2 + 3 log (XM1Fe -XM2Fe)opx + 2 log

(12)

where T is the temperature in K and P is the pressure in bars. Each part of this equation can be understood in terms of Equation (2a). The activity of Fe 2 Si0 4 in olivine is modeled assuming random mixing over the two cation sites, and is represented by the term, " - 1 2 log X o l F e - (2620/ IXX^Mg)2", in which XolFe and XolMg represent the mole fractions of fayalite and forsterite endmembers, respectively, in olivine. The activity of FeSi0 3 in orthopyroxene is treated as an ideal two-site solution, represented by the term, "3 log (XM1Fe -XM2Fe)°px", in which XM1Fe and XM2Fe represent the atomic fractions of Fe on M l and M2, respectively. These are calculated as described by Wood (1990): "Al was added to Si to fix tetrahedral occupancy at 2.0 per 6 oxygens. The remaining Al (VI), Cr and Ti were placed in M l , while Ca and Mn were placed in M2. Fe and Mg were then evenly distributed between M l and M2 positions and atomic fraction of Fe on M l (XM1Fe) and M2 calculated." The last term involves the activity of the magnetite component in spinel (aspFe3o4); markedly, no expression is incorporated into the equation. The largest uncertainty in determining oxygen fugacity with this method derives from the uncertainty in aspFe3o4- For this reason there exist several equations for calculating aspFe3o4, outlined below. The remaining terms in Equation (12) reflect the FMQ buffer to which the results of the calculation are related. The term, "log/o 2 (FMQ)p r " is the oxygen fugacity of the FMQ buffer at the P and T of interest. Numerous formulations of the FMQ buffer exist (Table 1); therefore,

Basalts as Probes of Planetary Interior Redox State

539

it is especially important to note that the Myers and Eugster (1983) formulation was used in the derivation of Equation (12), which is: log (/b2) = - 2 4 4 4 1 . 9 / r + 8 . 2 9

(13)

The term, "220IT + 0.35 - 0.0369PIT' represents the pressure and temperature dependency of the difference between the FMQ buffer (Eqn. 13; Myers and Eugster 1983) and "FFM" (Eqn. 2), after Mattioli and Wood (1988; their Eqn. 32). The derivation of the pressure term in Equation (12) is not explicit in Wood (1990) or Wood (1991). Significantly, oxygen fugacity calculated using Equation (12) is de facto relative to the FMQ buffer of Myers and Eugster (1983). Absolute log ( f 0 l ) can be calculated by also calculating log f0l (FMQ) according to Equation (13). Temperature and pressure are required for the calculation. Pressures for planetary basalts are typically close to atmospheric (~1 bar). Temperature can be calculated using any of several formulations of the chromite-spinel geothermometer (e.g., Fabries 1979; Sack and Ghiorso 1991a). Wood (1991) estimates that oxygen fugacity can be calculated to within ± 0.5 log units using the olivine-pyroxene-spinel oxybarometer, assuming good analyses of all three phases are available. The largest contributor to the uncertainty is in the determination of aspFe3o4- Mattioli and Wood (1988) determined aspFe3o4 across the MgAl 2 0 4 -Fe 3 04 join, between 900 and 1000 °C at 1 atm (1 bar) pressure. They did not, however, account for the effects of the chromite (FeCr 2 0 4 ) component. Insights into order-disorder in spinel from models and electrochemical measurements by O'Neill and Wall (1987) and Nell and Wood (1991) provided updates of the model to account for Cr. The Nell-Wood equation for aspFe3o4 (see also Wood 1991) is applicable to XFe3o4 = 0.008 to 0.06 and T= 800 to 1400 °C. The quinary model for cubic oxides of Sack and Ghiorso (1991a,b) provides an alternative method of calculating aspFe3o4- This is accessible using the MELTS Supplemental Calculator (.http://melts.ofm-research.org/CalcForms/index.html), and requires that the user first calculate mole fractions of chromite, hercynite, magnetite, spinel and ulvospinel, which are then used to calculate thermodynamic properties at a chosen Tand P. The Sack and Ghiorso (1991 a,b) models are applicable to a range of spinel compositions, which is more desirable when using the olivine-pyroxene-spinel oxybarometer for the compositions typical of basaltic rocks. Ballhaus et al. (1991) provide an empirical calibration of the O'Neill and Wall (1987) olivine-pyroxene-spinel oxybarometer, using synthetic spinel harzburgite and lherzolite assemblages between 1040 and 1300 °C and 0.3 to 2.7 GPa. The advantage of the formulation is that it obviates the need for an explicit calculation of the activity of the magnetite component in spinel. However, the formulation is simplified by canceling orthopyroxene against the ideal part of the fayalite activity in olivine. This simplification cannot be expected to be valid at XFeo1 > 0.15. As such, its application is limited to Mg-rich upper mantle-derived rocks. Wood (1990) ran experiments equilibrating olivine, orthopyroxene and spinel at known oxygen fugacity (between FMQ and FMQ - 2) and temperature (1188 to 1205 °C), in order to test the Mattioli-Wood, O'Neill-Wall and Nell-Wood expressions for aspFe3o4 in the calculation of fo2 using Equation (12). He demonstrated that the Mattioli-Wood model was slightly dependent on Cr content, and that the Nell-Wood version more accurately reproduced the known fo 2 compared to the O'Neill-Wall version. For comparison, the calculation was reproduced for this work by using the MELTS calculator for magnetite activity. It more closely reproduced the known/o 2 than did the Nell-Wood version, as shown in Figure la. Whereas the Nell-Wood version underestimated the/o 2 by 0.3 log units, the present calculation underestimates fo 2 by only 0.07 log units, on average. The calculation was repeated using the Ballhaus et al. (1991) formulation (Fig. lb); this model underestimates/o 2 by an average of 0.45 log units.

540

Herd Wood and MELTS Version

1.00

o

U i -1-00

o

-2.00 0.00

i 0.10

i 0.20

i 0.30

i 0.40

i 0.50

0.60

i 0.40

i 0.50

0.60

Cr/(Cr+AI) spinel

Ballhaus Version

1.00

'S

Ö ) -1.00

O

-2.00 0.00

i 0.10

i 0.20

i 0.30

Cr/(Cr+AI) spinel Figure 1. (a) Comparison of oxygen fugacity calculated according to Equation (12), using the MELTS Supplemental Calculator for the activity of magnetite in spinel, applied to the experimental data of Wood (1990). The average of the calculations is offset f r o m the k n o w n oxygen fugacity by 0.07 log units. N o dependence on the Cr/(Cr + Al) ratio is observed, (b) Comparison of oxygen fugacity calculated according to the expression of Ballhaus et al. (1991). The average of the calculations is offset f r o m the k n o w n oxygen fugacity by 0.45 log units. N o dependence on the Cr/(Cr + Al) ratio is observed.

Therefore, the use of the MELTS Supplemental Calculator to calculate the activity of magnetite in spinel, in combination with Equation (12), is as good as other formulations for spinel peridotite compositions. The diversity of solid solutions used in the MELTS Supplemental Calculator allows for wider applicability. As an example, this method has been applied to olivine-phyric martian basalts (Herd et al. 2002; Goodrich et al. 2003; Herd 2003, 2006). In many cases, the results agree with other oxybarometers applied to the same rock (e.g., Goodrich et al. 2003; Herd 2006). The olivine-pyroxene-spinel oxybarometer has seen little use for low-pressure assemblages in terrestrial basaltic samples. Ballhaus et al. (1991) used their model to calculate fo 2 for midocean ridge basalts (MORB), island arc basalts (IAB) and ocean island basalts (OIB), but due to the inherent assumptions, the model is limited to mantle-derived primitive melts, and is not

Basalts as Probes of Planetary Interior Redox State

541

appropriate for more evolved basalts. The Wood - MELTS version of the oxybarometer is readily applicable to olivine-phyric terrestrial basalts. O'Neill and Wall (1987) used a slightly different approach to calculation of oxygen fugacity from olivine, pyroxene and spinel in mantle assemblages. Instead of "FFM", they used FMQ, i.e., (14)

(15)

(16)

(17) This method has not been tested for use with basaltic samples from planetary surfaces, although it is potentially applicable, with use of the MELTS Supplemental Calculator (or some other method) to calculate the activities of fayalite, forsterite, magnetite and enstatite. Multivalent trace elements The previous oxybarometers are based on the equilibria between iron-bearing minerals in planetary basalts. In essence, they depend on the partitioning of ferrous and ferric iron among phases. Similarly, there exist a number of other multivalent elements whose valence states can be used to determine oxygen fugacity. Those that have a range of valence states under the redox conditions of planetary basalts include some of the transition elements (Ti, V, and Cr) and Eu, a particularly useful rare earth element (REE). The main difference between these elements and iron is that they are present in trace concentrations in major minerals. Their concentrations are such that they can be analyzed by in situ microbeam methods, including Electron Microprobe (EMP) and Secondary Ion Mass Spectrometry (SIMS). Determining their valence states, however, is a challenge, and methods have been developed, or are presently in development, to overcome this hurdle. The multivalent trace elements of interest for planetary basalts include Eu 2 + , 3 + , y2+, 3+, 4+, 5+^ Q.2+, 3+^ a n ( j 4+ These are shown schematically (along with iron) in Figure 2, after Papike et al. (2005). This diagram shows the range of valence states of the multivalent trace elements, for comparison with the range off 0 l of planetary basalts. Each point represents the oxygen fugacity at which the oxidized and reduced species are present in approximately equal proportions in a basaltic melt. The diagram is a useful "roadmap" for selecting appropriate oxybarometers—it is readily seen that mineral equilibria oxybarometers involving Fe 2 + and Fe 3 + are most applicable to the range of f 0 l experienced by terrestrial and martian basalts. Of these elements, only vanadium exists in 4 valence states, covering the range of redox conditions of planetary basalts. For this reason, much recent effort has focused on the development of vanadium oxybarometers. Oxygen fugacity from multivalent trace elements Europium. Europium is the only REE whose geochemical behavior is significantly different from the rest of the REE in planetary magmas, due to its stability as Eu 2 + or Eu 3 + at fo2 < FMQ. The "Eu anomaly" that is often observed in chondrite-normalized REE patterns in minerals and bulk rocks is due to this effect, coupled with crystal chemistry. The valence state

542

Herd | Earth Moon | Wars IW-itolW IWtolWfi IWt-i to IWt-6 •

•

V4+ O Fe2+ V3+ Cr2+

Cr3+

Fe

Fe2+

Eu2+

tu3+

V5+ Fe3+

V3+

V2+ < = i O T'@+ C

o

i=> Ti4+ .

.

-G

.

J

-4 -2 0 2 4 L o g f O2 r e l a t i v e t o IW

10

Figure 2. Schematic diagram showing the dominant valence states of multivalent elements in planetary basalts, and their relationship to oxygen fugacity relative to the IW buffer, after Papike et al. (2005). See text for discussion.

of Eu in a silicate melt is related to oxygen fugacity by EuO (melt) + %0 2 (g) = EuO L5 (melt)

(18)

which is analogous to the relationship between ferrous and ferric iron and / 0 , . However, europium's status as a trace element requires a different approach to for determination of the relationship of Eu behavior to oxygen fugacity, which involves the partitioning of Eu between minerals or between mineral and melt. Equilibrium (18) can be rearranged to solve for oxygen fugacity as log/o, = - 4 log

[fl E u0(melt/flEu01.5(melt)]

- 4 log K

(19)

which demonstrates that at constant T and P, the ratio of the activities of Eu 2+ and Eu 3+ is a function of oxygen fugacity. As soon as minerals become involved, crystal chemical effects, especially crystal/liquid partitioning, must be taken into account. Thus, it is expected that the Eu 3+ /Eu 2+ ratio in a given mineral will be a function of the composition of the melt, the crystal chemistry of the mineral, and the oxygen fugacity at the time of crystallization. Recognizing the potential of Eu as an oxybarometer, J. A. Philpotts developed a method for calculating Eu 2+ and Eu 3+ concentrations in igneous phases (Philpotts 1970). His formulation addresses the fact that Eu 2+ and Eu 3+ cannot be directly measured. Instead, it uses partition coefficients, i.e., D E u X + p/ a = E u x + p / E u x + a

(20)

x+

where Eu j, is the concentration of Eu of some valence X in phase y. In each phase, the concentration of Eu can be expressed as Eu a = Eu 2+ a + Eu 3+ a

(21a)

2+

(21b)

E u p = Eu

p + Eu

3+

P

Basalts as Probes of Planetary Interior Redox State

543

Assuming two phases in equilibrium, for example, where a is the matrix (as a proxy for the melt) and (3 is a plagioclase phenocryst, Equations (20), (21a) and (21b) can be combined to solve for the concentration of one of the valence states of Eu in the matrix (a), i.e., Eu 3 + a = [Eu p -

DEu2+p/a- Eua]/[DEll3+p/a

-

DEu2+p/a]

(22a)

with the concentration of Eu 2 + in the matrix determined by Eu 2 + a = Eu a - E u 3 + a

(22b)

The same could be repeated if other phenocrysts are in equilibrium with the matrix, to check for internal consistency (Philpotts 1970). Note that, in order to apply Equation (22a), the concentrations of Eu in the matrix and phenocryst (Eu a and Eu p , respectively), and the partition coefficients for Eu 2 + and Eu 3 + for the phenocryst mineral must be known. Philpotts (1970) used analyses determined by mineral separation and stable isotope dilution mass spectrometry for the former (e.g., Schnetzler and Philpotts 1970); today, in situ microbeam techniques (e.g., SIMS, LA-ICPMS) would be used. The partition coefficients were estimated by Philpotts (1970), with D S r 2 + p / a used as a proxy for D E u 2 + p / a , and D E u 3 + p / a interpolated from a plot of D R E E 3 + p/a for the same phases. Since that time, experimental studies have addressed the partitioning of Eu in different phases under different redox conditions (e.g., Grutzeck et al. 1974; Sun et al. 1974; Drake 1975; Weill and McKay 1975; McCanta et al. 2004), and parameterized it using D Eu /D Gd or D E u /D S m , as described below. Philpotts (1970) applied Equations (22a) and (22b) to a suite of terrestrial and lunar samples, and to the eucrite meteorites Moore County and Juvinas (thought to be from the differentiated asteroid 4 Vesta; Drake 1979). He employed the Eu 2 + /Eu 3 + ratios of the sample matrices to calculate oxygen fugacity according to Equation (19). In spite of certain variations in temperature, pressure and bulk composition among the samples, he determined that lunar basalt crystallized under redox conditions 4 to 5 log units below that of terrestrial basalts, and that the Juvinas eucrite oxygen fugacity is an additional 2 or 3 log units more reduced than lunar basalts. Drake (1975) applied his experimental Eu partitioning data for plagioclase to the Philpotts (1970) Eu 2 + /Eu 3 + ratios and obtained f 0 l results consistent with those of Philpotts (1970). These results are broadly consistent with observations of lunar basalts, which are essentially at metal saturation; however, similarities in mineral compositions and experimental petrology results for lunar and eucrite basalts suggest virtually identical oxygen fugacities (e.g., Kesson and Lindsley 1976; Stolper 1977; Longhi 1992). Regardless, subsequent studies of planetary basalts (e.g., B V S P 1981; Wadhwa 2008) have corroborated this range of oxygen fugacity among the terrestrial planets. The approach of McKay (1989) and McKay et al. (1994) is worth highlighting, since it uses the Eu oxybarometer to constrain the oxygen fugacity of a planetary basalt, specifically the L E W 86010 angrite meteorite, to within one log unit. Using appropriate experimental partitioning data for Eu among A1-, Ti-rich (fassaitic) pyroxene, anorthite and melt at 1175 to 1210 °C and atmospheric pressure, McKay (1989) determined the relationship between D Eu /D Gd and oxygen fugacity for plagioclase and pyroxene. At high oxygen fugacity (~ FMQ), Eu 3 + is the dominant species and, having a smaller ionic radius (^Eu 3 " 1 " radius = 1.09 A; Shannon 1976) than the divalent cation (^Eu 2 " 1 " radius = 1.31 A; Shannon, 1976), it is more readily incorporated into the pyroxene M sites. Presumably the substitution of Eu 3 + occurs via a coupled substitution with Na + in the M2 site or Al 3 + in the tetrahedral (T) site. At low/o 2 (~ IW), Eu 2 + is favored in the large feldspar site, as a substitution for Ca 2 + or Na + ( [ v n i l Eu 2 + radius = 1.39 A; [ v l n i Eu 3 + radius = 1.21 A; Shannon 1976). The predicted relationship of D Eu as a function o f f 0 l is an S-shaped curve that asymptotically approaches the values of D E u 2 + at low fo2 and D E u 3 + at high f0r The observations follow the predicted relationship between D Eu and f 0 l on the basis of theory, as outlined by McKay et al. (1994). The advantage of the application of Eu in pyroxene and plagioclase for planetary basalts

544

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is that the greatest variation in DEu in each of these phases occurs over the range of f 0 l between IW-2 and FMQ, which covers the majority of the range of oxygen fugacity of planetary basalts. The contrasting behavior of DEupl and DEupx, with the former increasing and the latter decreasing with/o 2 , provides an additional advantage: assuming that the concentrations of Eu and Gd in pyroxene and plagioclase are known, the two phases are in equilibrium, and the experimental phase compositions and temperatures are similar to those of the rock, then the two curves can be combined to give a calibration curve for the Eu oxybarometer. This is done by determining the ratio of (DEupl/DGdpl)/(DEupx/DGdpx); because these are mineral-melt partition coefficients, the melt concentrations cancel out, and only the Eu and Gd concentrations in the two phases are required. Thus, a plot of log (Eu/Gd)pl/(Eu/Gd)px vs. log/o 2 will yield a linear relationship from which oxygen fugacity can be calculated (McKay et al. 1994). In the application of McKay et al. (1994), the Eu oxybarometer was used to determine that the LEW 86010 meteorite crystallized at an oxygen fugacity between IW and IW + 1. The primary advantage of the Eu oxybarometer is that the REE are relatively immobile elements (e.g., Van Orman et al. 2001) that are taken up by igneous phases and are present in concentrations that are measurable by available in situ methods. Thus, the magmatic Eu/ Gd (or Eu/Sm) ratio is likely to be preserved through subsolidus re-equilibration; assuming an appropriate calibration, the magmatic oxygen fugacity can be determined. As with other oxybarometers, the user should be aware of potential pitfalls. Ideally, the Eu oxybarometer is applied to samples in which pyroxene and plagioclase are in equilibrium on the liquidus. In reality, this does not often occur. The crystal chemistry of the REE in pyroxene and plagioclase is such that compositional effects on Eu partitioning may be significant. For example, the Si0 2 /Al 2 0 3 activity ratio in the melt may influence Eu partitioning in plagioclase (Morris and Haskin 1974; Drake 1975). Likewise, the uptake of Eu 3+ into pyroxene M sites requires a coupled substitution, and so is presumably influenced by the content of univalent cations (especially Na + ); hence the need for calibration curves using appropriate melt compositions. Recently, a variation on the Eu oxybarometer has been developed for martian meteorites involving Eu in pyroxene (Wadhwa 2001; Musselwhite and Jones 2003; McCanta et al. 2004). Pyroxene, especially low-Ca pyroxene, is a liquidus phase in these basalts, with plagioclase crystallizing later. This work has provided an important means of estimating the magmatic oxygen fugacity in martian basalts (Wadhwa 2008). The fractionation of Eu into pyroxene becomes small at higher oxygen fugacity, however, where Eu 3+ becomes dominant, resulting in greater uncertainty, up to ± 1 log unit at FMQ (McCanta et al. 2004). Therefore, this version of the oxybarometer has the greatest resolution at lower oxygen fugacity (< IW + 1). Vanadium. As shown in Figure 2, vanadium exists in four valence states. In theory, if the relative proportions of V species can be calibrated for oxygen fugacity, then it would provide an oxybarometer appropriate for all planetary basalts. This step has been accomplished for V in glass by Sutton et al. (2005) using vanadium K-edge X-ray Absorption Near-Edge Structure (XANES) spectroscopy, done with synchrotron light at the Advanced Photon Source (APS), Argonne National Laboratory in Argonne, IL. Vanadium K edge XANES spectra have a pronounced pre-edge feature whose intensity and energy increase systematically with valence state. Sutton et al. (2005) used a suite of synthetic basalt (forsterite-anorthite-silica and forsterite-anorthite-diopside) glasses in which vanadium valence state had been determined by titration. They then determined the relationship between the pre-edge peak intensity, I, and the effective vanadium valence, V*, I = - 153 + 199(V*) - 106(V*)2 + 22.4(V*)|3:

(23)

V * = 3f(V 3+ ) + 4f(V 4+ ) + 5f(V 5+ )

(24)

where V* is given by and f is the fractional content of each species. It is assumed that V

2+

has zero pre-edge peak

Basalts as Probes of Planetary Interior Redox State

545

intensity. Equation (23) was used to determine V* for 5 other suites of synthetic glasses; observations were consistent with predictions of the dominant species at different oxygen fugacity. The effects of temperature and melt structure are accounted for, but introduce little additional uncertainty (± 0.2 and 0.5 log units, respectively). A calibration curve of vanadium pre-edge peak intensity vs. log/o 2 was constructed from the data from all synthetic glasses. Due to uncertainties at very low f0l, and the dominance of V 5+ at the highest f0l (and a concomitant reduction in resolution) the oxybarometer can be effectively used for glasses between IW - 2 to IW + 6 (corrected to a temperature of 1400 °C); a range that is unparalleled in oxybarometry. Application of the V XANES oxybarometer to natural glasses from the Earth, Moon and Mars yields results that are broadly consistent with previous studies, over a range from IW - 2 (lunar) to IW + 4 (terrestrial) with an uncertainty of ± 0.2 log units (Karner et al. 2006). The method is non-destructive and can be used on traditional polished thin sections at micron-scale resolution; furthermore, it is sensitive to V concentrations at the ~100 ppm level (Sutton et al. 2005). The main disadvantage is the lack of preserved glasses in natural basaltic samples. For this reason, the application of K edge XANES spectrometry to V in minerals is under development. This application faces the dual challenges of orientation effects and crystal chemical controls on V species partitioning. Papike et al. (2005) developed a semi-quantitative oxybarometer for V in chromite. The basis of this oxybarometer is the crystal chemistry of spinel, which has an affinity for V 3+ over V 4+ . Canil (2002) recognized the dependence of D v s p on oxygen fugacity, noting that D v s p in high Cr/Al spinels (chromite) decreases by about an order of magnitude (from 32 to 5) between IW - 1 and IW + 4. Papike et al. (2005) note that V behaves differently in spinel from different planetary basalts, as evidenced by core-to-rim traverses across grains using the EMP. As illustrated in Figure 2, V 4+ dominates under terrestrial redox conditions, V 3+ dominates under lunar redox conditions, and martian basalts have subequal proportions of both. This is reflected in the core-to-rim patterns in spinel, which show compositional zonation from chromite cores to ulvospinel rims: in lunar basalts, V as predominantly V 3+ follows Cr (as Cr 3+ ), decreasing from core to rim; in terrestrial basalts, V as predominantly V 4+ follows Ti (as Ti4+), increasing from core to rim; in martian basalts, the trends vary, consistent with differences in oxygen fugacity between different samples (Papike et al. 2005). A comparison of the V contents of chromite cores among planetary basalts also reflects differences in oxygen fugacity; a plot of 100V/(Cr+Al) atomic with distance in chromite cores from lunar, martian and terrestrial basalts qualitatively differentiates the ranges of oxygen fugacity for the Moon, Mars and Earth (Papike et al. 2005). The method can be made quantitative if the V content of the parental melt is known, assuming equilibrium between the melt and chromite cores, by calculating oxygen fugacity from the relationship of D v s p to fo 2 of Canil (2002), log[rf(Vmdt/Vsp)-l] = b log/o2 + c

(25)

Where d, b and c are fit parameters dependent on melt composition and temperature. Papike et al. (2005) use the fit parameters for komatiite from Canil (1999), which are d = 32.8, b = 0.41, and c = 0.77. Other fit parameters for different compositions, including spinel with low Cr/Al, are provided in the Background Data Set that accompanies Canil (2002), http://www.elsevier. com/locate/epsl. In the absence of parental melt data, Papike et al. (2005) used whole-rock V concentrations to approximate the V contents of parental melts, and obtained results that are broadly consistent with those of other studies. As it currently stands, application of the V-in-chromite oxybarometer is limited to /o 2 's between IW - 1 and IW + 4, which is the range over which D v s p for chromite has been determined (Canil 2002). In addition, V 4+ becomes dominant at ~ FMQ (IW + 3.5), and the oxybarometer loses resolution due to the limits of the crystal chemistry of spinel. The relationship of D v s p

546

Herd

t° fo2 is dependent on melt composition, P and T, and the user should take care to choose appropriate fit parameters. Furthermore, application of the oxybarometer to spinels in which Cr and A1 vary significantly, or to ulvospinel-rich compositions, could yield spurious results. Regardless, this oxybarometer relies only on EMP data, and so is widely accessible. Perhaps the most powerful use of vanadium to determine oxygen fugacity relies on its geochemical behavior in combination with its existence in multiple valence states. The V/Sc ratio in terrestrial basalts and mantle xenoliths has been used to infer the oxygen fugacity of the primary magma or mantle source (Lee et al. 2003, 2005; Li and Lee 2004). Vanadium and scandium behave so similarly in magmatic systems that the V/Sc ratio will remain largely unaffected by olivine fractionation, cryptic metasomatism, crustal contamination, or degassing (Canil 2004; Lee et al. 2005). Oxygen fugacity will have the most significant effect on V/Sc; whereas V has variable valence, Sc exists only as Sc3+. Therefore, the V/Sc ratio will "see through" post-extraction processes and reflect the oxygen fugacity of magmagenesis. Lee et al. (2005) implemented the V/Sc oxybarometer by modeling the dependence of V/Sc on oxygen fugacity, under assumptions of isothermal (1410 °C) and isobaric (1.5 GPa) partial melting within the spinel stability field, and a fertile convecting mantle with a constant V/Sc ratio. Analyses of whole-rock V and Sc concentrations are all that are required; these values are compared to the modeled values to determine the primary oxygen fugacity. The results of Lee et al. (2005) have significant implications for the question of whether basalt oxygen fugacity reflects that of its mantle source. THE BASALT-MANTLE SOURCE REDOX RELATIONSHIP Is basalt oxygen fugacity reflective of the redox state of its mantle source? Whether the oxygen fugacity of a planetary basalt reflects the redox state of the mantle source from which it was derived is of fundamental importance in determining the redox states and histories of planetary interiors. In the absence of mantle samples from the other terrestrial planets, insights into the basalt-mantle source relationship can be gained from attempts to explain the diversity of basalt oxygen fugacity on the Earth. The relationship between oxygen fugacity and FeO and Fe 2 0 3 in silicate liquids (Eqn. 8) is extended by Kress and Carmichael (1991) to quantify the effect of pressure. The Fe 3+ / EFe ratio of a melt closed to oxygen during its ascent will change such that the oxygen fugacity will be broadly parallel to FMQ; i.e., the relative oxygen fugacity is nearly independent of pressure. Pressure affects the solid oxygen buffers differently, with the FMQ buffer changing by -0.17 log units/GPa and the NNO buffer by -0.51 log units/GPa (Kress and Carmichael 1991). The implication is that a silicate liquid that is closed to oxygen during ascent will retain a record of its oxygen fugacity relative to FMQ to within a fraction of a log unit; therefore, glassy lavas can be used as probes of planetary interior redox state. Given that there is a 7-log unit range in oxygen fugacity in basic lavas on the Earth, it follows that there must exist mantle sources with an equivalent range of oxygen fugacity (Carmichael 1991). Other workers hold a different view, in which the C-H-O volatile species play a more significant role in buffering basalt mantle sources, or ascending magmas, and in determining the oxygen fugacity of the erupted basalt (Mathez 1984; Blundy et al. 1991; Ballhaus and Frost 1994). In the Ballhaus and Frost (1994) model, the range of oxygen fugacity of basalts is explained by variation in mantle source redox state combined with pressure effects, and the redox states of mantle sources are constrained to a much narrower range than proposed by Carmichael (1991). Ballhaus and Frost (1994) envision decompression melting, along an adiabatic ascent path, and assume that the melt remains in major element equilibrium with the crystalline phases of the mantle material until fairly low pressure. They choose an arbitrary

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initial oxygen fugacity of FMQ - 4 at a depth of 4 - 5 GPa; at these conditions, the system is buffered by ferrous-ferric equilibria. Although this is within graphite-water-methane stability, the oxygen fugacity, and hence C 0 2 activity, are too low to be controlled by C-H-O equilibria. Given that the relative change in oxygen fugacity with increasing pressure is approximately - 0 . 6 log units/GPa (Ballhaus 1995), upwelling asthenosphere will experience oxidation (relative to FMQ, for example) with decreasing pressure. This results in oxidation of graphite, an increase in the activity of C 0 2 , and a shift in buffering from ferrous-ferric to C-H-O; that is, buffering by the presence of graphite and the activity of C 0 2 . At some point, the adiabatic ascent path intersects the dry solidus of the asthenosphere, causing a rapid advance in the degree of partial melting. Eventually, the melt reaches a critical pressure and oxygen fugacity interval where the solubility of carbon in the melt as C 0 2 exceeds the amount of elemental carbon present in the residue; graphite is eliminated as a phase, and the buffering switches to ferrous-ferric equilibria, with an increase in relative oxygen fugacity with further decompression. Mathez (1984) outlines a model in which degassing of C-rich vapor species play a significant role in buffering f0r An example is provided in which a C-supersaturated, reduced magma degasses in a shallow magma chamber (depths of < 0.3 GPa), exsolving a CO-rich gas and liberating 0 2 , which oxidizes iron in the melt according to equation (7). Assuming slow, continuous and infinitesimal exsolution of the vapor, an initially reduced (~IW) magma can be oxidized such that the erupted melt approaches FMQ (Mathez 1984). However, this model assumes a magma that is C-supersaturated and reduced; contrast this with the Ballhaus and Frost (1994) model, in which the magma is no longer buffered by graphite by the time it has reached shallow depths. The Ballhaus and Frost (1994) model requires much less variation of redox state among basalt mantle sources. Instead of an intrinsic variation in redox state of mantle sources, the oxygen fugacity of the basalt at the surface will depend on the depth at which buffering switches from C-H-O to ferrous-ferric: the greater the depth at which graphite becomes eliminated in the residue, the more oxidized the melt will be at the surface. The model is used to explain the higher oxygen fugacity of OIB relative to M O R B ; because MORBs are typically derived from shallow mantle sources, their recorded oxygen fugacity is only slightly higher than that at which major melting (and separation from graphite buffering) occurred. Ocean island basalts, on the other hand, are more oxidized because melting occurs at greater depths and the melt experiences more relative oxidation during ascent. The much higher oxygen fugacity of island arc basalt relative to MORB and OIB can be explained by a source that is intrinsically more oxidized than graphite stability, due to the oxidation of the mantle wedge by slab-derived fluids. As such, the IAB source may be around FMQ, and these basalts have high oxygen fugacities as a result of relatively deep major melting. In contrast to Carmichael (1991), this model implies that the oxygen fugacities of MORB and OIB can be produced from the same mantle, by simply changing the depth of first major melting; the Mathez (1984) model implies that oxidation can occur by stalling and degassing the magma in a crustal reservoir. Such models highlight a potential incongruity between the oxygen fugacity of the basalt collected at the surface, and the redox state of the mantle from which the basaltic melt was derived. Insight into which of these models is more accurate may be derived from the use of V/ Sc ratios to infer the oxygen fugacity of melting for terrestrial basalts and mantle xenoliths (Lee et al. 2005). Using new V and Sc measurements, and data from the literature, Lee et al. (2005) obtain oxygen fugacity results of FMQ - 2 to FMQ for peridotites. This is in good agreement with the results from ferrous-ferric mineral equilibria (e.g., olivine-pyroxene spinel oxybarometry) for suboceanic abyssal peridotites and undepleted, fertile subcontinental mantle xenoliths. V/Sc results for MORB are self-consistent with the results for peridotites, falling within FMQ - 1.25 and FMQ + 0.25 (Lee et al. 2005), which agree with previous studies of MORB oxygen fugacity (Christie et al. 1986; Wood et al. 1990). Remarkably, V/Sc ratios for IAB overlap those from MORB, implying that the IAB and MORB mantle sources have the same oxygen fugacity, between FMQ - 1.25 and FMQ + 0.25. This is in contrast to

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results from the application of equation (8) to glassy arc lavas, which yield FMQ to FMQ + 6 (Carmichael 1991). The implication of the V/Sc results is that the oxygen fugacity of the Earth's asthenosphere is buffered, perhaps in a manner suggested by Blundy et al. (1991) or Ballhaus and Frost (1994). It is notable that in the range of 1 to 3 GPa, the CCO buffer curve falls between FMQ - 2 and FMQ - 1 (Ballhaus and Frost 1994). Studies of V and Cr in basalts and mantle xenoliths with ages ranging from Archean to the present demonstrate that the oxygen fugacity of the Earth's asthenosphere has remained buffered to within a log unit of FMQ throughout its history (Delano 2001; Canil 2002; Lee et al. 2003; Li and Lee 2004). The implications of the V/Sc results for the interpretation of oxygen fugacity results from planetary basalts are profound. The observation that the oxygen fugacity of IAB (as derived from mineral equilibria and ferrous-ferric ratios in glass) is several log units higher than that of its source implies that post-extraction processes, such as fractional crystallization, dissociation of volatiles and degassing, auto-oxidation (e.g., Holloway 2004), and hydrothermal alteration may significantly affect the oxygen fugacities recorded by most oxybarometers. Furthermore, it calls into question whether models of basaltic magma evolution in which the system is closed to oxygen (e.g., Kress and Carmichael 1991; Ghiorso 1997) are applicable to most natural systems. Implications for understanding the redox states of planetary interiors There are several methods now available for determining the oxygen fugacities of planetary basalts; many of these methods have been applied to samples from the Moon, Mars and asteroids, as summarized by Wadhwa (2008). The debate and uncertainty regarding the cause of the variation in oxygen fugacity of terrestrial basaltic samples is instructive - in the absence of mantle xenoliths, what can be said about the redox state of the interiors of the other terrestrial planets? The observation from V/Sc ratios that the mantle source of arc basalts is reduced, while the corresponding eruptives are up to several log units more oxidized, indicates that basalts from arc environments are open systems whose oxygen fugacity reflects post-extraction processes such as differentiation, degassing and assimilation; the same might be true for basalts from other tectonic environments. Although plate tectonics is removed as a complicating factor when discussing other terrestrial planets, the question of whether planetary basalts are open to oxygen during ascension and eruption is equivocal. Variations in oxygen fugacity, such as those observed for Mars (Wadhwa 2001; Herd 2003), need to be interpreted in the context of other indicators of mantle source characteristics. In spite of the lack of mantle xenoliths among our sample suites from the other terrestrial planets, insights into redox states of planetary interiors can be gained through the consideration of the relative roles of volatiles and Fe-bearing mineral equilibria, by analogy with the Earth. High-pressure experiments on mantle or primitive basalt compositions can assist in elucidating the relative change in oxygen fugacity with depth, and determine whether processes such as iron disproportionation have influenced the redox states of the lower mantles of the larger bodies (e.g., Venus, Mars). Comparative studies are particularly relevant and lead to new avenues of research. For example, given the large differences in the C-H-O budgets of the Earth, Moon, Mars and asteroids, what factors in the formation and geologic evolution of the planetary body have the greatest influence on whether mantle source redox state is dominated by volatiles or Febearing mineral equilibria? Given the similarities in C-H-O budgets of the Earth and Mars, what is the role of volatiles in controlling basalt oxygen fugacity on Mars? How can the observation of methane in the martian atmosphere (Formisano et al. 2004) be interpreted in this context? The relatively high (~FMQ) and constant redox state of the Earth's asthenosphere since the Archean has implications for the speciation of volatiles that were available during the development of life (Delano 2001; Canil 2002). Likewise, insights into the redox evolution of Mars will influence our understanding of the conditions on early Mars and contribute to the assessment of whether the environment of early Mars was conducive to life.

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ACKNOWLEDGMENTS The author thanks all of the participants of the Oxygen in the Terrestrial Planets Workshop for contributing to the discussion of redox conditions in planetary samples. This paper benefited from discussion with Bob Luth, Karlis Muehlenbachs, Tom Chacko and Thomas Stachel. Thanks to Cin-Ty Lee for the fugacity poetry. A thorough review by John Longhi and the editorship of Steve Simon are gratefully acknowledged.

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Reviews in Mineralogy & Geochemistry Vol. 68, pp. 555-569, 2008 Copyright © Mineralogical Society of America

Rheological Consequences of Redox State Stephen Mackwell Lunar

and Planetary Institute 3600 Bay Area Blvd Houston, Texas 77058, U.S.A. mackwell @ Ipi. usra. edu

ABSTRACT Within the upper mantle of Earth, there is a gradient from a relatively more oxidizing near-surface region, with oxygen fugacities near those of the fayalite-magnetite-quartz buffer (FMQ), to more reducing conditions at depth near the iron-wustite buffer (IW). Oxygen fugacity appears to vary laterally, as well as vertically, by as much as a factor of 10 4 . As flow within the interior of the Earth and other terrestrial planets occurs due to the (mostly) solidstate deformation of rocks, an understanding of the effect of oxygen fugacity on creep is critical in modeling planetary interior dynamical behavior. This is especially important for the asthenosphere, that anomalously weak region in the uppermost mantle that accommodates isostasy and largely decouples mantle convection from plate motions. Experimental studies of the rheological behavior of iron-bearing minerals have demonstrated that oxygen fugacity can play an important role in deformation. We have shown that olivine rich rocks deformed near F M Q deform in the dislocation creep regime about a factor of 6 faster when buffered near F M Q than at IW. Experiments on olivine single crystals and aggregates indicate that this difference in behavior results from an increase in the concentration of silicon vacancies under more oxidizing conditions, as dislocation creep is rate-limited by the climb of dislocations, which is controlled by diffusion of silicon defects. Although fewer data are available for the effects of oxygen fugacity on pyroxene deformation, clinopyroxene appears to be stronger under more oxidizing conditions, while the data on orthopyroxene deformation show no dependence on oxygen fugacity. These results indicate that vertical and lateral variations in oxygen fugacity may result in, at most, an order of magnitude difference in viscosity, while other factors, such as water fugacity and lithology may be more significant.

INTRODUCTION E x p e r i m e n t a l s t u d i e s of t h e r h e o l o g i c a l b e h a v i o r of i r o n - b e a r i n g m i n e r a l s h a v e d e m o n s t r a t e d t h a t o x y g e n f u g a c i t y c a n p l a y an i m p o r t a n t r o l e in d e f o r m a t i o n . A s flow w i t h i n t h e i n t e r i o r of t h e E a r t h a n d o t h e r t e r r e s t r i a l p l a n e t s o c c u r s d u e to t h e ( m o s t l y ) s o l i d - s t a t e d e f o r m a t i o n of r o c k s c o m p o s e d of s u c h m i n e r a l s , an u n d e r s t a n d i n g of t h e e f f e c t of o x y g e n f u g a c i t y o n c r e e p is critical in m o d e l i n g p l a n e t a r y i n t e r i o r d y n a m i c a l b e h a v i o r . W i t h i n t h e E a r t h , t h e r e is a g r a d i e n t f r o m a r e l a t i v e l y m o r e o x i d i z i n g n e a r - s u r f a c e r e g i o n , w i t h o x y g e n f u g a c i t i e s n e a r t h o s e of t h e f a y a l i t e - m a g n e t i t e - q u a r t z b u f f e r ( F M Q ) , to m o r e r e d u c i n g c o n d i t i o n s at d e p t h ( f o r r e c e n t r e v i e w s , s e e M c C a m m o n 2 0 0 5 , H e r d 2 0 0 8 ) . It is g e n e r a l l y a c c e p t e d t h a t t h e o x y g e n f u g a c i t y in t h e l o w e r m a n t l e is b u f f e r e d n e a r t h e i r o n - w i i s t i t e b u f f e r ( I W ) . M e a s u r e m e n t s of o x y g e n fugacity f r o m rocks derived f r o m the u p p e r mantle suggest that o x y g e n fugacity varies significantly both vertically and laterally ( M c C a m m o n 2 0 0 5 ; Herd 2008). S u c h variability m a y c o r r e s p o n d to as m u c h as a f a c t o r of 10 4 d e c r e a s e in o x y g e n f u g a c i t y at a s t h e n o s p h e r i c d e p t h s . A s t h e a s t h e n o s p h e r e is t h e a n o m a l o u s l y w e a k r e g i o n in t h e u p p e r m o s t m a n t l e t h a t a c c o m m o d a t e s i s o s t a s y a n d l a r g e l y d e c o u p l e s m a n t l e c o n v e c t i o n f r o m p l a t e m o t i o n s , it is 1529-6466/08/0068-0020$05.00

DOI: 10.2138/rmg.2008.68.20

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important to quantify the difference in mechanical behavior of mantle rocks as a function of oxygen fugacity. The following sections describe the current understanding of the effect of oxygen fugacity on deformation of mantle minerals of the terrestrial planets, including some new data from our current research. DEFORMATION OF OLIVINE Olivine (Mg x Fe!_ x ) 2 Si0 4 is the most abundant mineral in the upper mantle of Earth (with x~0.9) and, likely, those of the other terrestrial planets (x~0.7-0.9). Based on experimental measurements and observations from ophiolites and other olivine-bearing rocks exposed at the Earth's surface, it is also believed to be the weakest mineral in the upper mantle. Thus, the mechanical behavior of the upper mantle is approximated well by the rheological behavior of olivine aggregates (dunites). There have been numerous studies of the deformation behavior of olivine, dating back to the earliest rheological experiments on olivine aggregates (e.g., Raleigh 1968; Carter and Ave Lallemant 1970; Raleigh and Kirby 1970; Kirby and Raleigh 1973; Post 1977) and single crystals (e.g., Blacic 1972; Phakey et al. 1972). These early studies made no attempt to control, least of all vary, oxygen fugacity, however; only in the last three decades has there been a systematic attempt to constrain the role of oxygen fugacity in deformation experiments on olivine single crystals and aggregates. Olivine single crystal studies Although the earliest work on deformation of olivine single crystals dates to the pioneering work of Phakey et al. (1972) and Blacic (1972) using solid-medium deformation apparatus, rigorous control of oxygen fugacity was first attempted by Kohlstedt and Goetze (1974) on unoriented samples of single-crystal olivine, followed by the work of Durham and Goetze (1977) on single crystals oriented to favor slip on the easiest slip systems. Both of these studies used a fixed mixture of C 0 2 and H 2 to control oxygen fugacity. While neither study attempted to vary the oxygen fugacity during the experiment, they both acknowledged that iron-bearing natural olivine is only stable under conditions significantly more reducing than in air. In addition, the study reported in Durham and Goetze (1977) and Durham et al. (1977) demonstrated that, at high temperatures in the range 1080 to 1575 °C, significant differences in dislocation micro structure accompany the different slip systems. Interestingly, these authors observed the dislocation microstructures optically in thin sections, using an oxidation technique (Kohlstedt et al. 1976) to decorate dislocations by briefly heating the thin sections in air. The first studies to investigate the dependence of creep of olivine single crystals on oxygen fugacity were performed by Hornack (1978), Jaoul et al. (1980) and Kohlstedt and Hornack (1981). Working with natural olivine single crystals that were unbuffered in silica activity and using mixed C 0 - C 0 2 gases to control oxygen fugacity, they demonstrated a positive dependence of creep rate on oxygen fugacity, with the strain rate approximately proportional to/o 2 1/6 , consistent with point defect models (e.g., Stacker 1978a, 1978b; Stacker and Smyth 1978). In a subsequent study, Ricoult and Kohlstedt (1985) deformed olivine single crystals buffered against either orthopyroxene (Mg,Fe)Si0 3 or magnesiowiistite (Mg,Fe)0 to control the silica activity. Although the oxygen fugacity was nominally varied over a range of conditions using mixtures of CO and C 0 2 gases, a subsequent investigation by Bai et al. (1991) demonstrated that the tungsten/molybdenum furnace and foils used in the Ricoult and Kohlstedt (1985) study effectively buffered the oxygen fugacity at a fixed, much more reducing condition than suggested by the gas mixtures, so that no range in oxygen fugacity was actually investigated. Bai et al. (1991) used a new room-pressure apparatus with internal components that did not react with the mixed C 0 - C 0 2 gases, and with a downline zirconia oxygen sensor that monitored the oxygen fugacity of the gas mixture. They performed a detailed study of the

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creep behavior of natural olivine single crystals, investigating three orientations that favor the four easiest slip systems, varying temperature from 1200 to 1525 °C, oxygen fugacity from 10~12 to 10~3 atm, and compressive stresses between 14 and 180 MPa to yield creep rates between 10"4 and 10~6,7 s -1 . Samples were buffered in silica activity using either orthopyroxene or magnesiowiistite buffers. While the dependence of creep rate on applied stress (a) was remarkably constant, with creep rate proportional to a 3,5 , they noted significant complexity in creep behavior with variation in oxygen fugacity, silica buffer and temperature. Figure 1 (after Fig. 5 of Bai et al. 1991) illustrates this complexity for samples deformed favoring the easiest high-temperature slip system in olivine (010)[100], involving motion of [100] Burgers vector dislocations on the (010) planes. The dependence of creep rate on oxygen fugacity changes as oxygen fugacity is increased at lower temperatures, and a third dependence on oxygen fugacity becomes apparent at higher temperatures. Thus, creep of olivine single crystals favoring this slip system shows changes in dependence on oxygen fugacity and temperature as oxygen fugacity and temperature are varied, requiring complex formalisms (see equations in Table 4 of Bai et al. 1991) to describe the high-temperature deformation for each slip system. In a subsequent analysis of the sample microstructures using optical and electron microscope techniques (Bai and Kohlstedt 1992), the dislocation substructures were shown to be distinctly different between many of the deformation fields that possess unique fugacity dependencies and activation energies. These observations suggest that extrapolation of experimentally determined rheologies to model the behavior of the Earth or other planets must be done cautiously in order to ensure that the constitutive parameters are appropriate to the deformation mechanism operative in that body (Mackwell et al. 1990). Recent unpublished work by S. Mackwell, G. Hirth and D. Kohlstedt on deformation of olivine single crystals in high-pressure gas-medium deformation apparatus demonstrates a

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Figure 1. Plot of strain rate versus oxygen fugacity from experiments by Bai et al. (1991) on single crystals of olivine buffered by orthopyroxene and oriented to favor slip on (010)[100], the easiest high-temperature slip system in olivine. The symbols represent strain rate measurements at a variety of temperatures for an applied stress of 30 MPa. The solid lines are fits to the experimental data, which have been decomposed into individual creep laws illustrated by the dashed lines. The dotted lines show the oxygen fugacities for iron-wiistite (IW) and nickel-nickel oxide (NNO) for each temperature. /

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similar dependence of creep rate on oxygen fugacity as the room pressure studies. In this work, the oxygen fugacity was controlled using either a Ni jacket with a NiO powder coating the samples (NNO buffer), or an Fe jacket with an FeO coating (IW buffer). It is worth noting that the NNO buffer is ~1 order of magnitude in oxygen fugacity more oxidizing than the fayalitequartz-magnetite (FMQ) buffer (Fig. 2). Figure 3 shows that, in general, samples buffered at IW have lower strain rates, and are therefore stronger than, samples buffered at NNO under both wet and dry deformation conditions. Olivine aggregate studies As noted above, dunites (olivine aggregates) have been the focus of experimental deformation studies for more than forty years. Early pioneering studies by Raleigh (1968), Carter and Ave Lallemant (1970), Raleigh and Kirby (1970), Kirby and Raleigh (1973), and Post (1977) mapped out the deformation behavior of olivine aggregates using solid-medium, high-pressure deformation apparatus. These studies established that the behavior followed a power-law behavior, where the strain rate is proportional to a", where n lies in the range 3.5 to 5.5, consistent with deformation by dislocation creep. While these studies were pivotal in establishing rheologies applicable to lithospheric and asthenospheric processes, they were not well constrained in terms of the chemical state of the samples during deformation, notably in terms of the oxygen and water fugacity. While it is often assumed that the talc solid-medium assemblies approximately buffer the samples near FMQ, oxygen fugacity was not rigorously constrained. Also, the dehydration of hydrous minerals that are ubiquitous in natural dunites likely resulted in some water-weakening of the samples. The presence of some orthopyroxene in dunite probably buffered the silica activity. In later experiments on dunites in the solid-medium apparatus, Zeuch and Green (1984) attempted to minimize a number of these problems. They used samples synthesized from olivine sands and made improvements in the solid-medium apparatus that allowed an increase in stress resolution and decrease in

T (°C) 1600 1400

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Figure 2. Plot of oxygen fugacity versus inverse temperature illustrating the phase boundaries for the iron-wiistite (IW), wiistite-magnetite (WM) and magnetite-hematite (MH), NNO, and fayalite (Fe 2 Si0 4 ) systems at room pressure. The subvertical solid line represents the melting of fayalite at room pressure. IQF represents the phase boundary between fayalite and quartz/iron, and F M Q represents that between fayalite and quartz/magnetite.

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