Earth's Early Atmosphere and Surface Environment [504] 9780813725048, 0813725046

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Earth’s Early Atmosphere and Surface Environment

edited by George H. Shaw Geology Department Union College Schenectady, New York 12308, USA

Special Paper 504 3300 Penrose Place, P.O. Box 9140

Boulder, Colorado 80301-9140, USA

2014

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Copyright © 2014, The Geological Society of America (GSA), Inc. All rights reserved. Copyright is not claimed on content prepared wholly by U.S. government employees within the scope of their employment. Individual scientists are hereby granted permission, without fees or further requests to GSA, to use a single figure, a single table, and/or a brief paragraph of text in other subsequent works and to make unlimited photocopies of items in this volume for noncommercial use in classrooms to further education and science. Permission is also granted to authors to post the abstracts only of their articles on their own or their organization’s Web site providing that the posting cites the GSA publication in which the material appears and the citation includes the address line: “Geological Society of America, P.O. Box 9140, Boulder, CO 80301-9140 USA (http://www.geosociety.org),” and also providing that the abstract as posted is identical to that which appears in the GSA publication. In addition, an author has the right to use his or her article or a portion of the article in a thesis or dissertation without requesting permission from GSA, provided that the bibliographic citation and the GSA copyright credit line are given on the appropriate pages. For any other form of capture, reproduction, and/or distribution of any item in this volume by any means, contact Permissions, GSA, 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA; fax +1-303-357-1073; [email protected]. GSA provides this and other forums for the presentation of diverse opinions and positions by scientists worldwide, regardless of their race, citizenship, gender, religion, sexual orientation, or political viewpoint. Opinions presented in this publication do not reflect official positions of the Society. Published by The Geological Society of America, Inc. 3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301-9140, USA www.geosociety.org Printed in U.S.A. GSA Books Science Editors: Kent Condie and F. Edwin Harvey Library of Congress Cataloging-in-Publication Data Earth’s early atmosphere and surface environment / edited by George H. Shaw, Geology Department, Union College, Schenectady, New York 12308, USA. pages cm. — (Special paper ; 504) Includes bibliographical references. Summary: “This volume explores the range of conditions and compositions that have been proposed for Earth’s early surface and atmosphere with a collection of papers presented at a 2011 GSA Pardee Symposium. Each chapter is accompanied by a commentary and followed by a transcript of the ensuing discussion at the meeting”— Provided by publisher. ISBN 978-0-8137-2504-8 (pbk.) 1. Earth (Planet)—Origin—Congresses. 2. Atmospheric chemistry—Congresses. 3. Physical geography—Congresses. I. Shaw, George H., 1945–. II. Geological Society of America. Meeting (2011: Minneapolis, Minn.) III. Pardee Keynote Symposium (2011 : Minneapolis, Minn.) QC879.6.E27 2014 551.51′109012—dc23 2014001614 Cover: (Front) The first unequivocal evidence of life on Earth, a 3.5 Ga stromatolite, lies nestled in a remote hillside in the Australian Pilbara’s North Pole Dome. Photo by Lev Horodyskyj. (Back) Ferruginous stromatolites from the Biwabik Iron Formation, Minnesota (ca. 1.8 Ga). The stromatolite columns are approximately 1 cm across. Photo by E. Calvin Alexander Jr. of a sample from the collections in the Earth Sciences Department, University of Minnesota.

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Contents

Introduction—Earth’s early atmosphere and surface environment George H. Shaw . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . v

SHAW 1. Evidence and arguments for methane and ammonia in Earth’s earliest atmosphere and an organic compound–rich early ocean . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1 George H. Shaw 2. How low can you go? Maximum constraints on hydrogen concentrations prior to the Great Oxidation Event . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 Shawn Domagal-Goldman 3. Discussion of “Evidence and arguments for methane and ammonia in Earth’s earliest atmosphere and an organic compound–rich early ocean” (Shaw) . . . . . . . . . . . . . . . . . . . . . . . . 15

KASTING 4. Atmospheric composition of Hadean–early Archean Earth: The importance of CO . . . . . . . . . . 19 James F. Kasting 5. Atmospheric composition of Hadean–early Archean Earth: The importance of CO: Comment . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 Laura Schaefer and Bruce Fegley Jr. 6. Discussion of “Atmospheric composition of Hadean–early Archean Earth: The importance of CO” (Kasting) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33

ZAHNLE and CATLING 7. Waiting for O2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 Kevin Zahnle and David Catling 8. Discussion of “Waiting for O2” (Zahnle and Catling) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49

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Contents OHMOTO et al. 9. Oxygen, iron, and sulfur geochemical cycles on early Earth: Paradigms and contradictions. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 Hiroshi Ohmoto, Yumiko Watanabe, Antonio C. Lasaga, Hiroshi Naraoka, Ian Johnson, Jamie Brainard, and Andrew Chorney 10. The upside-down biosphere: “Evidence for the partially oxygenated oceans during the Archean Eon” . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 97 Shawn Domagal-Goldman 11. Discussion of “Oxygen, iron, and sulfur geochemical cycles on early Earth: Paradigms and contradictions” (Ohmoto et al.) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 101

REVIEW AND CONCLUSIONS 12. Earth’s early atmosphere and surface environments: A review . . . . . . . . . . . . . . . . . . . . . . . . . . 105 Martin J. Van Kranendonk 13. Concluding comments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 131 George H. Shaw

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The Geological Society of America Special Paper 504 2014

Introduction: Earth’s early atmosphere and surface environment Investigations of the chemical state of Earth’s early surface and atmosphere have been guided by geological evidence, cosmochemical analysis, and comparisons to other terrestrial bodies. A Pardee Symposium was held at the Geological Society of America Annual Meeting in Minneapolis, Minnesota, on 12 October 2011 with the goal of covering the broadest possible range of ideas that had then been developed addressing this important, if obscure, period in Earth’s history. The symposium provided for the presentation and discussion of several, often contradictory, models for the early Earth. The emphasis was on the later Hadean (post–moon-forming impact) to the late Archean (Great Oxidation Event), that is, from about 4.3 Ga to 2.1 Ga. The symposium used an unusual format with the intention of stimulating the maximum amount of discussion of the various views of the nature of Earth’s early environment. There were four invited papers, with one hour devoted to each topic. The presenters were asked to submit in advance a written paper of about 3000 words. Each paper was then sent to a “commentator” who prepared a statement of about 1000–1500 words, which was sent back to the presenter.1 The commentary was not intended or expected to be (solely) a critique of the original paper, but rather to include additional thoughts based on the content of the paper, in order to further stimulate discussion. At the meeting, each presenter delivered the paper as originally conceived, not in response to the commentary, in about 25 minutes. The commentator then delivered the commentary in about 10 minutes, followed by a brief response from the presenter if desired. The remainder of the time was available for discussion. A recording was made of the entire session in order to prepare a transcript of the discussion. In most cases it was possible to identify those who made comments or asked questions at the meeting, and they are so indicated in the discussion sections. Very minor edits and changes for clarity have been made to the discussion segments. This volume is the tangible result of the symposium. It incorporates the presentations, commentaries, and discussions in the order presented at the meeting. The review process for the presentations resulted in some modifications to the various texts, but the goal was to preserve, as much as possible, the flavor of the symposium. In addition, the inevitable passage of time for preparation of the volume led to some additional text and references from the past couple years. I believe I have been reasonably true to the essence of the symposium even with these modifications, and they have undoubtedly resulted in improvements to the various texts. With the hope of further increasing the value of this volume, I also commissioned an audience participant to prepare a paper covering the symposium topic as a whole and summarizing, to the extent possible, the current state of the field (at least from his viewpoint). One of the goals was to demonstrate a somewhat novel session format that might be used in the future, and this volume provides a record that might be useful for someone contemplating this. The success of this effort can perhaps best be suggested by the comment of one audience participant. After the session, he described his skepticism upon being informed at the beginning, of the format to be used. He then stated that he felt it was one of the best sessions at the meeting. It should be clear that this format is probably not amenable to all topics, and may be best suited for those where there is a

1

Although the paper by K. Zahnle and D. Catling (Chapter 7) was sent to a commentator, the commentator withdrew at a point in time too late to obtain a replacement. Shaw, G.H., 2014, Introduction: Earth’s early atmosphere and surface environment, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. v–vi, doi:10.1130/2014.2504(00). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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G.H. Shaw

broad range of views in a focused area, with some degree of controversy implicit in the range. It is our hope that this effort could be the stimulus for someone to try this again at a future GSA meeting. As for the substance of the topic discussed, it is clear that an extremely broad range of viewpoints is held by geologists, geochemists, atmospheric chemists, climate scientists, and various others, as informed by their own perspectives and the data they feel are most critical to discussion. In some cases, there are disparate inferences and conclusions drawn from more or less the same basic data. While this may seem surprising, it is, perhaps, a consequence of the nature of such an ancient and often skimpy geologic record. Although it is unlikely that all (if any) of the views presented can be correct, there appears at least some possibility that some (variable) fraction of each may have some validity in forming a picture of an important period of Earth’s history, one in which it is highly probable that life began its long journey. It is our hope that the discussions presented herein will help in the development of a self-consistent model of the early Earth. The underlying theme of the symposium centers on the question of the oxidation state of the atmosphere, and by extension, the general surface environment. The sequence of presentations was chosen to proceed from consideration of a highly reduced initial surface (G.H. Shaw, Chapter 1), through what is seemingly the generally accepted view of a more or less neutral (in the sense of oxidation state) N2-CO2 (w/CO?) atmosphere (J.F. Kasting, Chapter 4), and the processes by which it ultimately became strongly oxidizing and with free O2 (K. Zahnle and D. Catling, Chapter 7), to the rather controversial view that free oxygen has been present from very early (perhaps since 3.5 Ga) in Earth’s history (H. Ohmoto et al., Chapter 9). This scheme to some extent follows the historical development of ideas concerning the atmosphere, but also “bookends” the more conventional view that N2-CO2 “defines” the chemical state of the atmosphere and surface, with more radical views on both the more reducing and oxidizing ends of the spectrum. The student of Earth’s early atmosphere and/or surface will find here the widest range of thought available on this aspect of a critical and extensive part of Earth’s history I would like to thank GSA and the Pardee Symposia organizers for sponsoring this symposium, allowing us to test a novel (for us) format. I would also like to thank all of the participants for their thoughtful contributions, and the symposium attendees for their energetic discussions. Robert Pepin was gracious enough to be a co-convener and his participation before, during, and after the symposium is much appreciated. George H. Shaw Geology Department, Union College, Schenectady, New York 12308, USA [email protected]

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The Geological Society of America Special Paper 504 2014

Introduction: Earth’s early atmosphere and surface environment Investigations of the chemical state of Earth’s early surface and atmosphere have been guided by geological evidence, cosmochemical analysis, and comparisons to other terrestrial bodies. A Pardee Symposium was held at the Geological Society of America Annual Meeting in Minneapolis, Minnesota, on 12 October 2011 with the goal of covering the broadest possible range of ideas that had then been developed addressing this important, if obscure, period in Earth’s history. The symposium provided for the presentation and discussion of several, often contradictory, models for the early Earth. The emphasis was on the later Hadean (post–moon-forming impact) to the late Archean (Great Oxidation Event), that is, from about 4.3 Ga to 2.1 Ga. The symposium used an unusual format with the intention of stimulating the maximum amount of discussion of the various views of the nature of Earth’s early environment. There were four invited papers, with one hour devoted to each topic. The presenters were asked to submit in advance a written paper of about 3000 words. Each paper was then sent to a “commentator” who prepared a statement of about 1000–1500 words, which was sent back to the presenter.1 The commentary was not intended or expected to be (solely) a critique of the original paper, but rather to include additional thoughts based on the content of the paper, in order to further stimulate discussion. At the meeting, each presenter delivered the paper as originally conceived, not in response to the commentary, in about 25 minutes. The commentator then delivered the commentary in about 10 minutes, followed by a brief response from the presenter if desired. The remainder of the time was available for discussion. A recording was made of the entire session in order to prepare a transcript of the discussion. In most cases it was possible to identify those who made comments or asked questions at the meeting, and they are so indicated in the discussion sections. Very minor edits and changes for clarity have been made to the discussion segments. This volume is the tangible result of the symposium. It incorporates the presentations, commentaries, and discussions in the order presented at the meeting. The review process for the presentations resulted in some modifications to the various texts, but the goal was to preserve, as much as possible, the flavor of the symposium. In addition, the inevitable passage of time for preparation of the volume led to some additional text and references from the past couple years. I believe I have been reasonably true to the essence of the symposium even with these modifications, and they have undoubtedly resulted in improvements to the various texts. With the hope of further increasing the value of this volume, I also commissioned an audience participant to prepare a paper covering the symposium topic as a whole and summarizing, to the extent possible, the current state of the field (at least from his viewpoint). One of the goals was to demonstrate a somewhat novel session format that might be used in the future, and this volume provides a record that might be useful for someone contemplating this. The success of this effort can perhaps best be suggested by the comment of one audience participant. After the session, he described his skepticism upon being informed at the beginning, of the format to be used. He then stated that he felt it was one of the best sessions at the meeting. It should be clear that this format is probably not amenable to all topics, and may be best suited for those where there is a

1

Although the paper by K. Zahnle and D. Catling (Chapter 7) was sent to a commentator, the commentator withdrew at a point in time too late to obtain a replacement. Shaw, G.H., 2014, Introduction: Earth’s early atmosphere and surface environment, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. v–vi, doi:10.1130/2014.2504(00). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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broad range of views in a focused area, with some degree of controversy implicit in the range. It is our hope that this effort could be the stimulus for someone to try this again at a future GSA meeting. As for the substance of the topic discussed, it is clear that an extremely broad range of viewpoints is held by geologists, geochemists, atmospheric chemists, climate scientists, and various others, as informed by their own perspectives and the data they feel are most critical to discussion. In some cases, there are disparate inferences and conclusions drawn from more or less the same basic data. While this may seem surprising, it is, perhaps, a consequence of the nature of such an ancient and often skimpy geologic record. Although it is unlikely that all (if any) of the views presented can be correct, there appears at least some possibility that some (variable) fraction of each may have some validity in forming a picture of an important period of Earth’s history, one in which it is highly probable that life began its long journey. It is our hope that the discussions presented herein will help in the development of a self-consistent model of the early Earth. The underlying theme of the symposium centers on the question of the oxidation state of the atmosphere, and by extension, the general surface environment. The sequence of presentations was chosen to proceed from consideration of a highly reduced initial surface (G.H. Shaw, Chapter 1), through what is seemingly the generally accepted view of a more or less neutral (in the sense of oxidation state) N2-CO2 (w/CO?) atmosphere (J.F. Kasting, Chapter 4), and the processes by which it ultimately became strongly oxidizing and with free O2 (K. Zahnle and D. Catling, Chapter 7), to the rather controversial view that free oxygen has been present from very early (perhaps since 3.5 Ga) in Earth’s history (H. Ohmoto et al., Chapter 9). This scheme to some extent follows the historical development of ideas concerning the atmosphere, but also “bookends” the more conventional view that N2-CO2 “defines” the chemical state of the atmosphere and surface, with more radical views on both the more reducing and oxidizing ends of the spectrum. The student of Earth’s early atmosphere and/or surface will find here the widest range of thought available on this aspect of a critical and extensive part of Earth’s history I would like to thank GSA and the Pardee Symposia organizers for sponsoring this symposium, allowing us to test a novel (for us) format. I would also like to thank all of the participants for their thoughtful contributions, and the symposium attendees for their energetic discussions. Robert Pepin was gracious enough to be a co-convener and his participation before, during, and after the symposium is much appreciated. George H. Shaw Geology Department, Union College, Schenectady, New York 12308, USA [email protected]

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The Geological Society of America Special Paper 504 2014

Evidence and arguments for methane and ammonia in Earth’s earliest atmosphere and an organic compound–rich early ocean George H. Shaw* Geology Department, Union College, Schenectady, New York 12308, USA

ABSTRACT The preponderance of geologic evidence does not support carbon dioxide as the main carbon species degassed from early Earth, nor a carbon dioxide–rich early atmosphere. In fact, there are several problems that cannot be addressed by assuming either of these facets of what has become conventional wisdom about the early atmosphere. A careful examination of the conditions that most likely accompanied late accretion, incorporating the most probable average composition of accreting materials, suggests an early atmosphere produced by degassing of reduced carbon and nitrogen species, followed by photochemical processing to yield a surface environment rich in organic compounds. Recycling of these organics through hydrothermal and volcanic systems would have maintained a level of reduced gases (photochemically unstable as they may be) in the early atmosphere for an extended period, accompanied by a growing carbon dioxide component derived from mantle magmatism. Such a model for the early atmosphere is not only consistent with geological data, it also solves many problems of the early history of Earth.

PROBLEMS WITH THE “STANDARD MODEL”

of times present atmospheric concentrations (Kasting, 1993), the aforementioned problems remain. Solving any one of these problems does not make the others go away. Earth’s atmosphere is the most changeable element in the Earth system, especially over geologic time scales. Earth’s atmosphere was significantly different in the geologic past, with virtually no free oxygen before the latest Archean (Cloud, 1968, 1972). The processes involved in atmosphere–ocean–solid Earth interactions also operated at significantly different rates during the Hadean–Archean compared to the present, given the more energetic state of early Earth. Volcanic exhalation and hydrothermal activity were probably greater, and sedimentation was mainly in deep water, given the early absence of stable continental masses. Only by taking into account reasonably probable

The problems with an early CO2-rich atmosphere have been discussed at length in Shaw (2008a). Briefly, these include the excessive delay in atmospheric oxidation following the advent of oxygenic photosynthesis (given an enormous supply of photosynthetically available CO2), difficulties in producing prebiotic organics, a lack in the geologic record of the requisite amounts of Archean carbonate sediments or their metamorphic equivalents, a lack of intense and pervasive weathering due to atmospheric acidity/reactivity, and difficulties with producing the early sedimentary carbon isotope record. In addition, even if (as has been suggested) the faint young sun paradox might be overcome with a CO2-rich atmosphere, perhaps at a level hundreds to thousands

*[email protected] Shaw, G.H., 2014, Evidence and arguments for methane and ammonia in Earth’s earliest atmosphere and an organic compound–rich early ocean, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 1–10, doi:10.1130/2014.2504(01). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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system changes through time can we obtain a picture of Earth’s earliest atmosphere and surface, and their evolution through time. Furthermore, whatever insights we obtain should be consistent with known geological data preserved in early rocks. In a recent review of the Hadean and Archean (Shaw, 2008a), I proposed a possible resolution for many of the problems of early atmospheric composition vis-à-vis the geologic record. A key element of this proposal was a reconsideration of the nature of Earth’s earliest atmosphere and surface environment. In particular, I proposed an overwhelming dominance of reduced carbon species as a result of degassing, and as components of the early oceans and sediments. This contrasts with the conventional view of oxidized carbon species, especially CO2 and/or CO, as the primary degassed carbon compounds and as the main reservoir of carbon (e.g., Rubey, 1951, 1955; Ahrens et al., 1989; Nisbet and Sleep, 2001; Zahnle, 2006). This paper examines details of the generation and regeneration of reduced gases and condensed carbon species as the principal phases at/ near the surface of early Earth. SOURCE(S) OF VOLATILES AND THEIR COMPOSITION(S) There is no doubt that Earth’s atmosphere is “secondary,” the result of degassing volatile phases from solids that accreted to form the planet, as attested by the low abundances of noble gases and hydrogen in the modern atmosphere compared to solar nebula compositions (Pepin, 1991, 2006; Humayun and Cassen, 2000). The noncondensable gases were separated from condensable solids during the cooling of the nebula and formation of the early, small planetesimals, which could not gravitationally retain even the heaviest gases. The two major reservoirs of volatilebearing solids in accreting objects were comets and silicate-iron bodies, i.e., smaller asteroids and meteorites. Although comets have been considered important sources for Earth’s volatiles at various times (e.g., Morbidelli et al., 2000), the isotopic signature of cometary material suggests a relatively low upper limit on the amount of cometary debris contributing to Earth’s atmosphere (Drake and Righter, 2002). The recent measurement of Earth-like oxygen isotope ratios in at least one comet (Hartogh et al., 2011) somewhat relaxes this restriction, but it remains likely that meteoritic material was the major source of Earth’s volatiles. The best estimate of the bulk composition of the meteoritic fertile volatile

source material comes from consideration of Earth’s bulk composition and from examination of the meteoritic material arriving on Earth at present. The former consideration has given rise to the so-called “chondritic Earth model” to explain the bulk silicate composition of the outer two thirds of the planet, with a suitable addition of metallic meteorites to produce the core (Taylor, 1964; Sharma, 2009). The chondritic Earth model may or may not be relevant to considerations of Earth’s volatile inventory. Antarctic meteorites can probably be considered a good approximation to what is currently being added to Earth, and perhaps an indication of the previously accreted matter as well. Table 1 shows the masses of various meteorites collected in Antarctica up to June 2011 (Meteoritical Bulletin Database, www .lpi.usra.edu/meteor/metbull.php). A previous summary made in 1994 (Grossman, 1994) is slightly different, mostly in having a larger fraction of irons. The total amount of metallic iron in this compilation is ~19%, if one adds the metallic fraction in the various chondrites and stony-irons to the iron meteorite mass. This is somewhat smaller than the ~33% of metal that makes up Earth’s core, but it is close enough to bulk Earth to suggest that accreting compositions have not changed dramatically, and that the earliest accreting material was probably not compositionally significantly different from current additions, although differences in oxidation state and/or volatile content may have occurred over time (e.g., Wanke and Dreibus, 1988). In any case, it is hardly likely that the last 1% of accretion was markedly different in composition from the modern meteorite mix. If we use the Antarctic meteorite data and available analyses of various meteorites to calculate a bulk composition for the putative volatile source material (Table 2), we can then calculate the potential volatile production from an average meteorite mix equivalent to Earth’s mass. Assuming no contribution from the iron meteorite component, the amount of water available (~1.25 × 1026 g) is about two orders of magnitude greater than the water in Earth’s near-surface environment (~1.55 × 1024 g; Walker, 1977), and similarly for C and N species. It is clear that Earth has either lost a major fraction of its degassed volatiles, or that significant volatiles remain hidden in the interior, or that much of the early accretion consisted of volatile-poor material, i.e., there was a significant degree of inhomogeneous accretion, at least in terms of minor phases. Large-scale degassing of volatiles is likely for energetic (thermal) reasons, while volatile loss is consistent with several

TABLE 1. ANTARCTIC METEORITE DATA Mass fraction of Metal fraction in class meteorites (Huss, 1987) Carbonaceo us chondrites 0. 017 0 H chondrites 0.269 0. 122 L chondrites 0.418 0. 064 LL chondrites 0.091 0.028 O t h e r s t o ne s 0 .0 5 4 0 Stony-iron 0.047 0.5 Irons 0 .105 1 Meteorite type

Total metal fraction in mix

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Metal fraction in meteorite mix 0 0. 033 0. 027 0.003 0 0 .0 23 0 .1 0 5 0.19

Primary paper | Methane and ammonia in Earth’s earliest atmosphere and a rich early ocean

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TABLE 2. VOLATILE COMPONENTS IN METEORITES (AFTER ANDERS AND EBIHARA, 1982; PALME AND O’NEILL, 2003; BRADLEY, 2004; SCHAEFER AND FEGLEY, 2007) Type Mass fraction in each type Mass fraction in “average mix” H C N H C N CI (Orguiel) 0.02 0.035 0.003 0.0038 0.0018 7E–05 H 0.0032 0.0012 0.00003 L 0.0034 0.0016 0.00003 Mass fraction in mix as compounds LL 0.006 0.0024 0.00005 H2O CO2 N2 0.034 0.0067 7E–05

aspects of accretional history, especially the highly energetic moon-forming impact at ~50 m.y. after the beginning of accretion (Safronov and Koslovskaya, 1977; Safronov, 1978; Nisbet and Sleep, 2001; Righter, 2007; Schoenbaechler et al., 2010). There are numerous proposed models for accretional history (e.g., Wetherill, 1976; Raymond et al., 2004; Koeberl, 2006; Schoenbaechler et al., 2010), but most allow or postulate a late “tail,” amounting to one to several percent of the total mass of Earth following the moon-forming impact as shown in Figure 1. There is considerable flexibility as to the details of late accretion, but a significant “late veneer” is probable. Table 3 shows the amounts of various volatiles that could be produced from degassing 1% of Antarctic meteorite mix, as well as the volatile inventory of the crust-ocean-atmosphere (Walker, 1977). The table also shows the relative amounts of volatiles. It is probably only a coincidence that the relative amounts are so similar, but this at least suggests that the source of Earth’s volatiles may well be a meteorite mix, and that degassing, if not quantitative, is approximately so, with roughly equal degassing efficiency for all components. Note that the Late Heavy Bombardment amounts to a small fraction of 1% of Earth’s mass (Gomes et al., 2005), though its impact on the Moon’s surface appearance was substantial. There is little doubt that some, maybe considerable, recycling occurs in the upper mantle–crust–ocean–atmosphere system, and

perhaps some net “reinjection” of surface volatiles into the upper (and lower?) mantle occurs (Varekamp et al., 1992). This is possible not only for water and carbon, but may also include nitrogen. Given 4.5 b.y. and a modern thermal structure not drastically different from the early Archean, it seems reasonable to assume that most of this recycling is in some kind of dynamic equilibrium, established long ago. This, of course, means that we should not be surprised to find samples in relatively recently (or ancient) mantle-derived rocks/minerals (e.g., diamonds) reflecting (isotopic) signatures almost certainly due to surface processes such as photosynthesis. The continued production/release of volatiles in upper-mantle–derived magmas demonstrates that upper-mantle conditions remain conducive to extraction of volatiles from the upper mantle, whether they are primitive or recycled (Table 3). There are two points to this discussion especially relevant to the composition of Earth’s earliest atmosphere and surface environment. First, the atmospheric-oceanic components that we see now were almost certainly added late in accretional history, after the most violent impacts (although many of the late impacts probably involved objects of considerable size.) Second, whatever degassing mechanisms were active almost certainly processed material very similar to this meteorite mix, the composition of which controlled the chemistry of the degassing process in important ways. ENERGETICS AND PROCESSES OF DEGASSING

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Degassing of a silicate-iron meteorite mix containing condensed volatile compounds requires energy input. There are three important sources of this energy: impact energy, volcanic heating driven by residual impact heating and radioactive sources, and hydrothermal systems associated with volcanism (but at lower temperatures, and probably more spatially widespread.) The vital

Fraction of Earth accreted

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Accretion time (m.y.) Figure 1. A variety of accretion models for Earth.

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TABLE 3. VOLATILES AVAILABLE FROM METEORITE MIX EQUAL TO 1% OF EARTH’S MASS Potential from Earth’s near-surface degassing inventory 24 24 H2O 1.25 × 10 g 1.55 × 10 g 23 23 CO2 2.7 × 10 g 2.6 × 10 g 21 21 3.1 × 10 g 5.4 × 10 g N2 5.1 6 H2O/CO2 CO2/CO2 1 1 N2/CO2 0.012 0.02

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question in terms of the gases most likely present in the earliest atmosphere is: What geochemical conditions were associated with each of these processes given the composition of the degassing solids and the pressure and temperature of degassing? Impact-derived volatiles may be the quantitatively most important source of the volatiles in Earth’s earliest atmosphere. This question has received some attention, with the usual conclusion being that the high temperatures associated with such a process would lead to evolution of CO2 (and maybe CO; Ahrens et al., 1989; Kasting, 1990, 1993; Fiske et al., 1995; Tyburczy et al., 2001; Ellwood Madden et al., 2006; Sleep and Zahnle, 2001; Zahnle and Sleep, 2002; Zahnle, 2006). This argument is similar to that associated with volcanogenic production of volatiles, with CO2 being the stable phase at magmatic temperatures, provided iron is absent (see following). It is not clear how metallic iron has been treated in these considerations of impact degassing, although the importance of the reducing capacity of accreted iron has been recognized as a potentially significant source of hydrogen (Wanke and Dreibus, 1988; Kasting, 1993). Numerical models of the moon-forming impact show a significant injection of core component into near-Earth space, which then rapidly re-accreted (Benz et al., 1989). We might take this as the most extreme case for late impacts, though following the moon-forming impact, it seems unlikely that any subsequent impacts would approach that magnitude without producing noticeable effects. It is difficult to imagine that all iron in the impactor would quickly disappear from the surface, and conservatively one might guess that much of the metal in an impactor would initially be present in the atmosphere as iron-vapor (for the largest impacts) or rain out as particulate matter, perhaps into a “steamy” ocean. Some number of late impacts may have been large enough to completely vaporize whatever oceans were present from previous degassing, and any large impactor would certainly vaporize considerable water. Objects of 500 km diameter are capable of this degree of heating, assuming the presence of water equivalent to the modern ocean (Zahnle and Sleep, 1996). A few hundred such objects can account for 1% of Earth’s mass. If there were a smaller number of objects with somewhat larger dimensions, they would still have been small compared to the moon-forming impactor. Surface conditions produced by a 500 km impactor have been estimated (Zahnle and Sleep, 1996). Their modeling, as depicted in Figure 2, yields a very hightemperature (rock vapor) atmosphere lasting for a few months, followed by continued presence of temperatures high enough to prevent the presence of liquid water for an additional 1000 yr or so. A new ocean would rain out over the following couple of thousand years. The essential point is that atmospheric temperatures quickly drop to less than ~600 K due to radiation from the high-temperature atmosphere. The shift in atmospheric chemistry shown schematically is in response to atmospheric cooling because atmospheric hydrogen pressure is high following reaction between iron and water, both in the atmosphere immediately following the impact and at the surface following rain out of both iron and water. Obviously, smaller impacts would produce

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even shorter-lived high-temperature atmospheres but would also involve significant hydrogen production (Table 4). Table 4 shows what might be expected as a result of complete vaporization of a 500-km-diameter object, as well as somewhat larger and smaller impactors. Table 4 assumes that approximately one third of the mass of the impactor is metal, which is about double the Antarctic meteorite mix, but approximately equal to the proportion in bulk Earth. It also assumes that metal (iron) entirely reacts with the steam atmosphere (or possibly, after/during rain out, with the ocean). The table also assumes a hydrogen loss rate from the atmosphere during atmospheric cooling of about a million times the present-day H2 loss rate, probably a substantial overestimate. Tian et al. (2005, 2006) calculated a hydrogen escape rate 1000 times the present rate for an early hydrogen-rich atmosphere. For impacts large enough to produce high-temperature conditions that would stabilize CO2, the simultaneous production of hydrogen is sufficient to maintain high levels of hydrogen in the atmosphere well after the atmosphere cools to the range where methane becomes the stable carbon species. In other words, there may be a brief period when CO2 is the dominant carbon species in the atmosphere, but this quickly changes to a situation in which CH4 dominates. The presence of iron in late accreting material is an important control on early atmospheric chemistry. Many of our ideas about the composition of the early atmosphere stem from considerations of the modern production of volcanic gases, an obvious source of volatiles emanating from Earth’s interior (Rubey, 1951, 1955). These analyses suggest that CO2 is (and was?) the dominant carbon species produced, with water the most abundant volatile. Thermodynamic considerations support this idea, as long as metallic iron is absent from the magmagenic zone (Holland, 1984). This condition is supported by the energetics and timing of accretion, which imply core separation early in accretionary history. However, it is unlikely that primordial degassing was as simple as modern mid-ocean ridge volcanism suggests. Degassing almost certainly began upon impact, as discussed already, which may have liberated anywhere from a few percent or less of meteoritic volatiles to nearly 100% for large impacts in which a high degree of impactor vaporization occurred. The meteoritic debris that was not devolatilized on impact would have “rained out” on the surface, probably into a shallow(?) primordial ocean. This debris must of necessity have included some fraction of meteoritic metal, mostly iron, some of which would be buried in sediment to provide a source of reducing capacity when the debris was eventually thermally processed to release volatiles. This thermal processing may have taken place either in a shallow hydrothermal setting (quite probably), or following deeper “subduction” (not necessarily mirroring modern subduction in style) during high-temperature magmagenesis. Each of these processes must be examined in detail as to the likely geochemical conditions accompanying them. It is important to keep in mind that late accretion will have occurred and continued long after the surface had solidified, and

Primary paper | Methane and ammonia in Earth’s earliest atmosphere and a rich early ocean oceans (if any) formed following the moon-forming impact, within a few million years (Nisbet and Sleep, 2001; Zahnle, 2006). There is no conceivable way that the fraction of the late accreting matter that was not vaporized on impact could have been selectively stripped of its metallic component before being hydrothermally or volcanically processed. Sedimentation of solid meteorite debris (probably into an essentially global ocean) would necessarily produce a mixture of meteoritic silicate, metal, and some carbonaceous material as a “sediment” veneer. Indeed, recent thermochemical calculations suggest that even simple

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H chondrite material would give rise to reduced carbon (and nitrogen) species at temperatures near magmagenic conditions, as shown in Figure 3 (Schaefer and Fegley, 2007; Zahnle et al., 2010). Similar results have been obtained for L and LL chondrites as well, thus encompassing the vast majority of meteoritic material (Schaefer and Fegley, 2010). Even at higher, clearly magmatic, temperatures, the unavoidable presence of iron during this initial processing would lead to production of largely reduced gases, including CO and H2. On the other hand, it seems probable that much, perhaps most, of this initial degassing occurred

Figure 2. Conditions following a 500-km-diameter impact (modified from Zahnle and Sleep, 2006).

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Impactor diameter (km)

1000 500 50

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24

1.6 × 10 23 2 × 10 20 2 × 10

TABLE 4. POST–LARGE-IMPACT ATMOSPHERE SCENARIOS Possible H2 H2 pressure Peak temperature Time until produced (bar) (K) temperature drops (g) to 650 K (yr) 22 1.6 × 10 5 2500+ 2500+ 21 2 × 10 0.6 2000+ ≤1000 18 2 × 10 0.0006 100s (?) 0–10

at shallower depths under hydrothermal conditions, which were almost certainly much more widespread. LOWER-TEMPERATURE DEGASSING Hydrothermal processing of meteoritic material may well have been the most important degassing mechanism, at least for meteoritic material that survived impact with its volatile inventory more or less intact. The initial production of water, even from a small part of the 1% meteorite mix (and some water might have been residual from earlier degassing), would give rise to a global (if shallow) ocean, and volcanic activity would undoubtedly have been accompanied by much more widespread hydrothermal activity, exposing newly accreted material to hydrothermal processing before subduction-like processes carried it into deeper magmagenic zones. Such hydrothermal processing would obviously be at lower temperatures much more conducive to the production of reduced carbon and nitrogen species, as is evident from thermochemical considerations (Seewald et al., 1990, 1994), experimental results (Seewald et al., 1990, 1994, 2006; Seewald, 2001), and field observations of modern hydrothermal systems (Seewald et al., 1994; Cruse and Seewald, 2006). Such hydrothermal processing is less dependent upon the presence of free iron, but some free iron is likely to have been present. It is

H2 loss while temperature drops (g) 17

10 + 16 4 × 10 14 4 × 10

difficult to see how freshly accreted meteoritic material would lack free iron. Figure 4 schematically shows the environments where late veneer material would be degassed. The precise nature of the volcanic activity, thickness of surface veneer, and “crustal” foundering is not important. In particular, it is not essential that any type of plate-tectonic processes were active. It is only relevant that volcanic activity would imply associated shallow hydrothermal systems, and that foundering material entering high-temperature (magmagenic) zones would initially contain reduced iron. During the first pass of surface material through the magmagenic zone, excess iron would be removed (while affecting gas chemistry) and sink to the core, leaving the upper mantle iron free, but not before affecting the redox conditions during magmagenesis. Note that general upper-mantle redox conditions can certainly reflect an iron-free state. ATMOSPHERIC PROCESSING, RECYCLING OF VOLATILES, AND MAINTENANCE OF A REDUCING ATMOSPHERE If the early degassing processes outlined here produced methane (and ammonia) as significant atmospheric constituents, it is clear that these reduced gases had a limited lifetime

Figure 3. Carbon and nitrogen species in equilibrium with chondritic solids in near-surface environments (after Schaefer and Fegley, 2007). The shaded regions roughly approximate conditions expected during degassing, ranging from magmatic to hydrothermal.

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Primary paper | Methane and ammonia in Earth’s earliest atmosphere and a rich early ocean

H2 to space N2, H2O, H2, CO2, CH4, NH3 Photolytically condensed organics

CH4, NH3, CO2

H2, CO, N2, (CH4, NH3?)

Dissolved Organics

Seawater

Shallo w Hyd CH4, rotherm al acti vity

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Fe-rich Magmagenesis Fe to core

Figure 4. Primordial degassing of late veneer.

due to photochemical (and atmospheric electric discharge) destruction (Kasting et al., 1983; Levine, 1985; Zahnle, 1986; Kasting and Brown, 1998). Indeed, there are plenty of experimental and theoretical results proving the rapid conversion of the reduced C and N components of such an atmosphere into more complex molecules and ultimately a variety of condensed phases, most of which end up in the ocean. One of the principal consequences of this proposed model for the early atmosphere is copious production of prebiotic molecules (e.g., Miller, 1953, 1955), and in quantities highly favorable to the emergence of life, considering the amounts of carbon and nitrogen involved. The difficulties of doing this starting from a CO2 atmosphere are well known (but see discussion of Tian et al., 2005, 2006). Assuming the atmosphere is stripped of most of the reduced carbon and nitrogen compounds, the oceans end up as a concentrated “soup,” probably with an immiscible organic phase floating on the surface, and/or sinking to the bottom, carried by inorganic sediment particles. Processing of this organic-rich reservoir through hydrothermal systems (Seewald et al., 1990, 1994, 2006; Seewald, 2001; Cruse and Seewald, 2006) either as infiltrating organic-rich water or by heating of organic-rich sediment, would recycle the organics back into the atmosphere as methane and ammonia. This would serve to maintain modest but significant atmospheric levels of CH4 and NH3 in spite of photochemical destruction. The constant recycling of organics may well have been conducive to continuous prebiochemical “experimentation,” favoring the emergence of life. In fact, this hydrothermal reprocessing is not very different from the initial hydrothermal release of methane and ammonia from the primary meteorite mix suggested earlier.

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The hydrothermal regeneration of methane and ammonia can be approximately quantified based on modern observations and experiments, and by making reasonable estimates of hydrothermal activity. Hydrothermal processing of organic-rich sediments (~2.5% organic C and 0.22% N) for periods of ~100 h at 375 °C releases ~5% of the carbon as methane and ~50% of the nitrogen as ammonia (Seewald et al., 1994). Longer reaction times result in higher release fractions. If we assume that condensation reactions in the early reducing atmosphere led to rain out of most of the atmospheric carbon and even a few percent of the atmospheric nitrogen into the ocean, the concentrations of organic compounds in solution (and presumably in seafloor sediments) were likely a minimum of a few percent. Circulating this solution through mid-ocean ridge hydrothermal systems is guaranteed to yield large amounts of methane and ammonia. Modern seawater hydrothermal circulation has been estimated at 2.5 × 1012 m3/yr (Sleep and Zahnle, 2001). It is likely that rates on young Earth were even higher, but even using the modern value, the rate of production of both methane and ammonia by hydrothermal activity would have exceeded modern biogenic rates (Kasting, 2005) by at least an order of magnitude, i.e., almost certainly high enough to maintain atmospheric concentrations at significant levels (Shaw, 2008a). I have described this earliest reducing atmosphere as one dominated by methane as the carbon species, which may or may not have been the actual situation. The initial state is one of high hydrogen pressure, which leads to methane (and ammonia) production under equilibrium conditions following rapid atmospheric cooling. On the other hand, it is not necessary to posit an intermediate state in which methane is the main carbon species. Tian et al. (2005, 2006) have described the importance of high H/C ratios in production of prebiotic compounds in an atmosphere where carbon dioxide is the main carbon species, but where the hydrogen fraction is high. In their case, they posit hydrogen production from the mantle, which, with a low hydrogen loss rate from the upper atmosphere, produces high H/C. In the present case, the hydrogen would be produced copiously and promptly in the postimpact environment, and the loss rate is not critical. The result, in terms of production of prebiotic compounds and an organic-compound–rich early ocean, is much the same as if the atmosphere went through a short-lived intermediate phase with high methane (and ammonia). Over the longer term, the hydrothermal processing of this organic reservoir does, in fact, result in production and maintenance of atmospheric methane and ammonia. ATMOSPHERIC AND SURFACE VOLATILE RESERVOIR EVOLUTION This raises the problem of how to evolve from an Earth with virtually all of the carbon in organic compounds to the modern situation in which most of the carbon is oxidized, and present mainly as carbonate rocks. There is a simple mechanism to explain this. Deposition of organic-rich sediment leads to subduction of carbon (compounds), with some fraction of the organic carbon carried

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into the magmagenic zone of the mantle. Because accreted iron will have been removed from the upper mantle (quite early, during the initial degassing), subsequent magmagenesis incorporating organic-rich sediment will produce magmatic gases dominated by oxidized carbon (CO2), with an admixture of water and hydrogen, consistent with modern systems. Over geologic time, accompanied by loss of the hydrogen to space (much more effective with a nonoxidizing atmosphere), the magmagenic CO2 is gradually added to the surface environment and eventually becomes incorporated in carbonate sediments. (Whether/how H2 loss over geologic time can lead to eventual surface oxidation is discussed by Zahnle and Catling [this volume].) Rates of volcanism are entirely consistent with producing the requisite amount of CO2 over a billion years or so (Walker, 1977; Hayes and Waldbauer, 2006; Williams et al., 1992; Shaw, 2008a). Figure 5 depicts volatile cycling after degassing is essentially complete. As in Figure 4, this is a schematic view carrying no implications of tectonic details. Once the late accreted material has been magmatically processed once, it has lost all free iron to the deep interior. Subsequent magmagenesis produces carbon dioxide as the main carbon species. This also means that the oxidation state of the upper mantle soon after final accretion/ degassing at ca. 4.4–4.45 Ga would have been be controlled by iron-free magmatism, consistent with observations on old zircons (Trail et al., 2011). However, the large surface reservoir of reduced carbon would result in long-term production and cycling of reduced gases into the atmosphere, maintaining its overall reducing state, perhaps until the late Archean.

H2 to space

N2, H2O, H2, CO2, CH4, NH3 Photolytically condensed organics

H2, CO, N2, (CH4, NH3?) CH4, NH3, CO2

Deposition of organic sediments

Dissolved Organics

Seawater

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H2, CO CO2 Fe-free Magmagenesis

Fe-free Magmagenesis

These mechanisms of degassing and early atmospheric evolution are clearly consistent with ideas of (re)cycling of carbon species into/through the mantle, as seems necessary in order to provide the source material for upper-mantle carbon minerals (e.g., diamonds) that have isotopic characteristics indicative of surface (photosynthetic) processes (Walter et al., 2011). It is most likely that such recycling reached a dynamic equilibrium between the mantle and surface very early in Earth’s history, given the heat available to drive convective processes responsible for the exchange. The implications for this alternative view of early degassing are profound (Shaw, 2008a). In addition to providing a surface environment highly favorable to prebiotic synthesis and the ultimate emergence of life, recycling of methane and ammonia into the early atmosphere would provide a means of offsetting the faint young sun. The limited amount of carbon dioxide available explains the absence of abundant carbonate rocks in metasedimentary units until later in the Archean, and the subsequent gradual evolution of the carbon reservoirs. Inclusion of continuous addition of mantle-derived CO2 provides a straightforward explanation for the evolution of carbon isotope systematics. In addition, applying a similar analysis to Mars could provide a resolution to the problem of the climate transition on Mars from warm and wet to cold and dry (Shaw, 2008a, 2008b). OTHER CONSEQUENCES Although it is difficult to formulate a “clinching” observation that would support this model for early degassing and surface chemistry, that is also a problem for the conventional model of CO2 degassing and an early CO2-rich atmosphere. Indeed, much of the available geologic evidence does not “prove” the conventional model but instead (as seen here) appears to be quite inconsistent with it. At the very least, the proposed highly reduced model is not inconsistent with the available geologic data. A potential “test” that might distinguish between these models may be found by looking at Mars. Application of the proposed model to Mars, and what is known of its history, suggests that the frozen ocean likely beneath the Northern Plains may well be rich in organic compounds left over from the warm-wet era. This would contrast with the CO2-rich ocean expected if Mars’ early degassing was in the form of CO2. The limited amount of carbonate found on Mars surface already suggests a problem with an early CO2-rich atmosphere. If the CO2 is not found in the frozen ocean, but organic compounds are abundant, the case for an early reduced surface environment on Mars would be hard to reject, and by analogy (perhaps dangerous when considering planets), the same would apply to early Earth. REFERENCES CITED

Figure 5. Carbon and nitrogen cycling on early Earth after primordial degassing. These conditions applied after ca. 4.5 Ga. The Late Heavy Bombardment at ca. 3.9 Ga may have contributed a small amount in the mode of Figure 4.

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Ahrens, T.J., O’Keefe, J.D., and Lange, M.A., 1989, Formation of atmospheres during accretion of the terrestrial planets, in Atreya, S.K., Pollack, J.B., and Matthews, M.S., eds., Origin and Evolution of Planetary and Satellite Atmospheres: Tucson, Arizona, University of Arizona Press, p. 328–385.

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The Geological Society of America Special Paper 504 2014

How low can you go? Maximum constraints on hydrogen concentrations prior to the Great Oxidation Event Shawn Domagal-Goldman* Research Space Scientist, Planetary Environments Laboratory, National Aeronautics and Space Administration Goddard Space Flight Center, 8800 Greenbelt Road, Greenbelt, Maryland 20771, USA

ABSTRACT Shaw postulates that Earth’s early atmosphere was rich in reducing gases such as hydrogen, brought to Earth via impact events. This commentary seeks to place constraints on this idea through a very brief review of existing geological and geochemical upper limits on the reducing power of Earth’s atmosphere prior to the rise of oxygen. While these constraints place tight limits on this idea for rocks younger than 3.8 Ga, few constraints exist prior to that time, due to a paucity of rocks of that age. The time prior to these constraints is also a time frame for which the proposal is most plausible, and for which it carries the greatest potential to explain other mysteries. Given this potential, several tests are suggested for the H2-rich early Earth hypothesis.

INTRODUCTION

and research have been focused. Past research, as evidenced by a later paper in this series by Ohmoto et al. (this volume), has debated the limits on the oxidizing extent of early Earth. However, significantly less consideration has been given to limitations on the reducing potential of early Earth’s atmosphere. This is where Shaw’s hypothesis comes into play.

Shaw (this volume) contends the pre-oxygen Earth was more reducing than previously thought. That suggestion can be considered in the context of past research on the redox state of the atmosphere using geochemical measurements and atmospheric models. The call for an H2-rich atmosphere is intriguing because it contains some explanatory power; however, we can limit this idea in scope, given knowledge obtained through past analyses of the Earth’s geological record. The “sweet spot” for this hypothesis is for the earliest portion of the planet’s history. This is a time when the rock record cannot exclude this as a possibility, when the impact-driven mechanism for this phenomenon was most prevalent, and when this idea has the greatest potential to solve perplexing riddles. This idea is intriguing because it probes the opposite end of the oxidizing–reducing scale on which much of the past debate

DISCUSSION To evaluate the possibility of a highly reducing atmosphere on early Earth, we consider geological constraints that place an upper limit on the reducing power of the atmosphere. These include data used to argue for a “whiff” of oxygen in the midArchean (e.g., Anbar et al., 2007). These conclusions—based on the concentrations of elements for which mobility is sensitive to the redox state of the atmosphere-ocean system—place periodic, yet strict, limits on how reducing the atmosphere could have been

*[email protected] Domagal-Goldman, S., 2014, How low can you go? Maximum constraints on hydrogen concentrations prior to the Great Oxidation Event, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 11–13, doi:10.1130/2014.2504(02). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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between ca. 2.8 and ca. 2.5 Ga. Unless some mechanism could have rapidly oxidized the atmosphere, and then rapidly returned it to a highly reducing state, we must therefore rule out an H2-rich atmosphere for this time period. Further constraints come from the presence of massindependent sulfur isotope fractionation (S-MIF). Highly reducing conditions can upset the exit channel balance between sulfates and sulfides, a balance that is needed to transfer atmospherically derived isotopic features to the rock record. Under extremely reducing conditions, all S in the atmosphere will leave in the form of sulfides, and this will eliminate any S-MIF created before it is deposited (Domagal-Goldman, et al., 2008). Extremely reducing conditions also would have caused the CH4/CO2 ratio in the atmosphere to approach unity. Had this occurred, a haze would have formed, blocking the ultraviolet (UV) wavelengths responsible for S-MIF production (Zerkle et al., 2012). Given the presence of S-MIF from 3.8 Ga through the rise of oxygen (Farquhar and Wing, 2003), we can eliminate the possibility of an H2-rich atmosphere for any time after 3.8 Ga. The potential for haze formation places another constraint on CH4 concentrations for a climatic reason: Hazes cause significant antigreenhouse effects. These effects will be much greater than the greenhouse effects from CH4 and other organic gases (Pavlov et al., 2001). For CH4/CO2 ratios significantly above 0.1, global glaciations would have been triggered. However, there is no evidence for global glaciations prior to 2.4 Ga, nor for any glaciations prior to ca. 2.8 Ga. This corroborates the limit stated above: CH4 concentrations must have been less than CO2 concentrations since at least 3.8 Ga. Such conditions are not consistent with an H2-rich atmosphere. Prior to 3.8 Ga, there is not much of a rock record, so these geological constraints do not exist. However, there are minerals that have been dated to be as old as 4.4 Ga, and analyses of these suggest that the redox state of the mantle has been relatively consistent throughout Earth’s history (Trail et al., 2011). In terms of Shaw’s hypothesis for an H2-rich atmosphere, this means that had such an atmosphere been in place prior to 3.8 Ga, it did not have a significant effect on the redox state of the mantle. Thus, the (admittedly sparse) geological data prior to 3.8 Ga cannot exclude the possibility of an H2-rich atmosphere—they can only limit the degree to which the mantle would have been affected by such an atmosphere. The time period prior to 3.8 Ga is also the time when Shaw’s hypothesis is most appropriately applied. The mechanism proposed to drive the atmosphere to such a state is the delivery of highly reducing extraterrestrial materials. The rate of delivery of this material would have significantly decreased with time, with perhaps a spike in delivery rates associated with the still controversial “Late Heavy Bombardment” at 3.8 Ga. Regardless, prior to 3.8 Ga, significantly more of this reducing material would have been delivered to Earth. Further, if the pre–3.8 Ga Earth had been more reduced, it would have allowed for greater abiotic production of compounds necessary for the origins of life, and for the buildup of extremely efficient greenhouse gases that could

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solve the “faint young sun paradox” during the time frame for which the Sun was at its faintest. If escape of H through the top of the atmosphere was limited by energy deposition to the atmosphere (as opposed to limited by diffusion of H into the upper atmosphere, as it is on modernday Earth), then escape of H would have been much slower. If this was case, then the “oxidant source” provided by escape of H would have been lower, and the atmosphere would have been much more reducing. Previous models of escape show that this could produce a dramatic effect (Tian et al., 2005). However, such models have been criticized for not being complete enough (Catling, 2006), and further study is warranted before this process can be considered a potential mechanism for maintaining a reducing atmosphere on early Earth. Future research could test this “H2-rich early Earth” hypothesis. First, one must determine if models of the impacts themselves that delivered the reducing material would have allowed this material to be partitioned into the atmosphere and crust without significant effects on the redox state of the mantle. Then, it must be demonstrated that an atmosphere such as this would have been capable of evolving into the considerably more H2-poor atmosphere that was in place since at least 3.8 Ga. Finally, models that reproduce this evolution should also be able to predict the resulting changes to noble gas and isotopic reservoirs, ultimately leading to “ground-testing” of the hypothesis with geochemical measurements of rocks deposited after this evolution was complete. Alternatively, should we be fortunate enough to uncover (meta-)sedimentary rocks older than 3.8 Ga, or develop the capability to find them on the Moon, we will be able to analyze them for some of the same geochemical constraints, such as the presence of S-MIF, that limit H2 concentrations after this time. ACNOWLEDGMENTS This work was performed as part of the NASA Astrobiology Institute’s Virtual Planetary Laboratory, supported by the National Aeronautics and Space Administration through the NASA Astrobiology Institute under solicitation No. NNH05ZDA001C; it was also supported by the Oak Ridge Associated Universities NASA Postdoctoral Program. REFERENCES CITED Anbar, A.D., Duan, Y., Lyons, T.W., Arnold, G.L., Kendall, B., Creaser, R.A., Kaufman, A.J., Gordon, G.W., Scott, C., Garvin, J., and Buick, R., 2007, A whiff of oxygen before the Great Oxidation Event?: Science, v. 317, no. 5846, p. 1903–1906, doi:10.1126/science.1140325. Catling, D.C., 2006, Comment on “A hydrogen-rich early Earth atmosphere”: Science, v. 311, no. 5757, p. 38, doi:10.1126/science.1117827. Domagal-Goldman, S.D., Kasting, J.F., Johnston, D.T., and Farquhar, J., 2008, Organic haze, glaciations and multiple sulfur isotopes in the mid-Archean Era: Earth and Planetary Science Letters, v. 269, p. 29–40, doi:10.1016/j .epsl.2008.01.040. Farquhar, J., and Wing, B.A., 2003, Multiple sulfur isotopes and the evolution of the atmosphere: Earth and Planetary Science Letters, v. 213, p. 1–13, doi:10.1016/S0012-821X(03)00296-6.

Commentary | Maximum constraints on hydrogen concentrations prior to the Great Oxidation Event Ohmoto, H., Watanabe, Y., Lasaga, A.C., Naraoka, H., Johnson, I., Brainard, J., and Chorney, A., 2014, this volume, Oxygen, iron, and sulfur geochemical cycles on early Earth: Paradigms and contradictions, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, doi:10.1130/2014.2504(09). Pavlov, A.A., Brown, L.L., and Kasting, J.F., 2001, UV shielding of NH3 and O2 by organic hazes in the Archean atmosphere: Journal of Geophysical Research–Planets, v. 106, p. 23,267–23,287, doi:10.1029/2000JE001448. Shaw, G.H., 2014, this volume, Evidence and arguments for methane and ammonia in Earth’s earliest atmosphere and an organic compound–rich early ocean, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, doi:10.1130/2014.2504(01).

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Tian, F., Toon, O.B., Pavlov, A.A., and De Sterck, H., 2005, A hydrogen-rich early Earth atmosphere: Science, v. 308, p. 1014–1017, doi:10.1126/ science.1106983. Trail, D., Watson, E.B., and Tailby, N.D., 2011, The oxidation state of Hadean magmas and implications for early Earth’s atmosphere: Nature, v. 480, p. 79–82, doi:10.1038/nature10655. Zerkle, A.L., Claire, M.W., Domagal-Goldman, S.D., Farquhar, J., and Poulton, S.W., 2012, A bistable organic-rich atmosphere on the Neoarchaean Earth: Nature Geoscience, v. 5, p. 359–363, doi:10.1038/ngeo1425.

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The Geological Society of America Special Paper 504 2014

Discussion of “Evidence and arguments for methane and ammonia in Earth’s earliest atmosphere and an organic compound–rich early ocean” (Shaw) George H. Shaw: There’s one thing I’d like to point out: I spent a lot of time emphasizing the methane and the ammonia. What I probably should make clear, I should have made clear, is that I don’t deny the possibility and the likelihood of existence of significant carbon dioxide in the atmosphere. In fact, there must have been carbon dioxide in the atmosphere because once the upper mantle loses its metallic material, once the surface crustal material loses its metallic material which eventually obviously ends up in the core, the conditions are much more like the modern Earth. And on the modern Earth magmatic processes produce carbon dioxide and water as a result of magma genesis. So this must have started happening very, very early. So the earliest metal-free magmatism must have appeared on the scene almost immediately and would have produced carbon dioxide. What this means is that we could easily imagine an atmosphere with modern levels of carbon dioxide, maybe higher, but modern is maybe good enough: 300 parts per million (and rising). And this means that if it’s, say, 300 parts per million and if your upper limit for methane/CO2 is one, you could still have an atmosphere with a few hundred parts per million methane, along with a few hundred parts per million CO2. And then some of the objections, maybe they don’t go away, but they’re much less serious constraints. So I think the early atmosphere may have been far more complicated than we’ve been giving it credit for, and my hope is that we can get a little more exploration of some of these possibilities which up until recently, and who knows, maybe into the future, have been considered absurd.

mantle dynamics, which [were] stressed in yesterday’s session. What we learned also from [the] paper by Kump and Barley1 is before 2.5 billion years ago dominant volcanism was submarine, and the degassing occurred by submarine hydrothermal fluid at temperature of less than 450 degrees [Centigrade], and under high water pressure. In those conditions gases in equilibrium with basalt could be methane, which absorbed the CO2; SO2 was nothing. H2S could be. But after 2.5 billion years dominant volcanism changed to subaerial volcanisms, where high temperature volcanic gas could come out at slow pressure. That’s the time SO2 becomes important. So [my] point is yes, I think I agree with George, how the atmosphere could be methane-rich, methane as well as CO2, and also ammonia as well as nitrogen. So I agree with you, Miller-Urey–type reaction could have occurred in the early atmosphere. It could be more reducing than many people have thought, based on the assumption volcanic gas is [the] same through the geologic time. But the important thing is you have to think about change in mantle dynamics where the gassing occurred. Shaw: Yes, I agree that a lot of early volcanic activity was probably submarine, and those conditions and especially hydrothermal areas associated with them probably do favor reducing gases; that’s certainly part of my point. I’d also, however, like to point out that it’s very likely that some kind of small, perhaps arc-like geologic units, if you will, were probably formed fairly early on. We know that the oldest continental crust might have been as much as 4.2 billion years old. There was probably some subaerial

Robert Pepin: This topic or combination of topics is now open for discussion. Excellent. The lineup has started. Yes, go ahead.

1 L.R. Kump and M.E. Barley, 2007, Increased subaerial volcanism and the rise of atmospheric oxygen 2.5 billion years ago: Nature, v. 448, p. 1033– 1036, doi:10.1038/nature06058.

Hiroshi Ohmoto: In thinking about [the] volcanic gas compositions idea, it is very important to think about the changing

Shaw, G.H., Domagal-Goldman, S., et al., 2014, Discussion of “Evidence and arguments for methane and ammonia in Earth’s earliest atmosphere and an organic compound–rich early ocean” (Shaw), in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 15–17, doi:10.1130/2014.2504(03). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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Shaw et al. volcanism, even very early on, although it was probably not the dominant mode of volcanic activity. But there’s no reason to think that there might not have been carbon dioxide produced under conditions very early in the Earth’s history as well.

James F. Kasting: I have a talk coming up so I’ll be brief. I agree with Hiroshi that you need to take into account the mode of degassing early on. I’d argue that it’s mostly impact degassing, which you mentioned in your talk, George, because most of the material coming in during the late stages of accretion and then during the heavy bombardment comes in in big impactors and it gets degassed directly into the atmosphere. So then it’s not really subaerial or submarine volcanism, it’s impact degassing, and that’s a different issue, which will be addressed here; I think Laura Schaefer’s going to talk about that. I wanted to raise one other point on your talk, George; you mentioned that you looked at modern-day meteorites and said these are strong—they contain iron and so this is strong evidence against inhomogeneous accretion of the Earth. I was at the Geochemical Society, the Goldschmidt meeting a couple months ago over in France, and Alessandro Morbidelli gave a talk in which he said there’s now strong evidence for inhomogeneous accretion of the Earth. And that’s based on these dynamical models of accretion that Morbidelli and others have been doing. When you put the planet together, the Earth accretes first material from around 1 AU and then more volatile-rich material and more oxidized material from farther out in the solar system. So I think it’s important to take these theoretical but very believable, to me at least, constraints into account when you’re thinking about the Hadean. Shaw: That’s a good point. There’s a lot of discussion about homogeneity, homogeneous versus inhomogeneous accretion. My point is the stuff we see coming in now looks like bulk Earth. And there may have been inhomogeneous accretion based on distance from the sun early in Earth’s history, but it doesn’t seem at all unlikely to me that the last veneer, 1 percent, 2 percent—it doesn’t seem likely to me that it would look enormously different from what’s coming in now. Obviously the rates of influx now are much, much lower, but you’d have to come up with some kind of mechanism to make a major change between the start of that last 1 or 2 percent and the present time. And I’m not sure that there’s anything in the modeling that would call for that kind of difference. Clearly there could be a major difference between the earliest accretion, the first 80 percent and what happened after that; that’s certainly possible. Kasting: Okay, well I think we’re in agreement on that. I think the bulk of the Earth could—probably did accrete

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from these rather dry and reduced materials that came from the 1 AU region. Aaron Cavosie: First [I’d] like to thank the speakers for interesting presentations. I’m not an atmospheric chemist so I think I’ll not address that issue and leave that to those who operate at those elevations. I would like to address—I’m going to pull out, excise surgically a single statement that was made that I have a difference of opinion in, and I think it’s relevant to this discussion. And that was the statement that there’s no evidence for severe weathering in the early Archean. In the last few years, some evidence is emerging for severe weathering, and it may not be readily apparent by the titles of the papers, perhaps. But this beckons to be a detrital zircon record, particularly the very early zircons that have been found in Western Australia, and it comes from two lines of evidence. One is well known by now but one is relatively new. The first, and I won’t take time to go into detail, is the oxygen isotope record preserved in the zircons. To summarize that: The zircons have a memory of the alteration of the parent material that melted to crystalize the zircon, the particular oxygen isotope ratio that’s indicative of alteration. There’s been a lot published on this but that’s the first point. And that record goes back to about 4.2 [billion]. So we have evidence, we think, for profound weathering at 4.2 billion years. But what I’d really like to bring to the attention of this room full of people, is a different system, and that is the lithium isotope ratios in zircon. This is a relatively new isotopic system that’s been applied to zircon. The first paper [to come] out on this was by Ushikubo et al. in Earth and Planetary Science Letters in 20082. And what Ushikubo et al. demonstrated was a profound fractionation of the lithium isotope system recorded in zircon and also the demonstration—there’s arguments presented on why that’s a reliable record, why it’s not demonstrating alteration after the fact. But the implication of the lithium data is even more profound weathering than the oxygen implicates. So I encourage people that are looking for evidence of profound alteration to check into those two systems. The punch line is the detrital zircons themselves may simply be that record of alteration as the rest of the rocks that incorporated them have simply turned to clays and washed away with time. So I would suggest or propose that there is a record of intensive alteration that’s recorded in a robust mineral as far back as 4.2 to 4.3 billion years ago. Can’t take it back further than that right now.

2 T. Ushikubo, N.T. Kita, A.J. Cavosie, S.A. Wilde, R.L. Rudnick, and J.W. Valley, 2008, Lithium in Jack Hills zircons: Evidence for extensive weathering of Earth’s earliest crust: Earth and Planetary Science Letters, v. 272, no. 3–4, p. 666–676, doi:10.1016/j.epsl.2008.05.032.

Discussion | Evidence and arguments for methane and ammonia in Earth’s earliest atmosphere Shaw: Thank you. I appreciate that. These are results that I certainly have not been aware of and I look forward to talking to you about that and getting some more information. It is certainly possible that there were intervals of intense weathering and relatively high CO2 in the atmosphere. That’s not an impossibility, even with a reducing atmosphere. In fact, in a sense it makes significant methane and ammonia concentrations more acceptable—if there’s higher CO2 you can also have higher methane and ammonia. There are these other problems such as where’s all the carbonate that should have been produced by this weathering. But these are old rocks, and sometimes they’re not even rocks. So there’s a lot of, shall we say, “wiggle room” in the first billion years of Earth’s history. But that’s an interesting and valuable point, I think. Pepin: Kevin? Kevin Zahnle: Yeah, I had actually a comment to Jim’s comment, which was that a dynamical theories of late accretion are not actually evidence; they’re more like fashions. You have to think in terms of what’s in favor today. When you talk about the Nice model, it’s whatever the Nice modelers say they’re thinking this year. That model has been evolving for six years now and now has no common element with the original Nice model but it’s still the Nice model. And that has to be borne in mind when you’re dealing with these dynamicists. They have many more theoretical tools—well you have a theoretical toolbox yourself! You know you can do this; you can make things happen. And I think that’s kind of where they are right now. To be sure but we’re not going to—I don’t really like calling theory “evidence.” I think of like lithium on a zircon as evidence. That’s kind of more of what I had in mind. I was going to say something about the CO2, but I’ve already forgotten it. Martin J. Van Kranendonk: Just getting back to the weathering aspect and the lack of, or the apparent lack of, carbonates in early Archean rocks, one factor has also maybe been overlooked and maybe not published significantly, but many of the best-preserved, oldest volcanic sequences are extensively altered by carbonate. They have up to 8 weight percent carbonate as an average. And actually it’s quite a significant sink for carbon ions in the ocean, carbonate drawdown

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in the volcanic rocks, and then you get these hydrothermal systems that actually exchange and bring silica back out of the oceans. Actually, one of the points that George raised was really interesting to me is that the oldest hydrothermal system in the world that’s preserved at North Pole has the silica veinlets with extensive evidence of methane in them. Uachiro Ueno3 has discovered these and done isotope anomalies investigations on them and there’s significant methane, both trapped as fluid inclusions and from evidence of the delta Protean Sea. So those are things to consider. Now actually I found what you were talking about, and maybe we’re getting to this idea that there maybe is a combination of high methane and high CO2; there’s certainly, I think, evidence in the rock record for high CO2 but not in the more modern analog of having thick deposits of carbonate rocks; it’s all actually sunk down into the substrate. So that’s just a comment to put into (the discussion?). Robert Pepin: Thank you. Shawn, you actually had a comment to make. Shawn Domagal-Goldman: Well, one thing that I think we could really use as a community is some constraint on the total atmospheric pressure at the time. Because it is tempting to rely on just sticking more things in the atmosphere to explain some of the things that we’re seeing in the rock record, whether that’s getting more greenhouse gases in there or more CO2 to promote weathering. And I think we have a freedom to do that right now because we have zero constraints on how much we can put into the atmosphere. I say this because I’m aware that there are some people that are working on this problem, specifically Sanjoy Som, who just finished his PhD with Roger Buick at the University of Washington, has been trying to get two separate constraints on paleopressure. The preliminary results suggest that it does limit things to a bar or few bars of total atmospheric pressure at sea level. Now that’s not published yet; they’re still trying to work out the kinks in their data analysis but that’s—I think those types of studies are going to be really useful as we go forward.4 And again, those aren’t models, right, these are guys that are in the rocks, looking at things that they can find in the rocks to constrain what was going on at the time.

3 U. Uno, Y. Hideyoshi, M. Shigenori, and Y. Isozaki, 2004, Carbon isotopes and petrography of kerogens in ~3.5-Ga hydrothermal silica dikes in the North Pole area, Western Australia: Geochimica et Cosmochimica Acta, v. 68, no. 3, p. 573–589. 4 Sanjoy Som has since published some of his results in S.M. Som, D.C. Catling, J.P. Harnmeijer, P.M. Polivka, and R. Buick, 2012, Air density 2.7 billion years ago limited to less than twice modern levels by fossil raindrop imprints: Nature, v. 484, p. 359–362, doi:10.1038/nature10890.

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The Geological Society of America Special Paper 504 2014

Atmospheric composition of Hadean–early Archean Earth: The importance of CO James F. Kasting* Department of Geosciences, Pennsylvania State University, University Park, Pennsylvania 16802, USA

ABSTRACT The mantle was probably oxidized early, during and shortly after accretion, and so the early atmosphere of Earth was likely dominated by CO2 and N2, not by CH4 and NH3. CO2 declined from multibar levels during the early Hadean to perhaps a few tenths of a bar by the mid- to late Archean. Published geochemical constraints on Archean CO2 concentrations from paleosols are highly uncertain, and those from banded iron formations are probably invalid. Thus, CO2 could have been sufficiently abundant during the Archean to have provided most of the greenhouse warming needed to offset the faint young Sun. H2 might have augmented this warming prior to the origin of methanogenic bacteria. Atmospheric CH4 concentrations increased from at most tens of parts per million (ppm) on prebiotic Earth to hundreds of parts per million once methanogens evolved. CO was an important trace gas on prebiotic Earth because of its high free energy and its ability to catalyze key reactions involved in prebiotic synthesis. Large impacts could have made the atmosphere transiently CO rich, and this may have played a role in the origin of life and in fueling early biological metabolisms.

tinue until sometime in the Archean, prior to the rise of O2 at 2.45 Ga (Farquhar et al., 2000). I will concentrate on the prebiotic atmosphere and the Hadean–early Archean time period, as this interval is the main focus of this volume. Some discussion of the later Archean is included, as the constraints from climate and geochemistry become much better during this time.

INTRODUCTION Earth’s atmospheric composition during the Hadean and Archean is still somewhat of a mystery because the data that bear on it are sparse, indirect, and occasionally conflicting. For this reason, it is necessary to rely on theory to fill in gaps in the data. The following is a brief attempt to paint a coherent picture of this time period. The emphasis here is on atmospheric composition, although other aspects of the Earth system, climate in particular, have strong connections with the topic under discussion. As other authors have done (Zahnle, 2006), I will begin my discussion following the Moon-forming impact, sometime before 4.44 Ga (the age of the oldest Moon rock), and I will con-

WAS THE EARLY ATMOSPHERE STRONGLY OR WEAKLY REDUCED? Zahnle (2006) has provided a well-reasoned, qualitative description of the way in which the atmosphere may have evolved in the several hundred million years following the

*[email protected] Kasting, J.F., 2014, Atmospheric composition of Hadean–early Archean Earth: The importance of CO, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 19–28, doi:10.1130/2014.2504(04). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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Moon-forming impact. Contrary to the expectations of early researchers like Oparin and Urey, the early atmosphere was probably dominated by CO2 and N2, rather than CH4 and NH3. A CO2-N2 atmosphere, supplemented by small amounts of H2 and CO, is said to be “weakly reduced” (Walker, 1977). That said, there could have been a period during the early Hadean, when highly reduced gases, especially CH4, were more abundant, for reasons described next. The relevant considerations are the following: (1) The early mantle was probably relatively oxidized, as it is today, with an effective oxygen fugacity, fO2, close to that of the QFM (quartz-fayalite-magnetite) synthetic buffer. Although some authors (e.g., Kasting et al., 1993; Holland, 2002) have proposed that the early mantle was more reduced, data from V (Canil, 1997, 2002; Li and Lee, 2004) and Cr (Delano, 2001) in ancient rocks indicate that this was not the case. More recently, analyses of Ce anomalies in zircons suggest that the upper mantle was already oxidized by ca. 4.4 Ga (Trail et al., 2011). Theory also supports an oxidized early mantle, at least after the first few hundred million years (Frost et al., 2004; Wade and Wood, 2005; Wood et al., 2006; Frost and McCammon, 2008). Large impacts, including the Moon-forming impact, would have created deep magma oceans at Earth’s surface. At the high pressures encountered near the base of such an ocean, ferrous iron should have disproportionated to ferric plus metallic iron 3Fe+2 → 2Fe+3 + Fe.

(1)

Some of the metallic iron segregated out and went into the core, perhaps aided by continued downward migration of molten iron from large impacts. The rest of the ferrous iron, along with virtually all of the ferric iron, remained behind in the mantle. The upper mantle was then oxidized as some of this ferric iron was lifted upward by convection. The time scale for upper-mantle oxidation depends on the convective overturn time, and the process requires whole mantle convection. This mantle oxidation process may have been supplemented by escape of hydrogen from impact-generated steam atmospheres, which could have produced as much as one third of the ferric iron now present in Earth’s upper mantle (Hamano et al., 2013). A strong point of the disproportionation model is that it accounts for the apparently heterogeneous oxidation state of the present upper mantle, because mixing may have been incomplete. This theory might allow some portions of the upper mantle to remain relatively reduced for as much as several hundred million years, if convective mixing was slow. An additional strong point of this hypothesis is that it can explain the reduced redox state of Mars’ mantle, which is thought to have an fO2 ~3–5 log units below QFM, based on analyses of Martian meteorites (Stanley et al., 2011; Grott et al., 2011). Mars, being smaller, never achieved such high pressures in its interior, and so its mantle remained closer to the initial, more reduced, iron-wüstite (IW) state. The upshot of this argument is that, after the first few hundred million years, volcanic gases should have been dominated by the types of compounds released from volcanoes today,

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namely, N2, CO2, SO2, and H2O (Holland, 1984, 2002). Smaller amounts of reduced compounds such as H2, CO, and H2S should also have been released (Holland, 1984, 2002). Virtually no CH4 or NH3 is expected to be released from high-temperature, subaerial volcanoes with magmas near QFM. Some CH4 and NH3 could have been released by submarine volcanism, although the dominant source of submarine CH4 would likely have been serpentinization, as discussed later herein. H2S should have been the dominant sulfur gas released by submarine volcanism, but little of it would have reached the atmosphere because most of it should have reacted with dissolved ferrous iron in the oceans to form insoluble iron sulfide, and eventually pyrite. (2) A second reason why highly reduced CH4-NH3 atmospheres are not expected to persist is that these compounds are both photolyzed by solar ultraviolet (UV) radiation, and the process is more or less irreversible. Photolysis of NH3 occurs at wavelengths below ~220 nm and is thus quite rapid (Kuhn and Atreya, 1979; Kasting, 1982). Photolysis of CH4 requires photons below ~145 nm and is much slower (Kasting et al., 1983; Zahnle, 1986; Pavlov et al., 2001). The word “slow,” however, is a relative term. In practical terms, this means that a CH4 flux equivalent to the modern biogenic CH4 flux of ~3 × 1013 mol yr–1 (Prentice et al., 2001) could have sustained an atmospheric CH4 mixing ratio of the order of a few hundred parts per million by volume (Pavlov et al., 2001; Kharecha et al., 2005). Prebiotic CH4 fluxes were probably considerably lower than this (see following), so this suggests that CH4 was at best a minor constituent of the prebiotic atmosphere. Once methanogens evolved, the biological methane flux should have been comparable to today (Kharecha et al., 2005), so CH4 should have been an important, but still relatively minor, component of the Archean atmosphere. Another way of looking at this problem is the following: Suppose that Schaefer and Fegley (2007) are correct and that Earth started out with a dense, CH4-dominated atmosphere as a consequence of its formation from carbonaceous chondritic material. Let’s take 10 bar of CH4 for the sake of concreteness. The column depth of this atmosphere is ~3 × 1026 CH4 molecules cm–2. CH4 is efficiently photolyzed by solar Lyman alpha (Ly α) radiation at 121.6 nm. The modern solar mean Ly α flux at 1 AU (astronomical unit) is ~3 × 1011 photons cm–2 s–1 (White et al., 1990), so the average flux received over Earth’s surface is one fourth that value, or 8 × 1010 photons cm–2 s–1. However, the Ly α flux early in Earth’s history was higher than this by a considerable amount. Ribas et al. (2005) gave the following formula for the Ly flux evolution F = 19.2τ−0.72.

(2)

Here, the solar age is in b.y., and F is expressed in ergs cm–2 s–1. The formula is valid for >0.1 b.y. If we integrate Equation 2 from 0.1 to 1 b.y., the average value of F compared to today is ~5; hence, the average solar Ly flux in photon units is (8 × 1010 photons cm–2 s–1) × 5 = 4 × 1011 photons cm–2 s–1. Finally, dividing the CH4 column depth by this number gives the CH4 lifetime as

Primary paper | Atmospheric composition of Hadean–early Archean Earth

tCH4 =

3 × 1026 cm−2 11

−2 −1

4 × 10 cm s

≈ 8 × 1014 s = 3 × 107 yr .

(3)

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space, making the entire process irreversible. The relevant reaction sequence is the following: H2O + hν → H + OH

In other words, even a 10 bar impact-produced CH4 atmosphere would have been destroyed by photolysis in only ~30 m.y. By “destroyed,” I mean “converted to CO2,” because that is where the carbon would have ended up after the hydrogen had escaped to space. The required oxygen atoms would come from H2O photolysis. So, unless life originated within the first 30 m.y. following the end of accretion, the atmosphere is not likely to have been highly reduced. (3) Finally, one should ask the question: What exactly would be the significance, or theoretical benefit, of a strongly reduced CH4-NH3 atmosphere? It might indeed aid in Miller-Urey synthesis of organic compounds needed for the origin of life, but other mechanisms for producing such compounds have been postulated, including synthesis within (off-axis) hydrothermal vents on the seafloor (Baross and Hoffman, 1985) and delivery of complex organic molecules by interplanetary dust particles of cometary or asteroidal origin (Chyba and Sagan, 1992). From a climate standpoint, Sagan and Mullen (1972) suggested that a CH4-NH3 atmosphere could have produced sufficient greenhouse effect to offset the faint young Sun. However, NH3 is difficult to keep around, as pointed out here, and CH4 in excess of CO2 photolyzes to produce organic haze—a phenomenon that was not accounted for in Sagan and Mullen’s model. Organic haze creates an anti-greenhouse effect that cools a planet’s surface (McKay et al., 1991; Haqq-Misra et al., 2008). Thus, there is no theoretical reason to believe that a Miller-Urey type atmosphere could have resolved the faint young Sun problem. A dense CO2 atmosphere does a much better job, as discussed below. Abundance of CO2: Constraints from Geochemistry and Climate The abundance of CO2 in the early atmosphere is a topic of considerable debate. Zahnle (2006) described the situation this way: Very high (tens of bar) CO2 levels are expected very early in Earth’s history as a consequence of volcanic outgassing and impact degassing of incoming planetesimals. Most of the mass of the impactors was in large, 100-km-diameter or more, planetesimals, and these objects would have released a large fraction of their volatiles directly into the atmosphere when they hit. Much of the carbon may have been released initially as CO rather than CO2 (Kasting, 1990). I shall come back to this thought later in this review. Eventually, almost all of the carbon would have been converted to CO2. The mechanism that accomplished this is the following: First, H2O would have been photolyzed by UV photons shortward of 240 nm. Such photons can penetrate all the way down to the surface unless organic haze was present or the CO2 concentration was very high. Then, the by-products of H2O photolysis would have reacted with CO, oxidizing it to CO2, and releasing H2 in the process. The H2 would then have escaped to

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CO + OH → CO2 + H H + H + M → Η2 + Μ Net: CO + H2O → CO2 + H2 .

(4)

Exactly what this implies for early atmospheric CO2 concentrations is unclear. As Zahnle (2006) pointed out (see also Sleep and Zahnle, 2001), CO2 would have been depleted rather quickly by weathering of the seafloor, followed by subduction of carbonatized seafloor into the mantle. How fast and how far CO2 levels would have been drawn down are difficult to determine from theory alone. Zahnle favors low CO2 levels and very cold conditions after the first few hundred million years. The Sun was less luminous by ~25% at 3.8 Ga (Gough, 1981), so the climate would have been extremely cold in the absence of enhanced concentrations of greenhouse gases. Assume for the moment that CH4 concentrations were low and that CO2 and H2O were the only important greenhouse gases. Then, roughly 0.2 bar of CO2 would have been required at 3.8 Ga just to keep Earth’s mean surface temperature above freezing (Kasting, 1987). It seems likely that CO2 concentrations were at least this high. If they were not, and if other factors did not compensate, Earth’s oceans would have frozen over entirely, and silicate weathering on the continents (whatever continents existed at that time) would have come to a virtual halt. Sleep and Zahnle (2001) argued that removal of CO2 by seafloor weathering would have continued; however, it is not clear that this argument is correct. The ice on the ocean surface would have been at least hundreds of meters thick, as it was limited only by geothermal heat flow. Under such conditions, transfer of CO2 from the atmosphere to the ocean would have been strongly inhibited; hence, any volcanic CO2 that was released directly into the atmosphere should have remained there. So, at steady state, it is difficult to see how CO2 could have stabilized at less than ~0.2 bar. Later in the Archean, constraints from the geologic record become available. The most convincing of these are the evidence for glaciation at 2.9 Ga (Young et al., 1998) and 2.2–2.45 Ga (Crowell, 1999). During these times, CO2 could not have been too high—otherwise, the glaciations would not have occurred. The upper limit on pCO2 at 2.2 Ga, based on climate calculations, is ~0.2 bar (Kasting, 1987). Geochemical indicators, however, suggest that pCO2 was much lower than this. Rye et al. (1995) suggested an upper limit of roughly 0.03 bar, or 100× present atmospheric level (PAL), at 2.8 Ga, based on the absence of the mineral siderite, FeCO3, in paleosols. Sheldon (2006) has criticized the Rye et al. analysis on various grounds. He pointed out that if the authors were to repeat their analysis with modern thermodynamic data, they would predict paleo-CO2 concentrations below that of today, which is unrealistic. Sheldon himself analyzed paleosol weathering from a kinetic standpoint and derived

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an upper limit of 0.009 bar at 2.2 Ga and, more recently, 0.015 bar at 2.69 Ga (Driese et al., 2011). This is just about the amount of CO2 needed to bring Earth’s mean surface temperature above freezing at this time (Haqq-Misra et al., 2008, their figure 10). Sheldon’s analysis can itself be criticized on several counts. His calculated pCO2 values come from an equation derived originally by Holland and Zbinden (1988): M

pCO2 = t

KCO2 · r 3

10

+κ

DCO2 · α

.

(5)

L

The first term in the bracketed expression accounts for transport of dissolved CO2 in rainwater; the second term accounts for gaseous diffusion of CO2 through the soil. Here, M is the total number of moles of CO2 cm–2 required to produce the observed weathering in the paleosol, t is the time required to form the soil, KCO is Henry’s law coefficient for CO2, r is the amount of rain2 fall that is absorbed by the soil, κ is a constant (= 1.43 × 103 s · cm3/[mol yr]), DCO is the diffusion coefficient for CO2 in air, α 2 is the ratio of the CO2 diffusion coefficient in air to that in soil, and L is the total depth of the soil horizon. Nominal values for these parameters (in cgs units) and associated uncertainties are listed in Sheldon’s paper. According to his discussion, the biggest uncertainty (a factor of 2 in either direction) is in t, the time required for soil formation. He allows α to vary from 0.1 by plus or minus 20%. However, Holland and Zbinden assigned a much larger uncertainty to α, a factor of 10 in either direction. Holland and Zbinden also assumed that r = 50 cm yr–1, or half the globally averaged modern rainfall rate, whereas Sheldon used r = 100 cm yr–1. This makes Sheldon’s estimate for pCO2 lower by a factor of 2 in cases where percolation of rainwater dominates. Holland and Zbinden themselves listed the uncertainty in calculated pCO2 as being a factor of 10 in either direction (their Eq. 19), and this is after arbitrarily reducing the uncertainties in α and L by a factor of 2 compared to their earlier analysis of O2 delivery to soils. Finally, by setting the flux of CO2 into the soil over time t equal to the observed weathering, M, all of these authors effectively assumed that weathering by CO2 is 100 percent efficient, that is, that all of the CO2 that makes it into the soil results in mineral dissolution. In reality, some of this CO2 probably does not react, and hence Equation 5 should yield only a lower limit on pCO2, and a highly uncertain one at that. Archean CO2 concentrations have also been estimated based on the mineralogy of banded iron formations (Rosing et al., 2010). These authors suggested that magnetite (Fe3O4) in banded iron formations would have been quantitatively converted to siderite (FeCO3) if pCO2 exceeded 0.001 bar, or 3 PAL. If correct, this limit would apply throughout the time that banded iron formations formed, i.e., the Archean–early Proterozoic. Their argument has been criticized on climatic grounds, as mentioned earlier (Goldblatt and Zahnle, 2011), and on geochemical grounds as well (Reinhard and Planavsky, 2011; Dauphas and Kasting,

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2011). Briefly, conversion of magnetite to siderite requires a reductant, probably organic carbon. If the supply of organic carbon to sediments was small, magnetite would have been preserved. So, the Rosing et al. limit on CO2 is probably invalid. N2 ABUNDANCE Early N2 partial pressures may well have been higher than today by a factor of 2–3 (but see counter-arguments below). Goldblatt et al. (2009) have presented a strong argument for this. I will reiterate their analysis briefly here. N2, like CO2, would have been released directly into the atmosphere by large impacts. (The nitrogen in impacting planetesimals was likely bound to organic compounds, but it would have been volatilized to N2 in impact plumes.) Venus has similar amounts of excess carbon as does Earth, but it has 2–3 times as much nitrogen as is present in Earth’s atmosphere. Earth probably has equivalent amounts of nitrogen, much of which is in the mantle today, but it was probably present in the atmosphere during the early Archean. The presence of extra N2 could have warmed Earth’s surface by 4°–5 °C by pressure-broadening the absorption lines of CO2 and H2O (Goldblatt et al., 2009). This could reduce the amount of CO2 needed to warm the early environment by roughly a factor of 2–3. Eventually, this excess N2 was removed by incorporation into organic sediments, followed by subduction down into the mantle. The time scale on which this happened is unclear. Recently, however, some quantitative evidence bearing on atmospheric N2 levels has emerged: Som et al. (2012) estimated that total atmospheric pressure at 2.7 Ga was no higher than twice the present value, based on raindrop impressions in ancient tuffs. This evidence is consistent with the Goldblatt et al. model, since the high pN2 values postulated by those authors probably applied prior to this time. But Marty et al. (2013) placed even lower limits on Archean pN2 based on the N2/36Ar ratio in fluid inclusions. If their analysis is correct, then Archean pN2 was no higher than today’s value. H2 ABUNDANCE Recently, Wordsworth and Pierrehumbert (2013) argued that H2 could have helped to warm early Earth if it was present in concentrations exceeding ~10% by volume. H2 turns out to be a good greenhouse gas, because collisions can allow it to absorb radiation across the thermal-infrared spectrum by exciting its pure rotational levels (Kasting, 2013). Whether it was really important for Earth’s early climate is uncertain, however. Most models of the early atmosphere (Kasting, 1993, and references therein) suggest that H2 mixing ratios were of the order of 0.1% by volume. So, to achieve an H2 mixing ratio of 10%, either volcanic outgassing rates would need to have been 100 times higher than today, or hydrogen must have escaped to space at less than the diffusion limit, or some combination of these factors must have applied. None of this is impossible, but it remains to be demonstrated that H2 would really have been abundant enough for its climatic effect

Primary paper | Atmospheric composition of Hadean–early Archean Earth to have been important. Furthermore, as pointed out by Kasting (2013), once methanogens evolved, they should have pulled atmospheric H2 mixing ratios down by catalyzing the reaction CO2 + 4H2 → CH4 + 2H2O.

(6)

Calculations by Kharecha et al. (2005), discussed in the next section, suggest that methanogens would have been highly efficient at converting H2 to CH4 by this mechanism. CH4 ABUNDANCE CH4 could have been generated by several processes on early Earth. As mentioned earlier, methanogens, once they evolved would have been the largest source. According to Kharecha et al. (2005), methanogens could have produced CH4 at a rate comparable to today’s biological production rate, 535 Tg(CH4) yr–1, or 3.3 × 1013 mol yr–1 (Prentice et al., 2001). In photochemist’s units, this is ~1.2 × 1011 cm–2 s–1. (The conversion factor is: 1 mol/yr = 0.00374 cm–2s–1.) I will henceforth ignore biology and concentrate on abiotic CH4 sources. The biological fluxes, though, are useful for comparison. CH4 could also have been produced by serpentinization of ultramafic rocks in mid-ocean-ridge hydrothermal systems. In this process, ferrous iron in the rock is oxidized to magnetite, and either H2O is reduced to H2 or CO2 is reduced to CH4. We can bound the methane flux by comparing it with the rate at which seafloor is oxidized by this process. The stoichiometry is CO2 + 2H2O ↔ CH4 + 2O2.

(7)

According to Sleep (2005), the rate of serpentinization of seafloor is ~0.2 × 1012 mol O2 equivalent per year, or approximately one-fifth of the total seafloor oxidation rate. Production of 1 mol of CH4 is equivalent to consumption of 2 moles of O2, according to Equation 7. So, if all of the serpentinization results in CH4 production, the rate of CH4 release should be ~1 × 1011 mol yr–1, or 4 × 108 cm–2 s–1. This is an upper limit on CH4 production today, because much of the serpentinization probably produces H2 instead. Even so, it is smaller than the biological flux by a factor of more than 300. A CH4 flux of this magnitude would produce a CH4 mixing ratio of roughly half a part per million, based on extrapolation from figure 1 of Pavlov et al. (2001). The rate of serpentinization could conceivably have been higher on early Earth if more ultramafic rock was exposed to interactions with warm seawater. Quantifying this source is difficult, though, because it requires understanding how plate tectonics in the distant past differed from plate tectonics today. CH4 could also have been produced from impacts via interactions with metallic iron particles in their plumes (Kress and McKay, 2004; Schaefer and Fegley, 2007). This source of CH4 could conceivably have been much larger, of the order of 1010–1011 cm–2 s–1. This number is derived by comparison to the numbers for impact production of CO given in the following, by

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allowing the impacts to generate CH4 instead of CO. A CH4 flux of 1011 cm–2 s–1 could produce an atmospheric CH4 mixing ratio of roughly 1000 ppmv, or 0.1% (Pavlov et al., 2001). This is enough to produce 10°–12 °C of greenhouse warming (Haqq-Misra et al., 2008, their figure 10), thereby further reducing the amount of CO2 needed to warm early Earth. So, a reasonable CO2 partial pressure for the late Archean is ~0.03 bar, or 100 PAL (Haqq-Misra et al., 2008). As pointed out earlier, however, CH4 has a double-edged effect on paleoclimate, because organic haze begins to form at CH4/CO2 ratios greater than ~0.1. This creates an anti-greenhouse effect that cools the climate, as happens on Saturn’s moon, Titan (McKay et al., 1991). So, a CH4-dominated early Earth, if it existed, would have been quite cold. I note parenthetically that the reason organic haze does not form on the CH4-rich giant planets themselves is that the polymerization process is poisoned by H2. However, terrestrial planets, and Titan, are expected to form hazes because their H2 is continually lost to space. IMPORTANCE OF CO Production of CO by Photochemistry A second gas that could have been generated by impacts is CO. Most of the carbon in the high-temperature portion of the impact plume would likely have been in this form (Kasting, 1990), as discussed further in the following. However, impacts are not required for CO to have been an important constituent of the early atmosphere, because it is also produced by photolysis of CO2: CO2 + hν → CO + O.

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A photochemical model calculation can provide some idea of how much CO should have been present. Figure 1 shows vertical profiles of selected atmospheric species for a hypothetical Archean atmosphere. The model used here is the one described in Segura et al. (2007). It tracks the photochemistry of H, O, C, N, and S compounds, along with hydrocarbons up through C1 (i.e., CH4). Parameters used in this calculation were the following: surface pressure—1 bar; CO2 mixing ratio—0.2; H2 outgassing rate—1 × 1010 cm–2s–1; SO2 outgassing rate—3.5 × 108 cm–2s–1, CO outgassing rate—0; CH4 outgassing rate—1 × 106 cm–2s–1 (i.e., purposely assumed negligible). A present-day solar EUV (extreme ultraviolet) flux was assumed in this calculation, along with a relatively cool surface temperature, 280 K. Hydrogen was assumed to escape to space at the diffusion-limited rate φesc(H2) ≅ 2.5 × 1013 ƒtot(H2),

(9)

where ƒtot(H2) is the total hydrogen mixing ratio in the stratosphere ƒtot(H2) = ƒ(H2) + ƒ(H2O) + 2ƒ(CH4) +… .

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Hydrogen is budgeted here, for convenience, in terms of H2 molecules rather than H atoms.

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J.F. Kasting very effective source of metabolic energy. Today, certain organisms, notably acetogens, consume CO by using the enzyme CO dehydrogenase. This enzyme is considered to be primitive and was likely present in the earliest prokaryotes (Ragsdale, 2004; Kharecha et al., 2005). An alternative way of saying the same thing is to calculate the free energy change for reaction when H2 and CO are set equal to their calculated values in the model atmosphere. Using the Nernst equation, this energy change is given by ΔG R = ΔG R0 + RT 1n

Figure 1. Vertical profiles of major atmospheric species for a typical early Archean model atmosphere. A surface pressure of 1 bar is assumed. Other parameters for the calculation are described in the text.

As shown in Figure 1, the volume mixing ratios of H2 and CO near the surface are 2.6 × 10−4 and 1.6 × 10−5, respectively. When combined with 0.2 bar of CO2, this mixture appears rather unfavorable to prebiotic synthesis of organic compounds. Spark discharge (i.e., lightning) would probably not result in good yields of organics in such a mixture. However, the apparently oxidized nature of this atmosphere is deceiving. In reality, such an atmosphere could be very good for driving primitive microbial metabolisms. Consider the equilibrium reaction

pCO2 · pH2 pCO

≅ −17 kJ/mol .

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For comparison, the free energy needed to synthesize adenosine triphosphate (ATP) is 35.6 kJ mol–1 (Kral et al., 1998), and some organisms are able to metabolize at free energies as low at 10–15 kJ mol–1 (Conrad, 1996). So, converting CO to H2 could have provided sufficient energy to power primitive metabolisms. Others have had this thought, as well. Huber and Wachtershauser (1997, 1998) have suggested that such CO-based metabolism may have been an early step in the origin of life. Their hypothesis is discussed further later herein. Let’s consider why this free energy would have been available on early Earth. The ultimate source of this free energy is the photolysis of CO2 (reaction), which creates the high-energy species CO and O. H2 is also a high-energy species compared to the H2O, from which it ultimately derives. However, in Earth’s atmosphere, H2 escaped to space whereas CO did not. That is why the free energy gradient exists in Equation 14. Addition of CO from Impacts

CO + H2O( ) ↔ CO2 + H2.

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The ‘ ’ means that the H2O is in liquid phase. According to the NIST Chemistry Webbook (2011; http://webbook.nist.gov/ chemistry/ [accessed November 2011]), the Gibbs free energy change for this reaction under standard conditions at 25 °C is: ΔG R0 = −20.0 kJ mol–1. The equilibrium constant is given by 0 Keq = pCO2 · pH2 = e−ΔGR /(RT) ≅ 3.2 × 103 atm . pCO

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In this expression, R is the universal gas constant, and pi represents the partial pressure of gas i (in units of atm). Rearranging equation and evaluating it at pCO2 = 0.2 bar gives Keq pH2 ≅ 1.6 × 104 . = pCO pCO2

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Now, compare this ratio to the H2:CO ratio in the model atmosphere, ~16. The calculated equilibrium ratio of these gases in water solution is 1000 times higher. Energetically, it is as if the effective pH2 in solution is ~0.26 atm instead of 2.6 × 10−4 atm. A primitive organism that could utilize CO would have a

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We have just begun to tap the chemical potential of CO. The previous analysis does not include any production of CO from volcanism or impacts. Of these two sources, impacts were probably the most important. The impact flux as a function of time is difficult to pin down, as it depends critically on formation models for the early solar system. The total integrated flux during the heavy bombardment period, 4.4–3.8 Ga, can be estimated from analysis of Moon rocks. It is equivalent to a layer of rock ~2.4 km deep on early Earth (Sleep et al., 1989; Kasting, 1990). Kasting (1990) assumed that this material accreted at an exponentially decreasing rate with a decay constant of 0.2 b.y. He estimated that the CO production rate at 3.8 Ga would have been 1.5 × 1010 cm–2 s–1 if the impactors consisted of comets, and approximately one third this rate if the impactors were similar to carbonaceous chrondrites, or approximately one fifth this rate if they consisted of ordinary chondrites. The average flux over the 0.6 b.y. heavy bombardment period would have been 3.8 times this value. Since that time, some planetary dynamicists have proposed the so-called “Nice” model, which revives the idea that the impacts were concentrated in a 0.1–0.2 b.y. pulse near 3.9 Ga (Gomes et al., 2005; Tsiganis et al., 2005). If all of the impactors arrived within a 200 m.y. time interval, then the CO production

Primary paper | Atmospheric composition of Hadean–early Archean Earth rate during that time would have been enhanced by an additional factor of 3. Thus, plausible rates of CO production could be as high as (3–15) × 1010 cm–2 s–1. The significance of these numbers is the following: The atmosphere has a finite ability to rid itself of CO (Kasting, 1990). The reaction sequence begins with the photolysis of H2O, and there are only so many photons available below 240 nm that are capable of doing that, especially as CO2 absorbs in this region as well, and as most of the water vapor is confined below 10 km. Hence, the removal rate of CO cannot exceed the columnintegrated rate of photolysis of H2O, designated here as ∫ JH2O. That number is of the order of 2 × 1010 cm–2 s–1 (Fig. 2). In Figure 2, ∫ JH2O is shown as a function of atmospheric CO mixing ratio (dotted curve). Also shown is the corresponding impact production rate of CO that would be required to sustain that CO mixing ratio. Both curves are quite flat at CO mixing ratios above 10−4. This illustrates a phenomenon that has been called “CO runaway” (Kasting et al., 1983; Zahnle, 1986). CO production rates higher than this critical value (2 × 1010 cm–2 s–1) result in accumulation of CO in the atmosphere. To give a concrete example, suppose CO was being produced at a rate of 3 × 1010 cm–2 s–1, i.e., the lower bound of the numbers given in the previous paragraph. Two thirds of this CO could have been consumed by the reaction sequence in Equation 4. The other one third would have accumulated in the atmosphere. If the accumulation continued for 200 m.y., then the total amount of CO built up would be 6 × 1025 molecules cm–2, equivalent to a 3 bar atmosphere of pure CO. The corresponding H2 partial pressure from Equation 11 is 5 × 104 bar. Alternatively, the free energy change from the Nernst equation is −48 kJ mol–1. This latter calculation takes into account the fact that the H2 mixing ratio builds up to ~10−3 under such circumstances (not shown). The calculation just described is meant to be illustrative, not realistic. Higher UV fluxes on early Earth (Ribas et al., 2005) would raise the H2O photolysis rate and make CO accumulation more difficult. The increase is ~50% for the Sun at 3.9 Ga, but it would be higher early on. Furthermore, hydration of CO to formate (HCOO–) in the oceans should have provided a sink for CO (Van Trump and Miller, 1973) CO + OH− → HCOO−.

the oceans more acidic, which would have slowed the reaction in Equation 15 by reducing [OH]. When all of the uncertainties are considered, it is difficult to make a reliable estimate of how high atmospheric CO levels might have gotten and how long they would have lasted. Still, this calculation does suggest that CO could have built up to significant concentrations in the immediate aftermath of large impacts, and that the chemical potential generated by this CO could have been a driving force for early biological metabolism. As mentioned earlier, Huber and Wachtershauser (1997, 1998) have suggested that CO was directly involved in life’s origin. Their chemoautotrophic model for life’s origin also requires H2S or CH3SH

(15)

The rate of this reaction is relatively slow at low temperatures; for surface temperatures near the modern value, its effect on atmospheric CO can be parameterized by assuming a deposition velocity of 10−8–10−9 cm s−1 (Kasting, 1990; Kharecha et al., 2005). The dashed curve in Figure 2A labeled “vdepnCO” shows the removal rate when the deposition velocity is 10−8 cm s−1. According to this figure, loss of CO by way of this mechanism would not have competed effectively with photochemical removal by the reaction sequence in Equation 4 until the CO mixing ratio approached 10% by volume. Higher atmospheric CO2 concentrations could have sped up this reaction by warming the surface; however, high CO2 may also have made

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Figure 2. Inverse calculations showing various fluxes as a function of atmospheric CO mixing ratio. (A) φCO is the CO flux from impacts needed to sustain the CO mixing ratio, ∫ JH2O is the columnintegrated H2O photolysis rate, and vdepnCO is the rate at which CO would flow into the ocean given a deposition velocity of 10−8 cm–2 s–1. (B) Remaining terms in the atmospheric hydrogen budget Equation 16. Solid curves, along with φCO from panel A, represent H2 sources; dashed curves represent H2 sinks. The surface deposition flux of CO, vdepnCO from panel A, has not been included in the budget shown. In reality, this would constitute an additional sink for hydrogen, which would need to be compensated by higher CO input. It was not included explicitly here because the lower boundary was used to parameterize CO production from impacts (i.e., CO flows upward at a rate φCO).

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(methyl mercaptan), and so they envision this process happening in a hydrothermal vent environment. However, the dissolved CO levels in vent fluids that they cite (0.26–0.36 cm3 kg–1), are equivalent to those that would be in equilibrium with atmospheric CO partial pressures of 0.01–0.016 atm. This is within the range of CO concentrations that might have been generated by impacts. So, impact-generated CO might have allowed early life to expand beyond the hydrothermal vent environment, if indeed that is where life originated. CO also helps to form peptide bonds between amino acids (Huber and Wachtershauser, 1998), and so early protein synthesis pathways might have relied on the presence of this compound. Before leaving this section, I should spend a few words describing panel B of Figure 2. This panel shows the remaining terms in the atmospheric hydrogen budget. The budget equation is φout(H2) + φCO + φoxi = φesc(H2) + φred .

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Here, φout(H2) (= 2 × 1010 cm–2 s–1) is the assumed outgassing rate of H2, φCO is the production rate of CO from impacts and volcanism, φoxi is the rainout + surface deposition rate of photochemical oxidants, φred is the same rate for photochemical reductants, and φesc(H2) is the hydrogen escape rate calculated from Equation 16. This equation represents conservation, not of hydrogen molecules, but of available electrons (Kasting and Brown, 1998). Whether or not this equation is balanced is a good check of whether one’s photochemical model is working properly. Our models are well balanced. The main photochemical oxidant removed from these model atmospheres is nitroxyl, HNO. This is an oxidant because H2 is required to convert the nitrogen back to its stable form of N2, 2 HNO + H2 → N2 + 2H2O.

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The main photochemical reductant in these model atmospheres is formaldehyde, H2CO. The importance of this molecule was originally pointed out by Pinto et al. (1980). Formaldehyde in solution hydrates to form methylene glycol, H2C(OH)2. In high concentrations, it can polymerize to form sugars. What would happen next is a matter of conjecture. Shaw (2008) has pointed out that organic compounds can be reduced to CH4 during passage through the mid-ocean-ridge hydrothermal vents. If this was the ultimate fate of formaldehyde, then it should have resulted in an additional CH4 source of a few times 109 cm–2 s–1. This is comparable to the production rate of CH4 from serpentinization estimated earlier, but it is still 30–100 times smaller than the present biological CH4 flux. Finally, to bring closure to this story, I should point out that dissolved CO and/or formate might also be reduced to CH4 during passage through the mid-ocean-ridge hydrothermal vents. The potential CH4 source can be estimated by taking the modern rate of seawater cycling through the vents, 2.5 × 1012 m3 yr–1 (Sleep and Zahnle, 2001; Shaw, 2008), and multiplying by the dissolved

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CO content of seawater. Assuming that atmospheric pCO = 0.1 bar, and using a Henry’s law constant of 1 × 10–3 mol/l/atm (at 25 °C), [CO] = 10−4 mol L–1, or 0.1 mol m–3. If this CO was quantitatively reduced to CH4 during passage through the vents, the rate of CH4 production would have been 2.5 × 1011 mol yr–1, or ~109 cm–2 s–1. Once again, this source is comparable to the CH4 source from serpentinization or from H2CO reduction. Like the serpentinization flux, though, this CH4 source could have been larger in the distant past if the rate of hydrothermal circulation was higher—which it almost certainly was—or if atmospheric pCO exceeded 0.1 bar. CONCLUSIONS The atmosphere during the Hadean and early Archean was probably dominated by CO2 and N2. Any dense CH4 atmosphere that might have been generated by impacts would have dissipated within ~30 m.y. Multibar levels of CO2 very early in Earth’s history probably gave way to CO2 concentrations of a few tenths of a bar or less by 2.8 Ga. N2 concentrations were probably 2–3 times higher than today during the Hadean and declined to modern values sometime after that. The faint young Sun problem was probably solved largely by CO2, with a little help from enhanced pressure broadening by N2. CH4 was probably a trace constituent during most of the Hadean, with concentrations of a few tens of parts per million. CH4 concentrations would have risen to several hundreds of parts per million once methanogens evolved, perhaps during the early Archean. At this point, CH4 would also have made a significant contribution (10°–12 °C of warming) toward solving the faint young Sun problem. Higher amounts of warming by CH4 would have been prevented by formation of organic haze. CO would have been an important constituent of the prebiotic atmosphere, particularly in the aftermath of large impacts. The free energy of CO is high and could have been tapped by early organisms to help drive their metabolisms. This should be considered when evaluating theories of life’s origin and early evolution. REFERENCES CITED Baross, J.A., and Hoffman, S.E., 1985, Submarine hydrothermal vents and associated gradient environments as sites for the origin and evolution of life: Origins of Life and Evolution of the Biosphere, v. 15, p. 327–345, doi:10.1007/BF01808177. Canil, D., 1997, Vanadium partitioning and the oxidation state of Archaean komatiite magmas: Nature, v. 389, p. 842–845, doi:10.1038/39860. Canil, D., 2002, Vanadium in peridotites, mantle redox and tectonic environments: Archean to present: Earth and Planetary Science Letters, v. 195, p. 75–90, doi:10.1016/S0012-821X(01)00582-9. Chyba, C., and Sagan, C., 1992, Endogenous production, exogenous delivery and impact-shock synthesis of organic molecules: An inventory for the origins of life: Nature, v. 355, p. 125–132, doi:10.1038/355125a0. Conrad, R., 1996, Soil microorganisms as controllers of atmospheric trace gases (H2, CO, CH4, OCS, N2O, and NO): Microbiological Reviews, v. 60, p. 609–640. Crowell, J.C., 1999, Pre-Mesozoic Ice Ages: Their Bearing on Understanding the Climate System: Geological Society of America Memoir 192, 106 p.

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Sheldon, N.D., 2006, Precambrian paleosols and atmospheric CO2 levels: Precambrian Research, v. 147, p. 148–155, doi:10.1016/j.precamres .2006.02.004. Sleep, N.H., 2005, Dioxygen over geologic time: Metal Ions in Biological Systems, v. 43, p. 49–73, doi:10.1201/9780824751999.ch3. Sleep, N.H., and Zahnle, K., 2001, Carbon dioxide cycling and implications for climate on ancient Earth: Journal of Geophysical Research, v. 106, p. 1373–1399, doi:10.1029/2000JE001247. Sleep, N.H., Zahnle, K.J., Kasting, J.F., and Morowitz, H.J., 1989, Annihilation of ecosystems by large asteroid impacts on the early Earth: Nature, v. 342, p. 139–142, doi:10.1038/342139a0. Som, S.M., Catling, D.C., Harnmeijer, J.P., Polivka, P.M., and Buick, R., 2012, Air density 2.7 billion years ago limited to less than twice modern levels by fossil raindrop imprints: Nature, v. 484, p. 359–362, doi:10.1038/ nature10890. Stanley, B.D., Hirschmann, M.M., and Withers, A.C., 2011, CO2 solubility in Martian basalts and Martian atmospheric evolution: Geochimica et Cosmochimica Acta, v. 75, p. 5987–6003. Trail, D., Watson, E.B., and Tailby, N.D., 2011, The oxidation state of Hadean magmas and implications for early Earth’s atmosphere: Nature, v. 480, p. 79–82, doi:10.1038/nature10655. Tsiganis, K., Gomes, R., Morbidelli, A., and Levison, H.F., 2005, Origin of the orbital architecture of the giant planets of the solar system: Nature, v. 435, p. 459–461, doi:10.1038/nature03539. Van Trump, J.E., and Miller, S.L., 1973, Carbon monoxide on the primitive Earth: Earth and Planetary Science Letters, v. 20, p. 145–150, doi:10.1016/0012-821X(73)90152-0.

Wade, J., and Wood, B.J., 2005, Core formation and the oxidation state of the Earth: Earth and Planetary Science Letters, v. 236, p. 78–95, doi:10.1016/j .epsl.2005.05.017. Walker, J.C.G., 1977, Evolution of the Atmosphere: New York, Macmillan, 318 p. White, O.R., Rottman, G.J., and Livingston, W.C., 1990, Estimation of the solar Lyman alpha flux from ground based measurements of the Ca II K line: Geophysical Research Letters, v. 17, p. 575–578, doi:10.1029/ GL017i005p00575. Wood, B.J., Walte, M.J., and Wade, J., 2006, Accretion of the Earth and segregation of its core: Nature, v. 441, p. 825–833, doi:10.1038/nature04763. Wordsworth, R., and Pierrehumbert, R., 2013, Hydrogen-nitrogen greenhouse warming in Earth’s early atmosphere: Science, v. 339, p. 64–67, doi:10.1126/science.1225759. Young, G.M., von Brunn, V., Gold, D.J.C., and Minter, W.E.L., 1998, Earth’s oldest reported glaciation; physical and chemical evidence from the Archean Mozaan Group (~2.9 Ga) of South Africa: The Journal of Geology, v. 106, p. 523–538, doi:10.1086/516039. Zahnle, K.J., 1986, Photochemistry of methane and the formation of hydrocyanic acid (HCN) in the Earth’s early atmosphere: Journal of Geophysical Research, v. 91, p. 2819–2834, doi:10.1029/JD091iD02p02819. Zahnle, K.J., 2006, Earth’s earliest atmosphere: Elements, v. 2, p. 217–222, doi:10.2113/gselements.2.4.217.

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The Geological Society of America Special Paper 504 2014

Atmospheric composition of Hadean–early Archean Earth: The importance of CO: Comment Laura Schaefer* Department of Astronomy, Harvard University, Cambridge, Massachusetts 02138, USA, and Planetary Chemistry Laboratory, Department of Earth and Planetary Sciences, McDonnell Center for the Space Sciences, Washington University, St. Louis, Missouri 63130, USA Bruce Fegley Jr. Planetary Chemistry Laboratory, Department of Earth and Planetary Sciences, McDonnell Center for the Space Sciences, Washington University, St. Louis, Missouri 63130, USA

INTRODUCTION

sis would lead to optically thick organic hazes in the atmosphere and therefore global cooling. New work by Wolf and Toon (2010) suggests that the hazes may instead have been optically thick in the ultraviolet spectrum while remaining relatively transparent in the visible spectrum. If this is the case, then a reducing atmosphere could exist as more than a transient phenomenon.

Kasting (this volume) describes the current state of knowledge about early Earth’s atmosphere starting with the Hadean (ca. 4.5 Ga) and going through the late Archean (ca. 2.6 Ga). In his paper, he avoids discussing the Hadean atmosphere prior to the Moon-forming impact, or shortly after, due to a lack of quantitative evidence for this time period. Work done by us at Washington University can be used to discuss the Hadean atmosphere in more detail. We have produced several models using thermochemical equilibrium to describe the atmospheres evolved from both passive planetary outgassing and impact degassing of different types of chondritic material. We will discuss both of these models and briefly describe how they may be used to further our understanding of the early atmosphere. The time line presented suggests that Earth’s atmosphere evolved from being a relatively thick CO2 atmosphere enriched in N2, with diminishing concentrations of CO2 during the Archean. The faint young Sun problem is partially mitigated by greenhouse warming with CO2, and pressure broadening due to the enhanced abundance of N2. Kasting states that CH4 was likely present at no more than a few tens of parts per million, but may have been larger at earlier periods. However, as described here, our models suggest that certain types of precursor materials for Earth could lead to relatively large CH4 abundances in the atmosphere. Large abundances of CH4 are considered unfavorable because photoly-

PLANETARY OUTGASSING Little is known about the Hadean atmosphere, except that it was likely the product of numerous processes, including planetary outgassing (i.e., volcanism), impact degassing (devolatilization of impactors), and photochemical processes. Our previous theoretical work may shed some light on two of these processes. In Schaefer and Fegley (2007), we described calculations of planetesimal outgassing for ordinary chondritic material. This work was motivated by the master’s thesis work of Dushan Bukvic in 1979 under John Lewis. In his thesis, Bukvic calculated that an early Earth composed of chondritic material (either all ordinary chondritic material or a mixture of ordinary plus CI [carbonaceous] chondritic material) would primarily outgas reduced species such as CH4, NH3, and H2 at temperatures up to ~1500 K. However, his work was largely ignored, due to the studies showing early oxidation of Earth’s mantle, suggesting that any outgassed material ought to have been oxidized (e.g., H2O, CO2, N2). On repeating his calculations with updated thermochemical data

*[email protected] Schaefer, L., and Fegley, B., Jr., 2014, Atmospheric composition of Hadean–early Archean Earth: The importance of CO: Comment, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 29–31, doi:10.1130/2014.2504(05). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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Schaefer and Fegley

Log mole fraction

and ordinary chondrite abundances, we were surprised to find that we did indeed produce significant amounts of reduced gases. We used ordinary chondrites, which are the most abundant type of meteorite, because models of the formation of Earth from mixing of meteorites suggest that Earth is primarily a mixture of ordinary and enstatite chondrites, with small amounts of carbonaceous material (e.g., Lodders, 2000). For a terrestrial temperature-pressure profile up to depths equivalent to 1500 K and 23 kbar, we found the gas from an ordinary chondrite to be composed primarily of CH4 and H2, with smaller amounts of H2O and NH3 (see figure 18 in Schaefer and Fegley, 2007). These results are very similar to those of Saxena and Fei (1988), who modeled mantle-fluid equilibria for a CI chondritic composition, and likewise found reducing fluids to be stable. In Figure 1, we show recent calculations of planetary outgassing using the bulk silicate Earth (BSE) composition of Kargel and Lewis (1993). The BSE is the composition of the mantle, crust, hydrosphere, and atmosphere (i.e., the primordial mantle). The total pressure is held constant at 1 bar, but results at higher pressures are similar. The BSE is more oxidized than H-chondrites, and it reflects the current-day mantle oxidation state. Outgassing from the BSE produces ~80% H2O and ~17% CO2 from ~600 to 1100 K. At higher temperatures, SO2 gas is more abundant than CO2 in the atmosphere. However, at temperatures below 600 K, CO2 is converted to CH4, with an abundance of ~0.25% at 1 bar total pressure. CH4 and NH3 are the only abundant gases produced at 300 K. Recent work by Frost et al. (2008) and Rubie et al. (2011) indicated that the primordial mantle was more reducing than the BSE, which would

increase the amount of reducing gases produced during outgassing in the Hadean. It appears that Earth’s mantle became more oxidized after planetary formation until it reached its presentday oxidation state around the quartz-fayalite-magnetite (QFM) buffer in the Archean (at or sometime before 3.8 Ga). This is ~700 m.y. after formation of Earth, which is comparable to the time between the present and the Cambrian era. IMPACT DEGASSING Schaefer and Fegley (2010) extended our earlier work and modeled impact degassing of carbonaceous, ordinary, and enstatite chondritic material. This paper assumed chemical equilibrium between the impacting body and its generated atmosphere, but it did not consider interaction of the degassed material with a preexisting atmosphere. (In other words, we looked at formation of the initial atmosphere.) We found that for a nominal model of 1500 K and 100 bars, CI and CM chondrites produced atmospheres composed primarily of H2O and CO2, whereas CV chondrites produced mostly CO2. Note that CI, CV and CM are subgroups of carbonaceous chondrites that are classified based on their compositional similarity to a representative meteorite (e.g., Ivuna [CI], Vigarano [CV], Mighei [CM]). Ordinary and enstatite chondrites, on the other hand, produced a combination of H2 and CO, with significant amounts of H2O. Kasting (this volume) also suggests that CO was a major component of early Earth’s atmosphere. He suggests that CO, which has a large free energy, may have served as a source of metabolic energy for early organisms. However, in a previous work, Fegley et al. (1986) showed that impacts into a preexisting atmosphere composed of N2 + CO produce substantial amounts of CO2, graphite, H2, CH4, NH3, and HCN. This agrees with Kasting’s (this volume) statement that impacts will convert CO into CO2, and it suggests that buildup of CO would not have been not possible during the Late Heavy Bombardment. The ordinary and enstatite chondrites also produced nontrivial amounts of methane (~0.7%) even at 1500 K. At temperatures below ~1100 K at 100 bars pressure, CH4 and N2 dominated in the gases generated by the ordinary and enstatite chondrites. Ammonia was less abundant, but it was still generated at levels of several 100 ppmv by the ordinary and enstatite chondrites at temperatures between ~500 and 1000 K. Wolf and Toon (2010) suggested that the organic hazes may have stabilized NH3 against ultraviolet photolysis at low altitudes. Ammonia is an important greenhouse gas and may further mitigate the faint young Sun problem. PHOTOCHEMICAL MODELS

Figure 1. Equilibrium composition of gases released from the bulk silicate Earth (BSE) (Kargel and Lewis, 1993) as a function of temperature at a total pressure of 1 bar.

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Based on the impact degassing and planetary outgassing studies outlined here, the early atmosphere seems to have been composed of mixtures containing CO2, H2O, H2, CO, N2, NH3, CH4, SO2, HCl, and HF. Results for initial compositions derived from the ordinary or enstatite chondrite outgassing data

Commentary | Atmospheric composition of Hadean–early Archean Earth (i.e., H2 + CO + CH4 mixtures) would be very interesting. For the sake of completeness, it would also be interesting if photochemical models included HCl and HF. Of the major gases predicted from outgassing models, these are the only gases not included in photochemical models of early Earth (e.g., Kasting, 1990; Zahnle, 1986). In the present day, only very longlived Cl and F compounds are of concern in photochemical models because HCl and HF are extremely soluble and have short rainout times. In the Hadean and Archean, they likely resided primarily in the primitive oceans, as they do today. However, continual production through increased volcanism (e.g., Martin et al., 2007) and large impacts, particularly into oceans (e.g., Genda and Abe, 2005), may have caused them to be injected into the atmosphere at higher rates than today. Although it is tempting to assume that their effect would be small, until they are included in calculations, their effect on the lifetimes of CO, CH4, or other major gases is uncertain. REFERENCES CITED Bukvic, D.S., 1979, Outgassing of Chondritic Planets [M.S. thesis]: Cambridge, Massachusetts, Massachusetts Institute of Technology, 80 p. Fegley, B., Jr., Prinn, R.G., Hartman, H., and Watkins, G.H., 1986, Chemical effects of large impacts on the Earth’s primitive atmosphere: Nature, v. 319, p. 305–308, doi:10.1038/319305a0. Frost, D.J., Mann, U., Asahara, Y., and Rubie, D.C., 2008, The redox state of the mantle during and just after core formation: Philosophical Transactions of the Royal Society, v. A-366, p. 4315–4337, doi:10.1098/rsta.2008.0147. Genda, H., and Abe, Y., 2005, Enhanced atmospheric loss on protoplanets at the giant impact phase in the presence of oceans: Nature, v. 433, p. 842–844, doi:10.1038/nature03360.

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Kargel, J.S., and Lewis, J.S., 1993, The composition and early evolution of Earth: Icarus, v. 105, p. 1–25, doi:10.1006/icar.1993.1108. Kasting, J., 1990, Bolide impacts and the oxidation state of carbon in the Earth’s early atmosphere: Origins of Life and Evolution of the Biosphere, v. 20, p. 199–231, doi:10.1007/BF01808105. Kasting, J.F., 2014, this volume, Atmospheric composition of Hadean–early Archean Earth: The importance of CO, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, doi:10.1130/2014.2504(04). Lodders, K., 2000, An oxygen isotope mixing model for the accretion and composition of rocky planets: Space Science Reviews, v. 92, p. 341–354, doi:10.1023/A:1005220003004. Martin, R.S., Mather, T.A., and Pyle, D.M., 2007, Volcanic emissions and the early Earth atmosphere: Geochimica et Cosmochimica Acta, v. 71, p. 3673–3685, doi:10.1016/j.gca.2007.04.035. Rubie, D.C., Frost, D.J., Mann, U., Asahara, Y., Nimmo, F., Tsuno, K., Kegler, P., Holzheid, A., and Palme, H., 2011, Heterogeneous accretion, composition and core-mantle differentiation of the Earth: Earth and Planetary Science Letters, v. 301, p. 31–42, doi:10.1016/j.epsl.2010.11.030. Saxena, S.K., and Fei, Y., 1988, Phase equilibrium in a system of chondritic composition: Implications for early mantle-fluid compositions: The Journal of Geology, v. 96, p. 601–607, doi:10.1086/629255. Schaefer, L., and Fegley, B., Jr., 2007, Outgassing of ordinary chondritic material and some of its implications for the chemistry of asteroids, planets, and satellites: Icarus, v. 186, p. 462–483, doi:10.1016/j.icarus.2006.09.002. Schaefer, L., and Fegley, B., Jr., 2010, Chemistry of atmospheres formed during accretion of the Earth and other terrestrial planets: Icarus, v. 208, p. 438– 448, doi:10.1016/j.icarus.2010.01.026. Wolf, E.T., and Toon, O.B., 2010, Fractal organic hazes provided an ultraviolet shield for early Earth: Science, v. 328, p. 1266–1268, doi:10.1126/ science.1183260. Zahnle, K., 1986, Photochemistry of methane and formation of hydrocyanic acid (HCN) in the Earth’s early atmosphere: Journal of Geophysical Research, v. 91, p. 2819–2834, doi:10.1029/JD091iD02p02819.

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The Geological Society of America Special Paper 504 2014

Discussion of “Atmospheric composition of Hadean– early Archean Earth: The importance of CO” (Kasting) James F. Kasting: Thanks. I’ll just take my one minute and just address that one question that you (Laura Schaefer) posed at the end [of the commentary]. I think you’re totally right, that when the atmosphere, in the immediate aftermath of a large impact; if all the water’s up there in the atmosphere there’s going to be a lot of chlorine there. The reason that chlorine is important on Venus and Mars is there’s no surface liquid water. So chlorine chemistry is very important. But on the early Earth once the ocean is there you’ll form HCl, the chlorine goes in and forms chloride in the ocean, you’ll get a little bit of chlorine in the atmosphere from sea salt, providing you modest amounts. But I’ve never really seen the reason to include this in our weakly reduced atmospheres. It would change kinetic rates in there a little bit, but the main thing that’s controlling the chemistry in there is the redox balance—it’s the outgassing rates of hydrogen and carbon dioxide and the escape rate to space. That’s what’s controlling the composition.

You can get a very reducing environment, there are lots of origin of life theories that involve the midocean ridge vents. Just one footnote on that: We had a talk by Jack Szostak at Penn State a couple of weeks ago. He’s of course a Nobel Prize–winner who has worked on the origin of life. He doesn’t think you can form life in the ocean because the salinity’s too high. So he’s worried about membrane formation; he argues that it has to happen in a freshwater environment. If he’s right about that, then you need a little bit of land area and some hydrothermal activity, perhaps on land. Frost: Well, the interesting thing about hydrothermal vents is that we have in our biochemical systems the participation of various transition metals: copper and stuff like that, which are much more abundant in a vent environment than they are floating around in the ocean. Kasting: So maybe this emphasis on Miller-Urey synthesis is misplaced.

Ron Frost: One of the overlying, shall we say, assumptions that I see here is that we need to talk about a reducing environment in the atmosphere to produce life in the Hadean or earliest Archean. And it’s quite clear that life could easily form today in an oxidized atmosphere in a serpentinization vent. Serpentinites are in equilibrium about one log unit above iron magnetite, and at that oxygen fugacity all organic hydrocarbons are stable. If you had a much more reducing environment but still, let’s say, a CO2 water environment instead of a methanehydrogen environment, you could still get that kind of an oxygen fugacity maybe over a broad area around serpentinization vents. And if you’re erupting a lot of komatiites at that time, serpentinization vents might be much more abundant than they are today.

Nathan Sheldon: So I just want to address a couple of things that Jim said about the paleosol data because I actually don’t think we can use a model to say that the data is wrong. So one of the things that you mentioned, that our error margins are smaller. There’s a simple reason for that, that analytical uncertainty of measurements that we did was about one-fifth of what it was __(?)__. So the shrink in the error margins is just due to the fact that we use better, more modern techniques for data measurement. So that’s one piece of it; it’s just the actual measurements of what was happening __(?)___. The second thing is, you know, the paleosol data, in terms of being reliable really only goes back to about that 2.7-billionyear-old point that we had last year. So I mean I would say, before that it becomes a lot more uncertain. Obviously, we’d like to fill that gap in but it’s not such a big gap now.

Kasting: Thank you for making that point. I actually had that point on one of my slides and I skipped over it in my haste.

Kasting, J.F., Schaefer, L., et al., 2014, Discussion of “Atmospheric composition of Hadean–early Archean Earth: The importance of CO” (Kasting), in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 33–35, doi:10.1130/2014.2504(06). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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Kasting et al. But at the CO2 levels that we reconstruct which I think again [are] something like 25 to maybe 40 times preindustrial levels—actually the model that I’m doing with Shawn [Domagal-Goldman] right now we can show that you actually can get, you know, habitable, equable conditions at the surface at these sort of CO2 levels with a reasonable CO2-methane mixing ratio. So we actually don’t need to have a CO2 level as high as 100 times; we can actually do it with 25 to 40. And I showed this at Goldschmidt [in 2011].

Kasting: Thank you for making those points. Actually there is a lot of uncertainty, though, in soil formation times, for instance. So there [are] some inherent uncertainties in paleosol analysis. And the other thing I failed to mention is that by including the extra nitrogen that Colin Goldblatt and Zahnle would argue was there, we can bring our estimate for CO2 down; we could make that fit your paleosol data. So there is a little bit of overlap, you know, even—we could fit your data if we have a little bit more nitrogen. Shawn Domagal-Goldman: So I just wanted to bring up one point with regards to one slide that Jim showed. And this is the one with regards to the fractal haze. And this, on the surface, seems like a very nice, elegant solution to how you would preserve certain gases like ammonia in an atmosphere or other greenhouse gases that might keep things warm. However, there is a problem with it that I don’t think has been discussed very much. We’re going to bring this up— there’s some work I’ve done with Mark Claire and Aubrey Zerkle and James Farquhar that’s going to hopefully come out in Nature Geosciences [sic] pretty soon.1 We see a number of things in that paper. One of them is that the haze is put in the atmosphere as a hypothesis of how you could protect some gases from photochemistry below the haze. The problem with that is the haze itself requires photochemical reactions to be formed. So it’s fundamentally, especially if it’s a fractal haze that’s very efficient at shutting off the photochemical reactions, it’s fundamentally self-limiting. As you increase the thickness of the haze you slow down the haze production rate in the atmosphere and so there’s only a certain amount of bang you can get out of that buck.

Domagal-Goldman: Yeah, I’m not saying that there’s no haze; I’m just saying that there’s a limit to its thickness. Kasting: You can extend the lifetime of ammonia, though, because that would arguably be contained below the haze. Domagal-Goldman: Perhaps. Harry Becker: I have a question for Laura. When you do [these] impact volatilization models do you see, or would you expect to see, some fractionation between the different elements, like when you look at the sulfur versus the hydrogen versus the nitrogen? Laura Schaefer: You’re asking if we have more volatilization of certain elements than others? Becker: I mean I would expect that solubility in the magma has an impact on what goes into the atmosphere and what remains in the magma and maybe, you know, solidifies or whatever. Schaefer: Right, right. So our models are looking at equilibrium between the impacting body itself and the gas that is evolving. So we’re not including, say, this impactor generates a magma ocean; we’re not including equilibrium with that. But yes, there is a certain amount, if I go back, maybe— okay for instance in this graph you can see where sulfur is becoming more abundant—maybe I can show it. So here, I mean there’s no sulfur, really, in the gas below this temperature, at temperatures below here. So if you’re not above this temperature you’re going to be producing a lot of SO2 or H2S in the atmosphere because it’s going to remain in the solid body. So yes, in that sense you are getting a certain amount of fractionation and it’s going to depend on the temperature of the impact. Becker: And would you expect that some of the volatized species are getting lost by hydrodynamic escape, for instance, and others not? Schaefer: I mean I’m personally not a dynamicist but I am aware there is a paper by—I think by Genda and Abe—in Nature in 20052 that shows that with large impacts, especially those into an ocean, you get a lot of atmospheric loss. So yes, I would expect you’re getting some amount of hydrodynamic escape.

Kasting: Well the methane will rise above the haze, actually, because if the haze is blocking out all those photons you have to photolyze methane to form the haze. So methane is going to get photolyzed.

Becker: But you cannot put any constraints on these kinds of models?

1 This has been published as A.L. Zerkle, M.W. Claire, S.D. DomagalGoldman, J. Farquhar, and S.W. Poulton, 2012, A bistable organic-rich atmosphere on the Neoarchaean Earth: Nature Geoscience, v. 5, p. 359–363, doi:10.1038/ngeo1425.

2 H. Genda and Y. Abe, 2005, Enhanced atmospheric loss on protoplanets at the giant impact phase in the presence of oceans: Nature, v. 433, p. 842–844, doi:10.1038/nature03360.

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Discussion | Atmospheric composition of Hadean–early Archean Earth Schaefer: It’s something that you could—you could take this atmosphere and put it into a dynamic model, but this does not involve any dynamics.

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CO, so [that] by the time the plume has expanded, and it’s cooled below 1000 [Kelvin], some of that CO is going to be converted to methane. But it’s a kinetic process; you actually have to have a model of that impact plume and how much particulate material is in there. It’s not that easy to get the composition. You can’t get it, the results, just from thermodynamic equilibrium; you have to think about the dynamics of the process.

Becker: Yeah, the reason why I’m asking these questions is that because if you look at the ground truth data of what we have on the estimates, what the abundances of the major volatiles in the Earth’s mantle, or the silicate Earth are, I mean if you exclude nitrogen, which there may be a problem because some of it might be hidden in the mantle. But all the other ones they have CI-like ratios within ± 30 percent or something like that, that may be a factor of two of order. But for sulfur or something like the sulfur-carbon ratio is like CI. So that would argue that there is not [a] huge loss of some kind of the species that you’re modeling. I thought you may __(?)__.

Martin J. Van Kranendonk: On one of your earlier slides you made some comparisons with the atmosphere of Venus. But Venus doesn’t have the moon like Earth does, and although they’re [a] similar size, now the mass may have been different. And from what George has been saying—and some of the planetary community—of a late addition of the volatiles to the Earth, is it a fair comparison in that scenario?

Kasting: I have some comments on that. I’m glad we went back to this slide because it actually shows that Laura and I are not in serious disagreement on this. If you look at this slide on the right that’s the more reduced material from which you might get a lot of methane. You get methane at temperatures below—in this case below about 1000 Kelvin, and you get CO for the carbon compound above 1000 Kelvin. So think about what happens in a big impact: You vaporize this thing, the initial temperatures are tens of thousands of Kelvins, right? So all the carbon starts out as—well it starts out as C, actually, then it goes to CO, and then that plume is expanding, though. It’s expanding and cooling, and there are particles in it, too, which condense out: the rock materials, some of which may contain some iron. This is the Cress and McCay paper—that catalyzes the reduction of

George H. Shaw: Well remember, Venus and—think about the CO2—Venus’s atmospheric CO2 is essentially equal to Earth’s surface CO2 in carbonates. This was Rubey’s excess volatiles—so you take Rubey’s excess volatiles for Earth— they’re the same for carbon for Venus and the Earth, so that would make it look like there—either you didn’t lose all the volatiles in the moon-forming impact or most of the volatiles were delivered after that. In homogeneous accretion, which I think had to be the case for theoretical reasons, that means that a lot of those volatiles arrive late in the process. The moon-forming impact was arguably the last impact of that size. But there were probably a lot of large impacts after that of material from the asteroid belt region. So there’s plenty of room to bring in carbon and water.

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The Geological Society of America Special Paper 504 2014

Waiting for O2 Kevin Zahnle* Space Science Division, National Aeronautics and Space Administration Ames Research Center, MS 245-3, Moffett Field, California 94035, USA David Catling Department of Earth and Space Sciences/Astrobiology Program, University of Washington, Seattle, Washington 98195, USA

ABSTRACT Oxygenic photosynthesis appears to be necessary for an oxygen-rich atmosphere like Earth’s. However, available geological and geochemical evidence suggests that at least 200 m.y., and possibly as many as 700 m.y., elapsed between the advent of oxygenic photosynthesis and the establishment of an oxygen atmosphere. The interregnum implies that at least one other necessary condition for O2 needed to be met. Here, we argue that the second condition was the oxidation of the surface and crust to the point where free O2 became more stable than competing reduced gases such as CH4, and that the cause of Earth’s surface oxidation was the same cause as it is for other planets with oxidized surfaces: hydrogen escape to space. The duration of the interregnum was determined by the rate of hydrogen escape and by the size of the reduced reservoir that needed to be oxidized before O2 became favored. We speculate that hydrogen escape determined the history of continental growth, and we are confident that hydrogen escape provided a progressive bias to biological evolution.

INTRODUCTION

atmosphere, a state that is more widespread in the solar system. Oxygen and oxidation are different things and reflect different processes acting on different time scales, although it is plausible that one is prerequisite to the other. It could be that it was free oxygen in the atmosphere that oxidized the surface, or it could be that oxidation of the surface allowed free oxygen to endure. Here, we presume that surface oxidation is a prerequisite to O2. It has long been considered probable from hints in the geological record that oxygenic photosynthesis appeared much earlier than widespread crustal oxidation (Holland, 1962; Buick, 2008), and thus that surface oxidation played a role in the rise of oxygen (Berkner and Marshall, 1965). Where we go beyond Berkner and Marshall is that we give a reason: Oxygenation is caused by the

This volume addresses Earth from its beginnings in the Hadean ca. 4.4 Ga to the rise of oxygen in the Paleoproterozoic a mere 2.2 b.y. ago. Two very interesting things happened on Earth during the first half of its history: (1) life began, and (2) later, and perhaps a bit less important, the atmosphere began to fill up with oxygen. From the perspective of its inhabitants, these may be the two most important events in Earth’s history. Neither is well understood. Oxygen raises two issues that are usefully separated. One is the matter of abundant O2, which is the distinctive feature of Earth’s atmosphere. The other is the oxidation of the surface and

*[email protected] Zahnle, K., and Catling, D., 2014, Waiting for O2, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 37–48, doi:10.1130/2014.2504(07). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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steady loss of hydrogen to space. In our story, the apparent delay between the origin of oxygenic photosynthesis and the establishment of an oxygenated atmosphere is explained by how long it took to oxidize the surface through the steady loss of hydrogen to space (Catling et al., 2001; Claire et al., 2006). This chapter is not intended to give an even-handed overview of the history of oxygen, if such were possible. The intent of the October 2011 Pardee symposium on which this volume is based was for the speakers to take clear and opposing points of view. This essay holds to the original intent. Our point of view is as stated earlier, but it harms nothing to state it again more forcefully: We think that oxygenation of Earth was caused by hydrogen escape. Hydrogen escape stepped Earth through a series of titratable reservoirs and provided the bias that drove biological evolution to develop the O2 atmosphere. Readers seeking a modern, broadly encompassing review of oxygen that is relatively free of overt biases would be well served to read Farquhar and Johnston (2008). OXIDATION More than 60 years ago, Harold C. Urey (1952, p. 352) wrote that the “highly oxidized condition is rare in the cosmos and exists in the surface regions of the Earth and probably only in the surface regions of Venus and Mars. Beyond these we know of no highly oxidized regions at all, although undoubtedly other localized regions of this kind exist.” His underlying interest in the matter was in the conditions pertinent to the spontaneous origin of life. Urey was greatly influenced by Oparin’s arguments in favor of an anaerobic origin of life. Oparin regarded a prebiotic source of organic molecules as essential to the phenomenon, and therefore concluded that Earth’s first atmosphere was strongly reducing (Urey, 1952). Urey argued that, because hydrogen is the most abundant element in the cosmos, a reduced atmosphere is to be expected at early times relevant to the origin of life, while the “highly oxidized condition of planetary surfaces” is an evolutionary result of hydrogen escape. In addition, Urey regarded the tension between the reduced interior and the oxidized surface as contributing to life’s subsequent evolution. This too is an important insight into the boundary conditions imposed by a planet on life and its evolution. It turns out that oxidation of planetary surfaces is not rare in the solar system. In all cases apart from Earth, the oxidation is clearly caused by hydrogen escape. For example, several icy satellites (Ganymede, Europa, Rhea) have extremely thin O2 atmospheres derived from splitting water molecules in ice by ultraviolet (UV) photons or energetic particle bombardment (Cruikshank, 2010). Hydrogen escapes easily, but oxygen, which is much heavier, does not. As a consequence, over time, the surface ice and any contaminants in the ice become highly oxidized. Mars provides a more Earth-like example. Mars is red because much of the iron at its surface is oxidized. On Earth, red surfaces like these first appear in the geological record during the Huronian glaciation ca. 2.32 Ga and are one of the classic

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indicators of the first appearance of an oxic atmosphere (Holland, 1999; Kump, 2008; Guo et al., 2009; Bekker and Holland, 2012). Mars’s surface is locally characterized by the presence of strong oxidants such as peroxides (seen by Viking; Hunten, 1979) and perchlorates (seen by Phoenix; Hecht et al., 2009), and on a global scale, there are extensive and locally thick sulfate deposits (encountered by the Mars exploration rovers on the ground and mapped by satellites from above). Hydrogen escape is fast enough on Mars to generate its modest atmospheric reservoir of O2 in just 105 yr (Nair et al., 1994; Zahnle et al., 2008). On Mars, the oxidation appears to be quite shallow, with no evidence that oxidation extends to the mantle or even to the deeper crust. If so, planetary oxidation would have been quick (Hartman and McKay, 1995). In contrast to the other worlds of the solar system, most accounts of the rise of oxygen on Earth either marginalize or ignore the role of hydrogen escape in oxidizing the surface. Rather, oxidation of the surface is usually ascribed to a shrinking influence of reduced volcanic gases from a reduced mantle compared to the accumulation of reduced carbon (from CO2) in continents, or to an enhanced role of continents versus the mantle in weathering, or both, although how the mechanism(s) might work is debated (Holland, 1962; Kasting et al., 1993; Kump et al., 2001; Holland, 2002, 2009; Kump and Barley, 2007; Gaillard et al., 2011). A variant that is easier to understand posits the preferential subduction of reduced matter by the mantle (Hayes and Waldbauer, 2006). The variant is easier to understand because, like H escape, it actually oxidizes the surface. What these theories have in common is the primacy they assign to secular cooling in Earth’s evolution. Secular cooling is undeniable and has been viewed as a major driver of planetary evolution since the nineteenth century, if not before. For example, Lowell (1908) provides a good general overview of the hypothesis. Secular cooling is obviously a major factor setting the direction and pace of planetary evolution. Its consequences are easy to see throughout the solar system. However, secular cooling does not by itself change the surface’s oxidation state. Hence, theories of oxygenation that start with secular cooling depend on second-order effects to produce oxidation. These are usually effects that decrease the influence of the mantle on the surface over time, either through changes in the rate or style of plate tectonics or in the quantity or quality of volcanic gases, or more indirectly through the consequences of continental growth (which, if a fact, might have something to do with cooling). On the face of it, the idea that the mantle oxidizes the surface is rather puzzling. The gases that come from the mantle are reducing and have always been reducing, because the mantle is reduced. That they might have been more or less reducing in the past does not make them oxidizing. The reasonable expectation is that the surface would be reduced by these gases. Why the sum of second-order effects should favor oxidizing the surface is a fair question. Net oxidation of the surface by the mantle can occur only if the surface exports more reductant to the mantle (via subduction) than the mantle exhales in volcanic gases. In

Primary paper | Waiting for O2 other words, convention makes the rather awkward argument that an oxidized surface oxidizes itself by further reducing an already reduced mantle. The Hayes and Walbauer mechanism is not immune from this criticism. By contrast, hydrogen escape acts directly on the redox state of the surface and is categorically oxidizing. OXYGENIC PHOTOSYNTHESIS There is at present little doubt that oxygenic photosynthesis is required to create an oxygen-rich atmosphere like Earth’s. Although it is possible to imagine atmospheres that might become oxygenated abiotically through some combination of vigorous hydrogen escape and high rates of stellar UV irradiation, to the best of our knowledge no one has yet succeeded in demonstrating this convincingly with an actual photochemical model. The fundamental reason for this is that thermo-dynamics of typical planetary materials favor O2 only at very high temperatures. Photosynthesis on Earth makes O2 from 5800 K sunlight. Even under Mars’s favorable circumstances (weak gravity, an oxidized surface, extremely limited weathering owing to extremely low water activities and low surface temperatures, and minimal volcanism and crustal recycling), abiotic processes produce an O2 partial pressure that is only 4 × 10−5 that of Earth. Detailed photochemical models show that for an Earth-like planet, the abiotic photochemical source of O2 would be overwhelmed by Earth’s volcanic gases, resulting in O2 partial pressures at the surface on the order of 10−12 bar (Kasting and Walker, 1981; Kasting, 1993; Haqq-Misra et al., 2011). Oxygenic photosynthesis uses sunlight to split water molecules into hydrogen and oxygen. The O2 is released to the atmosphere as a by-product. The hydrogen is used to reduce CO2 to water and organic matter. Under an oxygenated atmosphere, most of the organic matter is aerobically respired, closing the cycle, but a fraction of the organic matter is further reduced to inedible carbon that is ultimately buried. Methane is also a product, but it is mostly eaten (using O2, sulfate, nitrate, or even ferric iron; Bekker et al., 2010) before it reaches the atmosphere. On long time scales, the net effect of oxygenic photosynthesis today is approximately CO2 → C + O2. This is why C burial is often treated as the source of O2 (Hayes et al., 1983). Many models focus on biological and geological controls on C burial to reconstruct Phanerozoic histories of O2 (Berner, 2001). Under an anoxic atmosphere, aerobic respiration might still be possible locally where O2 is made, but we would expect a bigger role for anaerobic pathways and a higher fraction of the organic matter to be reduced, and a much higher fraction of the organic matter made by oxygenic photosynthesis to reach the atmosphere as methane (Walker, 1987). Most of the CH4 reaches the stratosphere, where it is oxidized with the help of solar UV. Most of the H in CH4 is destined to escape to space. The net is therefore H2O = ½O2 + H2 (to space). When oxygenic photosynthesis leads to methane, H escape becomes a source of O2 (Catling et al., 2001).

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Methane is much less reactive than O2. If both are vented to the atmosphere by an oxygenic photosynthetic ecosystem, the O2 and free radicals that are made from O2 will react with rocks and such (Berkner and Marshall, 1965), while the CH4 reacts only with free radicals generated from O2. As a result, one expects the atmosphere after the origin of oxygenic photosynthesis to be CH4 rich rather than O2 rich, and that this will remain the state of the atmosphere for as long as chemical sinks for O2 remain easily accessed and plentiful. Methane itself leaves few geological traces, but it can polymerize to make aerosols, possibly nitrogenous or sulfurous (Pavlov et al., 2001; Domagal-Goldman et al., 2008; Trainer et al., 2006; Zerkle et al., 2012), which should be isotopically and structurally distinctive. A methane-rich atmosphere also welcomes other more fragile biogenic reduced gases, like sulfides, and thus may reveal itself indirectly by enabling a richer sulfur cycle (Zahnle et al., 2006; Kurzweil et al., 2013). It is sometimes argued that the advent of oxygenic photosynthesis was not just a necessary, but also a sufficient condition to create an O2 atmosphere (e.g., Kopp et al., 2005). The appeal here is to brute force kinetics: Photosynthesis using H2O is presumed to be so much more biologically productive than photosynthesis reliant on other hydrogen donors that it ought to take over the world on a biological time scale. This is not an irrefutable notion: It is not obvious a priori that the hydrogen source was much more limiting than all other nutrients, such that by removing this limit, biology would pour forth like a flood after a dam burst (Sleep and Bird, 2008). A second presumption is that the flux of biogenic O2 into the atmosphere would be great enough to utterly overwhelm the capacity of weathering reactions or hungry mouths to consume it; this, too, is at least debatable, albeit there is some evidence in favor of an oxygen overshoot ca. 2.1 Ga (Bekker and Holland, 2012). A more serious problem with this hypothesis is that it has not proved possible to identify when exactly in the geological record the one true revolution took place (Des Marais, 2000; Buick, 2008; Guo et al., 2009). One extreme solution to this paradox is to argue that Earth’s atmosphere has always been oxygenated and that the origin of oxygenic photosynthesis predates the geological record (Ohmoto, 1999). However, the overwhelming consensus recognizes several substantial changes in redox conditions at Earth’s surface between 3.5 Ga and 2.1 Ga (to say nothing of the evolution of multicellular animals ca. 0.6 Ga, an innovation best explained as heralding the advent of breathable levels of O2; Knoll, 2004). The stepwise ratcheting up of oxygenation is inconsistent with its origin in a single event. WHERE DOES OXYGEN COME FROM? The atmosphere, ocean, and crust are oxidized. There are 0.37 × 1020 moles of O2 in the atmosphere. The current rate of C burial is estimated to be on the order of 1 × 1013 moles/yr (Holland, 2002). At this rate, it would take only 4 m.y. to build up the O2 in the modern atmosphere. The equivalent to another 5 × 1020 moles of O2 is stored on the continents as sulfates and sedimentary ferric iron (e.g., iron formations); these are obvious

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products of oxic weathering. It would take only 50 m.y. at current rates to build up these inventories. The buried reservoir of reduced carbon is bigger than the sedimentary reservoirs of oxidized iron and sulfur. A net accumulation of reduced materials in the crust is to be expected if the surface chemistry is dominated by input of reduced gases from the mantle and hydrogen escape is negligible. This is the relationship to be expected if reduced (organic) carbon burial is the reason the surface is oxidized. The 12 × 1020 moles of reduced carbon in the continents (Claire et al., 2006) correspond to 120 m.y. of accumulation. Evidently, oxic weathering of old continental reduced carbon back to CO2 must be a major part of the modern oxygen cycle. The final victory of O2 has sometimes been identified with a pronounced excursion of 13C in carbonates to high positive values ca. 2.22–2.06 Ga, called the Lomagundi event (Karhu and Holland, 1996; Bekker et al., 2006; Melezhik et al., 2007; Bekker and Holland, 2012). The isotopic excursion suggests a simultaneous massive burial of reduced carbon. The 120 m.y. time scale for reduced carbon to accumulate in continents corresponds to the duration of the Lomagundi event. Two bigger reservoirs of oxidized material in the crust are often neglected. Iron in aging continental basalts is generally more oxidized than the iron in freshly erupted basalts. The difference suggests that 20 × 1020 moles of O2 have gone into oxidizing continental basalts (Lecuyer and Ricard, 1999). Another way to see this is to compare the mantle, which has an Fe2O3 content of only 0.1–0.4 wt% (Canil et al., 1994), to average continents, which have 3 wt% Fe2O3 (Lecuyer and Ricard, 1999). The oxygen reservoir linked to continental ferric iron is bigger than the reduced carbon reservoir, which means that the continental crust as well as the surface are oxidized. This is the relationship to be expected if H escape is the reason the surface is oxidized. The mechanism may be abiotic (e.g., 3FeO + H2O = Fe3O4 + H2), but the result would be a net source of H2 to the atmosphere and biosphere while the continents accumulated. If this H2 passed directly to space without passing through life, it might be possible to keep continental oxidation separated from carbon burial (which would leave the hypothesis that O2 originates by burying carbon liberated from CO2 intact), but in reality, the H2 would be utilized biologically (Kharecha et al., 2005), thus entangling continental oxidation with the carbon cycle. The biggest oxidized crustal reservoir is CO2 in carbonates. Although CO2 ice can be abundant in comets, most carbon in asteroids, comets, and meteorites is reduced. Earth probably accreted most of its carbon in a reduced form. Substantial amounts of reduced carbon in a reduced early mantle cannot be ruled out (Hirschmann and Dasgupta, 2009). The 50 × 1020 moles of CO2 at Earth’s surface (and another >100 × 1020 moles of CO2 in the mantle) were probably (but not certainly) oxidized during accretion or early in the Hadean, and they may have been oxidized by processes taking place inside the mantle (McCammon, 2005), and if so would not count toward the oxidized inventories, but they do represent a significant reservoir that became oxidized on Earth. It is worth recalling that there is little evidence that

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much of the 50 × 1020 moles of CO2 now in the continents was present in the Archean (Holland, 2009). The crustal reservoirs are small compared to what could be lost to the mantle, and it is possible that too much has been made of them. Evidence that the upper mantle has not changed its oxidation state since 3.8 Ga (Canil, 2002; Delano, 2001; Trail et al., 2011) tells us little about the lower mantle. The effect of export to the mantle could go either way. Oxidation of the surface could have been accomplished by net export of reduced carbon (Hayes and Waldbauer, 2006) and sulfur rather than by H escape. On the other hand, oxidized iron could have been exported to the mantle. The excess oxygen in the crust and atmosphere must derive ultimately from splitting H2O followed by hydrogen escape to space or from splitting CO2 and subsequent burial of reduced carbon in the mantle. In other words, what is needed is to separate the oxygen from the reductant and then to keep them separated, forever. Hydrogen escape is clearly irreversible, while carbon burial just as clearly is not. Both are possible on a planet with oxygenic photosynthesis. In practice most discussions of the history of oxygen on Earth have focused on CO2 and carbon burial, because hydrogen escape today is negligible. WHEN DID OXYGEN FIRST APPEAR? The geological record of the anoxic-oxic transition between ca. 2.5 Ga and 2.1 Ga is often referred to as the Great Oxidation Event (GOE). Authorities differ on what the GOE means, on how broadly it should be defined, or on what geologic indicators should take precedence. Figure 1 provides a cartoon summary of some key events. The classical evidence is concisely summarized by Holland (1999). Changes that mark the GOE include the near disappearance of reduced detrital minerals such as sulfides, siderite, and uraninite from fluvial deposits (Rasmussen and Buick, 1999; Hofmann et al., 2009); changes in the weathering rates of redox-sensitive elements (Anbar et al., 2007; Frei et al., 2009; Reinhard et al., 2009); the appearance of oxidized sediments (red beds) in the place of comparable unoxidized sediments (gray beds; Holland, 1999); the first appearance of massive sulfate deposits (Melezhik et al., 2005); the changing redox state of soils revealed by redox-sensitive elements such as cerium and iron; the evolution and eventual disappearance of massive banded iron formations (this was drawn out well into the Proterozoic; Canfield, 2005; Bekker et al., 2010); fluctuations in iron isotopes that speak of biological meddling (Rouxel et al., 2005; Bekker et al., 2010); and looming above all, the end at 2.45 Ga of huge massindependent fractionations of sedimentary sulfur isotopes (MIFS), a signal for which the only plausible source is photochemistry in an anoxic atmosphere (Farquhar et al., 2000; Farquhar and Wing, 2003; Farquhar and Johnston, 2008). In Figure 1, we have put this event at 2.45 Ga. A more conservative assessment of the sparse temporal record brackets the event between 2.32 and 2.45 Ga (Bekker et al., 2004). Oxygenic photosynthesis appears to predate the GOE significantly, although the earliest evidence is highly circumstantial

Primary paper | Waiting for O2 (Eigenbrode and Freeman, 2006; Buick, 2008). Buick (2008) regarded thick and widespread black shales ca. 3.2 Ga as suggestive of aerobic photoautotrophy. The Pongola ice ages ca. 2.94 Ga (Young et al., 1998) could signal the disappearance of major greenhouse gases, probably CO2 or CH4. Cold climates can be interpreted as weak circumstantial evidence for oxygen, the enemy of CH4 (although not of CO2), or they could be explained by an anti-greenhouse organic haze (Domagal-Goldman et al., 2008) triggered by an increase of biogenic CH4. The massive Steep Rock carbonate reefs suggest a great leap forward in biological productivity, as if some limit had been breached (Wilks and Nisbet, 1988). They are older than 2.78 Ga and probably younger than 2.82 Ga (Fralick et al., 2008). Evidence that water was used as the hydrogen source in photosynthesis is found in 2.72 Ga lacustrine stromatolites that are big, biogenic, and “evidently phototrophic” (Buick, 2008). There is no evidence of any electron donor for photosynthesis apart from water (Buick, 2008). Although stealthy anoxygenic photosynthesis based on H2 cannot be ruled out (Kharecha et al., 2005), evidence of oxygen bubbles (Bosak et al., 2009) and structural features characteristic of cyanobacterial mats (Flannery and Walter, 2012) strengthen the case for oxygenic photosynthesis. Extremely low δ13C values suggest both methanogenesis and methanotrophy (Hayes, 1994), and the latter implies that O2 was consumed metabolically. A similar but more widespread prevalence of low δ13C values in deep-water carbonates paired with modern photosynthetic δ13C values in shallow water suggest a more far-reaching impact of aerobic ecosystems between 2.7 and 2.45 Ga, perhaps achieving global reach by 2.45 Ga (Eigenbrode and Freeman, 2006). Molecular fossils in ca. 2.72–2.56 Ga rocks are indicative of sterol synthesis (Brocks et al., 2003; Eigenbrode et al., 2008). Biological synthesis of the original molecules demands the pres-

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ence of O2 itself (Summons et al., 2006; Waldbauer et al., 2011), which would leave little room for quibbling, were there not significant doubts that the molecular fossils are as old as the rocks (Rasmussen et al., 2008). Between 2.7 Ga and 2.45 Ga, there are several reports of elements such as molybdenum and rhenium (Anbar et al., 2007) and chromium (Frei et al., 2009) and minerals such as pyrite (Stüeken et al., 2012), which are insoluble unless oxidized, being weathered from continents and washed into the sea. These occurrences have been given the memorable name “whiffs of oxygen” (Anbar et al., 2007). However, the reported signals, both elemental abundances and isotopic anomalies, are small compared to more modern undoubted products of weathering under an O2-rich atmosphere (Scott et al., 2008). Oxidation of Cr to the Cr+6 of soluble chromate requires O2 (Frei et al., 2009). Konhauser et al. (2011) suggested instead that conditions may have been acidic enough to mobilize chromium as Cr+3. Acid indirectly indicates oxidation. A plausible source of strong acid is oxidation of sulfur or sulfide, but it may not mean O2, although Konhauser et al. (2011) preferred both biology and O2, as today at Rio Tinto. Whatever the details, it seems clear that by 2.6 Ga at the latest, H2O was being used as the hydrogen source in photosynthesis, and the oxidized product was mobile and regionally abundant. Although O2 need not have been the product, toxic plumes of biogenic O2 might account for things nicely. The best date to divide the Archean from the Proterozoic is ca. 2.45 Ga, defined by the seemingly sudden end to an extraordinary isotopic fractionation history of sedimentary sulfur (Farquhar and Johnston, 2008). This was followed quickly by Earth’s plunge into the first of a series of stupendous ice ages (Papineau et al., 2007). There remains some uncertainty regarding when these events took place. We have placed the end of MIF-S at 2.45 Ga following Farquhar and Johnston (2008), but

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Figure 1. A cartoon history of the Great Oxidation Event. Hydrogen escape rates and the range of mass-independent fractionations of sedimentary isotopes (MIF-S) are indicative of general trends. Approximate tropospheric O2 levels are estimated but should not be taken too literally. Below ~10−5 present atmospheric levels (PAL), O2 levels are ill-defined because the molecule would not in general be well mixed in the troposphere. Key dates are taken from Papineau et al. (2007), Scott et al. (2008), and Guo et al. (2009). “Whiff” [of O2] refers to geochemical hints that free O2 was present in the late Archean environment (Anbar et al., 2007).

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disappear ca. 2.48 Ga, with more modest but still significant Δ33S persisting until ca. 2.38 Ga (Farquhar and Johnston, 2008; James Farquhar, 2013, personal commun.). The only known way to get MIF-S big enough to give the observed Δ33S signal invokes atmospheric photochemistry driven by solar UV (Farquhar et al., 2000; Farquhar and Johnston, 2008). An atmospheric source has two requirements. First, the UV cannot be absorbed by ozone. Ozone becomes abundant when the atmosphere contains more than 0.1% O2. Thus, there cannot have been much O2 in the atmosphere. Second, there have to be at least two channels available to remove sulfur from the atmosphere (Pavlov and Kasting, 2002), in order to keep the two

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SULFUR Sulfur is particularly important to oxygen because it is, with carbon and iron, one of the three abundant elements that can easily change oxidation state at ordinary conditions. The record of sulfur’s ordinary mass fractionation (denoted δ34S) is consistent with increasing oxidation of the surface beginning in the late Archean (Fig. 2). The widening envelope bounding the scattered data indicates that greater amounts of sulfur were becoming biologically available, presumably as soluble sulfate in the seas, which afforded biology the opportunity to be more choosy about which sulfur isotopes they used (Canfield, 2005). The result is a progressively increasing scatter in δ34S from 2.7 Ga to 2.1 Ga. The extraordinary fractionations that characterize the Archean are recorded with sulfur’s third isotope. This fractionation, denoted Δ33S (often called mass-independent fractionation of sulfur, or MIF-S for short), is the difference between the massdependent fractionation of 33S that would be expected from the δ34S (were this mass dependent in the ordinary sense) and what is actually observed. Very large nonzero values of Δ33S seem to

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dates as early as 2.48 Ga and as late as 2.38 Ga are possible (James Farquhar, 2013, personal commun.). Both events signal profound changes in the composition of the atmosphere that are almost certainly linked to the atmosphere’s oxygenation. It is notable that these dates record the disappearance of things that were remarkable about the Archean environment—the mechanism for fractionating sulfur isotopes, thermal blanketing by unidentified greenhouse gases—rather than the advent of abundant O2 in the atmosphere. Red beds may be the highest fidelity and most directly pertinent of the classic geological indicators of the atmosphere’s changing redox state (Kump, 2008). Red beds record conditions at or near the surface that oxidized sediments in fluvial deposits. It is inferred that the oxygen came from the atmosphere. Figure 1 places the gray bed–red bed transition at 2.32 Ga and after the second of the three Paleoproterozoic ice ages that began ca. 2.45 Ga following Papineau et al. (2007) and Guo et al. (2009). The latest date that has been considered for the GOE, ca. 2.22 Ga, refers to the onset of the Lomagundi δ13C carbonate isotope excursion (Karhu and Holland, 1996; Bekker et al., 2006; Melezhik et al., 2007; Bekker and Holland, 2012) and to the massive Hotazel manganese deposit that followed the last of the great Paleoproterozoic ice ages. This—the last possible moment—is where Kopp et al. (2005) placed the origin of oxygenic photosynthesis. This is also around the time that the first massive sulfate deposits appear (Holland, 2002; Schroeder et al., 2008). The Lomagundi isotope excursion suggests that a great deal of isotopically light reduced carbon was buried, which on the modern Earth would imply that a great deal of O2 was left behind in the atmosphere, in obedience to the net reaction CO2 = C + O2 (Bekker and Holland, 2012). Kopp et al. (2005) suggested that the third ice age was the biggest and baddest and therefore the best ice age to link to the origin of oxygenic photosynthesis.

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Time [Ga] Figure 2. (Top) Carbon isotope history in carbonates (δ13C ≈ 0‰) and in reduced carbon (δ13C ≈ −30‰). (Bottom) The widening dispersion of ordinary mass-dependent fractionation δ34S as a function of time on Earth suggests that sulfur becomes progressively more abundant and less dear to life (Canfield, 2005). (Middle) Mass independent S fractionation Δ33S (Farquhar and Johnston, 2008) appears to increase as the biologically accessible sulfur pool grows until the signal vanishes ca. 2.45 Ga, to be followed by a series of major ice ages. The late Archean marks the time when biogenic reduced gases such as CH4 and H2S were abundant and biogenic O2 suppressed. This figure uses databases compiled by R. Buick, E. Stüecken, and J. Farquhar. Use of their data here does not imply their support for our interpretations of these data.

Primary paper | Waiting for O2 complementary Δ33S flavors separate (if, as today, nearly all S left the atmosphere as sulfate aerosols, the Δ33S would sum to naught). Laboratory experiments (Farquhar et al., 2001; DeWitt et al., 2010) show and photochemical models predict (Pavlov and Kasting, 2002; Zahnle et al., 2006; Kurzweil et al., 2013) that both sulfate aerosols and elemental sulfur (S8) aerosols should be generated photochemically in an anoxic atmosphere, and that elemental sulfur does not form in an oxic atmosphere. This mechanism could have generated significant nonzero Δ33S values if the atmosphere were deeply anoxic (2.7 Ga to ca. 2.2 Ga) for the atmospheric pO2 to increase from ~10−5 PAL to ~1% PAL and another ~1.6 b.y. (from ca. 2.2 Ga to ca. 0.6 Ga) to further increase to >50% PAL. It also implies that no mechanism has existed to regulate the atmospheric pO2 level, and that the level may continue to increase in the future. In contrast, the D-O model postulates: (1) the very rapid and early development of the fully oxygenated world (i.e., stage IV) by ca. 3.5 Ga; and (2) the regulation of atmospheric pO2 by the coupling of two negative feedback mechanisms: One is the pO2 dependence of O2 production by the burial of organic carbon (kerogen) in sediments, and the other is the pO2 dependence of O2 consumption by the oxidation of soil kerogen (Lasaga and Ohmoto, 2002). In discussions regarding the evolution of the atmosphere and biosphere, the most direct evidence should be the gases (e.g., O2, H2, CO2) trapped in fluid inclusions of ancient seawater in sedimentary rocks (e.g., cherts, carbonates, evaporites), and microfossils and the molecular biomarkers of various organisms (e.g., cyanobacteria, sulfate-reducing bacteria, methanogens, Eukarya) from various parts of the ancient oceans. However, due to diagenesis and metamorphism, primary fluid inclusions have not been preserved in Archean-aged sedimentary rocks. Despite a very large number of reports of Archean microfossils, paleontologists

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pO2 ≥ 15% P.A.L.

G.O.E.

Paleosols Williams Mt. Roe Denison Hekpoort Flin Flon & Wolhaarkop

Red beds

Sturgeon Falls

Jatulian

Uranium ores Witwatersrand

Blind River

Oklo

Figure 2. Lines of evidence summarized by Holland (1994) for the Great Oxidation Event (G.O.E.) at ca. 2.2 Ga. PAL—present atmospheric level.

Athabasca

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and organic geochemists have not been able to agree whether a particular “fossil-like” feature in an Archean sedimentary rock can be linked to a specific organism, and whether or not certain molecular biomarkers are indigenous. Therefore, we will limit our discussion on the chemical evolution of the atmosphere and oceans in this paper to geological, mineralogical, and geochemical lines of evidence. Holland (1994) summarized the following lines of evidence in pre–2.2 Ga rocks (Fig. 2) to support the C-W-H-K model: 1. loss of Fe from paleosols; 2. absence of red beds; 3. presence of detrital grains of uraninite and pyrite; 4. absence of groundwater-type U deposits; and 5. presence of banded iron formations (BIFs). Subsequent researchers have added the following lines of evidence in pre–2.2 Ga rocks to strengthen the C-W-H-K model: 6. absence of large variations in the δ34S values of sulfides and sulfates (Canfield and Raiswell, 1999); 7. presence of AIF-S, commonly known as the “massindependent fractionated sulfur isotopes (MIF-S),” in sulfides and sulfates in sedimentary rocks (Farquhar et al., 2000); 8. absence of U enrichment in black shales (Tribovillard et al., 2006); and 9. absence of Mo enrichment and Mo isotope fractionation in black shales (Anbar et al., 2007). Here, we will focus on the Archean geochemical cycles of O, Fe, and S, as they address the validity of items 1–9 as evidence for a reducing atmosphere. Our analyses of early Earth geochemical cycles of other redox-sensitive elements (U, Mo, Cr, Ce, N, and C) will be presented elsewhere.

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Paradigms and Contradictions Here, we use the term “paradigm” to mean the “standard model (theory) that is currently accepted by an overwhelming majority of scientists in a particular discipline (e.g., geoscience, astrobiology).” There are six major paradigms for the oxygen, iron, and sulfur geochemical cycles on early Earth (see Fig. 3):

No O2 & O3 UV SO42-

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Figure 3. A schematic illustration of the current paradigms for the geochemical cycles of O, Fe, and S on early Earth: The atmosphere and oceans were anoxic, the ocean was rich in Fe2+ supplied by mid-ocean-ridge hydrothermal fluids, but poor in SO42–; and ultraviolet (UV) photolysis of volcanic SO2 provided native S and SO42– to the ocean, which were used in the formation of pyrite and sulfate minerals in sedimentary rocks.

Primary paper | Oxygen, iron, and sulfur geochemical cycles on early Earth (I) The Great Oxidation Event: The change from a reducing atmosphere (pH2 > 10−6 atm > pO2) to an oxidizing atmosphere (pH2 < 10−6 atm < pO2) occurred at ca. 2.45 Ga, and the pO2 increased to >1% PAL by ca. 2.2 Ga. (II) Fe-rich Archean oceans: The oceans contained uniformly high (~3–10 ppm) ∑Fe2+ until ca. 1.85 Ga, compared to ~2 ppb in today’s oceans. (III) Hydrothermal origin for the oceanic Fe2+: Submarine hydrothermal fluids that were discharged at mid-ocean ridges (MORs) provided the oceanic Fe2+. (IV) SO42–-poor Archean oceans: The SO42– content of the Archean oceans was less than 1/200 of the present value (i.e., 2000 m) oceans (Ohmoto et al., 2006b). Since submarine hydrothermal fluids developed in such environments have salinity similar to normal seawater, the heated hydrothermal fluids would

Shale Deposition Stage clastic sediment atmosphere oxic zone

(goethite & silica nucleation) euxinic sea

Fe- & SiO2-rich brine pool BIF hydrothermal fluids

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organic & pyrite

shale

Figure 12. Proposed model (the Red and Black Seas model) for the environments and processes that resulted in the accumulation of banded iron formations (BIFs) and interbedded organic C–rich black shales in the Hamersley Basin during the period 2.7–2.4 Ga. Fe- and silica-rich brine pools were created by locally discharged submarine hydrothermal fluids during the BIF-forming stages. Organic-rich black shales accumulated during quiescent periods.

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A A. BIFs B. VMS C. Greenstones (Submarine volcanism)

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Mechanisms of Iron-Oxide Formation in BIFs

Labrador Trough Circum-Superior Belt Animikie Rift Lake Superior Wopmay Orogen Allik G. and Dubawnt G. Belt-Purcell-Wernecke Basins Grenville Supercrustals N. American Midcontinent Rift System Mckenzie Mt. SG.; Rae G. Windermere; Ekwi SG.

N. America 3.5

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A compilation of the temporal distributions of BIFs and rift systems (Fig. 13B) suggests that the scarcity of BIFs during the 1.5–ca. 0.6 Ga period was due to the paucity of rift basins that matured enough to achieve deep-marine conditions, not due to a change in the atmospheric chemistry. There is a serious misconception in the literature that sedimentary iron ores of Phanerozoic age all belong to the Minnetteand Clinton-type ironstones (Meyer, 1988). These ironstones are characterized by the general absence of chert, and oolitic instead of banded ore textures, as well as by their Fe mineralogy. The iron minerals are characterized by hematite and goethite, chamosite, and siderite, while the magnetite and low-Al iron silicates, such as greenalite or minnesotaite, are scarce (Maynard, 1983). These characteristics suggest that the ironstones were formed by a low-temperature groundwater that acquired aqueous Fe2+ from FeIII-enriched soils (e.g., laterites) by organic acid. The Fe2+ was then precipitated as FeIII-hydroxides when the reduced groundwater was discharged into oxidized environments, such as lakes, seas, and rivers. In contrast, the mineralogy of the Phanerozoic Algoma-type BIFs, especially the abundance of quartz with hematite, as that of the Archean and Proterozoic ones, clearly indicates that they were formed by mixing of locally discharged submarine hydrothermal fluids with O2-rich seawater (Ohmoto et al., 2006b).

0

Figure 13. (A) Temporal distributions of banded iron formations (BIFs), volcanogenic massive sulfide deposits (VMS), and greenstones (submarine volcanism) throughout geologic history (Ohmoto et al., 2006b). (B) Temporal distributions of marine (thick bars) and subaerial (thin bars) rift systems, BIFs (vertical arrows), and red beds (dots). Data from Green (1992). G.—Group; SG.—Supergroup.

have density ~60 °C under H2O-saturated conditions (Ohmoto et al., 2006b). Although Ahn and Buseck (1990) showed that FeIII-hydroxide minerals, and not magnetite (or green rusts), were primary minerals, we cannot determine whether the nucleation of Fe(OH)3 in the Mara Mamba Iron Formation occurred with or without molecular O2. However, during rapid mixing of hydrothermal fluid with seawater at depths below the photic zone, Fe III-(hydr) oxides could only form under conditions when the seawater contained appreciable concentrations (more than ~10 µm/kg H2O) of dissolved O2 (Ohmoto et al., 2006b). Indeed, conducting a detailed investigation of a 3.46 Ga jasper formation (i.e., a low-grade oxide BIF) that most likely formed at seawater depths greater than ~500 m, Hoashi et al. (2009) provided unequivocal evidence (Fig. 15) that the submicron-sized particles of hematite crystals in the jasper nucleated directly from solution, indicating their formation at T > 60 °C by the following reaction: 2Fe2+(hyd.) + 1/2O2(seawater) + 2H2O → [Fe2O3] + 4H+. (R17) Hoashi et al. (2009) also provided photographic evidence (Fig. 16) that magnetite and siderite (more abundant than hematite crystals in all BIFs) did not form during the mixing of hydrothermal fluids and seawater, but by reactions between the earlier hematite crystals and Fe2+-rich fluids in unconsolidated silica- and hematite-rich sediments by reaction Reaction R15. These findings illustrate that, contrary to a popular theory linking the presence of magnetite and siderite in BIFs to a very low-pO2 atmosphere (e.g., Rosing et al., 2010), magnetite and siderite are useful only to indicate the redox state of subseafloor sediments. Basalts associated with the 3.46 Ga jasper beds also exhibit strong enrichment of hematite and increased FeIII/∑Fe ratios, much like those of altered submarine basalts on midocean ridges (Fig. 17). The Re/Os age of 2.76 Ga for pyritebearing veins that crosscut (i.e., younger than) hematite aggregates (Kato et al., 2009) suggest that formation of hematite in these basalts was most likely caused when Fe2+-rich hydrothermal fluids and dissolved O2 in seawater reacted near the seafloor (i.e., Reaction R17). The primary formation of FeIII(hydr)oxides via Reaction R17, which requires the presence

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Figure 14. Transmission electron microscope (TEM) images of a hematite crystal from the 2.6 Ga Marra Mamba banded iron formation (BIF), Hamersley Basin, Western Australia (Ahn and Buseck, 1990). The photos show that a large (~10 µm) hematite crystal is composed of nano-sized hematite crystals and the voids among them, suggesting that the nano-sized hematite crystals were transformed during diagenesis at T > ~60 °C from ferrihydrate (and/ or goethite) crystals that nucleated at T ≈ 25 °C. xls—crystals.

Figure 15. Transmission electron microscope (TEM) images of a single crystal of hematite from the 3.46 Ga Marble Bar Jasper Formation, Pilbara, Australia (Hoashi et al., 2009). The photos show that the hematite is a single crystal, suggesting its nucleation at T > 60 °C during the mixing of Fe2+-rich submarine hydrothermal fluid (T > 100 °C) and O2-bearing deep-ocean water.

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Primary paper | Oxygen, iron, and sulfur geochemical cycles on early Earth

A

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M S 1 µm

C

M

1 µm

D

M H

S

H

S H

Figure 16. (A–B) Scanning electron microscope (SEM) images of hematite (H), magnetite (M), and siderite (S) crystals in the 3.46 Ga Marble Bar Jasper Formation. (C) SEM images of hematite and magnetite crystals after etching with dilute HF solution. Dissolution of magnetite crystals by HF revealed the presence of hematite crystals inside magnetite crystals, indicating that the hematite crystals formed before the magnetite crystals. (D) Transmission microscope photo of siderite crystals with hematite inclusions, indicating the hematite crystals formed before the siderite crystals (Hoashi et al., 2009).

1 µm

of molecular O2 in deep ocean water, therefore indicates that atmospheric pO2 was greater than 50% PAL (Lasaga and Ohmoto, 2002) at 3.46 Ga. Formation of oxide mesobands in the Algoma-type BIFs, ranging in age from 3.8 Ga to the present, is strong evidence for the atmospheric pO2 to have been greater than ~50% PAL since ca. 3.8 Ga.

-600

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Figure 17. (A) Increases in the Fe3+/∑Fe ratios of submarine basalts from a drill core at Ocean Drilling Program (ODP) Site 735, SW Indian Ridge (Bach et al., 2001), indicating oxidation of the submarine basalts due to a reaction with O2- and SO42–-rich deep-sea water (mbsf— meters below seafloor). (B) Fe3+/∑Fe values of 3.46 Ga submarine basalts and jasper beds in the ABDP #1 drill core, Pilbara, Australia (Ohmoto, unpublished data). The data indicate the abundant formation of hematite due to the reaction of Fe2+-rich submarine hydrothermal fluids with O2- and SO42–-rich deep-sea water.

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Linking Siderite to Oceanic and Atmospheric Chemistry The CO2 content of the early atmosphere has been a major research topic of geoscientists and astrobiologists, because CO2, as well as H2O, is essential for the generation of life and maintenance of habitable environments. Based on climate modeling of early Earth when the solar luminosity was presumably much less than today (i.e., the faint young Sun paradox), Owen et al. (1979) and Kasting (1987) suggested that the pCO2 of the early atmosphere was much higher than today (~300 ppm or 10–3.5 atm for the preindustrial atmosphere), such as ~1 atm (~3000 PAL) at 4.5 Ga, ~0.1 atm (~300 PAL) at 2.5 Ga, and ~10–2 atm (~30 PAL) at 1.0 Ga. However, based on the absence of siderite in pre–2.2 Ga paleosols and the application of thermodynamic data for reactions involving siderite, Rye et al. (1995) suggested that the pCO2 of the pre–2.2 Ga atmosphere was only about ~10 PAL, and the remaining greenhouse gas, necessary to counter the faint young Sun paradox, was CH4. Kasting and his associates (Kasting and Siefert, 2002; Pavlov et al., 2001) suggested levels of ~1000 ppm CH4 and ~3000 ppm CO2 (~10 PAL) for the pre–2.2 Ga atmosphere. The atmosphere cannot contain high concentrations of both O2 and CH4 because they react to form CO2. Therefore, the idea of a methane-rich Archean atmosphere was thought to be another line of evidence to support paradox I (i.e., a reducing atmosphere before ca. 2.2 Ga). However, subsequent researchers (Haqq-Misra et al., 2008) have recognized that the greenhouse effect of CH4 is much smaller than previously thought: At the same concentration level, CH4 is found to be only ~3 times more effective as a greenhouse gas than CO2, whereas it was previously thought to be ~20 times more effective than CO2. Increasing the CH4 concentration to levels

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necessary to counter the faint young Sun paradox would have created a haze, which would block sunlight and cool Earth’s surface. As a result, the idea of a CH4-rich Archean atmosphere has been subdued. Paradigm II-D: Siderite as an Indicator of Fe-Rich Archean Oceans Siderite, as well as magnetite, in BIFs are primary minerals, indicating Fe-rich oceans prior to ca. 1.85 Ga. Contradiction II-D.1: Abundance of Fe-Poor Archean Carbonates If the Archean oceans were Fe2+ rich (i.e., >10−4 m) as postulated by the current paradigm, we would expect all carbonates formed in Archean oceans to be Fe rich. Yet, Holland (1984) has recognized that many Archean carbonates are Fe poor, just like younger carbonates. For example, the average FeO content and (FeII/Ca) mole ratio of seven Archean limestone samples in the Russian Shield are 2.73 wt% and 0.07, respectively (Ronov and Migdisov, 1971). Based on the analyses of nine samples of stromatolite-bearing limestones of the 2.74 Ga Tumbiana Formation in Pilbara, Western Australia, we have estimated the carbonates have an average FeO content of 1.0 wt% and an Fe/Ca mole ratio of 0.02 (Fig. 18). These values of Archean carbonates are essentially the same as those of Neoproterozoic limestones in the Russian Platform (FeO = 1.63 wt%; Fe/Ca = 0.05) and the average composition of >1000 samples of Phanerozoic-aged limestones from the Russian and North American Platforms (FeO = ~0.5 wt%; Fe/Ca = ~0.01; Ronov and Migdisov, 1971). These data suggest that the formation of siderite (Fe/Ca > 1) was prob-

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ABDP#10 y = -3.4018x + 46.244 R = 0.9212

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1.0 0 0

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Figure 18. Plots of FeO and CaO content versus Al2O3 content of carbonate samples from the 2.74 Ga Tumbiana Formation, Pilbara, Western Australia (Ohmoto, unpublished data). These samples contain silica and clay minerals, as well as calcite. Therefore, the intersection of the linear regression lines at Al2O3 = 0 indicates the calcite in these samples has compositions of FeO = 1.03 wt%, CaO = 46.24 wt%, and an Fe/Ca mole ratio of 0.02.

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ably restricted to special environments with Fe2+-rich solutions, such as in/under Fe2+-rich brine pools. Contradiction II-D.2: Formation of Siderite during Diagenesis of Sediments While many researchers (e.g., Rosing et al., 2010) have assumed that siderite and magnetite in BIFs formed in equilibrium with the contemporaneous atmosphere, no one has presented petrographical or geochemical evidence to support such an assumption. In an earlier section, we presented a line of evidence suggesting that magnetite in BIFs was formed by a reaction between the primary FeIII-bearing minerals and later Fe2+rich solutions within sediments; therefore, magnetite was not in equilibrium with the contemporaneous atmosphere. Here, we present a line of evidence indicating that siderite in massive carbonate beds is also a diagenetic product; thus, it cannot be linked directly to atmospheric chemistry. The 2.6 Ga Wittenoon Formation in the Hamersley Basin, Australia, and the 1.9 Ga Gunflint Formation in Ontario, Canada, are two of the best-known examples of carbonate formations that were deposited in BIF-hosting basins. The former is a 150–200-m-thick predominantly dolomite unit that occurs immediately above the Mara Mamba Iron Formation and below the Mount Sylvia Formation (mostly black shale) throughout the Hamersley Basin (Fig. 10). The latter is a representative of carbonate-BIF, which may have formed nearly contemporaneously with the oxide-type BIFs in Minnesota and Michigan (Ohmoto et al., 2006b). It is a 100–220-m-thick unit, overlain by a thick black shale formation (the Robe Formation), extending over a 100,000 km2 area. We have carried out microscale analyses of the chemical and isotopic compositions of >300 samples from these two carbonate formations (Figs. 19A–19C). An important finding from our study is that these carbonate formations are very heterogeneous in their chemical (Fe/Ca/ Mg/Mn ratios) and isotopic compositions (δ13C and δ18O) on an approximately millimeter scale. The finding indicates that not all carbonate grains precipitated directly from a chemically homogeneous water body; therefore, their bulk-rock compositions do not represent those of the ocean water. Microscopic investigation of these samples (Kumazawa, 1997) has revealed that Fe-rich carbonates (siderite and ankerite) generally formed after the precipitation of Fe-poor carbonates (calcite and dolomite) during diagenesis. The large magnitudes of microscale variations in the chemical and isotopic compositions of the carbonates suggest that Fe, Mg, Ca, and Mn contents of the diagenetic pore fluids also varied widely on a millimeter scale due to reactions with the precursor carbonates (calcite and dolomite). A probable source of these Fe- and Mn-rich diagenetic fluids was the overlying brine pools that most likely existed in the BIF-forming basins (see Fig. 12B). The δ13C values of kerogen in the Wittenoon and Gunflint carbonates fall within a range of −30‰ ± 5‰ (Kumazawa, 1997). Therefore, the broad, but general trend of decreasing δ13C values, from ~0‰ to −15‰, with increasing Fe content (Fig. 19B) could

Primary paper | Oxygen, iron, and sulfur geochemical cycles on early Earth indicate that the carbonate C came from two sources. The first was normal seawater HCO3– with δ13C ≈ 0‰. The second was HCO3– with a δ13C ≈ −30‰ that was derived from the oxidation of organic carbon, probably by SRB (see Reaction R18), during the early diagenesis of the carbonate-rich sediments (Ohmoto et al., 2004): SO42– + 2H+ + 2(NRC-CH2O) → H2S + 2CO2 + 2H2O + 2NRC,

A

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in which NRC refers to nonreactive C, and CH2O is reactive organic matter. The H2S then reacts with reactive Fe (probably Fe2+-rich brine fluids) to form pyrite (FeS2), or is incorporated in organic matter as organo-S compounds, or diffuses out from the sediments. Reaction R18 is the reason for the existence of positive correlations between the residual organic C contents (i.e., NRC) and pyrite contents found in many C-rich black shales of all geologic age (Ohmoto and Goldhaber, 1997).

100 1.9 Ga Gunflint 2.6 Ga Wittenoon 2.6 Ga Munjina Bee Gorge Rio Tinto, Hamersley

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Figure 19. (A) Chemical compositions of carbonates from the 2.6 Ga Wittenoon Formation in the Hamersley Basin, Australia, and the 1.9 Ga Gunflint Formation, Ontario, Canada. (B) C isotopic compositions and Fe contents of carbonates in A. The CO32– in the carbonates is interpreted to represent various mixtures of normal marine HCO3– (δ13C = 0‰) and HCO3– from the oxidation of organic C (δ13C = 0‰); the percentages of organic-derived HCO3– are indicated for 10%, 20%, and 40%. (C) C and O isotopic compositions of the carbonates. The δ18O values of calcite in equilibrium with seawater bearing δ18O = 0‰ are indicated at T = 25 °C and 75 °C. Data are from supplementary information of Ohmoto et al. (2004).

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Decomposition of organic matter by fermenting bacteria (e.g., methanogens) would produce CH4 as well as CO2 through: CH2O → CH4 + CO2.

(R19)

CH4 produced by this reaction is isotopically very light (δ13C ≈ −60‰), and CO2 is very heavy (δ13C ≈ +10‰; Ohmoto et al., 2004). The addition of the heavy CO2 (and HCO3–) to normal seawater HCO3– would have caused the pore-water HCO3– to have positive δ13C values. Therefore, the negative δ13C values of these carbonates (Figs. 19B and 19C) are evidence that the seawater was SO42– rich and that SRB played an important role in the formation of Archean siderites during sediment diagenesis. We conclude that siderite in Archean sedimentary rocks, as well as in younger rocks, was not in chemical equilibrium with the contemporaneous atmosphere. Paradigm II-E: The Absence of Siderite in Paleosols Defines the Maximum pCO2 Siderite would have formed in pre–2.2 Ga paleosols if the atmospheric pCO2 was higher than ~10 PAL. Contradiction II-E: Thermodynamic Analyses of Multicomponent Systems and Kinetics of Redox Reactions In order to constrain the chemistry of fluid by applying thermodynamic data to natural rocks, it is essential to first determine

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log pCO2 (atm) Figure 20. Stability fields of Fe-rich minerals as a function of pCO2 and pO2 at T = 25 °C and SiO2(aq) = 10–3.4 moles/kg H2O (24 ppm) (solid lines). Dashed lines are the stability fields of minerals in an H2O-free system. Note that the stability fields of fayalite and magnetite disappear and the siderite field shrinks with the addition of H2O and SiO2(aq). The greenalite/siderite metastable boundary occurs at log pCO2 about −2.5. Based on the pCO2 value (>100 PAL) and carbon isotope data of the diagenetic fluids of banded iron formations (BIFs), we estimate that the pre–1.8 Ga atmospheric pCO2 was greater than ~80 PAL. Figure is modified after Ohmoto et al. (2004).

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an

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Archean atmosphere

log pO2 (atm)

which parameters (e.g., T, pH, fO2, fCO2, concentrations of aqueous species, and mineral phases) are important, which chemical reactions are likely to attain equilibrium, and which are not. For example, Rye et al. (1995) assumed, without presenting any justification, that only T, pCO2, and aSiO2(aq) determined the presence or absence of siderite in paleosols. However, if we also include pO2, pH, and aFe2+ as other important variables, it becomes obvious that siderite becomes unstable at pO2 > 10−60 atm, regardless of the pCO2 value (Fig. 20), as well as in the pH and aFe2+ range of the Archean surface waters, regardless of the pO2 and pCO2 values (Fig. 21). That is, the absence of siderite in paleosols cannot constrain the atmospheric pCO2 value, whereas the presence of siderite can constrain the diagenetic fluid composition; the pCO2 of a diagenetic fluid can be used to reduce the atmospheric pCO2 by considering the carbon isotopic composition of siderite. Rosing et al. (2010) constrained the CO2/H2 ratio of the Archean atmosphere by assuming chemical equilibrium among CO2, CH4, H2, and H2O at ambient temperatures:

Log(aCa2+)

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=1

00

PA L

0P AL

Modern Rivers Modern Oceans

-10 4

5

6

7

8

9

pH Figure 21. Solubility diagrams for (A) calcite and (B) siderite as a function of pCO2, pH, aCa2+, and aFe2+ at 25 °C. We estimate that before ca.1.8 Ga, the atmospheric pCO2 was 100–1000 PAL; the resulting pH values were 4.0–4.5 for rainwater, 4.5–6.0 for river water, and 6.5–7.5 for ocean water. The Archean ocean water was saturated with calcite but undersaturated with siderite. The estimated concentrations in the pre–1.8 Ga ocean water were 10−5–10–2 m for Ca2+ and 0.03 atm or >~100 PAL) was of atmospheric origin, because atmospheric pCO2 is related to the HCO3– con-

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tent of normal seawater by the solubility constant. That is, the pCO2 value for the pre–1.8 Ga atmosphere was greater than ~100 PAL. Our conclusion suggests that CO2 was the main greenhouse gas acting to counter the faint young Sun paradox in the Archean (Ohmoto et al., 2004). The pH and Fe2+ Content of Archean Oceans The pH and concentrations of Ca2+ and Fe2+ in early oceans may be constrained from evaluations of the following carbonate equilibria: Ca2+ + CO2(g) + H2O = CaCO3(calcite) + 2H+

(R23)

Fe2+ + CO2(g) + H2O = FeCO3(siderite) + 2H+.

(R24)

and

These reactions indicate that the concentration of Ca2+ (or Fe ) in solution in equilibrium with calcite (or siderite) is a function of pH and pCO2. Figure 21A shows that: (1) modern river water (pH = 5.5–7; mCa2+ = ~10−4) has an increased pH (i.e., decreased concentration of H+) relative to rainwater (pH = ~5.5) and increased solute concentrations due to water-rock interaction during soil formation; (2) modern rainwater, soil water, and river waters are undersaturated with respect to calcite; (3) the Ca2+ content of ocean water is ~30 times higher than that of river water due to the hydrological cycle of H2O involving evaporation of ocean water, condensation of water vapor to form rain, and reactions between rainwater and rock; and (4) modern ocean water (pH = 8.1 and mCa2+ = 0.01) is slightly oversaturated with respect to calcite. The common occurrence of limestone in Archean sedimentary formations suggests that the pH–mCa2+ values of Archean ocean water were near the calcite saturation line at pCO2 = 100– 1000 PAL (Fig. 21A). If the pCO2 of the Archean atmosphere was 100–1000 PAL, the pH of the rainwater would have been 4.0–4.5 (Ohmoto et al., 2004); the pH of Archean river water was possibly in the range of 4.5–6.0. Lasaga and Ohmoto (2002) have proposed that the average concentration of a non-redox-sensitive element (Ci) in river waters at a particular geologic time can be related to the atmospheric pCO2 by: 2+

(Cit / Ci0) = (pCO2t/ pCO20)0.25,

(Eq. 1)

where superscripts t and 0 refer, respectively, to the values at time t and today. It follows that the average concentration of Ca2+ in Archean rivers would have been ~3–6 times higher than the present value: therefore ~(1–3) × 10−3 moles/kg H2O. Assuming the mCa2+(seawater)/mCa2+(river water) ratio in the Archean was between 10 and 100, we can estimate that the pH of the Archean ocean was probably within a range of 7.0 ± 0.5 (Fig. 21A).

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If the Archean Earth had been an entirely anoxic world, reactions Reactions R2 and R3 would not have been important. Therefore, during the weathering of rocks under an anoxic atmosphere, Fe would have dissolved in soil waters similar to the way in which Mg does (Holland, 1984). The average Mg content of river water is ~4 ppm (Holland, 1984). Considering the FeO/ MgO ratios of average crustal rocks, Holland (1984) estimated that the average Fe2+ concentration of Archean river waters was ~7 ppm. Applying Equation 1, we can estimate the average Fe2+ content of anoxic Archean river waters to be ~20–60 ppm (i.e., ~10–3.2 ± 0.2 m). Even with such a high concentration of Fe2+, the solution would be undersaturated with siderite because of low pH (Fig. 21B). If the Archean atmosphere was oxygenated, then the Fe2+ contents of the surface water would have been ~10−7 m, which is below the siderite saturation. Therefore, irrespective of the pO2 value of the Archean atmosphere, siderite would not have been stable in surface waters (see Fig. 21B). The formation of low-Fe carbonates from normal Archean ocean water (see previous) suggests that the Archean oceans were undersaturated with respect to siderite in a pH range of 7.0 ± 0.5. This would place the maximum Fe2+ concentration of the Archean oceans to be ~10−4 m, i.e., less than ~6 ppm.

cal sedimentary rocks (e.g., carbonates, cherts, phosphates) depend on the δ18O of Archean seawater. The δ18O values of carbonates from the Wittenoon and Gunflint Formations generally fall in a range of −20‰ and −25‰ (Fig. 19C), suggesting that these carbonates attained an isotopic equilibrium with the water of δ18O = ~0‰ at T = ~30–60 °C. These temperature values are reasonable for diagenetic/hydrothermal conditions. Our data, therefore, support the theory (Muehlenbachs and Clayton, 1976) that the δ18O of ocean water has remained essentially constant by high-temperature (200–300 °C) seawater–rock interaction at mid-ocean ridges throughout geologic history. These discussions indicate that siderite in BIF-forming environments most likely formed during the diagenetic stage, possibly involving Fe-bearing submarine hydrothermal fluids. The pO2 values of the diagenetic/hydrothermal fluids were decreased dramatically from the atmospheric value by the consumption of dissolved O2 and the production of H2 by reactions with organic matter and Fe2+-rich rocks. However, the pCO2 of the diagenetic fluids increased above the atmospheric value by 10%–30%. SULFUR CYCLE AND FORMATION OF PYRITE

Oxygen Isotopic Composition of Archean Seawater

Geochemical Cycle of Sulfur and Formation of Pyrite during the Phanerozoic Eon

There has been a debate about whether the δ18O values of the Archean oceans were essentially the same as today (Muehlenbachs and Clayton, 1976; Knauth, 2005) or much lighter, such as ~-20‰ (Kasting et al., 2006). This question is important because the ocean temperatures estimated from the δ18O values of chemi-

Figure 22 shows the Phanerozoic geochemical cycle of sulfur through the atmosphere–continental crust–oceans–oceanic crust– mantle reservoirs. Sulfur in the crust occurs mostly as sulfates (mostly gypsum/anhydrite) in evaporites (250 × 1018 moles), sulfides (mostly pyrite) in sedimentary rocks (250 × 1018 moles), and

Mass: in unit of 1018 moles Flux: in unit of 1012 moles/yr

SO2 (Vol. Gas)

Fv

, SO

FeS2

2

= 0.3

Fppt, py = 0.5

(Sed. Rocks)

MS = 250

Fw, py = 0.5

Fw, gyp = 1.0

Oceans SO42(900 ppm)

(Sed. Rocks)

Fppt, gyp = 1.0

MS = 40 FeS2 (Ign. Rocks)

FOC, py = 0.3

MS = 300 Upper Mantle FeS

FOC, py = 0.3

Oceanic Crust FeS2

(200 ppm)

(1000 ppm)

MS = 4000

MS = 200

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CaSO4 MS = 250

Figure 22. Modern geochemical cycles of sulfur. The flux values are from Lasaga et al. (1985); the S contents and masses for the reservoirs are estimated from the data in Henderson and Henderson (2009).

Primary paper | Oxygen, iron, and sulfur geochemical cycles on early Earth sulfides (mostly pyrite and pyrrhotite) in igneous rocks (~300 × 1018 moles; Lasaga et al., 1985; Ohmoto and Goldhaber, 1997). Modern oceans are rich in SO42– (28 mM or 900 ppm; ΣS = 38 × 1018 moles) because of the input from the oxidative weathering of sulfides (Fpy,weathering = 0.5 × 1012 moles/yr), weathering of gypsum/anhydrite (Fgyp,weathering = 1 × 1012 moles/yr), and the oxidation of volcanic SO2 gas (FSO ,v = 0.3 × 1012 moles/yr; Lasaga et 2 al., 1985). Seawater sulfate is in turn removed as biogenic pyrite in marine sediments (most importantly in shales), gypsum in evaporating basins, and as pyrite in mid-ocean ridges; their fluxes are estimated to be the same as the weathering flux of pyrite, the weathering flux of evaporite, and the volcanic flux of SO2, respectively (Lasaga et al., 1985). The residence time of SO42– in the oceans is ~10 m.y. (Henderson and Henderson, 2009). The following three components are required for the formation of syngenetic and diagenetic pyrite in Phanerozoic-aged sedimentary rocks: (1) SO42–, (2) reactive organic matter, and (3) reactive Fe. Items 1 and 2 are used to generate H2S by BSR (i.e., R16). The H2S then reacts with reactive Fe to form pyrite (FeS2). However, the reaction pathways are complicated. In most Phanerozoic sediments, reactive Fe resides almost exclusively in FeIII-(hydr)oxides that formed during oxidative weathering of FeII silicates on land (i.e., R2). Because the dissolution rates of FeII silicates in seawater (pH ≈ 8) are much slower than those in rainwater (pH ≈ 5.5), FeII silicates are not primary sources of reactive Fe for the formation of pyrite (Raiswell and Canfield, 1996). Instead, FeIII-(hydr)oxides decompose to generate Fe2+ in sediment pore water by FeIII-reducing bacteria and/or by organic acids generated from organic matter in the sediments (cf. R3). The first mineral phase that forms from the reaction between Fe2+ and H2S (or HS–) is generally termed “iron monosulfides (FeS),” Fe2+ + H2S → “FeS” + 2H+,

(R25)

where “FeS” includes mackinawite (Fe1+xS), pyrrhotite (Fe1–xS), and greigite (Fe3S4). “FeS” is then converted to FeS2 by: (1) reaction with native S (S8, simplified as S0) that was formed by the oxidation of H2S by O2 in the overlying seawater (Reactions R26 and R27) or (2) a reaction with additional H2S and the loss of H2 (Reaction R28): H2S + 1/2O2 → S0 + H2O.

(R26)

“FeS” + S0 → [FeS2].

(R27)

“FeS” + H2S → [FeS2] + H2↑.

(R28)

The overall reaction involving reactive FeIII–(hydr)oxides and H2S is therefore [FeIIIOOH] + 2H2S → [FeS2] + 2H2O + 1/2H2↑. (R29)

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Reactions R28 and R29 are promoted by the consumption of H2 by O2 in the overlying water column, or by methanogens and other chemolithotrophic microbes in anoxic water/sediments, such as: CO2 + 2H2 → CH2O + H2O.

(R30)

Reactions R26–R30 indicate that the formation of pyrite, which occurs in anoxic water bodies and sediments, is favored near the redox boundary with an overlying oxygenated water body. Furthermore, because the reactive FeIII-hydroxides are formed under an oxygenated atmosphere, we can conclude that the high abundance of pyrite in Phanerozoic-aged sediments is a reflection of an O2-rich atmosphere. Geochemical Cycle of Sulfur and Formation of Pyrite during the Archean Eon Reactions R25 and R28 indicate that pyrite could have formed in a totally anaerobic world (i.e., stage I in Fig. 1) if sources of reactive Fe and H2S were available. For example, reactive Fe could have been Fe2+-rich fluids (e.g., hydrothermal fluid) rather than FeIII-(hydr)oxides. H2S could have been hydrothermally derived, generated in situ by BSR or thermochemical sulfate reduction (TSR) of seawater SO42–, or generated from atmospherically derived S0. An important question is which processes formed pyrite in Archean-aged sedimentary rocks. Paradigm IV: SO42–-Poor Archean Oceans Another paradigm is that the SO42– contents of the Archean oceans were less than 1/200 of the present level (i.e., +6‰ for FeS2

nd wa

A

δ56Fe (‰)

The discoveries of AIF-S–bearing pyrite from the hydrothermally altered, organic C–rich 1.8 Ga black shales from Outokumpu, Finland (Young et al., 2011), and of the Paleoproterozoic-aged Mississippi Valley–type deposits (Kim et al., 2009) support our hypothesis linking AIF-S signatures to reactions involving SO42–-rich hydrothermal solutions and in situ solid/amorphous organic matter during sediment diagenesis, rather than to UV photolysis of volcanic SO2 in a reducing atmosphere. However, except for these samples from hydrothermal deposits, AIF-S signatures in submarine hydrothermal terrains younger than ca. 1.8 Ga in age have not been reported. An important question is why? An obvious answer is that we have not yet examined enough samples from different types of hydrothermal deposits of various ages. Another may be related to the evolution of the core-mantle-crust geodynamics (see following). The Hamersley Province was a part of perhaps the world’s largest and longest-lived igneous province in geologic history. The igneous activity began with the eruptions of a series of large subaerial flood basalts including the Mount Roe Basalt (ca. 2.775 Ga), the Kylena Formation (2.740 Ga), and the Maddina Formation (2.717 Ga), all of which were followed by the deepening of the rift basin and the accumulation of an ~2-kmthick succession of shales and BIFs between 2.69 Ga to 2.41 Ga (Fig. 10). The presence of numerous dolerite sills in the shale/ BIF succession (Trendall and Blockey, 1970) indicates that the Hamersley Basin was the site of intermittent igneous and hydrothermal activity for more than 300 m.y. No other place in the world is known to have had such a long history of a high thermal anomaly. Such a large-scale and long-lived igneous province may have been the result of thicker lithosphere and slower plate movements during the Archean, as suggested by the Korenaga model for early Earth (Korenaga, 2008). Thus, a long-term thermal anomaly may have been required to produce such a large AIF-S signature by hydrothermal-diagenetic processes, rather than SO2 photolysis in an anoxic Archean atmosphere.

0.1

1

10

100

1000

Fe (ppm)

Figure 30. (A) Equilibrium Fe isotope fractionation factors (δ56Fe) between aqueous Fe2+ and Fe-bearing minerals at 25 °C. See text for the data source. (B) Correlation between the concentration and isotopic composition of aqueous Fe2+ in natural waters, modified from C.M. Johnson et al. (2008).

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concentrations of Fe2+ (Fig. 30B). A similar, but more pronounced trend of decreasing δ56Fe values with decreasing Fe2+ concentrations is found in groundwaters (Fig. 30B). These trends indicate that a major process for decreasing δ56Fe values of Fe2+ in natural solutions is the Rayleigh isotope fractionation caused by removal of isotopically heavier FeIII-(hydr)oxides from solution. If the atmospheric pO2 is decreased to below ~50% of the present atmospheric level, ocean waters below the photic zone will become anoxic (Lasaga and Ohmoto, 2002). If there were no dissolved O2 in the Archean deep-oceans, no FeIII-(hydr)oxides would have formed near deep-sea submarine hydrothermal vents. Consequently, the δ56Fe values of Fe2+ in the discharging submarine hydrothermal fluids and the Archean oceans should have been very close to 0‰. If the FeIII-oxides, siderite, and pyrite in BIFs formed in basins that were open to the ocean, and if the ocean water had constantly supplied Fe2+ to form Fe-rich miner-

als, as proposed by the current paradigms for BIFs, then the δ56Fe values of the oxides, carbonates, and pyrite that formed in the basin should have been very constant at ~+1‰, ~+0.5‰, and ~+0.3‰, respectively (see Fig. 30A). Therefore, the current BIF paradigms cannot explain the observe large variation in the δ56Fe values with distinct trends toward negative values in BIF-forming environments (Fig. 29A). The observed Fe isotope data of BIFs, however, provide strong support for the Red and Black Seas model for the formation of the Hamersley-Superior–type BIFs (Figs. 12 and 31). Precipitation of FeIII-oxides from a brine pool during a high-temperature stage would have caused the δ56Fe value of residual Fe2+ in the brine pool to become progressively more negative with time (Fig. 29A). Thus, the siderite and pyrite that formed in organic-rich sediments during the lower-temperature stage of the brine pools would have acquired very negative δ56Fe values (Fig. 29A).

Figure 31. Ohmoto model for the environments and processes for the formation of volcanogenic massive sulfide (VMS) deposits and banded iron formations (BIFs) through geologic time (from Ohmoto et al., 2006b). MOR—mid-ocean ridge; IF—iron formation.

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Primary paper | Oxygen, iron, and sulfur geochemical cycles on early Earth Judging by the sample descriptions, we interpret that the post–2.0 Ga pyrite samples (δ56Fe = 0‰ to +1.0‰) analyzed by Rouxel et al. (2005) utilized FeIII-(hydr)oxides that formed through oxidative weathering of the source rocks. In contrast, the 2.4–2.6 Ga pyrite samples they analyzed are predominantly pyrite nodules like those in Figure 11, which most likely formed by reactions between Fe2+-rich hydrothermal fluid and H2S generated by TSR. Therefore, neither of the δ56Fe values of pre– and post–2.2 Ga pyrites in Figure 29A can be linked to that of Fe2+ in normal ocean water. Contradiction III-B.2: Similarity in Fe Isotopic Composition between Archean and Post-Archean Shales By analyzing the chemical composition and Fe isotope ratios of 120 sediment samples of 3.3–2.2 Ga in age, Yamaguchi et al. (2005) determined that the bulk-rock δ56Fe values of Archean rocks are close to 0‰, like those of Phanerozoic shales (Fig. 29A). Based on this finding, they have concluded that the Fe in these rocks has behaved conservatively during soil formation, sediment transport, diagenesis, and lithification since the Archean. This is strong evidence supporting the theory that the Fe geochemical cycle has remained the same as today since at least 3.3 Ga (Fig. 4). SUMMARY AND CONCLUSIONS Recent literature regarding the evolution of the atmosphere, hydrosphere, and biosphere on early Earth has been strongly influenced by six major paradigms for the early Earth: (I) the Great Oxidation Event at ca. 2.4 Ga; (II) Fe-rich oceans; (III) hydrothermal origin for the oceanic Fe; (IV) SO42–-poor oceans; (V) atmospheric origin for the oceanic sulfur species; and (VI) sulfidic Proterozoic oceans. At first glance, these six paradigms appear to be supportive with each other. However, they are actually contradictory with each other. For example, paradigm IV was developed from an interpretation of the δ34S relationships between barite and pyrite in the oceans with a fundamental assumption that pyrite formed by utilizing H2S generated by bacterial sulfate reduction (BSR) of seawater SO42–. Conversely, paradigm V was developed under the assumption that the S isotopic characteristics of Archean pyrites were inherited from native sulfur created in the atmosphere by UV photolysis of volcanic SO2. If paradigms I, II, and V were true, Fe-rich minerals (especially pyrite and siderite) should be found abundantly in all types of sedimentary rocks, especially those formed in shallow seas (e.g., carbonates, evaporites, cherts). However, these Archean-aged chemical sediments are poor in Fe-rich minerals, much like those of younger age. Paradigm I has also been reinforced by other paradigms for the Archean systems, concerning: (1) the Fe geochemistry of paleosols; (2) the temporal distribution of red beds; (3) the Fe geochemistry of shales; (4) the temporal distribution, geological environments, and geochemical processes behind

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the formation of Fe-oxides, Fe-carbonates, and pyrite in BIFs; (5) the only known process of AIF-S generation (until recently) in geologic samples; and (6) the processes behind barite mineralization. However, we have recognized that all these paradigms have serious flaws because they contradict with the geological, mineralogical, and geochemical data on natural samples of Archean and younger rocks, and/or experimental and theoretical data on the solubility, stability, and isotopic fractionations of S- and Febearing compounds (e.g., siderite, polysulfides, monosulfides, pyrite, organic compounds, etc.). In contrast, the mineralogical and geochemical data on natural samples of Archean age can be better explained by the Dimroth-Ohmoto model for the evolution of the atmosphere, hydrosphere, and biosphere. This model postulates the development of a fully oxygenated world and present-day geochemical cycles of Fe, S, C, and other redox-sensitive elements by 3.5 Ga. These data include, but are not restricted to, the findings of the following mineralogical and geochemical characteristics in Archean- and younger-aged rocks: 1. the scarcity of primary siderite and pyrite in chemical sediments formed in shallow oceans (e.g., limestone, dolomite, chert, evaporites); 2. the scarcity of detrital grains of pyrite, siderite, and uraninite in normal marine sedimentary rocks (carbonates, chert, evaporites, shale, sandstone); 3. the scarcity of detrital grains of kerogen in normal marine sedimentary rocks; 4. the enrichment (and depletion) of FeIII in paleosols; 5. the depletions of pyrite, siderite, and uraninite in paleosols; 6. the presence of red beds; 7. the enrichment (and depletion) of FeIII in shales; 8. the abundance of primary (syngenetic) FeIII-(hydr) oxides in deep-sea, as well as in shallow-sea submarine hydrothermal deposits, including the Algoma- and Superior-type BIFs, and submarine hydrothermal alteration zones; 9. the abundance of syngenetic/diagenetic siderite in organic C–rich sedimentary rocks; 10. the abundance of syngenetic/diagenetic pyrite in organic C–rich shales; 11. the presence of AIF-S signatures in organic C–rich sedimentary rocks; 12. the large variations in δ56Fe values of oxides, siderite, and pyrite; and 13. the correlation between the magnitude of AIF-S signatures in pyrite and the magnitude of hydrothermal effects on organic C–rich shales, which suggests the involvement of SO42–-rich hydrothermal fluids in the generation of AIF-S by TSR. Other lines of evidence supporting the D-O model, which were not included in this paper are: 14. the abundance of barite in sedimentary rocks and in submarine hydrothermal systems (Ohmoto, 1992);

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Ohmoto et al. 15. the enrichment of U with FeIII-oxides in submarine hydrothermal systems, indicating the enrichment of U in the Archean oceans due to the oxidative weathering of rocks (Kerrich and Said, 2011; Manikyamba et al., 2012; Ohmoto et al., 2011); 16. the enrichment of uranogenic Pb isotopes relative to thoriogenic Pb isotopes in hydrothermally altered submarine volcanic rocks, suggesting the enrichment of U in the Archean oceans (Ohmoto et al., 2011); 17. the enrichment of Mo with FeIII-oxides in submarine hydrothermal systems, indicating the enrichment of Mo in the Archean oceans due to the oxidative weathering of rocks (Ohmoto et al., 2010, 2011); 18. the variation in Cr isotopes associated with FeIII-oxides in BIFs, indicating the enrichment of Cr in the Archean oceans due to the oxidative weathering of rocks (Frei et al., 2009); 19. the ubiquitous presence of positive and negative Ce anomalies in oxide-type BIFs and hydrothermally altered submarine volcanic rocks, indicating oxygenated oceans (Kato et al., 2006; Kerrich and Said, 2011); 20. the similarities in the δ13C values of carbonates (~0‰) and organic C (~–30‰) in Archean sedimentary rocks with those of younger age, suggesting the early development of the C geochemical cycle that was linked with an O2-rich atmosphere (Schidlowski and Aharon, 1992); and 21. the similarity in the δ15N values of kerogen (~0‰ to ~+10‰) in Archean sedimentary rocks with those of younger ages, suggesting the early development of the N geochemical cycle that was linked with an O2-rich atmosphere (Pinti et al., 2007). Other important suggestions made in this chapter are: (A) The geochemical cycles of O, Fe, and S (and other redoxsensitive elements) through the atmosphere–ocean–oceanic crust–mantle–continental crust have been basically the same as today since ca. 3.5 Ga. (B) The terrestrial biosphere, as well the marine biosphere, developed by ca. 3.5 Ga, influencing the geochemical cycles of nutrients and other elements. (C) The geochemistry of sedimentary rocks, including shales, red beds, cherts, evaporites, and carbonates, has been basically the same throughout geologic history. (D) FeIII-oxides in BIFs were formed by reactions between locally discharged Fe2+- and silica-rich submarine hydrothermal fluids and O2-rich deep seawater (Fig. 31). (E) Magnetite in BIFs was formed during the hightemperature diagenetic stages of BIFs through reactions between primary goethite/hematite and Fe2+-rich hydrothermal fluids. (F) BIFs were formed throughout geologic history. (G) Sulfidic oceans (i.e., the “Canfield ocean”) did not exist during the Proterozoic Eon. However, regional sulfidic seas, like the Black Sea, have existed in globally oxygenated oceans throughout geologic history.

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(H) The primary carbonate in Archean oceans, as in younger oceans, was Fe-poor calcite. (I) The pre–1.8 Ga atmosphere was CO2 rich with the pCO2 level greater than ~100 PAL. CO2 alone provided the greenhouse effect necessary to compensate for the young Sun’s lower luminosity. (J) The Archean pH values were 4.0–4.5 for rainwater, 4.5– 6.0 for river water, and 7.0 ± 0.5 for ocean water. The oceans were saturated with calcite but undersaturated with siderite. (K) The δ18O of Archean oceans was ~0‰, as today. (L) Fe-rich carbonates (siderite, ankerite) have formed during the diagenesis of sediments throughout geologic history by reactions between the primary calcite and Fe2+-rich solutions, either hydrothermal solutions or those derived from biological/abiological dissolution of FeIII-(hydr) oxides within the sediments. (M) The ranges of δ34S values of pyrite and sulfates in Archean sedimentary rocks are comparable with those of Proterozoic age. (N) Pyrites in organic C–rich black shales associated with BIFs were formed by a reaction between Fe2+- and SO42–-rich hydrothermal solutions and organic C–rich shales during early diagenetic stages of the host sediments. This reaction also created anomalous isotopic fractionation of sulfur (AIF-S) in the pyrite and the residual SO42–. (O) The AIF-S signatures in Archean and younger rocks were not created by the UV photolysis of volcanic SO2 in a reducing atmosphere. AIF-S signatures are not evidence for a reducing atmosphere. (P) Contrary to a popular belief that AIF-S–forming events ceased at ca. 2.45 Ga, AIF-S was also formed at later geologic times. (Q) The presence of AIF-S in some pre–2.4 Ga rocks, but the lower abundance of AIF-S in post–2.4 Ga rocks, may reflect changes in the core-mantle-crust dynamics, including changes in the thickness and movements of oceanic lithosphere. In an effort to link the geochemical data from Archean rocks to the (bio)geochemical nature of the Archean world, we have learned several important lessons: 1. The geochemical data from a natural sample must be interpreted within the context of its geologic settings. For example, in order to relate the S and Fe isotopic compositions of pyrite in a siltstone sample to ocean-water chemistry, we must first understand whether that siltstone was deposited in an open sea, a closed marine basin, a freshwater lake, or a fluvial channel, as well as determine petrographically whether the pyrite crystals in the sample formed before, during, or after sedimentation. 2. There is no silver bullet (or smoking gun) that provides unequivocal evidence for a particular paradigm. For example, presence or absence of AIF-S in a rock does not

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Primary paper | Oxygen, iron, and sulfur geochemical cycles on early Earth Thode, H.G., Ding, T., and Crocket, J.H., 1991, Sulphur-isotope and elemental geochemistry studies of the Hemlo gold mineralization, Ontario: Sources of sulphur and implications for the mineralization processes: Canadian Journal of Earth Sciences, v. 28, p. 13–25, doi:10.1139/e91-002. Trendall, A.F., and Blockley, J.G., 1970, The Iron Formations of the Precambrian Hamersley Group, Western Australia: Geological Survey of Western Australia Bulletin 119, 366 p. Tribovillar, N., Algeo, T.J., Lyons, T., and Riboulleau, A., 2006, Trace metals as paleoredox and paleoproductivity proxies: An update: Chemical Geology, v. 232, p. 12–32, doi:10.1016/j.chemgeo.2006.02.012. Ueno, Y., 2008, Mass dependent and non-mass dependent sulfate influx into ca. 3.5 Ga ocean: Astrobiology, v. 8, Appendix, p. BAR-21. Van Kranendonk, M.J., Hickman, A.H., Smithies, R.H., Williams, I.R., Bagas, L., and Farrell, T.R., 2006, Revised Lithostratigraphy of Archean Supracrustal and Intrusive Rocks in the Northern Pilbara Craton, Western Australia: Geological Survey of Western Australia Record 2006/15, 57 p. Vogel, D.E., 1975, Precambrian weathering in acid metavolcanic rocks from the Superior Province, Villebon Township, south-central Quebec: Canadian Journal of Earth Sciences, v. 12, p. 2080–2085, doi:10.1139/e75-183. Walker, J.C.G., and Brimblecome, P., 1985, Iron and sulfur in the pre-biologic ocean: Precambrian Research, v. 28, p. 205–222, doi:10.1016/0301 -9268(85)90031-2. Watanabe, K., Naraoka, H., Wronkiewicz, D., Condie, K., and Ohmoto, H., 1997, Carbon, nitrogen, and sulfur geochemistry of Archean and Proterozoic shales from the Kaapvaal craton, South Africa: Geochimica et Cosmochimica Acta, v. 61, no. 16, p. 3441–3459, doi:10.1016/S0016 -7037(97)00164-6. Watanabe, Y., Martini, E.J., and Ohmoto, H., 2000, Geochemical evidence for terrestrial ecosystems 2.6 billion years in age: Nature, v. 41, p. 574–578, doi: 10.1038/35046052.

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The Geological Society of America Special Paper 504 2014

The upside-down biosphere: “Evidence for the partially oxygenated oceans during the Archean Eon” Shawn Domagal-Goldman* Research Space Scientist, Planetary Environments Laboratory, National Aeronautics and Space Administration Goddard Space Flight Center, 8800 Greenbelt Road, Greenbelt, Maryland 20771, USA

ABSTRACT This is a commentary on the preceding chapter by Ohmoto et al., in which it is suggested that oxygen concentrations have been high throughout Earth history. This is a contentious suggestion at odds with the prevailing view in the field, which contends that atmospheric oxygen concentrations rose from trace levels to a few percent of modern-day levels around 2.5 b.y. ago. This comment notes that many of the data sets cited by Ohmoto et al. as evidence for a relatively oxidized environment come from deep-ocean settings. This presents a possibility to reconcile some of these data and suggestions with the overwhelming evidence for an atmosphere free of oxygen at that time. Specifically, it is possible that deep-ocean waters were relatively oxidized with respect to certain redox pairs. These deep-ocean waters would have been more oxidized than surface waters, thus representing an “upside-down biosphere,” as originally proposed 25 years ago by Jim Walker.

these existing debates, and in the interest of fostering a constructive conversation, this commentary explores a hypothesis that is partially consistent with Ohmoto et al.’s arguments, but that is also consistent with the conventional wisdom that molecular oxygen (O2) was only a trace gas prior to 2.45 Ga. The hypothesis to be explored is the one originally proposed by Walker (1987), that of the “upside-down Archean biosphere.” This proposes that Archean oceans were relatively oxidized at depth, yet still anoxic at the surface; this is “upside-down” from modern oceans, which are more anoxic with depth. Such a scenario would be consistent with many of the interpretations in Ohmoto et al. Limiting these conditions to deep-ocean waters also allows for consistency with much of the “standard interpretation” that the Archean atmosphere was anoxic (a viewpoint to which Ohmoto is opposed). The “link” between the anoxic

INTRODUCTION This manuscript is a commentary on the preceding paper by Ohmoto et al. (Chapter 9, this volume), originally presented as part of the Geological Society of America Pardee Keynote Symposium in 2011. The title of that presentation, “Evidence for the fully oxygenated oceans and atmosphere during the Archean Eon,” nicely summarizes Ohmoto et al.’s position, which is contrary to that of most experts in the field. The main thesis presented by Ohmoto et al. is that there was not a “Great Oxygenation Event” (GOE) at ca. 2.4 Ga. While I disagree with this overall conclusion—as well as the majority of the analyses presented in that manuscript—those viewpoints have been repeatedly voiced in prior conferences and publications. They have also been effectively countered when previously presented. To avoid rehashing

*[email protected] Domagal-Goldman, S., 2014, The upside-down biosphere: “Evidence for the partially oxygenated oceans during the Archean Eon,” in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 97–99, doi:10.1130/2014.2504(10). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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DISCUSSION The upside-down biosphere is consistent with the strongest lines of evidence from Ohmoto et al. Of the 11 separate pieces of evidence presented for oxidized surface conditions during the Archean, 10 are sensitive to the redox conditions of bottom waters. The strongest positive line of evidence presented, “formation of primary hematite crystals in deep-sea hydrothermal sediments and volcanic rocks,” is directly tied to deep-water redox state. The upside-down biosphere is also consistent with the information most directly tied to Archean atmospheric redox conditions. Specifically, wide variations in mass-independent fractionation of sulfur (MIF-S) values, the loss of Fe in paleosols, and the presence of detrital uraninite and pyrite in sediments all place tight constraints on the oxygen concentrations in the Archean atmosphere. The other lines of evidence for anoxic conditions in the Archean are related to the mobility of certain redox-sensitive elements. Such arguments would not apply to deep waters, so long as the peaks of mid-ocean ridges—through which most of the redox-sensitive elements flow—are above the depths for which conditions are oxidizing enough to affect metal solubility. It is reasonable to question whether an anoxic atmosphere could coexist with oxic deep-ocean waters. At first glance, such a scenario seems dramatically inconsistent with what we know of modern-day redox profiles. Such profiles show more reducing conditions with depth, whereas the scenario proposed here requires more oxidizing conditions with depth. However, the cause of Earth’s modern-day redox profiles might not apply to a biosphere with different primary producers. The increase in reducing conditions with depth in modernday waters is a result of rain out of organic carbon from oxygenic photosynthesis, followed by consumption of both organic carbon and dissolved oxidants in deeper waters by heterotrophs. The heterotrophic activity draws down dissolved oxygen concentrations in deeper waters. The redox state of deep waters is then a balance between fluxes of the reduced particulate matter and the dissolved oxidized species. When organic carbon fluxes outpace the fluxes of dissolved oxygen, then all the oxygen can be consumed. This allows for the stability of organic carbon within microenvironments that are significantly out of equilibrium with the modern-day atmosphere. In the context of a different biosphere, this same principle may lead to dramatically different results. Consider a view of the Archean biosphere in which primary production is dominated by iron-oxidizing heterotrophs. The oceans on such a world would receive a flux of particulate oxidants (Fe-oxyhydroxides and sul-

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log f H2(g)

atmosphere and the more oxidized deep-ocean waters would be provided through rain out of iron oxides from surface waters, which would have been consumed at depth by iron-reducing bacteria. These organisms would have removed reductants from deep water, making this part of the ocean more oxidized than the surface.

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pH Figure 1. Stability diagram for various Fe phases: hematite, magnetite, and siderite minerals as well as Fe2+ and Fe3+ in solution. These are shown as a function of pH and the dissolved H2 concentration. Lower H2 values represent less-reducing/more-oxidizing conditions. This figure was generated using Geochemists’ Workbench (Bethke, 2002).

fur particles) and dissolved reductants (hydrogen and methane). In this case, the rain out of these oxidized particles would lead to consumption of dissolved reduced species at depth, and more oxidizing conditions in deeper waters. The redox state of deep waters would then be a balance between fluxes of the oxidized particulate matter and the dissolved reduced species. If mixing in the system is poor, or if productivity levels are high, then the rain out from primary producers will outpace the mixing of dissolved species. In this case, the fluxes of Fe-oxyhydroxides dominate, and the reduced species are consumed first. This creates a microenvironment that allows for the preservation of Feoxyhydroxides. However, the presence of this microenvironment does not mean the entire surface must have been oxidized. Walker (1987) originally presented this model over 25 yr ago, and the few studies that have tested it since then have all been consistent with it. Most notably, series of iron isotope analyses suggest the presence of bacterial Fe reduction in the Archean (nicely summarized by Johnson et al., 2008). Many of these Fe isotope studies also show covariance in the fractionations of carbon, sulfur, and oxygen isotopes (e.g., Czaja et al., 2010). These studies suggest the presence of complex ecosystems in the Archean in which Fe oxidation and Fe reduction were both biologically mediated. More tests of this conceptual model are required. Numerical simulations showing which boundary conditions (if any) could produce this upside-down biosphere are needed. These models

Commentary | Upside-down biosphere: “Evidence for the partially oxygenated oceans during the Archean Eon” must include a flexible ecosystem model that can predict the productivity of anoxygenic phototrophs, and the heterotrophs that would have consumed the by-products of these metabolisms. More importantly, further geological and geochemical tests are needed. One particular test that could resolve this would involve proxies for the paleodepths of various samples. Such measurements, combined with measurements of the local redox environments, would present the ultimate test for this hypothesis. For example, the amount of reductant consumption would determine which Fe mineral phase will be stable (see Fig. 1) in the microenvironment. At increasing H2 drawdown levels, siderite, then magnetite, and finally hematite will be preserved. If a quantitative proxy for paleodepth is made available, then a test for this hypothesis would be a comparison of depths of these various mineralogies: Deeper waters should be associated with more oxidized conditions and a greater presence of hematite. ACNOWLEDGMENTS This work was performed as part of the NASA Astrobiology Institute’s Virtual Planetary Laboratory, supported by the National Aeronautics and Space Administration through the

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NASA Astrobiology Institute under solicitation No. NNH05ZDA001C; it was also supported by the Oak Ridge Associated Universities NASA Postdoctoral Program. REFERENCES CITED Bethke, C.M., 2002, The Geochemist’s Workbench Release 4.0: A User’s Guide to Rxn, Act2, Tact, React, and Gtplot: Urbana, Illinois, University of Illinois, 236 p. Czaja, A.D., Johnson, C.M., Beard, B.L., Eigenbrode, J.L., Freeman, K.H., and Yamaguchi, K.E., 2010, Iron and carbon isotope evidence for ecosystem and environmental diversity in the ~2.7 to 2.5 Ga Hamersley Province, Western Australia: Earth and Planetary Science Letters, v. 292, p. 170– 180, doi:10.1016/j.epsl.2010.01.032. Johnson, C.M., Beard, B.L., and Roden, E.E., 2008, The iron isotope fingerprints of redox and biogeochemical cycling in modern and ancient Earth: Annual Review of Earth and Planetary Sciences, v. 36, no. 1, p. 457–493, doi:10.1146/annurev.earth.36.031207.124139. Ohmoto, H., Watanabe, Y., Lasaga, A.C., Naraoka, H., Johnson, I., Brainard, J., and Chorney, A., 2014, this volume, Oxygen, iron, and sulfur geochemical cycles on early Earth: Paradigms and contradictions, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, doi:10.1130/2014.2504(09). Walker, J.C.G., 1987, Was the Archaean biosphere upside down?: Nature, v. 329, p. 710–712, doi:10.1038/329710a0.

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The Geological Society of America Special Paper 504 2014

Discussion of “Oxygen, iron, and sulfur geochemical cycles on early Earth: Paradigms and contradictions” (Ohmoto et al.) Hiroshi Ohmoto: Thanks, Shawn [Domagal-Goldman]. A number of problems exist in your oxygen oasis model. You could argue for an oxygen oasis for a shallow, near-coastal water body. But you cannot argue for the existence of a local oxygen oasis in a reducing atmosphere during the formation of oxidized paleosol. You don’t call it an oxygen oasis if a deep sea [in the 3.5 Ga Marble Bar region], over a 100,000 square kilometers area, was oxygenated. It was unlikely that ferric hydroxides [in the 3.5 Ga Marble Bar region] formed in the surface zone, rained down [to the seafloor], and were converted to ferric oxide (hematite). [This is because our electron microscope study revealed that the hematite crystals in the Marble Bar jasper were not transformed from low temperature ferric hydroxides. Rather they nucleated directly as hematite at T > 60 °C at near hydrothermal vents.] If the hematite crystals had formed initially as ferric hydrates in a shallow zone and deposited in a deep sea, you would see very dispersed and homogeneous layers of hematite, not specifically confined to submarine hydrothermal fluid discharge areas as we see. So you have to go beyond the model, and look for geologic evidence. Geology is the key. The correlation between iron and sulfur isotope variations you mentioned, that’s a characteristic in the Hamersley Basin. And what I said was this [i.e., Fe-S correlation] was a diagenetic problem. Diagenetic reactions occurred among iron, sulfur, and carbon in the sediments under hydrothermal conditions. That is why you get correlations among them. If they are not linked together you don’t get any correlations. Ron Frost: This is a minor point from Hiroshi’s talk, and several other people have made the point that relatively oxidized calcalkalic magmas indicate that there must have been an oxidized atmosphere, and that I think is totally false. The reason why you get really reduced granites in, let’s say, Yellowstone rhyolites is because they’re dry and you differentiate an oxide, mostly olivine and pyroxenes, and you’re driving

the iron into oxides, making the pyroxenes more iron-rich; you make very reduced rock. In calcalkalic rocks you’re differentiating biotite and hornblende, both of which have ferric iron, and neither of them get very iron-rich and you can demonstrate pretty simply the differentiation of biotite’s going to make it relatively oxidized. These aren’t very oxidized rocks; they’re only two log units above FMQ1 and they probably carry no necessary indication of an oxidized atmosphere. Ohmoto: Well you can do it perhaps one or two above FMQ but these Kaapvaal craton granites you are talking about four orders of magnitude above. Simple differentiation of iron is not going to make a ferric, high-end ferric content. You need to start with source rocks which are enriched in ferric iron. Frost: Four log units is probably a little too high but you get minettes that are coming out of rocks which are probably not necessarily seeing oxygen that are 6 log units above FMQ. Ohmoto: How do you know [that ferric-enrichments in the source rocks were not caused by reactions with O2-rich atmosphere or ocean water]? Frost: I don’t know but I certainly trust Ian Carmichael, who measured them. In Mexico today early minettes with very, very small amounts of melting can come out with lots of ferric iron and can be very, very oxidized and not necessarily have anything to do with the oxygen atmosphere today. Ohmoto: Do you know these mantle edges, where these magmas were generated, are highly oxidized because of the influence from subducting, oxygenated oceanic crust? I think I have many data to support such a model.

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FMQ stands for fayalite-magnetite-quartz.

Ohmoto, H., Domagal-Goldman, S., et al., 2014, Discussion of “Oxygen, iron, and sulfur geochemical cycles on early Earth: Paradigms and contradictions” (Ohmoto et al.), in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 101–104, doi:10.1130/2014.2504(11). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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Frost: Okay that’s an assumption. Ohmoto: It’s an assumption, yes, at least to fit to various geochemical data, trace element data indicating oceanic crust, oxygenated oceanic crust. So I think I have all these data. Frost: Yeah, but my point is a lot of the oxidation and reduction that we see in the Earth’s crustal rocks today can be driven abiotically and have nothing to do with subducted oxygen. So what you have to do is you have to prove that you cannot produce these things by abiotic emissions, or by processes without adding oxygen to [?]. Ohmoto: I think you misunderstood me. What I’m saying is the creations of oxidized crusts (oceanic and continental), not necessarily having free-O2 present in these midoceanic region basalts. Basalts in midoceanic ridges become highly oxidized, highly increased in ferric/ferrous ratio, due to reactions with oxygen- and sulfate-rich seawater. This high-ferric oceanic crust, when it is subducted, releases water and SO2. They in turn oxidize the mantle wedges. And if you melt those mantle wedges, you produce highly oxidized granites, which make up most of the continental crust. I’m not talking about a survival of oxygen as O2 molecules during the passage through the oceanic crust, upper mantle, and continental crust. Most of the O2 molecules that existed in the seawater go through this passage as oxidized forms of iron and sulfur. Frost: Well I think I made my point, essentially. A lot of this can be explained without necessarily having to oxidize the wedge [?] by simple oxidation reduction reactions between melt, oxides and silicates in the melt. Kasting: Hi, Hiroshi. I think one of the most interesting things that you and your colleagues have done are these experiments on the sulfur isotope fractionation. Yumiko [Watanabe]’s experiments where she’s getting positive Δ33 from reacting solid glycine with sulfate, and then these new data, which you showed from your Japanese colleagues who were getting up to plus 13. Ohmoto: The plus 13‰ Δ33 value was obtained by Henry Oduro of James Farquhar’s group. Kasting: So I have a question: Have you measured or have they measured Δ36 for those products? Ohmoto: Oduro’s sample had a Δ36 value close to 0‰. This is a mystery because we found a significant Δ36 variation [ranging from –1.5 to +1.5‰] in our experimental products. So I think at least two types of isotope fractionation mechanisms play important roles during reactions between solid organic matter and aqueous sulfate: surface chemical reactions [i.e.,

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chemisorption isotope effect] and radical formation reactions, which produce magnetic isotope effect. Kasting: So you’re suggesting surface reactions, but the reason the 36 is important is that if you look at Archean rocks there are well-known correlations between Δ33 and Δ36, which says it is not a magnetic isotope effect because to get a magnetic isotope effect you have to have a nuclear spin and you need an odd number of nucleus, so 32, 34, and 36 don’t have nuclear spin; 33 does. There is a known mechanism for producing magnetic isotope effects in liquid solution, and if that’s what’s going on in your experiments then you could have a magnetic isotope effect that’s fractioning sulfur 33 and you should not see a strong fractionation in 36. Ohmoto: But in our experiments we did find significant fractionations of both 36S and 33S; some of the data are already published. Kasting: Right. But the other thing that goes along with that, I’ve just been learning some of this, is not every deviation from zero means that it’s a mass independent effect because the mass dependent fractionation line is not that well defined, or it’s defined for a particular reaction. And so you can get small deviations from the MIFs2 that look like MIF without a mass-independent process. Ohmoto: I think a problem, like many definitions, is a tendency for researchers to relate a particular observation (e.g., a Δ33-Δ36 relationship of a sample) to a specific process (e.g., photochemical process). Definitions of mass independent isotope fractionation are simply based on the magnitudes of deviations of δ values of a sample from more common values. [Our experimental samples definitely show MIF characteristics regardless of how you define the δ, Δ33, and Δ36 values, either as natural logarithmic functions or as normal arithmetic functions.] Whether the isotope fractionation was related to equilibrium, kinetic, magnetic, surface chemisorption, or photochemical reaction—that’s a separate problem. Kasting: So sulfur 36 is the way to distinguish between these different mechanisms? Shawn Domagal-Goldman: I want to ask actually a follow-up to that which is I think I heard there, that the largest MIF you’re seeing from the Farquhar experiments didn’t have the 36 fraction ___(?)__? Ohmoto: That’s what I heard, right.

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MIF stands for mass-independent fractionation.

Discussion | Oxygen, iron, and sulfur geochemical cycles on early Earth Domagal-Goldman: But the ones that you did that don’t have quite as large a magnitude do have a 36 __(?)___? Ohmoto: Correct. [Our Δ33 values range from 0 to +2.3‰ and the Δ36 values from –1.5 to +1.5‰. These Δ33 and Δ36 values overlap with more than 70% of those of Archean sedimentary rocks.] How large the Δ33-Δ36 box will be expanded for products of thermochemical reactions between organic matter and sulfate is an important topic for future research. Johan Varekamp: If all this early volcanism indeed was submarine and let’s assume that the magma production rates were somewhat higher and that the submarine hydrothermal systems were not complete analogs with the current ones, one could imagine that the SO2 ultimately reaches that seawater and that you have disproportionation of SO2 and elemental sulfur and sulfate. The elemental sulfur would ultimately sink to the bottom of the ocean and you (have) the question [of] what has ultimately happened to that elemental sulfur is something else. But you would build up some steadystate inventory of sulfate in that very early ocean. This is not totally a pipe dream; this happens every day in many volcanic crater lakes where we have large layers of elemental sulfur sitting at the bottom of the lake and fairly rich, sulfaterich waters that also happen to be very acidic at the top. One way to check this is that that sulfate is actually isotopically very heavy, because a light elemental sulfur is separated. And I was wondering if anyone has thought about that. And my personal view, since I’ve worked from crater lakes half of my career, these are the ideal current analogs of pre-Cambrian oceans if you want. Ohmoto: Yes. Subaerial volcanism may produce SO2-dominant volcanic gas, especially when water pressures of eruptions were less than 10 bar [i.e., shallow volcanic eruptions] and under high temperature [e.g., T > 900 °C] and normal pO2 conditions [FMQ ± 2 log units]. When the volcanic gas cools down to below ~400 °C, SO2 dissociates to sulfate and H2S. H2S may be oxidized to native sulfur if O2 is present. [When the water pressure was greater than ~100 bar, such as for typical submarine volcanism, volcanic gas (hydrothermal fluid) typically remains H2S-rich and SO2-poor throughout the temperature range from ~1000 °C to < 200 °C. Therefore, simple cooling of such magmatic fluids would not produce sulfate or native sulfur. However, today’s submarine hydrothermal fluids typically contain some sulfate of seawater origin.] So you have to think about high temperature, shallow subaerial volcanism to produce SO2, H2S, sulfate, and native sulfur. Varekamp: So if you do this is in some early ocean at a relatively shallow water depth that would be one way to getting sulfur in the water. Or some kind of interface that’s warm enough where this could happen.

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Ohmoto: Yeah. The question is what were the temperatures of volcanic gas exiting out of ocean water. If the volcanic gas exited from the ocean at T > 400 °C, it would not have generated sulfate and native sulfur in the shallow sea. However, if volcanic gas from a shallow magma exited from the oceans at T < 400 °C, the volcanic gas could have generated sulfate and native sulfur before exiting the ocean. Varekamp: It’s just another idea. Ohmoto: Thank you. Harry Becker: I would like to make another comment. If you want to have an oxidizing environment, continental environment in the Archean you also have to oxidize elements like rhenium and osmium, and of course you should see a trace of them somewhere. And there is the decay system of rhenium-187 to osmium-187, and because the crust has a high rhenium-osmium ratio they’ll develop radiogenic osmium isotope as it comes in a very short period of time, 200 million years. And so you should see these radiogenic osmium isotopes in marine sediments at that time, like shales or whatever, but there’s not a lot of data over its history in terms of initial osmium isotopic composition in marine sediments but there are some Archean sands and this is __(?)__ by chance. If you go back into the Archean the values are chondritic; the initial isotopic compounds there are two data points—no actually—yeah, two data points, from 2.5 billion and 2.7 billion, and they are within the contradicted values. And it’s only until, well, 1.8 billion or so, or 1.5 billion that the values really take off. That suggests that there’s no radiogenic osmium from the continents introduced into the oceans. There’s not a lot of data but, you know, it’s quite convincing to me there was no oxidated [?]. Ohmoto: I’m not familiar with the rhenium-osmium system, and I should be. Okay, it’d be interesting. Martin J. Van Kranendonk: Hiroshi, again, as Shawn said with deep respect for you as a person I have many disagreements with your model, and I think we’ve talked about those. But I do want to make the point to the audience because there are some aspects, I think, that have been glossed over very quickly in your presentation, and I think you present comparisons which are not valid. Shawn picked up on many of the points that relate to the deep ocean, but you also mention many features that should be seen on the surface. And you mentioned things like red beds, and the fact that prior to 2.5 [billion] there are no continental deposits, or very few. But actually there are very well-preserved thick continental deposits and there are no red beds. The other thing you mentioned, and one of them is about the uranium ore deposits, detrital uraninite, in the

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Ohmoto et al. modern examples you show those are extremely immature sediments that come directly from high mountain belts whereas the detrital uraninite in pre-Cambrian and Archean rocks are from highly mature quartzites that have been washed back and forth within the ambient atmosphere for long periods of time. These are ultramature orthoquartzites; the two are not comparable. The third point, and we’ve had these discussions before, is about the paleosol that you quote in the Pilbara; that has not been documented as a real paleosol, and I have a completely different interpretation. Those are metamorphic, hydrothermal mineral assemblages related to intrusions in that area that are uncomformably overlain by the rocks above it. So I don’t think it’s a good argument to bring that into the mix. And then the other one I had is a question for you: If your model for the generation of the sulfur MIF anomalies is solely related to cooking of organic-rich shales with sulfate reducers, why don’t we see that signature all the way through the rest of the pre-Cambrian record up to today? I mean we have very high metamorphic rate and low metamorphic grade thick sections of organic-rich shales; we do not have the sulfur MIF signature. And I just wanted to clarify from your presentation: Are you actually saying that you don’t believe the mechanism that James Farquhar has presented for the sulfur MIF signature? Because we know from modern examples of very large, volcanic eruptions that have emitted out into the stratosphere that they do preserve that sulfur MIF signature. So I don’t quite understand the way you compare one thing with another and try to make a single model; to me those just do not fit.

Ohmoto: Okay, yeah. I was expecting all the comments from you, Martin. We’ll start with the last one. I believe some photochemical reactions involving SO2 produce MIF sig-

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natures on the reaction products, but I do not believe photochemical reactions were responsible for the MIF signatures we see in sediments. Okay? I am not arguing about the mechanism. However, the MIF signatures from photochemical experiments [especially those obtained by using broad-band UV lamps], do not match well with MIF signatures in natural rocks. Back to the MIF signatures, whether or not we find them in younger organic carbon-rich sediments—We do see those in the 1.8-billion-year-old Outokumpu Formation [and in some Phanerozoic-aged sediments]. In general, what we need are: (1) abundant SO4 (e.g., sulfate-rich seawater and evaporite), (2) an environment to accumulate a large amount of organic matter, and (3) a very large-scale heat source for hydrothermal systems that lasted for a period over 10s of million years. These requirements were met in the Hamersley Basin. Today where do we go [to find such a geologic environment]? We don’t find that. So it’s no mystery to me at all that we don’t see many MIF signatures in younger sediments. How do you explain the MIF signatures in the 1.8 Ga and also in some younger sediments? Van Kranendonk: Well I think it’s not surprising to find the MIF signature coming back at 1.8 because you actually have a reappearance of standard iron formations in the rock record; immediately prior to that you have komatiites erupting again at 2.1 to 1.9 billion years and you actually have the formation of granite greenstone-type crust in that time period too. You actually seem to be returning to a phase of a more Archean-like Earth for a period. So actually that doesn’t surprise me. But I think you still have to explain 1.8 billion years of history that doesn’t record what should be a prevalent, widespread phenomenon if it’s just merely from cooking of organic shales. I think that’s a real issue.

The Geological Society of America Special Paper 504 2014

Earth’s early atmosphere and surface environments: A review Martin J. Van Kranendonk* School of Biological, Earth and Environmental Sciences and Australian Centre for Astrobiology, University of New South Wales, Kensington, NSW 2052, Australia

ABSTRACT This review summarizes and assesses a series of papers presented at the Geological Society of America Annual Meeting in 2011 on the changing composition of Earth’s early atmosphere. This is a developing field, with differing views, due largely to the facts of an incomplete rock record and negligible preservation of the gaseous components of the atmosphere. Nevertheless, there are constraints, available through geological proxies in the form of chemical sedimentary rocks that reflect the composition of the hydrosphere, and because the two are directly linked, the atmosphere. A review of the geological constraints on atmospheric conditions is presented for early Earth, from its formation at 4.56 Ga up to 1.8 Ga, followed by a developing model that links changing atmosphere/hydrosphere conditions and biosphere evolution to changes in planetary tectonics, including the evolving supercontinent cycle.

INTRODUCTION The Pardee keynote symposium on “Earth’s Early Atmosphere and Surface Environment” at the 2011 Geological Society of America meeting in Minneapolis, Minnesota, provided an opportunity for specialists and others to hear and discuss the latest ideas regarding the early evolution of the outer shells of our planet. The format of the session was structured so that rather than a high number of speakers presenting for a short time (12 min), more limited numbers of invited speakers were allotted longer time slots to present overviews of their work (25 min), with time allotted after each presentation for discussion. Comments and their replies were taped so that a full record of the session could be preserved. The results of this symposium are presented in this Special Paper. Most presentations focused on the composition of the early atmosphere and the redox state of the early oceans, with little attention paid to surface environments. Three chapters by Jim

Kasting, by George Shaw, and by Laura Schaefer and Bruce Fegley Jr. discuss the composition of the atmosphere in the earliest part of Earth history (the Hadean and earliest Archean). Shawn Domagal-Goldman (this volume, Chapter 2) provides a commentary on Shaw’s paper. Two chapters, by Hiroshi Ohmoto et al. (this volume) and by Kevin Zahnle and David Catling (this volume), concentrate on the rise of atmospheric oxygen through the latter part of early Earth history (the latest Archean and earliest Proterozoic). Shawn Domagal-Goldman (this volume, Chapter 10) provides another commentary (this time on the contribution by Ohmoto et al., this volume), in the process reviving a model of an “upside-down” biosphere first proposed by Walker (1987). As a nonspecialist regarding the composition of Earth’s earliest atmosphere and oceans, I can only attempt to summarize the different views presented in this Special Paper. My immediate impression was that this is an exciting field, in the early stages of development, with a number of contrasting hypotheses and more than a few controversies. However, as someone with a

*[email protected] Van Kranendonk, M.J., 2014, Earth’s early atmosphere and surface environments: A review, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 105–130, doi:10.1130/2014.2504(12). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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degree of familiarity with early Earth, I can try to supply some of the context that may reflect on the issues discussed in this special volume. To this end, I present a summary of papers presented in this volume, discuss the sources of controversy related to these topics, and then present a review of Precambrian Earth evolution and provide a geological template for events affecting the atmosphere/hydrosphere. I wrap up the conclusions and then lastly discuss how these may influence our thinking with regard to the models/ideas on atmospheric evolution presented here. PAPERS PRESENTED IN THIS SPECIAL PAPER An Early, Reducing Atmosphere? It is widely believed that conversion of latent energy of accreted planetesimals into heat led to impact outgassing of volatiles during the accretion of Earth (e.g., Tyburczy et al., 1986; Ahrens et al., 1989). Because H2O and CO2 were released from these authors’ study of carbonaceous chondrites, this has given rise to the widespread belief that an atmosphere arising from impact origins was a H2O-CO2 steam atmosphere. However, more recent studies have shown that only carbonaceous chondrites produced a steam atmosphere, whereas ordinary chondrites (including enstatite chondrites) and high-Fe chondrites produced H2-rich atmospheres (Schaefer and Fegley, 2007, 2010). Modeling of oxygen isotope compositions suggest that Earth was primarily accreted from enstatite (70%) and ordinary chondrites (21%), with only a subordinate contribution from carbonaceous chondrites (10−5 PAL, is confirmed by more and more data from a variety of detailed studies. The model of H escape to space as the cause for the rise in atmospheric oxygen presented by Zahnle and Catling (this volume) certainly touches on one important aspect of oxygen’s rise, but a several other factors come into play at this time, potentially the most important of which is a near-global shutdown in magmatic and orogenic activity and a resultant dramatic decrease in volcanic degassing. Indeed, a compilation of geoscience information through the Precambrian suggests that tectonics—driven by heat loss from the core-mantle system—was a principal driver of atmospheric change throughout the Precambrian, at least since 3.2 Ga, when the onset of modern-style plate tectonics is considered to have commenced. A comparison with the mid-Cretaceous superplume event provides a modern analogue for past events, with changes driven by changing mantle dynamics. These changes are cyclical over time and vary from warm mantle to cool mantle states that

Earth’s early atmosphere and surface environments: A review are linked to the supercontinent cycle. These deep-level drivers of planetary change are upwardly linked to the outer shells of the planet through changes in the rates of volcanism and tectonism, which cause variations in the degree of continental emergence and weathering, which in turn drive changes in atmosphere/ hydrosphere compositions, as well as changes in the biosphere (e.g., through C-availability and nutrient supply). The links and feedbacks among the core-mantle-crusthydrosphere-atmosphere provide a rich field for future research and provide a rewarding, though challenging, means to a better understanding of whole-Earth evolution through time. ACKNOWLEDGMENTS I am grateful to George Shaw for inviting me to present this overview, and I wish to acknowledge a host of colleagues who helped shape this contribution, particularly Aivo Lepland and Chris Kirkland. The Geological Survey of Western Australia provided me with the opportunity to attend the Geological Society of America Annual Meeting in the United States. I thank Roger Buick for his thorough review. This is Australian Research Council Centre of Excellence for Core to Crust Fluid Systems contribution number 252 (www.ccfs.mq.edu.au). REFERENCES CITED Abbott, D.H., and Isley, A.E., 2002, The intensity, occurrence, and duration of superplume events and eras over time: Journal of Geodynamics, v. 34, p. 265–307, doi:10.1016/S0264-3707(02)00024-8. Ahrens, T.J., O’Keefe, J.D., and Lange, M.A., 1989, Formation of atmospheres during accretion of the terrestrial planets, in Atreya, S.K., Pollack, J.B., and Matthews, M.S., eds., Origin and Evolution of Planetary and Satellite Atmospheres: Tucson, Arizona, University of Arizona Press, p. 328–385. Akhmedov, A.M., Belova, M.Y., Krupenik, V.A., and Sidorova, I.N., 2000, Fungous microfossils from Paleoproterozoic black shales of the Pechenga Complex in the Kola Peninsula: Doklady Earth Sciences, v. 373, p. 782–785. Albarède, F., and Blichert-Toft, J., 2007, The split fate of the early Earth, Mars, Venus, and Moon: Comptes Rendus Geoscience, v. 339, p. 917–927, doi:10.1016/j.crte.2007.09.006. Allègre, C.J., Manhès, G., and Göpel, C., 1995, The age of the Earth: Geochimica et Cosmochimica Acta, v. 59, p. 1445–1456, doi:10.1016/0016 -7037(95)00054-4. Allwood, A.C., Walter, M.R., Kamber, B.S., Marshall, C.P., and Burch, I.W., 2006, Stromatolite reef from the Early Archaean era of Australia: Nature, v. 441, p. 714–718, doi:10.1038/nature04764. Allwood, A.C., Walter, M.R., Burch, I.W., and Kamber, B.S., 2007, 3.43 billionyear-old stromatolite reef from the Pilbara craton of Western Australia: Ecosystem-scale insights to early life on Earth: Precambrian Research, v. 158, p. 198–227, doi:10.1016/j.precamres.2007.04.013. Altermann, W., and Schopf, J.W., 1995, Microfossils from the Neoarchean Campbell Group, Griqualand West Sequence of the Transvaal Supergroup, and their paleoenvironmental and evolutionary implications: Precambrian Research, v. 75, p. 65–90, doi:10.1016/0301-9268(95)00018-Z. Anbar, A.D., Duan, Y., Lyons, T.W., Arnold, G.L., Kendall, B., Creaser, R., and Kaufman, A.J., 2007, A whiff of oxygen before the Great Oxidation Event?: Science, v. 317, p. 1903–1906, doi:10.1126/science.1140325. Andreasen, R., and Sharma, M., 2009, Comment on “Neodymium-142 evidence for Hadean mafic crust”: Science, v. 325, p. 267, doi:10.1126/ science.1169604. Arculus, R.J., and Delano, J.W., 1980, Implications for the primitive atmosphere of the oxidation state of Earth’s upper mantle: Nature, v. 288, p. 72–74, doi:10.1038/288072a0.

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The Geological Society of America Special Paper 504 2014

Concluding comments George H. Shaw Geology Department, Union College, Schenectady, New York 12308, USA

Even a casual reading of the chapters in this volume will suggest dramatically different views of Earth’s early atmosphere, from shortly after the Moon-forming impact (MFI) to the end of the Archean. It might appear that many, perhaps most, of the ideas presented must necessarily be incorrect because of mutual contradiction. Indeed, it is fair to say that some of what is presented here must categorically be wrong. That said, it is also clear to me that some apparent contradictions are just that, apparent. As one who has proposed to revive an old view of early atmospheric composition, and with little success in convincing others, I want to suggest ways in which the disparate views may at least partially be rationalized. To be sure, I am at the opposite end of the compositional spectrum from H. Ohmoto et al. (this volume), though I am willing to accept that some of their ideas should receive careful analytical attention (as they have). Because my own ideas, while rooted in much older notions about the early (prebiotic) atmosphere, have only been recently reintroduced, I will suggest ways in which they can be made consistent with the prevailing view(s). It is important to keep in mind the time line for various events, conditions, and lines of evidence relevant to understanding the state of Earth’s atmosphere and surface. For example, conditions implied by geological evidence dated at 3.8 Ga may not apply to earlier times (as indicated by S. Domagal-Goldman [this volume]). It is also true that conditions for one part of Earth, say the upper mantle, may or may not be indicators for other locations, e.g., the surface conditions. Also, it is important to recognize that some geologic evidence may be so fragmentary that it, by chance, may be telling us more about a micro-environment than about global conditions. With these caveats in mind, I suggest the following “resolutions.” First, my analysis of degassing applies to conditions at the earliest phases of the post-MFI Earth, perhaps from 4.5 to 4.2 Ga. This 300 m.y. interval probably saw the accretion of much of the last 1+% of Earth’s mass. It is clear (to me) that the

Late Heavy Bombardment, which has come under some scrutiny of late, represents a very small fraction of the total of this late accretion. The degassing I imagine, almost certainly largely driven by impact energy, must have produced an abundance of reduced compounds both in the atmosphere (initially) and as rain-out in the earliest ocean. This process had very little to do with mantle chemistry, which, as the result of earlier core separation, was largely iron free. Whatever magmas were produced by the mantle, they would certainly have been consistent with measurements of mantle oxidation state indicative of conditions close to the modern state. This magmatism (and evolution of oxidized gases) in no way contradicts a reduced surface environment, or (as I suggest) massive amounts of reduced compounds in the primordial ocean, and some in the atmosphere. The subsequent addition of carbon dioxide (and carbon monoxide, following J.F. Kasting’s suggestion [this volume]) would clearly affect surface conditions over time and was almost certainly vital in regulating the state of the early atmosphere and climate. Second, an atmosphere containing a significant fraction of reduced gases, such as methane (and ammonia to a much lesser extent), provides the conditions (probably) necessary for prolonged and substantial hydrogen escape as per K. Zahnle and D. Catling’s (this volume) ideas. It is very difficult to imagine hydrogen escape necessary to eventually lead to oxidized surface conditions with an atmosphere composed almost solely of nitrogen, carbon dioxide (and monoxide), and water. Zahnle and Catling make a very persuasive case for this hydrogen loss, and a process for “pumping” CH4 (and NH3) into the atmosphere by hydrothermal processing of organic compounds can only help. Third, assuming a fairly early advent of oxygenic photosynthesis, probably no later than 2.7 Ga, and perhaps somewhat earlier, the delay in atmospheric oxygenation can readily be explained by the necessity not only for hydrogen escape, but for processing (oxidizing) the large reduced carbon pool at the surface. The hydrothermal processing of the reduced carbon pool is

Shaw, G.H., 2014, Concluding comments, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, p. 131–132, doi:10.1130/2014.2504(13). For permission to copy, contact [email protected]. © 2014 The Geological Society of America. All rights reserved.

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just the sort of mechanism to get the reduced compounds into the atmosphere where hydrogen escape can do the necessary work of oxidizing the surface. Finally, although I clearly would not agree with most of the conclusions of Ohmoto et al., I can certainly envisage local conditions of high oxidation state, especially once oxygenic photosynthesizers are on the scene. Such conditions could certainly yield geologic evidence that might well be preserved, albeit even if the conditions were highly localized. One need only consider the current surface of Earth, where an oxygen-rich atmosphere is clearly accompanied by local (even fairly widespread) highly reducing environments. Also, we should always keep in mind that geologists have a special knack for finding the unusual when looking for materials to study! I am not aware of any substantial efforts to attempt atmospheric modeling that would address the conditions I am proposing for Earth’s earliest atmosphere and surface. To do this in a meaningful way would require making realistic estimates of processes not yet included in many models, especially hydrothermal processing of a large organic-rich pool. Modern methanogenesis and serpentinization reactions tell only part of the story, and biological methanogenesis was clearly absent during the earliest times when hydrothermal processing would probably have been the main source of reduced compounds in the atmosphere. Until this kind of modeling takes into account the widest range of possible mechanisms, I would have little confidence in the results. I will make one last suggestion that I hope will pique the interest of the modelers. An important problem with any kind of global model is the issue of homeostasis, especially in the climatic sense. There have been many efforts to explain climate stability (and the serious deviations from it) based on considerations of atmospheric chemistry and changes through time. Current knowledge about early Earth conditions (ca. 4.2 Ga or so) suggests temperatures that were not so high as to boil the oceans or so low as to produce globally extensive glaciations. The atmospheric/surface model I propose should presumably result in conditions including some kind of feedback to achieve not only warm enough conditions, but reasonably equable temperatures. The proposals which state that this can be accomplished using high levels of carbon dioxide leave me cold (pun intended), because the reactivity of carbon dioxide with silicates is a fairly efficient route for its removal from the atmosphere. Assuming that volcanic-hydrothermal recycling of carbon dioxide into the atmosphere could maintain the huge levels necessary for a mild climate, such conditions must necessarily have persisted well into the late Archean. The geological effects of this later Archean high-CO2 atmosphere should be more than apparent, as first pointed out by Rubey (1955). The presence of carbonatized oceanic crust may be one such indicator, but it is hardly

definitive, even at the levels seen in some locations. After all, the modern ocean crust exhibits some carbonatization, even under a 350 ppm CO2 atmosphere. The degree of weathering one might expect from an atmosphere with hundreds to thousands of times as much atmospheric CO2 should be both pervasive and striking, and would have persisted until late in the Archean. On the other hand, if climate stability were achieved in a nitrogen–carbon dioxide(monoxide)–water–methane–ammonia atmosphere over an organic-compound–rich ocean, I can imagine a feedback mechanism that could maintain equable conditions, though I am not the one to try modeling it. Imagine that hydrothermal activity generates enough methane and ammonia to moderate the climate, in conjunction with a substantial amount of carbon dioxide (and water vapor). What happens if extra methane is added and the climate heats up? There is a feedback that accentuates the amount of reduced gas in the atmosphere as the ocean heats. The hotter ocean will release more methane, but it will probably also release other volatile (low-molecular-weight) organic compounds, providing an enhanced source of feedback. This extra organic loading to the atmosphere will lead to the wellknown production of an anti-greenhouse haze, which then cools Earth. As temperature falls, the volatile organics will be retained in the cooler ocean, reducing the load to the atmosphere and allowing “normal” greenhouse conditions to return, still provided by hydrothermal production of methane and ammonia (to add to the carbon dioxide, which largely stays in the atmosphere). Modeling this kind of system is probably possible (though not by me!) and could provide some further insights into Earth’s earliest era. REFERENCES CITED Domagal-Goldman, S., 2014, this volume How low can you go? Maximum constraints on hydrogen concentrations prior to the Great Oxidation Event, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, doi:10.1130/2014.2504(02). Kasting, J.F., 2014, this volume, Atmospheric composition on the Hadean–early Archean Earth: The importance of CO, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, doi:10.1130/2014.2504(04). Ohmoto, H., Watanabe, Y., Lasaga, A.C., Naraoka, H., Johnson, I., Brainard, J., and Chorney, A., 2014, this volume, Oxygen, iron, and sulfur geochemical cycles on early Earth: Paradigms and contradictions, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, doi:10.1130/2014.2504(09). Rubey, W.W., 1955, Development of the hydrosphere and atmosphere, with special reference to probable composition of the early atmosphere, in Poldervaart, A., ed., Crust of the Earth: Geological Society of America Special Paper 62, p. 631–664, doi:10.1130/SPE62-p631. Zahnle, K., and Catling, D., 2014, this volume, Waiting for O2, in Shaw, G.H., ed., Earth’s Early Atmosphere and Surface Environment: Geological Society of America Special Paper 504, doi:10.1130/2014.2504(07). MANUSCRIPT ACCEPTED BY THE SOCIETY 8 NOVEMBER 2013

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