Johnson Et Al 2013 GSA Spec Pap

spe 500-08 1st pgs page 1

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The Geological Society of AmericaSpecial Paper 500

2013

Smaller, better, more: Five decades of advances in geochemistry

Clark M. Johnson*Department of Geoscience, University of Wisconsin, Madison, Wisconsin 53706, USA, and National Aeronautics and

Space Administration (NASA) Astrobiology Institute, Wisconsin Team

Scott M. McLennanDepartment of Geoscience, State University of New York at Stony Brook, Stony Brook, New York 11794, USA

Harry Y. McSweenDepartment of Earth and Planetary Sciences, University of Tennessee, Knoxville, Tennessee 37996, USA

Roger E. SummonsDepartment of Earth, Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology, Cambridge,

Massachusetts 02139, USA, and National Aeronautics and Space Administration (NASA) Astrobiology Institute, MIT Team

ABSTRACT

Many of the discoveries made in geochemistry over the last 50 yr have been driven by technological advances that have allowed analysis of smaller samples, attainment of better instrumental precision and accuracy or computational capability, and automation that has provided many more data. These advances occurred during development of revolutionary concepts, such as plate tectonics, which has provided an overarching framework for interpreting many geochemical studies. Also, spacecraft exploration of other planetary bodies, including analyses of returned lunar samples and remote sensing of Mars, has added an additional dimension to geochemistry.

Determinations of elemental compositions of minerals and rocks, either through in situ analysis by various techniques (e.g., electron microprobe, secondary ion mass spectrometry [SIMS], synchrotron X-ray fl uorescence [XRF], laser ablation) or bulk analysis (e.g., XRF, inductively coupled plasma–atomic emission spectrometry [ICP-AES], inductively coupled plasma–mass spectrometry [ICP-MS]), have become essential approaches to many geochemical studies at levels of sensitivity and spatial resolution undreamed of fi ve decades ago. Although major-element distributions in igneous rocks have been understood at a basic level for some time, advances using major-element abundances to understand sedimentary provenance and processes have been especially noteworthy during the past half-century. The great diversity of trace elements in terms of geochemical behavior (e.g., lithophile, siderophile, etc.) has made them invaluable to many studies, providing unique constraints on redox conditions, mineral-melt and mineral-fl uid reactions, and planetary differentiation.

*[email protected]

Johnson, C.M., McLennan, S.M., McSween, H.Y., and Summons, R.E., 2013, Smaller, better, more: Five decades of advances in geochemistry, in Bickford, M.E., ed., The Web of Geological Sciences: Advances, Impacts, and Interactions: Geological Society of America Special Paper 500, p. 1–44, doi:10.1130/2013.2500(08). For permission to copy, contact [email protected]. © 2013 The Geological Society of America. All rights reserved.

CELEBRATING ADVANCES IN GEOSCIENCE

1888 20138 02

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Signifi cant advances in microanalytical techniques have markedly improved experimental determinations of trace-element partitioning among phases and in characterizing elemental distributions in rocks and minerals using two- dimensional (2-D) and three-dimensional (3-D) mapping. Rare earth elements, in particular, have proved to be invaluable tracers of magmatic, sedimentary, aqueous, redox, and cos-mochemical processes, and siderophile trace elements form a basis for modeling many aspects of planetary accretion and early evolution. An anomalous amount of iridium at Mesozoic-Cenozoic boundary has revolutionized our view of one of Earth’s most important biologic extinctions.

Isotopic variations, whether produced by stable or radiogenic isotopes, provide a third dimension to the Periodic Table of Elements, and tremendous advances in instru-mentation since the early 1960s have greatly broadened this fi eld of geochemistry. Early work outlined the stable H and O isotope fi ngerprints of natural waters and water-rock interactions, and stable C and S isotope studies defi ned the biological frac-tionations that occur by photosynthesis and microbial sulfate reduction, respectively, topics that have since been applied to problems relating to the evolution of life and Earth’s atmosphere. Recent work on stable O isotopes has documented the likelihood that liquid water existed >4 b.y. ago on Earth, which profoundly affects our view of Earth’s evolution. New work on “nontraditional” stable isotopes has investigated redox cycling over Earth’s history, as has study of non-mass-dependent stable iso-tope variations. New approaches using stable isotopes as paleothermometers include exploiting the unique energetics of bonds between rare stable isotopes. Early work on the radiogenic Rb-Sr and U-Th-Pb isotope systems documented the key distinc-tions between continental crust and mantle, setting the stage for later tracing of mass fl uxes via plate tectonics, as well as documenting the great antiquity of continental crust formation and mantle differentiation on Earth. The Sm-Nd and Lu-Hf isotope systems provided a temporal context for earlier studies of rare earth element varia-tions in nature, including new constraints on crustal growth rates and mechanisms extending back earlier than 4 Ga. The siderophile Re-Os isotope system has been used to study the accretion of planetary bodies, core-mantle interaction, and the nature of the ancient lithospheric mantle.

The branch of geochemistry that deals with fossilized organic molecules had its origins in elucidating the processes and pathways that led to petroleum formation. As awareness of the richness and diversity of organic compounds that can be preserved in sedimentary rocks grew, this gave way to the broader endeavor of molecular paleo-biology. Despite great challenges in tying specifi c biomolecules to groups of organisms, or to metabolic processes, as well as issues of preservation mechanisms, molecular paleobiology remains a prime approach for studying the history of microorganisms, which have been the dominant life form for most of Earth’s history and yet are rarely preserved in the fossil record. Work on molecular biomarkers has produced numer-ous paleoenvironmental proxies for the chemistry and redox state (euxinia, anoxic, oxic) of the ancient oceans, as well as new paleoclimate records. The biochemical diversity of relatively simple life forms, including bacteria and archaea, has provided a wealth of lipid biomarkers that inform us about the evolution of metabolisms over Earth history, including oxygenic and anoxygenic photosynthesis, methanogenesis, and methanotrophy, and these records have been tied into stable isotope variations of many individual chemical elements (C, H, N, O, S, Fe, Mo, etc.), which provide a broad view of the biogeochemical evolution and biologically catalyzed redox cycling of Earth, and, potentially, other planetary bodies.

Although many geochemists focus exclusively on terrestrial problems, research over the past fi ve decades has been intimately linked to the chemistry of other solar system bodies and the universe beyond. We routinely rely on meteorite falls, inter-planetary dust particles, and Moon rocks for a baseline for comparison to Earth, which has been extensively differentiated and repeatedly resurfaced. Sophisticated

Five decades of advances in geochemistry 3


INTRODUCTION

Fifty years ago, in the early 1960s, modern geochemistry had begun to take shape, and from efforts during the fi rst half of the twentieth century, many of its fundamental questions had come into focus. Even biogeochemistry, arguably the most mod-ern branch of geochemistry, had been pioneered years earlier by Vernadsky and Baas Becking. The fi rst journal specifi cally devoted to the fi eld, Geochimica et Cosmochimica Acta, began publication in 1950. Two years later, the fi rst English-language geochemistry text was published by Mason (1952), building on the pioneering efforts of his mentor Goldschmidt (whose classic tome was published posthumously two years later; Goldschmidt, 1954). The year 1952 also witnessed Nobel Prize–winning chem-ist Harold Urey publish his seminal book, The Planets, widely considered a pivotal event in what would become the fi eld of cosmochemistry. In 1961, the American president John F. Ken-nedy laid down the challenge of sending humans to the Moon, and within a decade, intense competition for the opportunity to study returned lunar samples would result in major advances in geochemical techniques and thought.

On the other hand, the technical aspects of modern geochem-istry were very much in their infancy 50 yr ago. Three examples illustrate this in the period leading up to our 50 yr review time frame. High-pressure, high-temperature experimental methods, pioneered at the Geophysical Laboratory by Norman Bowen dur-ing the fi rst half of the twentieth century, were still only capable of exploring conditions relevant to the upper few kilometers of Earth by the early 1960s. In the early 1950s, distribution of the fi rst international geochemical rock standards, G-1 and W-1,

revealed a disturbing lack of agreement among the major geo-chemical laboratories of the world (e.g., Fairbairn et al., 1951), and analytical geochemists continued to struggle with data qual-ity until the demanding standards of lunar sample analysis seeped through the broader geochemical community. Eventually, major strides were made as geochemical analyses moved away from techniques based on wet chemistry to those based on X-ray or mass spectra, allowing a marked increase in sample throughput. The use of chemical modeling in geochemistry has taken place almost entirely within the past 50 yr (Bethke, 2008). Garrels and Thompson (1962) fi rst modeled the speciation of seawater, but the use of computers in geochemical modeling was not intro-duced until the work of Helgeson (1968).

Accordingly, our judgment is that developments in analyti-cal, experimental, and modeling methods have largely facilitated the major advances of modern geochemistry during the last fi ve decades. These developments have followed three simple but fun-damental themes: smaller, better, and more. By smaller, we mean the capability to analyze increasingly smaller samples and thus to evaluate geochemical problems at increasingly fi ner scales. For example, analytical capabilities now often achieve subfemto-gram (10-15 g) sensitivities, permitting study of vanishingly small samples, such as individual interstellar dust particles, or isotopic analysis of micron-sized regions of individual mineral grains. By better, we mean that modern geochemical methods permit ever-increasing precision, accuracy, and resolution, greater control on experimental conditions, and more stringent constraints on geo-chemical modeling. Thus, it is possible to measure many isotopic ratios with precisions approaching parts per million, and Moore’s law has witnessed >106 improvement in computing capabilities

remote-sensing capabilities based on past and current spacecraft missions are enabling active study of other planetary bodies such as the Moon, Mercury, and Mars. Ideas about nucleosynthesis within stars are tested by reference to the measured isotopic compositions of tiny presolar grains extracted from chondrites. Short-lived radio-nuclides in meteorites provide a detailed record of the condensation, mixing, and differentiation history of the earliest solar system. Mass-independent oxygen isotope fractionation in extraterrestrial samples may identify photochemical processes in the early solar nebula. More broadly, the temperature stabilities of elements and minerals constrain the sequence of nebular condensation, which provides a fi rst-order explana-tion for the bulk composition of the terrestrial planets relative to the planets of the outer solar system. Organic compounds from space inform us on the delivery of com-plex organic molecules to the early Earth, which likely infl uenced the earliest organic chemistry reactions, which in turn must have affected the origin and evolution of life. Chemical characterizations of samples of the Moon from the Apollo missions have provided the key data to recognize the Moon’s formation by impact of a Mars-size object with Earth and the likelihood that both bodies solidifi ed from magma oceans.

The individual subfi elds in geochemistry are becoming increasingly integrated, where systems are now viewed in a more holistic fashion, such as multi-element or multi-isotopic studies of biogeochemical cycles. Such approaches seem likely to con-tinue in the future, and they offer a comprehensive way to test multiple hypotheses and address geologic questions that continue to be important as we use geochemistry to better understand the geologic history of Earth and the solar system.

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over the past 50 yr. By more, we mean that with modern methods, it is possible to study increasing numbers of samples and ever-widening geological conditions. For example, high-throughput mass spectrometers allow for construction of high-resolution paleoclimate proxy records; improvements in high-pressure experimental equipment (e.g., diamond anvils) allow for experi-ments to be carried out at conditions equivalent to the centers of terrestrial planets; improvements in chemical models permit evaluation of increasingly high ionic strength aqueous condi-tions. Importantly, these advances in instrumentation occurred during, and after, the emergence of plate tectonics, and in terms of providing a framework for terrestrial geochemical studies, plate tectonics permeated work in the fi eld. Another conceptual advance in terrestrial geochemistry has been examination of Earth within a planetary perspective, with much insight coming from studies of meteorites and rocks from the Moon and Mars. Exploration of the solar system by spacecraft also occurred dur-ing this time, prodding the expansion of geochemical methods into remote sensing.

Our charge has been to summarize the remarkable progress that has been made in geochemistry during the last 50 yr in a journal-length chapter. We therefore begin with an apology to our colleagues for the large body of important work that must be omitted by the constraints imposed. Here, we are able to highlight only a small sample of the remarkable advances in geochemistry research in the last fi ve decades. This contribution is intended to be useful for geochemists who desire a glimpse of progress in areas of research beyond their own, as well as generalists who are curious about what has been going on in those laboratories down the hall.

GEOCHEMISTRY OF THE ELEMENTS

Major- and trace-element compositions of rocks, minerals, and natural fl uids have long been understood to be among the most fundamental data of geochemistry. Systematic evaluations of the abundances of the individual elements, and the basic laws governing their distributions in rocks and minerals were largely established by the 1960s, pioneered by the likes of Frank Clarke, Victor Goldschmidt, Louis Ahrens, and Ted Ringwood. Early analytical methods relied mainly on classical gravimetry and X-ray spectrographs to determine element compositions (Mason, 1992). During the 1950s, these laborious techniques began to give way to rapid and sensitive spectrophotometric methods, with the development and availability of instruments such as arc-source optical emission spectrographs and fl ame photometry, and chem-ical complexing agents (e.g., EDTA, ethylenediaminetetraace-tic acid) that could be used for colorimetry. Advances in bulk chemical analyses over the past 50 yr have witnessed astonishing developments of high-sensitivity instrumentation and methods (e.g., X-ray fl uorescence, plasma-emission spectroscopy, neu-tron activation, chromatography, thermal ionization and plasma-source mass spectrometry), allowing for increasingly rapid and precise measurements of increasingly smaller samples (e.g., Gill,

1997; Sutton et al., 2006). A variety of microbeam methods fur-ther allows for determination of major- and trace-element (and isotope) compositions of very small volumes of minerals, and to spatially accumulate such data into two-dimensional (2-D) maps and three-dimensional (3-D) tomographic images of the compo-sitions of rocks and minerals at micron to tens of micron scales of resolution.

Major-Element Geochemistry

Although early chemical analyses were time consuming, by the mid-twentieth century enough major-element data had been accumulated from igneous rocks to provide a basic appreciation of the ways in which fundamental magmatic processes (e.g., partial melting, equilibrium, and fractional crystallization) con-trolled observed variations (e.g., Harker and alkali-Fe-Mg [or AFM] diagrams). Because igneous and metamorphic processes are largely controlled by equilibrium, as shown, for example, by Norman Bowen’s studies of igneous systems, they are amenable to high-temperature, high-pressure experimental investigations, and experimental petrology has provided the fundamental back-ground for interpreting chemical compositions. This work con-tinues with the ability to carry out experiments at increasingly extreme conditions, for example, using diamond anvils that can attain pressures approaching the center of Earth (>300 GPa; Mao et al., 1990).

One area where fundamental advances in major-element geochemistry have been made over the past 50 yr is in micro-analysis. The electron microprobe was developed during the fi rst half of the twentieth century, but commercial probes only became available in the 1960s (Long, 1995). In order for these instru-ments to provide quantitative analyses of complex rock-forming minerals, improved understanding of the infl uences of atomic number, X-ray absorption and fl uorescence (so-called ZAF cor-rections), and various matrix corrections (e.g., Bence and Albee, 1968; Reed, 1995) was required. Applications of rapid quanti-tative analysis of minerals are legion. Among the pioneering studies were evaluation of phase relations in high-pressure, high-temperature experiments, studies of diffusion from elemental profi les through minerals, quantitative study of coexisting min-eral equilibria to constrain geothermometry, geobarometry, and oxygen fugacity (e.g., Andersen et al., 1993), and determination of mineral-melt partition coeffi cients in rocks and experiments allowing for quantitative geochemical modeling (see following).

In addition to the electron microprobe, other microbeam techniques have improved spatial resolution and sensitivity (i.e., detection limits) and allowed for elemental mapping (Jan-sen and Slaughter, 1982), including 3-D tomographic imaging (Jerram and Higgins, 2007). Examples include analytical elec-tron microscopy, secondary ionization mass spectrometry, laser-ablation sources for emission and mass spectrometers, and syn-chrotron X-ray fl uorescence microprobes. Figure 1 provides one example of modern geochemical mapping using a synchrotron X-ray fl uorescence microprobe.



Sedimentary Geochemistry and the Chemical Index of Alteration (CIA) Concept

Unlike igneous systems, where major-element distributions have been studied in great detail for decades, interpretations of the elemental geochemistry of sedimentary systems until recently lagged far behind. The main reason for this is that siliciclastic sediments (sandstones, shales) are mostly physical mixtures and have been affected by multiple episodes of kinetically dominated water-rock interaction prior to, during, and after sedimentation. Accordingly, there was no simple measure, analogous to partition coeffi cients or Harker diagrams, that quantitatively linked bulk chemistry to process. Well i nto the 1970s, textbooks routinely tabulated chemical analyses of sedimentary rocks (e.g., Pettijohn, 1975) but provided little in the way of quantitative interpretation. Weathering processes were understood but focused primarily on gains/losses in soil profi les to evaluate weathering intensity as an index of paleoclimates. Experiments and fi eld investiga-tions made signifi cant progress in evaluating the kinetics and time scales of weathering (Brantley and Lebedeva, 2011; White and Brantley, 1995). Garrels and Mackenzie (1971) fi rst began to describe the relations between low-temperature aqueous geo-chemistry (e.g., mineral aqueous stability diagrams) and the bulk composition of sedimentary rocks.

Accordingly, a major advance in sedimentary geochemistry was development of the so-called chemical index of alteration (CIA) concepts during the 1980s and 1990s, in a series of papers by Wayne Nesbitt and coworkers (summarized in Nesbitt, 2003).

Figure 1. Synchrotron X-ray fl uorescence continu-ous scan mapping of a Jurassic crinoid stem taken at the Brookhaven National Synchrotron Light Source Beamline X-26A. Image is 6.1 mm by 5.3 mm, col-lected with a pixel size of 10 mm at a 100 ms/pixel collection rate. Three maps of Ca (green), Fe (blue), and Sr (red) are superimposed, with green being suppressed in this image. Areas of high strontium content (red) are magnesian calcite, whereas the blue areas are ferroan calcite infi ll. Image is courte-sy of Steve Sutton and Tony Lanzirotti (University of Chicago/Brookhaven) and Troy Rasbury (State University of New York at Stony Brook).

Developed fi rst to evaluate weathering in soil and weathering profi les (Nesbitt and Young, 1984), this approach also allowed for quantitative, petrologically based understanding of the major-element composition of siliciclastic sediments. Over the past sev-eral decades, few papers reporting the chemical composition of siliciclastics have omitted geochemical concepts in interpreting the data.

One especially infl uential diagram is the “feldspar ternary” diagram (in mole fraction), Al

2O

3–(CaO* + Na

2O)–K

2O, or

A-CN-K (Fig. 2), where CaO* is CaO in silicate minerals only (i.e., corrected for carbonates and phosphates). The A-CN-K ter-nary diagram captures most of the major-element (and mineral-ogical) changes observed in weathering of igneous rocks (i.e., alteration of feldspar and glass to clay minerals). The CIA scale (CIA = 100 × Al

2O

3/[Al

2O

3 + CaO* + Na

2O + K

2O]) is shown

on the left side of Figure 2. This diagram also proves useful for interpreting the geochemistry of sedimentary rocks for several reasons: Minerals most relevant to sedimentary rocks plot at well-separated locations on the apices or along joins, and two-component mixing/unmixing relationships plot as straight lines and follow the lever rule. Accordingly, these relations have been used to quantify or constrain (1) the degree of weather-ing affecting sedimentary source regions, which typically has paleoclimatological implications; (2) mineral sorting and simple two-component mixing; (3) average provenance and mixing of provenance/mineral components; and (4) diagenesis. Variants on this diagram are the A-CNK-FM and A-CNKM-F diagrams (F-FeO

T; M-MgO), which have less thermodynamic/kinetic

foundation (Nesbitt and Young, 1984) but are especially useful in basaltic systems (Nesbitt and Wilson, 1992), and thus have been used in planetary applications (Hurowitz and McLennan, 2007; McSween and Keil, 2000).

Trace-Element Geochemistry

During the past fi ve decades, some of the most signifi cant advances in elemental chemistry have had to do with quantitative interpretations of trace-element distributions in rocks and min-erals (as pioneered by Louis Ahrens, Paul Gast, Larry Haskin, Denis Shaw, and Ross Taylor), and these advances have naturally mirrored developments in analytical geochemistry.

Partition Coeffi cientsOne such advance was the determination of partition coef-

fi cients (solid-liquid, solid-gas; expressed as Kd) from experi-

ments and natural rock systems over a broad range of pressure-temperature-composition, beginning in earnest during the 1970s and 1980s. These data provided the foundation to the fi eld of trace-element modeling of igneous processes (see reviews in Allègre and Minster [1978] and Green [1994]). Early studies were limited by the detection limits of the electron microprobe used to measure trace-element abundances in minerals and melts. Experiments were highly constrained by the necessary balancing act between having trace-element abundances high enough to be

6 Johnson et al.


detected by the microprobe but not exceeding the Henry’s law limit of remaining a trace element within the system (i.e., a

i =

kxi as x

i ⇒ 0 in some phase where i = element of interest, a =

activity, x = molar concentration, and k = Henry’s law constant). The development of highly sensitive modern microbeam tech-niques (e.g., ion microprobe, synchrotron X-ray fl uorescence, laser-ablation mass spectrometry) now allows for precise deter-mination of a wide range of partition coeffi cients, including those for highly incompatible elements (i.e., K

dmineral/melt < 10-3), under

conditions that represent nature (e.g., Frei et al., 2009a).

Rare Earth Elements The rare earth elements (REEs; La–Lu, Y1) are the most stud-

ied and infl uential group of trace elements, and they have provided crucial evidence for a wide variety of geochemical and cosmo-chemical processes. In addition to elemental abundances, the REEs are used for several important radiogenic isotope systems (e.g.,

Figure 2. Ternary diagrams plotting molar proportions of Al2O

3–(CaO* + Na

2O)–K

2O, or A-CN-K. CaO* refers to cal-

cium in the silicate components only (i.e., corrected for carbonates and phosphates). The chemical index of alteration (CIA) scale is shown on the left side of the diagram. Also plotted are selected igneous and sedimentary minerals, average compositions of major igneous rock types, average shale, and the range for most river waters. The horizontal line con-necting plagioclase and K-feldspar separates the lower part of the diagram dominated by primary igneous minerals (i.e., unweathered) from the upper part dominated by clay minerals (i.e., weathered). (A) Typical weathering trends where the arrows schematically represent the general pathways observed for increasing degrees of weathering for various rock types, based on fi eld studies of weathering profi les and on thermodynamic-kinetic modeling. (B) Plot of suspended sedi-ment from many of the world’s major rivers, illustrating how the overall effects of weathering are preserved within the composition of sediments.

Sm-Nd, Lu-Hf), discussed later herein. REEs have been used to address problems ranging from constraining the earliest history of the solar nebula based on variations in condensation temperatures, coupled with measurements of their abundances in Ca-Al–rich inclusions and minerals in meteorites (Mason and Taylor, 1982), to tracing the movement of water masses through the oceans based on their short residence times (Piepgras and Wasserburg, 1980). The past fi ve decades have also witnessed developments in the use of REEs as components for a variety of high-technology applications, such that these elements are now considered to be strategic met-als, thus providing further impetus for future geochemical research (Haxel et al., 2005). Modern research on the REEs dates from the development of effi cient separation methods and high-precision instrumental analytical techniques in the early 1960s that resolved, in the affi rmative, the long-standing question of whether or not these elements could be fractionated during formation of Earth’s crust (Haskin and Gehl, 1962).

By geochemical standards, REEs are an extremely coherent group in terms of size (ionic radius), charge, mineral cation site coordination, lithophile behavior, and aqueous complexing and speciation characteristics. From a planetary perspective, REEs

1Scandium may also be considered a REE but typically is not included with the others in geochemical discussions due to its smaller size and differing partition-ing behavior.



do not fractionate signifi cantly during planetary formation from the solar nebula, and, accordingly, average chondritic meteorite abundances serve as a useful reference for examining planetary and geological processes. Under geological conditions, REEs are trivalent, except for the distinctive redox chemistries of europium (Eu3+ or Eu2+) and cerium (Ce3+ or Ce4+), which result in unique insights into magmatic and aqueous processes, respectively. Reduction of Eu (17% increase in ionic radius) occurs under highly reducing conditions, which, with rare exception, exist only within magmatic environments. On the other hand, oxidation of Ce (15% decrease in ionic radius) is common under surfi cial aqueous conditions, such as those encountered during weathering.

The REEs have proven particularly valuable for constraining magmatic processes because their mineral/melt partition coeffi -cients, which (apart from Eu2+) vary smoothly as a function of atomic number, vary over several orders of magnitude among the common rock-forming minerals (Fig. 3). For example, the ionic radius of Eu2+ is virtually identical to that of Sr2+ and readily substitutes for Ca2+ in plagioclase. Accordingly, the presence of Eu anomalies in magmatic rocks commonly results from frac-tionation of plagioclase during partial melting or crystallization. Since plagioclase is stable only up to ~1 GPa pressure (~40 km depth on Earth), the presence of Eu enrichments or depletions in magmatic rocks generally indicates relatively shallow condi-tions. The ubiquitous presence of chondrite-normalized negative Eu anomalies in sedimentary rocks is thus interpreted to indicate that intracrustal igneous differentiation processes dominated for-mation of the upper continental crust (the source of sedimentary REEs). In another example, the presence of very steep REE pat-terns that are depleted in heavy REEs (high Gd/Yb) in magmatic rocks is taken to indicate fractionation of garnet, a mineral only stable at pressures higher than ~1 GPa in ultramafi c systems and thus indicative of mantle origins. Accordingly, the origin of steep REE patterns in the ubiquitous Archean tonalite-trondhjemite-granodiorite suites has been central to development of models for the origin of Archean continental crust, and determination of whether or not the modern style of plate tectonics was operating at that time (Taylor and McLennan, 2009).

Although all REEs are cosmochemically refractory elements (50% T

condensation > 1300 K at 10 Pa), slight differences in their

condensation temperatures lead to remarkably complex REE patterns in 4.567 Ga refractory Ca-Al–rich inclusions, the oldest material preserved in certain chondritic meteorites. In addition to unusual and highly variable overall shapes, the REE patterns of these objects also have both positive and negative anomalies involving the least refractory REEs (Ce, Eu, and Yb). These pat-terns thus provide persuasive evidence for very complex but rela-tively localized evaporation-condensation processes in the early solar nebula (Mason and Taylor, 1982).

The REEs have proven to be extremely useful tracers for understanding a wide range of aqueous processes (Byrne and Sholkovitz, 1996), primarily because they tend not to partition into the aqueous fl uid (Dfl uid/solid <<1) during fl uid-rock interac-tions. Accordingly, REEs have been shown to be resistant to

Figure 3. Plot of mineral-melt partition coeffi cients (Kd)

for the rare earth elements (REEs) in major minerals in equilibrium with basaltic melts. Data are from a compila-tion in White (2012).

remobilization beyond the mineralogical scale during weather-ing, diagenesis, and metamorphism, except under conditions of very high fl uid/rock ratios. REEs are particle-reactive and tend to be effi ciently scavenged from seawater by sediment particles, resulting in very low residence times, ranging from ~50 yr (Ce) to ~2900 yr (Lu). In seawater, Ce is readily oxidized on Mn-oxide particle surfaces, and insoluble ceric oxides and hydroxides are then preferentially removed from seawater compared to the other trivalent REEs, resulting in extreme negative Ce anomalies. The stability constants for most REE complexes tend to increase with increasing atomic number, resulting in the light REEs (La–Sm) being preferentially scavenged over Gd–Lu. The overall effect of these characteristics is that REEs in natural fl uids vary by many orders of magnitude in absolute abundances, with highly variable chondrite-normalized patterns, providing a very sensitive tracer.

8 Johnson et al.


Siderophile and Platinum Group ElementsA second group of trace elements that has proven especially

useful for evaluating geochemical phenomena is the siderophile trace elements in general, including the refractory and highly siderophile platinum group elements (PGEs = Ru, Rh, Pd, Os, Ir, Pt). The major value of highly siderophile elements is their strong partitioning into metal phases during metal-silicate equi-librium (Dmetal liquid/silicate liquid >> 103; Righter, 2003). Accordingly, their distributions provide insight into the nature of core-forming processes in planets and in the parent bodies of the various mete-orite classes, and into the evolution of planetary mantles after extraction of cores.

One key discovery, based on PGE abundances, was that the mass extinction at the ca. 65 Ma Mesozoic-Cenozoic ( Cretaceous-Paleogene; K-Pg) boundary was related to the impact of an asteroid, thus opening the door to an entirely new way to consider the relations between biologic evolution and geo-logic history. Alvarez et al. (1980) determined that clay-rich sedi-ment deposited exactly at the K-Pg boundary was highly elevated in iridium, typically by ~20–200 times the levels in the enclosing sediments (Fig. 4). Since PGEs (including Ir) occur at very low levels in Earth’s mantle and crust due to sequestration into the metallic core during planetary differentiation, they argued that the likely source of globally distributed Ir was from impact of a large meteorite. Mass-balance calculations indicated the impactor, assuming a CI chondrite composition, would have been ~10 km in diameter. Much subsequent work, including the presence within the K-Pg sediment of additional PGE enrichments, as well as Re-Os isotope data, in addition to the presence of soot, shocked quartz, stishovite, and impact glass spherules, reinforced this model, and discovery of the buried ~200-km-diameter Chicxulub crater in Mexico with a 65.0 Ma age (e.g., Swisher et al., 1992), has confi rmed that a major impact took place at that time.

Siderophile element distributions are also important in cos-mochemistry (McSween and Huss, 2010). In addition to elemental abundances, siderophile elements possess important radiogenic isotope systems (e.g., Re-Os, Pt-Os, Hf-W) that provide details about planetary differentiation time scales (Kleine et al., 2009; Shirey and Walker, 1998). The family of siderophile elements encompasses a broad range of siderophile tendencies (as measured by metal-silicate partition coeffi cients under different pressure [P], temperature [T], and oxygen fugacity [f

O2] conditions), volatilities

(condensation temperatures), and incompatibilities (mineral/sili-cate melt partition coeffi cients). Accordingly, they provide insights into core-mantle differentiation and silicate mantle melting in the terrestrial planets, Moon, and meteorite parent bodies (Righter, 2003; Righter and Drake, 1996; Walker, 2009).

ISOTOPE GEOCHEMISTRY

The fi eld of isotope geochemistry offers the opportunity to look at the Periodic Table of Elements in a “third dimension,” where isotopic variations may result from stable isotope frac-tionation or radioactive decay. Isotope geochemistry provides

an important means with which to trace mass fl uxes among Earth’s reservoirs that are independent of concentration effects in the sense of activity coeffi cients and the thermodynamics of mixing. Historically, stable and radiogenic isotopes were studied by different laboratories, refl ecting the distinct instrumentation required by various techniques, but such boundaries have now blurred. Our theme of “smaller, better, more” is embodied in the history of development of mass spectrometers. Modern mass spectrometers capable of high-precision isotopic measurements are based on the Nier (1940) geometry, and by the 1960s, the isotope ratio mass spectrometers (IRMS) available to the stable isotope community had a well-established dual-inlet system for comparing sample and standard gases, as well as simultane-ous collection of two isotopes. Later, continuous-fl ow systems would come online (Hayes et al., 1990), and after that, online laser fl uorination systems (Sharp, 1990). The thermal ionization mass spectrometers (TIMS) available to the radiogenic isotope community 50 yr ago, however, could not match the precision of dual-inlet, double-collector IRMS instruments. As electron-ics continued to improve, as well as the addition of multicol-lection, precisions attainable using TIMS instruments increased markedly in the 1980s (Thirlwall, 1991). Negative ion-capable TIMS later provided a breakthrough for the Re-Os isotope sys-tem (Creaser et al., 1991). A major innovation for both stable and radiogenic isotope geochemistry was development of a multicol-lector, magnetic-sector–based inductively coupled plasma–mass spectrometer (MC-ICP-MS) in the 1990s (Halliday et al., 1998). Finally, efforts have focused on in situ isotopic analysis at the micron scale from the 1990s to present. This includes second-ary ion mass spectrometry (SIMS, or ion microprobe), where the current state-of-the-art, large-radius, multicollector instru-ments can determine d18O precisions, for example, of ±0.3‰ on ~10 µm spots (Valley and Kita, 2009). Laser-ablation (LA) coupled to MC-ICP-MS has also emerged as an important tech-nique for in situ isotopic analysis, a technique widely applied, for example, to Hf isotope analysis (Griffi n et al., 2000), and new research is currently investigating ultrafast (femtosecond; 10-15 s) lasers (Poitrasson et al., 2003).

Next, we touch on some of the major stable and radiogenic isotope systems used in geochemistry in the last fi ve decades. Space limitations prevent us from covering some important isoto-pic systems, such as rare gas isotopes, which have seen important applications, for example, to mantle evolution using 3He/4He ratios (Kurz et al., 1982). Other important isotopic systems we have omitted include cosmogenic radionuclides, which, in solid Earth geochemistry, have provided unique constraints on sediment sub-duction using isotopes such as 10Be (Tera et al., 1986).

Stable Isotopes

Broadly, stable isotope variations refl ect partitioning between phases that results from differences in zero-point ener-gies for isotopically substituted species, although in detail, com-plexities exist that are mass-independent or refl ect nonstochastic



distribution of rare isotopes. The basic thermodynamic frame-work for stable isotope geochemistry was laid out by Bigeleisen and Mayer (1947) and Urey (1947). Here, we fi rst focus on the main players that existed 50 yr ago, H, C, O, and S, followed by a discussion of newer stable isotope systems that have been developed since then.

Hydrogen and Oxygen IsotopesThe importance of water in the geologic cycle logically

placed H and O isotopes as early targets in stable isotope geo-chemistry (Fig. 5). Signifi cant accomplishments in the 1960s were numerous. The canonical relation between D/H and 18O/16O ratios of meteoric waters was determined (Craig, 1961), defi ning the meteoric water line, a relation that is a key factor to understanding the origin of fl uids and fl uid-rock interactions. The kinetic and equilibrium isotopic fractionations associated with meteoric waters, as well as temperature, latitude, and oro-graphic effects, were established (Craig et al., 1963; Friedman

Figure 4. Stratigraphic variations in iridium (Ir) concentrations (parts per billion by mass) in sedimentary rocks at the Cretaceous- Paleogene (Tertiary in older nomenclature) boundary in the vicinity of Gubbio, Italy. Analyses are for 2 molar nitric acid insoluble resi-dues of limestones and boundary clay. The highly elevated Ir abun-dances at this stratigraphic boundary were the fi rst direct evidence that a major meteorite impact was involved with the mass extinctions at the 65 Ma Mesozoic-Cenozoic boundary. Figure is adapted from Alvarez et al. (1980).

Figure 5. δD-d18O variations for waters and rocks. Meteoric water line (MWL) and standard mean ocean water (SMOW) are shown for refer-ence, as is box for magmatic waters. Hydrothermal fl uids are shown for Salton Sea (Craig, 1966) and Lassen (Janik et al., 1983), which are shifted from the MWL, where the arrows mark increasing rock infl uence (low water/rock ratios). Sedimentary formation waters may lie along a slope lower than that of the MWL, refl ecting kinetic effects of evaporation (fi eld shown for oil-fi eld brines from California; see Sheppard, 1986). Hydrothermally altered igneous rocks have distinct δD-d18O trends compared to waters, and data shown for the Idaho batholith (Criss and Taylor, 1983); arrow marks direction of increas-ing water infl uence (high water/rock ratios). Diagram is adapted from those in Criss (1999).

et al., 1964), and especially large seasonal and temperature effects were documented in the d18O values of precipitation in polar regions (Dansgaard et al., 1969; Epstein et al., 1963), set-ting the stage for use of oxygen isotopes from ice cores to infer paleoclimate. The δD-d18O relations of hydrothermal waters that underwent fl uid-rock interaction were recognized to be different from those produced by evaporation of meteoric waters (Craig, 1966). Initial efforts began to study the d18O values of the ancient oceans when Perry (1967) suggested that the Precambrian oceans had much lower d18O values than the modern ocean based on analysis of cherts. Muehlenbachs and Clayton (1976) proposed the important concept that the d18O value of seawater was buff-ered near zero by extensive water-rock interactions at mid-ocean ridges (MORs), which would imply that the d18O value of ancient seawater would be roughly invariant as long as plate tectonics operated. A landmark study of the Skaergaard intrusion com-bined detailed fi eld studies with numerical modeling of convec-tive heat transport to quantify water-rock interaction at a fossil hydrothermal system (Norton and Knight, 1977; Norton and Taylor, 1979). Studies of ophiolites (Gregory and Taylor, 1981) confi rmed that extensive hydrothermal interaction at MORs has

10 Johnson et al.


been responsible for fi xing the d18O value of seawater. If the d18O value of seawater was relatively constant, the low-d18O values of Precambrian cherts could be interpreted to refl ect very high ocean temperatures (Knauth and Epstein, 1976), although the possibility exists that very old samples have later exchanged with fl uids that lowered their d18O values after deposition.

Oxygen isotope studies of igneous and metamorphic rocks in the 1970s and 1980s focused on (1) use of d18O values as a tracer of the sources of magmas, and (2) application of O isotope thermometry to metamorphism (Valley, 1986) and hydrother-mal alteration (Gregory and Criss, 1986). The large decreases in d18O values that occur from meteoric hydrothermal alteration at elevated temperature (e.g., Criss and Taylor, 1983), or the large increases in d18O that occur during weathering (e.g., Savin and Epstein, 1970) produce a wide range in d18O values for upper-crustal rocks that may be incorporated into magmas. Elevated d18O values are characteristic of “S-type” granites, a term used for granites that contain a sedimentary component, and con-fi rmed by O isotopes (O’Neil et al., 1977). In contrast, Friedman et al. (1974) discovered that magmas could attain low-d18O val-ues through interaction with meteoric waters directly, or through assimilation of hydrothermally altered country rocks. Collec-tively, this discussion points to one of the most important contri-butions of stable O isotope geochemistry: A rock that has a d18O value that is distinct from that of the mantle probably contains O that was cycled near the surface of Earth in the presence of water (Taylor and Sheppard, 1986).

One of the most profound demonstrations of the use of O isotopes as a “water tracer” is found in the >4 Ga zircons from the Jack Hills, Western Australia, which have elevated d18O val-ues, providing compelling evidence that liquid water existed on Earth over 4 b.y. ago (Cavosie et al., 2005; Mojzsis et al., 2001; Peck et al., 2001). This work, in fact, highlights the importance in isotope geochemistry of moving away from bulk sample analy-sis to in situ approaches (e.g., Valley and Kita, 2009). Without the ability to make precise, and accurate, O isotope analyses of micron-sized spots, the elevated d18O values found in the Jack Hills zircons would not have been discovered, and the concept of a “cool early Earth” (Valley et al., 2002) would not have arisen.

Carbon IsotopesBy 1960, it was already well established that photosynthetic

fi xation of CO2 into organic carbon produced a decrease in d13C

values by ~25‰–30‰ (Park and Epstein, 1960). In the follow-ing years, the fi rst surveys of d13C values for marine carbonates and organic C confi rmed that the overall isotopic fractionation between these C reservoirs can be found in natural samples, including those from Precambrian rocks (Hoefs and Schid-lowski, 1967; Keith and Weber, 1964). As the database for C isotope compositions of carbonates greatly expanded in the fol-lowing decade, it became clear that the vast majority of Ca-Mg marine carbonates had a restricted range in d13C values near zero for most of Earth history, which was interpreted to refl ect a relatively constant balance between the organic and inorganic

C reservoirs (Schidlowski et al., 1975). An important exception was Paleoproterozoic carbonates, which had unusually positive d13C values, fi rst documented in the Lomagundi Group, Rhodesia (Schidlowski et al., 1975), and later shown to be global and cor-relative with a major rise in atmospheric oxygen, likely, at least in part, refl ecting organic C burial (Karhu and Holland, 1996). In contrast to the generally zero d13C values for Ca-Mg carbonates, Fe-rich carbonates from Precambrian banded iron formations (BIFs) have signifi cantly negative d13C values. Early workers suggested microbial oxidation of organic carbon as an explana-tion (Becker and Clayton, 1972), but later work favored verti-cal zonation in d13C values for dissolved inorganic carbon (DIC) in the oceans (Beukes et al., 1990; Winter and Knauth, 1992). More recently, geochemical modeling and studies of shelf-to-basin transects suggest that vertical zonation in d13C values for DIC is unlikely, swinging the interpretation for negative d13C val-ues back to microbial respiration (Beukes and Gutzmer, 2008; Fischer et al., 2009).

By the 1970s, there was already a substantial database for C isotope compositions of kerogen from Precambrian shales as old as 3.4 Ga, documenting that d13C values for organic carbon generally lay between -25‰ and -35‰ (Oehler et al., 1972). An important exception was discovery of highly negative d13C val-ues for organic carbon in sedimentary rocks between ca. 2.8 and 2.6 Ga in age, down to −60‰. Such low values are accepted to refl ect a role for involvement of methane, and Hayes (1983) spec-ulated that aerobic methanotrophy might have been responsible. Follow-up work has confi rmed these unusually low d13C values and documented correlations with sedimentary facies that sug-gest a transition from an anaerobic ecosystem to one supported by oxygenic photosynthesis at the end of the Archean (Eigen-brode and Freeman, 2006).

Looking at the oldest known sedimentary rocks, Schid-lowski et al. (1979) noted that graphite in metasedimentary rocks from the Isua belt, SW Greenland, did not generally reach the low-d13C values that are characteristically thought to refl ect pho-tosynthetically fi xed C, and this was interpreted to refl ect the effects of metamorphism (Schidlowski, 1987). The importance of obtaining “primary” d13C values from C from Early Archean rocks, despite their commonly high grade of metamorphism, has generated numerous studies. Mojzsis et al. (1996) docu-mented d13C values for graphite from Akilia Island, SW Green-land, obtained via in situ SIMS analysis, that were signifi cantly lower than those measured in bulk samples, but this work has come under criticism on a number of fronts, including arguments (1) that the graphite is not photosynthetic in origin, but instead a breakdown product of Fe-bearing carbonates (van Zuilen et al., 2002), (2) that the sample analyzed was not a sedimentary rock but a metasomatic dike (Fedo and Whitehouse, 2002), and (3) that the graphite analyzed was not enclosed in apatite and therefore isolated from the effects of metamorphism (Lepland et al., 2005). The question of when the fi rst C isotope compositions that suggest photosynthesis appeared on Earth remains an impor-tant one, although there is no consensus on the answer.



Recent work in C isotope geochemistry has expanded into in situ isotopic analyses, including studies of organic carbon and individual microfossils (House et al., 2000). Increasingly, it is rec-ognized that the range in d13C values for individual microfossils, or kerogen on submillimeter scales, is much larger than would be suggested by bulk C isotope analysis. Such an approach has affected the range in C isotope fractionations inferred between organic and inorganic C, which in turn may constrain atmo-spheric CO

2 contents (Kaufman and Xiao, 2003), test the indig-

enous nature of molecular biomarkers (Rasmussen et al., 2008), or distinguish between eukaryotic and cyanobacterial photosyn-thesis (Williford et al., 2013).

Sulfur IsotopesEarly studies of S isotope variations in the laboratory and

natural environments had documented a large fractionation in 34S/32S ratios during microbial sulfate reduction that was depen-dent upon the rate of reduction and the abundance of sulfate, and similar ranges in isotopic compositions had been documented in modern marine sediments (Kaplan et al., 1963; Thode et al., 1961). In contrast to the strongly positive d34S values measured for Phanerozoic sulfates, Perry et al. (1971) found that older than 3 Ga barites from South Africa had only slightly positive d34S values, which they interpreted to refl ect very low seawa-ter sulfate contents, suggesting very low atmospheric O

2 con-

tents. Detailed experimental work on microbial reduction and S disproportionation in the 1990s demonstrated that extremely large 34S/32S fractionations could be produced during microbial S cycling that involved S species of intermediate oxidation state such as sulfi te and thiosulfate (Canfi eld and Thamdrup, 1994; Canfi eld and Teske, 1996). This in turn suggested that the excep-tionally low d34S values in sedimentary sulfi des of younger than 1 Ga age refl ected an increase in the oxidative component of the S cycle, which in turn provided strong evidence for a major rise in atmospheric O

2 contents in the Neoproterozoic. The decrease

in d34S values of sulfi des in the Neoproterozoic, relative to ear-lier time periods, was accompanied by an increase in the esti-mated d34S values for seawater sulfate (Canfi eld, 2001).

Burdett et al. (1989) proposed that the isotopic composi-tions of seawater sulfate may be determined through analysis of carbonate-associated sulfate (CAS), an approach that greatly extends the lithologies that may be used to infer ancient seawater S. The CAS proxy proved particularly valuable for the ancient rock record, where the d34S values of ancient seawater sulfate had been previously inferred indirectly based on the maximum values obtained in a suite of sedimentary sulfi des. Application of the CAS proxy to the Proterozoic has confi rmed expectations that seawater sulfate contents were very low in the Paleoproterozoic, before or immediately after the Great Oxidation Event (GOE), but rose substantially in the Neoproterozoic at the time of the second increase in atmospheric oxygen (Kah et al., 2004).

The last decade has seen a rapid increase in in-situ S isotope studies, primarily focused on pyrite. These efforts have shown that detrital, authigenic, and hydrothermal components may

be recognized in the same single pyrite grain (Williford et al., 2011). Work on ca. 3.4 Ga rocks from the Pilbara craton, Aus-tralia, has documented signifi cant S isotope variations on the micron scale, which have been generally interpreted to refl ect biological cycling of S, including S0 disproportionation and sul-fate reduction (Philippot et al., 2007; Wacey et al., 2011). A very large range in d34S values, >35‰, was observed in pillow lava textures in the Barberton greenstone belt, which were interpreted to refl ect microbial microboring, and which McLoughlin et al. (2012) suggested provides evidence for a Paleoarchean subsea-fl oor biosphere.

Mass-Independent Stable IsotopesMost stable isotope variations in terrestrial systems fraction-

ate in a mass-dependent manner. Mass-independent fraction-ations (MIF) for O isotopes (and others) are commonly observed in extraterrestrial samples (Birck, 2004), and will be discussed later in this review. Sulfur MIF (commonly termed “S-MIF”) has been studied extensively in ancient sedimentary rocks as a tracer of past atmospheric O

2 contents. The most common mechanism

called upon to produce S-MIF is photolysis reactions involving SO

2 and H

2S in the upper atmosphere, based on ultraviolet (UV)

radiation experiments (Farquhar et al., 2001, 2000b). Because ozone is a strong absorbent of UV in the atmosphere, its pres-ence would greatly decrease S-MIF in aerosols, which sug-gests that the S-MIF recorded in Archean and Paleoproterozoic rocks indicates very low atmospheric O

2 contents (Farquhar et

al., 2000a). Photochemical modeling suggests that transport of S-MIF to the sedimentary cycle requires ambient atmospheric oxygen contents less than 10−5 of present day (Pavlov and Kast-ing, 2002). The specifi c mechanisms and pathways responsible for creating S-MIF in the ancient rock record, however, remain unclear, where initial UV experiments may not have adequately modeled the full spectrum of UV radiation, and the roles of other gases, including methane and inert gases, are issues that have been raised (e.g., Domagal-Goldman et al., 2008; Lyons, 2009; Masterson et al., 2011; Zahnle et al., 2006). Further complex-ity arises from the fi nding that thermochemical sulfate reduction by organics can produce mild S-MIF (Watanabe et al., 2009). Nevertheless, most workers accept S-MIF in Archean and Paleo-proterozoic sedimentary sulfi des as indicating essentially anoxic conditions in Earth’s atmosphere at this time, although research continues on the ancient rock record, as well as the mechanisms to produce MIF.

Nontraditional Stable IsotopesNearly three quarters of the elements on the Periodic Table

of Elements have two or more stable isotopes, but beyond the stable isotope systems discussed here, and a few others, most elements have remained relatively unexplored due to analytical barriers or the opinion that the range of isotopic variations in nature is too small to be worth the trouble. Vanguard efforts in the “nontraditional” stable isotopes include early studies of Li (Chan, 1987), Si (Douthitt, 1982), and Ca (Russell et al., 1978)

12 Johnson et al.


isotopes. The largest isotopic variations of the “nontraditional” stable isotopes are observed for Li, where the range in 7Li/6Li in nature exceeds 75‰ (Tomascak, 2004). Smaller variations are found for intermediate-mass elements, where, for example, 26Mg/24Mg, 44Ca/40Ca, 53Cr/52Cr, 56Fe/54Fe, 65Cu/63Cu, 80Se/76Se, and 98Mo/95Mo ratios vary by ~5‰–10‰ in natural samples (e.g., Albarède, 2004; Anbar, 2004; Beard and Johnson, 2004a; DePaolo, 2004; Johnson and Bullen, 2004; Young and Galy, 2004). Results from additional stable isotope systems are being regularly reported at meetings and in publications, extending all the way up to mass U (e.g., Weyer et al., 2008). The fi eld is changing so rapidly that the fi rst review of the subject in 2004 is now signifi cantly out of date (Johnson et al., 2004).

One of the most intensively studied “nontraditional” stable isotope systems has been Fe. The large bonding changes that occur between Fe3+ and Fe2+ species in solutions and minerals were recognized early as the major driving force for producing Fe isotope fractionations (Beard et al., 1999; Bullen et al., 2001; Johnson et al., 2002; Polyakov and Mineev, 2000). The vast majority of Fe in the crust has a δ56Fe value near zero, includ-ing low-C, low-S sedimentary rocks (Beard et al., 2003a, 2003b) and most igneous rocks, although there are small variations in high-temperature rocks (Beard and Johnson, 2004b). The largest range in δ56Fe values, however, is restricted to organic-rich shales and BIFs, which vary from −4‰ to +2‰ (Johnson et al., 2008a; Rouxel et al., 2005; Yamaguchi et al., 2005). The origin of such large variations is debated, where some workers argue that they refl ect variable extents of oxidation of aqueous Fe2+ (Anbar and Rouxel, 2007), and others interpret such variations, particularly in Fe-rich rocks, to refl ect microbial Fe3+ reduction (Johnson et al., 2008b). Recently, multiple stable isotope systems have been used to study redox-driven geochemical cycling, greatly expand-ing the utility of stable Fe isotopes, including Fe and S isotopes (Archer and Vance, 2006), Cr isotopes and Fe redox changes (Frei et al., 2009b), Fe and C isotopes, including carbonate C (Heimann et al., 2010) and organic C (Czaja et al., 2010), and Fe and Mo isotopes (Czaja et al., 2012). In summary, the fi eld of nontraditional stable isotopes is growing rapidly, although progress is currently hampered by the relative paucity of stable isotope fractionation factors as compared to other stable isotope systems that have been studied for several decades.

Clumped Stable IsotopesStable isotope compositions generally consider substitution

of a single minor isotope, such as 13C16O16O or 12C18O16O in car-bon dioxide because multiply substituted minor isotopes, such as 13C18O16O, are very low in abundance in nature. Recently, how-ever, the unique aspects of multiply substituted minor isotopes have been exploited (Wang et al., 2004). Now termed “clumped-isotope” geochemistry, to describe, for example, the enhanced stability of 13C-18O bonds in CO

2 and carbonates relative to a

random or stochastic distribution of minor isotopes, this fi eld of stable isotope geochemistry promises great breakthroughs in paleoclimate studies, which require accurate determination

of paleotemperatures, as well as the study of atmospheric gases (Eiler, 2007). For CO

2 and carbonate, the measured enrichment

in 13C-18O bonds relative to that expected for a random distribu-tion of 13C-18O bonds is defi ned as Δ

47, refl ecting nominal mass

47 for 13C18O16O, and this has been shown to be an inverse func-tion of temperature (Ghosh et al., 2006). Because the enhanced stability, or “clumping,” of 13C-18O bonds, under equilibrium con-ditions, refl ects internal, homogeneous equilibrium that is inde-pendent of the bulk d13C or d18O values, Δ

47 provides an “inter-

nal” thermometer that does not require knowledge of the isotopic composition of the fl uid from which the carbonate precipitated. This critical aspect of “clumped-isotope” geochemistry obviates the need to know, or assume, the d18O or d13C value of ancient seawater when extracting paleotemperatures from marine car-bonates. Clumped stable isotopes therefore hold great promise for resolving debates such as the temperatures of the Archean oceans (Kasting et al., 2006; Knauth, 2005). An important com-ponent to using clumped isotope thermometry is the preliminary observation that the “vital” effects that are known to fractionate 13C/12C and 18O/16O ratios during biologically catalyzed carbon-ate formation in marine environments do not apparently affect Δ

47 values (Thiagarajan et al., 2011; Tripati et al., 2010). There

seems little doubt that as the very demanding analytical issues of clumped-isotope geochemistry are tackled by more laboratories, and refi ned with improvements and new directions in instrumen-tation and experimental studies, this area of stable isotope geo-chemistry will expand.

Radiogenic Isotopes

Radiogenic isotopes may be used for geochronologic infor-mation, or as a genetic tracer. Here, we will generally cover the latter, as the subject of geochronology is covered in another chapter in this series, but it should be recognized that radiogenic isotope systems often provide both types of information. The fol-lowing discussion is grouped by isotope system, as is tradition-ally done, but we note that modern radiogenic isotope studies commonly combine multiple isotopic systems, blurring such grouping of subjects.

Rb-SrThe isotope 87Rb decays to 87Sr with a half-life (t

1/2) of 49 b.y.

Work in the 1960s established that Precambrian continental crust had signifi cantly higher 87Sr/86Sr ratios than mantle-derived rocks, refl ecting the higher time-integrated 87Rb/86Sr ratios of continental crust (Faure and Hurley, 1963). Early studies of the mantle documented Sr isotope heterogeneity across tholeiitic and alkaline basalts, but they documented a generally nonradiogenic (low 87Sr/86Sr) isotopic composition that was similar to achondrite meteorites (Engel et al., 1965). In contrast, the fi rst studies of gra-nitic batholiths showed intermediate 87Sr/86Sr ratios, suggesting they were composed of mixtures of mantle and crustal material (Hurley et al., 1965). A landmark paper by Kistler and Peterman (1973) on the Sierra Nevada batholith documented across-arc



variations in 87Sr/86Sr ratios that correlated with outcrop patterns of Precambrian rocks. Following contouring of the data, they proposed that a line of initial 87Sr/86Sr = 0.706 marked the bound-ary of the continental crust. The 1980s saw an increased focus on interaction between arc magmas and continental crust, including many studies that combined O and Sr isotopes. Taylor (1980) and DePaolo (1981b) advanced the idea of coupled assimilation and fractional crystallization, drawing upon heat balance arguments put forth decades earlier by Norman Bowen. Numerous stud-ies on the covariation of O and Sr isotopes in granitic batholiths provided constraints on the crustal lithologies that were blended in arc magmas, as well as insights into deep crustal architecture (e.g., Fleck and Criss, 1985; Solomon and Taylor, 1989).

There has long been an interest in the Sr isotope composition of seawater, and Faure et al. (1965) made one of the fi rst measure-ments of modern seawater, found that it was isotopically homo-geneous, and proposed that the isotopic composition refl ected a mixture of weathering inputs from continental and young vol-canic components. This homogeneity was confi rmed by later studies, leading to the conclusion that the residence time of Sr in the oceans far exceeds the time scales of water mass mixing. The fi rst detailed survey of marine carbonates of Phanerozoic age sketched out the broad outlines of seawater 87Sr/86Sr variations, identifying periods of radiogenic (high 87Sr/86Sr) compositions in the Carboniferous and late Cenozoic, and relatively nonradio-genic (low 87Sr/86Sr) compositions in the Cretaceous (Peterman et al., 1970). Archean-age carbonates were found to be very nonra-diogenic, which was interpreted to indicate minimal input from continental crust, possibly refl ecting small continental exposure (Veizer and Compston, 1976). Spooner (1976) made the break-through that the nonradiogenic Sr fl ux was more likely to refl ect MOR hydrothermal input rather than erosion of exposed young volcanic rocks. As mass spectrometry precision improved, both broad (Burke et al., 1982; Veizer et al., 1983) and detailed (Hess et al., 1986) Sr isotope studies of carbonates, including shells and foraminifera, provided a highly resolved Sr seawater curve. This curve has been used for both stratigraphic chronology and to infer changes due to sea-level variations, tectonic activity, weathering, and continental-scale glaciations in the Phanerozoic (DePaolo, 1986; Elderfi eld, 1986; Raymo et al., 1988).

The Rb-Sr isotope research outlined here provided a broad view of Sr isotope variations in the crust and mantle, and in the last decade, there has been an increase in work on determining 87Sr/86Sr variations in individual minerals, rather that bulk sam-ples, using in situ approaches, which have included microdrill-ing, SIMS, and LA-ICP-MS. For example, studies on feldspar phenocrysts in igneous rocks have been used to determine open- versus closed-system behavior in magmatic systems, where both magma recharge and crustal assimilation have been documented from intracrystal variations in 87Sr/86Sr ratios (Davidson et al., 2001; Knesel et al., 1999). In addition, Sr isotope profi les in phenocrysts potentially provide information on crystal residence times when coupled with diffusion modeling across discontinui-ties (Davidson et al., 2007). In addition to analysis of individual

minerals, Sr isotope measurements of melt inclusions document very large 87Sr/86Sr variations in olivine-hosted melt inclusions from oceanic basalts, which Jackson and Hart (2006) interpreted to refl ect mixing of primitive magmas that originated from enriched and depleted mantle reservoirs.

U-Th-PbThe U-Th-Pb system involves decay of three parent isotopes

to distinct Pb isotopes: 238U-206Pb (t1/2

= 4.5 b.y.), 235U-207Pb (t1/2

= 0.7 b.y.), and 232Th-208Pb (t

1/2 = 14 b.y.). Early efforts on the

U-Th-Pb isotope system focused on determining the evolution of the continental crust and mantle, given the framework that Patterson’s (1956) Pb-Pb age of Earth provided. Broad surveys of galena of various ages confi rmed that a single-stage growth curve for the crust can explain some data (Kanasewich and Far-quhar, 1965). In contrast, Patterson and Tatsumoto (1964) stud-ied detrital feldspars as a widespread measure of North American continental crust and found that the relatively high abundance of 207Pb, which can only have been produced early in Earth’s his-tory, required a two-stage growth curve that included an early U/Pb differentiation event between 3.5 and 2.5 Ga, which they interpreted to be the primary formation age for the crust. Arm-strong (1968) recognized, as did many others, that the average isotopic composition of modern Pb plotted to the right of the geo-chron (the Pb-Pb array of the solar system, inferred to include the terrestrial planets), that is, at high 206Pb/204Pb on a 206Pb/204Pb-207Pb/204Pb diagram; this was commonly referred to at the time as “anomalous” or “future” Pb, and it was recognized that this likely required a multistage history of U/Pb evolution. The scope of the Pb isotope mass-balance problem (sometimes referred to as the “Pb paradox”; Allègre, 1968) became clear in studies of basaltic rocks from the ocean basins, where most oceanic basalts lay to the right of the geochron, requiring some sort of multistage his-tory (Gast et al., 1964; Tatsumoto, 1966). An important attempt at solving the “Pb paradox” and the failure of single-stage growth curves came from the “Stacey-Kramers” average crustal growth curve, used to this day, which is a two-stage growth curve that involved a global U/Pb enrichment event at 3.7 Ga, possibly refl ecting the fi rst major differentiation event on Earth (Stacey and Kramers, 1975). The “plumbotectonics” model put forth in the late 1970s, and refi ned in following years, attempted to explain the transport of Pb between the major reservoirs of Earth, and offered a possible explanation to the Pb paradox (Doe and Zartman, 1979; Zartman and Doe, 1981). Moving forward, Pb isotope studies, as well as Sr isotopes, became increasingly inte-grated with Sm-Nd isotope studies, which provided important breakthroughs in understanding of crust-mantle evolution; these are discussed in the next section.

Recent work on in situ Pb isotope analysis by SIMS and LA-ICP-MS has been applied to subjects ranging from sedimen-tary provenance to melt inclusions. Pb isotope analyses of single detrital feldspars in both modern (Alizai et al., 2011) and ancient (Tyrrell et al., 2007) sediments have been used to infer paleodrain-age confi gurations, the extent of sediment recycling, and mineral

14 Johnson et al.


diagenesis. In situ Pb isotope analyses of feldspar phenocrysts in large caldera-related volcanic systems have shown correlations between the proportion of mantle and crustal components in phe-nocryst cargos and eruptive frequency and volume (Simon et al., 2007). Initial work on olivine-hosted melt inclusions in oceanic basalt demonstrated very large ranges in isotopic compositions that greatly exceed those measured in bulk samples, and such ranges have been inferred to refl ect blending of mantle melts from distinct mantle reservoirs, in addition to wall-rock interac-tions in the plumbing system (Saal et al., 1998). Later work has suggested less extreme ranges in the isotopic compositions of melt inclusions, but it has also highlighted the utility of in situ isotopic analysis in distinguishing between mantle and crustal processes (Paul et al., 2011).

Sm-NdArrival of the Sm-Nd isotope system on the scene in the 1970s

opened new research avenues due to the relatively restricted REE variations in a large variety of rocks, as compared to Rb/Sr and U/Pb. The most widely used decay system has been 147Sm-143Nd (t

1/2 = 106 b.y.), although the 146Sm-142Nd (t

1/2 = 0.1 b.y.) decay

system has also been used to trace early solar system processes. A powerful component of the Sm-Nd isotope system was the uni-form nature of chondrite meteorites, which refl ected generally limited Sm/Nd fractionation (Jacobsen and Wasserburg, 1980), and this led to a reference reservoir, the chondritic uniform reser-voir, or “CHUR” (DePaolo and Wasserburg, 1976), which would be used extensively in studies of terrestrial rocks for inferring the age of differentiation events. Early work documented a broad anticorrelation between 87Sr/86Sr and 143Nd/144Nd ratios in ter-restrial samples (DePaolo and Wasserburg, 1976; O’Nions et al., 1977; Richard et al., 1976), which was recognized as refl ecting the distinct time-integrated Rb/Sr and Sm/Nd ratios of “depleted” (high 143Nd/144Nd) and “enriched” (low 143Nd/144Nd) components. For oceanic lavas, this relation was initially termed “the mantle array.” In the three components of Sr, Nd, and Pb isotopes, a “mantle plane” was proposed (Zindler et al., 1982), later modifi ed in multiple component space by the landmark paper by Zindler and Hart (1986), who defi ned four principal mantle components: a depleted mantle (DMM) anchored by mid-ocean-ridge basalt (MORB), a high 238U/204Pb (μ) and high 206Pb/204Pb component (HIMU), defi ned by ocean-island basalts (OIBs) such as St. Hel-ena, a moderate 87Sr/86Sr and low 143Nd/144Nd enriched component (EM I) represented by Walvis Ridge, and a high 87Sr/86Sr and low 143Nd/144Nd enriched component (EM II), represented by Samoa. It is generally accepted that the DMM component refl ects deple-tion of the mantle through long-term production of continental crust, and the three “enriched” components (HIMU, EM I, and EM II) are generally thought to refl ect lithospheric recycling (Fig. 6). Mantle xenoliths from the subcontinental lithospheric mantle demonstrated that the DMM, EM I, and EM II components can be identifi ed in the subcontinental lithospheric mantle for some isotopic systems such as Sr and Nd (Hawkesworth et al., 1990; Menzies, 1989).

Figure 6. 143Nd/144Nd-87Sr/86Sr variations in the mantle as inferred from oceanic mafi c lavas and mantle-derived xenoliths. Box for mid-ocean-ridge basalts (MORB) defi nes a depleted mantle (DMM) com-ponent, and ocean-island basalts (OIB) extend from MORB to lower 143Nd/144Nd and higher 87Sr/86Sr ratios toward two enriched mantle (EM) components: EM I (low 143Nd/144Nd, moderate 87Sr/86Sr) and EM II (moderate 143Nd/144Nd and high 87Sr/86Sr). A high U/Pb compo-nent (HIMU) is subtly defi ned for Nd-Sr isotopes, but it is prominent for Pb isotopes. These mantle components were discussed in detail in Zindler and Hart (1986) and Hart (1988). Nd-Sr isotope varia-tions for xenoliths derived from the subcontinental mantle greatly extend the range observed for oceanic basalts, refl ecting isolation from the asthenosphere, but the data are consistent with the DMM, HIMU, EMI, and EM II components (Hawkesworth et al., 1990). Subduction-related volcanic arcs may deviate from the OIB fi eld, sometimes toward relatively high 87Sr/86Sr ratios, possibly refl ecting a component from hydrothermally altered oceanic crust (McCulloch et al., 1981). Subduction-related volcanic rocks may also extend to very low 143Nd/144Nd and high 87Sr/86Sr ratios (data shown for Martinique; Davidson, 1983). Diagram is adapted from those in Dickin (1995).

Covariation of Sr and Nd isotopes in orogenic arcs showed that many arc lavas were shifted to high 87Sr/86Sr ratios relative to the Sr-Nd mantle array (Hawkesworth et al., 1979). Coupled with new Sr and Nd isotope data from ophiolites that showed strong shifts in 87Sr/86Sr ratios but invariant 143Nd/144Nd during hydrothermal alteration at MORs (McCulloch et al., 1981), preferential shifts in 87Sr/86Sr in orogenic arcs were interpreted to refl ect a subduction component derived from the altered slab. A landmark study by Hildreth and Moorbath (1988) proposed that mixing, assimilation, storage, and homogenization of arc magmas extensively occurred in the lower crust during pond-ing of mantle-derived basaltic magmas near the Moho, a model known as “MASH,” and one that still provides a framework for arc magmatism. The relative coherence of Sm/Nd ratios dur-ing melting provided a means of “seeing through” recent mag-matic events to infer the time when Nd was extracted from the mantle, an approach that led to the concept of “Nd model ages” (DePaolo, 1981a). In western North America, Nd model age



provinces correlated with Archean, Proterozoic, and Phanero-zoic crustal boundaries and allowed identifi cation of older Nd components in granitic batholiths (Bennett and DePaolo, 1987; Farmer and DePaolo, 1983). Neodymium isotope data from large caldera complexes from North America, however, identi-fi ed large proportions of mantle, indicating that such complexes represented new periods of net crustal growth (Johnson, 1991), suggesting that ignimbrites may contain a larger proportion of mantle than granitic rocks, and hence periods of net crustal growth (Johnson, 1993).

The Nd model age concept has been extensively applied to fi ne-grained sedimentary rocks, and for samples that had deposi-tional ages younger than 2 Ga, Nd model ages were usually older than 2 Ga, indicating that most Phanerozoic- and Proterozoic-age sedimentary rocks have been recycled from older sources (Allègre and Rousseau, 1984; Goldstein et al., 1984; O’Nions et al., 1983). An exception is sedimentary rocks that were depos-ited at the same time as orogenic episodes, which were found to contain a larger proportion of mantle Nd than nonorogenic sedi-ments, refl ecting net crustal growth (Michard et al., 1985). Mass-age distributions for the continents calculated from Nd crustal residence times using sedimentary rocks indicated that ~40% of the present-day continental mass formed by 3.8 Ga (Jacobsen, 1988), a conclusion that is similar to that inferred from Pb isotope compositions of detrital feldspars, as discussed earlier (Patterson and Tatsumoto, 1964). A critical component to extracting crustal growth curves from Nd model ages, however, is quantifying the extent of crustal recycling in sediments (Allègre and Rousseau, 1984); as will be seen later herein, signifi cantly different crustal growth curves have been inferred using Hf isotope data.

In contrast to Sr, the residence time of Nd in seawater is very short, on the order of 102 as, and early researchers in Sm-Nd iso-topes recognized that this should make the isotopic composition of Nd in seawater a sensitive indictor of local input (Piepgras and Wasserburg, 1980; Piepgras et al., 1979). This early work, which confi rmed the isotopic provinciality of Nd in seawater, and hence its use as a sensitive tracer of water masses, was later extended back in time through analysis of Fe-Mn crusts and authigenic marine sediments. Examples include tracing the distinct evolu-tion of the Pacifi c-Panthalassa and Iapetus Oceans from the Neo-proterozoic to present (Keto and Jacobsen, 1988), and closure of the Central American isthmus that restricted water mass com-munication between the Pacifi c and North Atlantic Oceans (Bur-ton et al., 1997). In addition to Nd, Pb isotope studies of Fe-Mn sediments demonstrated that Pb isotopes were also provincial, commensurate with the short residence time for Pb in seawater (Abouchami and Goldstein, 1995).

Advances in mass spectrometry in the 1990s allowed appli-cation of the 146Sm-142Nd isotope system, which requires mea-surement of 142Nd/144Nd ratios to a very high precision of several parts per million. The short half-life of the 146Sm-142Nd system provides a sensitive tracer of processes in the fi rst few hundred million years of solar system history. Harper and Jacobsen (1992) demonstrated an average 142Nd/144Nd enrichment of 32 ppm for

rocks from the 3.8 Ga Isua supracrustal belt, Greenland, relative to average terrestrial rocks. Consideration of coupled 142Nd/144Nd and 143Nd/144Nd variations suggested that the 142Nd enrichment recorded Sm/Nd fractionation in the source reservoir(s) of the Isua rocks likely occurred between 4.55 and 4.45 Ga. These results were confi rmed by Bennett et al. (2007), who interpreted the 142Nd enrichment to record early mantle differentiation in the fi rst 30–75 m.y. of Earth’s history. Boyet and Carlson (2005, 2006) com-pared high-precision 142Nd/144Nd data from meteorites and a wide variety of terrestrial mantle–derived rocks and documented that chondrite meteorites have 142Nd depletions on the order of 20 ppm relative to the average of terrestrial rocks, and they suggested that this records a major mantle differentiation event ~30 m.y. after Earth formation. Moreover, Boyet and Carlson proposed that there is a missing low-142Nd/144Nd reservoir in Earth, possibly located at the base of the mantle, assuming that Earth formed with a chondritic Sm/Nd ratio. Carlson and Boyet (2008) suggested that the missing reservoir, which should have enriched incompat-ible trace-element contents, refl ects high-density components that settled into the deep mantle from an early terrestrial magma ocean. Evidence for an incompatible element–enriched mafi c crust very early in Earth history lies in rocks from the Nuvvuagittuq green-stone belt, Canada, which has relatively low 142Nd/144Nd ratios that overlap those of chondrites, and which has a 146Sm-142Nd “age” of ca. 4.3 Ga (O’Neil et al., 2008). A contrasting view of the “miss-ing 142Nd reservoir” is offered by Jacobsen et al. (2008), who argued that the difference in 142Nd/144Nd ratios between terrestrial rocks and chondrite meteorites may be explained by isotopic het-erogeneity during Earth’s accretion, rather than invoking a miss-ing reservoir in Earth’s mantle.

Lu-HfThe isotope 176Lu decays to 176Hf with a half-life of 37 b.y. The

large Lu/Hf fractionations that are produced by garnet, as well as the high abundance of Hf in zircon, have long spurred interest in this analytically diffi cult isotope system. A series of pioneering papers in the 1980s outlined the range in Hf isotope composi-tions of meteorites, mantle-derived rocks, crustal zircons, and sedimentary rocks (Patchett et al., 1981; Patchett and Tatsumoto, 1980; Patchett et al., 1984). Hafnium and Nd isotope composi-tions for OIB are broadly correlated, but Hf and Nd isotopes for MORB are not correlated. Early work on zircons from igneous and metamorphic rocks as old as 3.7 Ga suggested a chondritic Hf evolution for the crust until ca. 2.8 Ga, at which time high initial 176Hf/177Hf ratios appeared, indicating the presence of a globally differentiated mantle (Patchett et al., 1981). Later work on detrital zircons indicated the presence of a depleted mantle by 3.0 Ga (Stevenson and Patchett, 1990). Detailed studies of the Hf isotope compositions for the oceanic mantle identifi ed HIMU, EM I, and EM II components defi ned by Sr-Nd-Pb isotope stud-ies (Salters and Hart, 1991), but Hf isotopes require two DMM components, one of which has high 176Hf/177Hf ratios that requires garnet as an important ancient component in the source regions of MORB (Salters, 1996).

16 Johnson et al.


The late 1990s saw a marked increase in Lu-Hf research as MC-ICP-MS instruments became widespread. High-precision Hf isotope measurements of juvenile crystalline rocks, as well as ancient sediments, initially documented a limited range in initial 176Hf/177Hf ratios prior to 3.0 Ga, which stood in contrast with Sm-Nd isotope data that suggested very large ranges in initial 143Nd/144Nd early in Earth’s history (Vervoort and Blichert-Toft, 1999; Vervoort et al., 1999), although later work, based on detri-tal zircons (see following), has since expanded the Hf isotope database for Precambrian rocks by several orders of magnitude and documented a much larger range in initial 176Hf/177Hf ratios. High-quality MC-ICP-MS data for OIBs identifi ed subtle varia-tions in Hf-Nd isotope arrays that provided strong support for an ancient subduction of pelagic sediment that was previously hypothesized but diffi cult to confi rm using TIMS data (Blichert-Toft et al., 1999).

A major effort in combined U/Pb geochronology and Hf iso-tope studies of detrital zircons, using in situ analysis methods, began in the early 2000s and has continued to the present. Much of this work has been aimed at addressing the growth and evolu-tion of continental crust, given the wide variety of crustal growth histories that have been proposed (e.g., Allègre and Rousseau, 1984; Armstrong, 1981; Condie, 1998; Hurley and Rand, 1969; Taylor and McLennan, 1985). U/Pb zircon geochronology has long highlighted major peaks in crystallization ages at ca. 2.7 and ca. 1.8 Ga, as well as other peaks in more detail (e.g., Condie, 1998), leading a number of workers to conclude that crustal growth has been episodic (e.g., Rino et al., 2004). This in turn has led to models where catastrophic events in the mantle have been invoked to explain punctuated periods of very rapid crustal growth. In contrast, the distributions of Hf isotope model ages for detri-tal zircons do not show the strong peaks recorded in U/Pb ages, which suggests that the U/Pb age record may refl ect biases in preservation rather than periods of episodic crustal growth (e.g., Belousova et al., 2010; Hawkesworth et al., 2009, 2010; Voice et al., 2011). Recognizing that Hf isotope model ages in detri-tal zircons may refl ect mixtures of mantle-derived and recycled components, Kemp et al. (2006) combined U/Pb geochronology and O and Hf isotopes to separate a recycled high-d18O sedimen-tary component from juvenile, mantle-derived components that refl ect net addition to the continents. This approach has forced a major change in thinking about crustal growth rates based on detrital zircons, and it suggests that as much as 70% of current continental crustal volume was produced by 3 Ga, followed by slower, but relatively uniform, rates of crustal growth since that time (Dhuime et al., 2012).

Looking at the oldest detrital zircons, the Archean and Hadean zircons from the Jack Hills, Australia, have been a logi-cal target of Lu-Hf studies. Early work on mineral separates by Amelin et al. (2000, 1999) highlighted the diffi culty in inter-preting both Lu-Hf and U-Pb data from complex zircons, and they found little evidence for enriched or depleted components at 3.8 Ga, suggesting minimal differentiation on Earth by this time. Later work by Harrison et al. (2005) reported remarkably

positive εHf

values up to +13 at 4.3 Ga, obtained by in situ meth-ods using LA-MC-ICP-MS, and they argued that an active plate-tectonic system existed at this time. Valley et al. (2006) noted that signifi cant errors could be introduced in calculating initial ε

Hf values given the complexity of the zircons and the disparate

volumes involved in U-Pb (SIMS) and Lu-Hf (LA-MC-ICP-MS) measurements. Harrison et al. (2008) combined Pb and Lu-Hf analysis to constrain 207Pb/206Pb ages to the same ablated volume analyzed for Lu-Hf, and this revised approach found only nega-tive ε

Hf values down to −5 at 4.2 Ga. Recently, Kemp et al. (2010)

reported LA-MC-ICP-MS Pb and Hf isotope analyses that pro-duced negative ε

Hf values between 3.9 and 4.3 Ga in age (Fig.

7), which they argued refl ects only small domains of enriched (low Lu/Hf) components in primitive crust, and not widespread melting and depletion of the mantle, nor an active plate-tectonic system prior to 4 Ga.

Re-Pt-OsRadiogenic Os isotope studies have generally focused on the

187Re-187Os system (t1/2

= 42 b.y.), although the 190Pt-186Os system (t

1/2 = 489 b.y.) has also been explored. The Re-Pt-Os isotope sys-

tem differs greatly from the radiogenic systems discussed previ-ously in that these elements are siderophile/chalcophile (Shirey

Figure 7. εHf

(t) vs. age relations for zircons from the Jack Hills, Austra-lia, the oldest known terrestrial samples. The ε

Hf(t) value was calculat-

ed from initial 176Hf/177Hf ratios and 207Pb/206Pb ages (U-Pb zircon geo-chronology). Chondrite uniform reservoir (CHUR) reference is shown, as well as the fi eld for depleted mantle inferred from mid-ocean-ridge basalts (MORB-DM), which spans a range in present-day ε

Hf values

from CHUR to highly positive values (Salters and Hart, 1991). The limiting line for an enriched reservoir, such as continental crust, is shown at Lu/Hf = 0. Some previous studies reported a limited number of positive ε

Hf(t) values (Blichert-Toft and Albarède, 2008; Harrison

et al., 2005), which would indicate widespread mantle depletion, but other studies found only negative ε

Hf(t) values (Amelin et al., 1999;

Harrison et al., 2008; Kemp et al., 2010); such values indicate the pres-ence of an enriched component, possibly tied to crust formation, but this does not provide conclusive proof of widespread mantle depletion and continental crust formation.



and Walker, 1998), and the major inventories exist in Earth’s core, distantly followed by the mantle (see previous discussion). Work on this analytically challenging system, which began in the 1980s, showed that the mantle has remained relatively nonradio-genic, lying within the range of chondritic meteorites, but the very high Re/Os ratios of continental crust produced very high 187Os/186Os ratios over time (Allègre and Luck, 1980; Luck and Allègre, 1983). Correlations between Os and O isotopes in OIBs have been interpreted to refl ect ancient subducted sediments (Lassiter and Hauri, 1998), or assimilation of hydrothermally altered oceanic crust during ascent of magmas to the surface (Gaffney et al., 2005). In contrast, Escrig et al. (2005) interpreted high 187Os/187Os ratios at Fogo Island to record assimilation of lower continental crust during opening of the Atlantic Ocean.

The large Re/Os fractionations that are produced dur-ing melting of peridotite have been used to trace the evolution of the lithospheric mantle. Walker et al. (1989) fi rst proposed this approach for cratonic peridotite xenoliths, where they cal-culated Os model ages that refl ect Re depletion “events” that may record the time of stabilization of the lithospheric mantle. Compilations of Os isotope data from both whole-rock samples and sulfi des from cratonic peridotite xenoliths commonly show very nonradiogenic 187Os/188Os ratios that may be interpreted as Re depletion ages between ca. 2.5 and 3.0 Ga (Carlson et al., 2005; Pearson and Wittig, 2008). Although Os model ages pro-vide insight into the evolution of the lithospheric mantle not available by other radiogenic isotope systems, complexities such as sulfi de breakdown, metasomatism, and mantle-melt interac-tions can cloud interpretations (Rudnick and Walker, 2009). In contrast to the nonradiogenic Os isotope compositions that have been used to infer Re depletion histories, Shirey and Richardson (2011) recently reported high 187Os/188Os ratios in sulfi de inclu-sions in diamonds of 3 Ga age or younger, which they interpreted to refl ect eclogites that originated as subducted oceanic crust that became incorporated in the subcontinental lithospheric mantle.

Radiogenic (high) 187Os/188Os ratios in orogenic arc lavas have been interpreted in two ways. Some workers have inferred such compositions to refl ect subducted sediments (Alves et al., 1999, 2002), although other workers have argued that the low Os content of crustal material is unlikely to shift the Os isotope com-positions of the subarc mantle, instead interpreting the radiogenic Os isotope compositions as refl ecting assimilation of young, mafi c lower crust (Chesley and Ruiz, 1998; Hart et al., 2003; Jicha et al., 2009). The importance of the second interpretation lies in the potential ability of the Re-Os isotope system to trace interaction with arc crust, which is essentially invisible using other isotopic systems such as O, Sr, Nd, Hf, and Pb isotopes, and which has implications for estimating net crustal growth in orogenic arcs.

Turning to the very long-lived 190Pt-186Os system, Walker et al. (1995) fi rst proposed that core material might be identifi ed in OIBs, often suggested to refl ect mantle plumes that originate at the core-mantle boundary (e.g., Hawkesworth and Schersten, 2007). Brandon et al. (1999, 2000) observed correlated enrich-

ments in 186Os/188Os and 187Os/188Os in modern OIBs that they interpreted to represent a high-190Pt component from the core, and later work on Archean komatiites identifi ed similar fea-tures (Puchtel et al., 2005). Alternative proposals for correlated 187Os/188Os and 186Os/188Os in OIBs include Pt- and Re-rich mate-rials that could refl ect ancient crustal components, or ancient pyroxenite in the mantle, although Brandon and Walker (2005) suggested that such materials are unlikely to be common compo-nents in OIBs. Recently, Luguet et al. (2008) documented 186Os enrichments in pyroxenites, eclogites, sulfi des, and Pt-rich alloys in peridotites, providing support for the argument that 186Os may not be a unique tracer of core components in OIBs.

ORGANIC GEOCHEMISTRY: INVESTIGATING CARBON COMPOUNDS PRESERVED IN ROCKS

The past 50 yr have witnessed exponential growth in our appreciation of the ubiquity and diversity of organic compounds that are preserved in sediments and sedimentary rocks. A number of these substances, or their derivatives formed through diage-netic alteration processes, were discovered in sediments before they were recognized as natural products in living organisms (Ourisson and Albrecht, 1992; Ourisson et al., 1979). Techno-logical innovation combined with the quest for an understanding of fossil fuel composition, formation, and occurrence were the initial drivers of research on sedimentary organic matter in the 1960s and 1970s (Kvenvolden, 2006). As the possibility of gain-ing evolutionary insights by way of “molecular paleobiology” became more clear, the driving force for organic geochemistry shifted (Eglinton, 1970; Peterson et al., 2007). In more recent times, it has been recognized that organic molecules in the air, in water, and in sediments carry in their chemical structures and stable isotopic compositions a myriad of environmental and cli-mate signals. Investigating these records on both human and geo-logic time scales has now become the prime focus of research in organic geochemistry (Eglinton and Eglinton, 2008; Gaines et al., 2009).

The Quest to Understand the Origins of Petroleum and Other Fossil Fuels

The invention of computerized mass spectrometers, in com-bination with high-resolution capillary gas chromatography, was a critical early technical development, and this led to the discovery that petroleum was composed of thousands of dis-crete organic compounds. In the 1960s, geochemists working in the fossil fuel industry, in collaboration with academic scien-tists, also demonstrated that sedimentary kerogen and bitumen, the precursors of petroleum, could be correlated to commercial deposits of oil and gas using combinations of organic compound distributions and isotopic patterns (Hunt et al., 2002; Peters et al., 2005). They also convincingly demonstrated that buried organic matter was predominantly of biological origin (e.g., Hills and Whitehead, 1966), and that progressive changes in its chemical

18 Johnson et al.


composition correlated with burial depth and geothermal heating rates, thereby providing a tool with which to assess the nature and distribution of coal, petroleum source rocks, and natural gas (Seifert and Moldowan, 1980, 1978). Transport of organic matter in clays opened a new mechanistic understanding to its preser-vation and accumulation in source rocks, and a way in which organic carbon burial could be understood in the context of the principles of sequence stratigraphy (Creaney and Passe, 1993; Hedges and Keil, 1995).

Diagenetic processes, mediated by microbes, water, and reduced sulfur compounds, render complex biomolecules into both simpler hydrocarbons and much more complex material (kerogen) through polymerization and cross-linking (Adam et al., 1993; Kohnen et al., 1989). Integration of geochemical understanding with sedimentary geology and basin modeling ultimately resulted in a more informed understanding of hydro-carbon systems, including resource and risk assessment (Hunt, 1996; Tissot and Welte, 1978).

Development of Molecular Paleontology

Paleontologists have long recognized that microbial life has dominated life on Earth for most of its history, yet microscopic fossils comprise an infi nitesimally small fraction of the biomass that has been present on Earth, and the vast majority of microbes leave no visible trace at all. It was recognized early, however, that microbial lipids derived from their cell walls, membranes, and pigments are preserved, and, in fact, form the dominant organic molecules that can be found preserved in sedimentary rocks. In Figure 8, we summarize some of the important organic mole-cules used in geochemistry, organized according to contemporary molecular phylogeny. The concept of “biomarker chemostratig-raphy” has emerged as a means for using these molecules to chart evolutionary innovation, mass extinctions and radiations, chemi-cal events in the ocean, and climate change (Gaines et al., 2009). For example, the rise of complex multicellular life in the Neopro-terozoic, the transitions in ocean plankton through the Phanero-zoic, and the advent and radiation of land plants are among a few of the biological innovations that have been, and continue to be studied using biomarkers.

After several decades of study, it is now clear that major paleontological mileposts are accompanied by corresponding biosynthetic innovations that are recorded in molecular fos-sils. Colonization of the continents by plants, for example, was enabled by invention of a range of structural biopolymers, includ-ing lignin and cellulose. Plants built defenses against desiccation and predation utilizing cuticles made of waxy hydrocarbons for the former, and an array of sesqui- (C

15), di- (C

20), and tri-

terpenoids (C30

) and resinitic compounds for signaling and for warding off insect predators. Distinct differences in the chemi-cal defenses utilized by conifers (Gymnosperms) and fl owering plants (Angiosperms) can be recognized through the presence of distinctive diterpenoids in Paleozoic and early Mesozoic rocks and petroleums (e.g., Noble et al., 1985). An increase in the prev-

alence of triterpenoids is now known to be associated with the increased abundance of fl owering plants in the late Mesozoic and Cenozoic (Moldowan et al., 1994; Murray et al., 1998; Simoneit et al., 1986). Recently, it has become clear that biogeographical aspects of plant occurrence and dispersal are refl ected in sedi-mentary hydrocarbons (Dutta et al., 2011). In the marine realm, micropaleontology informs us that the Paleozoic dominance of green and red clades of planktonic algae gave way to the modern dominance of chlorophyll a + c plankton, an observation that is directly correlated with the prevalence of their respective steroi-dal and acyclic isoprenoidal hydrocarbons in marine sediments and oils (e.g., Knoll et al., 2007; Rampen et al., 2007; Sinninghe Damsté et al., 2004; Volkman, 2003). A practical application of the evolution of complex organisms and changes in the composi-tion of ocean plankton through time has been development of age-diagnostic biomarkers that can place constraints on the tim-ing of deposition of petroleum source rocks (e.g., Grantham and Wakefi eld, 1988; Holba et al., 1998).

As with studies of algae, sensitive GS-MS (gas chromatography–mass spectrometry) methods developed in the last few decades have provided the means with which to study biomolecules that are unique to animal phyla. Cholesterol is the predominant membrane sterol of animals, and there is little else in the way of other preservable molecules that could constitute a robust “metazoan” biomarker. Basal invertebrate phyla, defi ned as those least removed from their single-celled ancestors, include sponges, cnidarians, and echinoderms, and these are the excep-tions in that they have a wealth of complex biochemistries, possi-bly refl ecting defenses against predation. Demosponges are well known for their biosynthetic capacity to produce distinctive ter-penoids, including sterols that have a rare 24- isopropycholestane skeleton (Bergquist et al., 1991). 24-Isopropylcholesterols have been found in the calcisponges, hexactinellid sponges, or choano-fl agellates, and these form a unicellular sister group to sponges (Love et al., 2009). Hydrocarbon derivatives of the distinctive C

30

sponge sterols are prevalent in sedimentary rocks and petroleums of Cryogenian (Neoproterozoic) to Early Cambrian age, and they are thought to refl ect an acme in the abundance of demosponges in sedimentary environments of this time (Love et al., 2009; McCaffrey et al., 1994). The molecular fossil record suggests that demosponges made their fi rst appearance in the Cryogenian, and this is concordant with molecular clock estimates of metazoan divergence times (Peterson et al., 2007).

It is now recognized that biomarker hydrocarbons can be diagnostic for certain kinds of paleoenvironmental circum-stances (Fig. 9). Water-column redox stratifi cation, elevated salinity, marine versus nonmarine sedimentation, and clas-tic versus carbonate environments are some of the conditions that can be encoded into fossil hydrocarbon distribution pat-terns. A prime example lies in the insights provided by organic geochemical proxies into biogeochemical processes occurring during oceanic anoxic events (OAEs), and especially those associated with mass extinction events. The work of Holser and others (Holser, 1977; Holser et al., 1989) fi rst recognized



shifts in the isotopic compositions of C, O, and S in marine sediments, interpreted to record the disruption of ocean chem-istry and consequent biological mass extinction. Biogeochem-ists have subsequently observed concomitant signals in organic molecules, redox-sensitive trace elements, and other geochemi-cal proxies. One recurring trend is seen in the prevalence of biomarkers derived from the green sulfur bacteria (Chlorobi) in sediments that were deposited during OAEs. Chlorobi are

anoxygenic phototrophs that use sulfi de as an electron donor for photosynthesis, and these produce distinctive light-harvesting carotenoid and chlorophyll pigments that may be recorded in preservable structural features (Grice et al., 1996; Sinninghe Damsté and Koopmans, 1997; Summons and Powell, 1987). Diagenetic reduction and stabilization of these compounds are enhanced under the strongly reducing (euxinic) conditions that are favored by the green sulfur bacteria.

Figure 8. A “tree of life,” based loosely on the nucleotide sequences of small subunit ribosomal ribonucleic acid (RNA), to illustrate life’s three domains and the diverse structures of lipids that are characteristic of organisms at the domain level and, in some cases, phylum level. For example, most eukaryotes utilize sterols, and, in a few cases such as diatoms, dinofl agellates, and demosponges, there are particular sterols that are very specifi c. The membrane lipids of bacteria and eukaryotes incorporate hydrocarbon chains that are linear or branched in relatively simple ways, and these are mostly linked to glycerol via ester linkages. Archaea, on the other hand, build equivalent structures using isoprenoidal chains, and they are linked to glycerol with ether bonds. Many of these features are preservable and can be observed in organic matter extracted from sedimentary rocks. Improved understanding of the way in which lipid chemistry relates to physiology, phylogeny, and environment, and deduction of the pathways by which organic molecules become preserved in sediments have been prime accomplishments of organic geochemists in the past 50 yr. GDGT—glycerol dibiphytanyl glycerol tetraethers.

20 Johnson et al.


Geologists and paleontologists have long debated the origin of the Permian-Triassic extinction, both the cause and the bio-logical response, and work in molecular paleontology in the last decade has provided important insights. Studies of a core drilled into the end-Permian type section at Meishan, China, show that Chlorobi carotenoids are particularly abundant throughout the last few million years of the Permian and into the Early Triassic, implying that shallow-water euxinic conditions were protracted at the type section of the greatest mass extinction of the geologi-cal record (Cao et al., 2009; Grice et al., 2005). Examination of other Permian-Triassic transition sections confi rms the pattern and suggests that euxinic conditions were pervasive globally, as far as can be discerned from continental margin sediments of the Tethys, Panthalassic, and Boreal Oceans (Hays et al., 2007, 2012). These observations are consistent with other isotopic and inorganic proxies for a global and intense oceanic anoxic event during the Permian-Triassic transition. Independent modeling also suggests that toxic hydrogen sulfi de would have been upwell-ing from the deep ocean onto continental shelves and entering the atmosphere (Kump et al., 2005), implying that sulfi de toxicity

was a contributing factor to the biological mass extinction in the marine realm as well as on land. Co-occurring with the evidence for euxinia, molecular fossils shed light on other aspects of this event, including transitions in ocean plankton (Cao et al., 2009), and a massive terrestrial weathering event (French et al., 2012; Sephton et al., 2005; Xie et al., 2007).

Although many theories abound, no consensus has been reached on the processes that instigated the end-Permian mass extinction (Erwin, 2006). Evidence that the event had its roots in ocean chemistry is consistent: Anoxia with accompanying effects of hypercapnia, ocean acidifi cation, and sulfi de toxicity could all have contributed to the loss of marine life, and, arguably, to loss of life on land (Holser, 1977; Knoll et al., 1996; Twitchett et al., 2001; Wignall and Hallam, 1992). Numerous researchers also point to volatile release during emplacement of the Siberian Traps lavas and intrusions (Payne and Clapham, 2012). These authors argued that gases liberated during the volcanism itself, augmented by additional C and S volatiles released from affected sediments, can account for most of the end-Permian paleonto-logical and geochemical observations, and this is consistent with

Figure 9. An example of the way in which marine water-column redox structure can be deduced from the preserved remains of photosynthetic pigments. The carotenoid β-carotene is an accessory pigment common to oxygen-producing phototrophs such as algae and bacteria and is, there-fore, produced in, and diagnostic for, oxygenated environments. In contrast, isorenieratene is produced by brown strains of the Chlorobi, which are obligately anaerobic photoautotrophic bacteria and which thrive in low-light, sulfi dic environments, typically where the oxycline is as deep as 60 m. Chlorobactene is produced by the “green” forms of the Chlorobi, which require higher light intensities and are generally found where the oxycline shoals to 20–40 m depth. Okenone is characteristic of the purple sulfur bacteria (Chromatiaceae), which require sulfi de and even higher light intensities. This pigment can be diagnostic for highly restricted environments such as hypersaline lagoons and where oxyclines are at 20 m or less. All these pigments can also be found in redox-structured microbial mats. It is probably no coincidence that the hydrocarbon derivatives of these pigments abound in sediments and petroleum from the Mesozoic and other warm intervals of Earth’s history and particularly when narrow seaways and lack of ice resulted in sluggish circulation and ready maintenance of deep-ocean anoxia.



the timing and tempo indicated by the most recent geochrono-logic studies (Kamo et al., 2003; Mundil et al., 2010; Shen et al., 2011). Despite continued debate on the ultimate trigger, most authors agree that the geochemical and geologic evidence for a bolide impact (Becker et al., 2001) is unclear (Farley and Muk-hopadhyay, 2001). In this context, it is also noteworthy that the organic geochemical signals that accompany the Cretaceous-Paleogene impact and mass extinction event bear no resemblance to those seen at OAEs (Sepulveda et al., 2009).

There are other examples where biomarker research in the last 20 yr has helped us to understand oceanic anoxia. The Chlo-robi carotenoid signal is one of several organic geochemical fea-tures that are consistently observed during intervals of enhanced black shale deposition during the Paleozoic and Mesozoic Eras (Koopmans et al., 1996; Pancost et al., 2004). Isorenieratane, chlorobactane, and the aryl isoprenoids derived from them, along with C, S, and N isotope and other biomarker anomalies that are indicative of anoxia, euxinia, and water-column stratifi cation, comprise a recurring signal that is seen in both clastic and car-bonate sediments deposited during Mesozoic OAEs (Kuypers et al., 2004; Schouten et al., 2000b). Many of these sediments, such as the Kimmeridge Clay Formation, are also prolifi c source rocks for petroleum deposits that carry the same signal (Van Kaam-Peters et al., 1998a, 1998b). The prevalence of organic-rich and petroleum-prone sediments that were deposited during the Mesozoic Era comprises the bulk of the world’s oil-in-place inventory, refl ecting pervasive development of narrow rift basins and seaways, greenhouse conditions, and sluggish ocean circula-tion during breakup of the Pangean supercontinent (Klemme and Ulmishek, 1991).

Studies of molecular fossils in Precambrian sedimentary rocks have been pursued in earnest since the 1980s, and this work has shown that such rocks exhibit an array of novel molecular and isotopic features that are not seen in younger sedimentary sequences (Grantham and Wakefi eld, 1988; Grosjean et al., 2009; Klomp, 1986; Summons et al., 1988). Unprecedented distribu-tions of steroids, for example, can be traced to the proliferation of green and red algae, as well as sponges, and this accords with other geochemical evidence for the Neoproterozoic oxygenation of the atmosphere and ocean (Canfi eld et al., 2007; Fike et al., 2006) and paleontological evidence of the proliferation of com-plex life (Erwin et al., 2011; Knoll et al., 2004). Euxinic condi-tions have been inferred from geochemical studies of Mesopro-terozoic strata, and this is refl ected in carotenoid molecular fossil distributions in the same rocks (Brocks et al., 2005). Steroid and triterpenoid hydrocarbons have also been recovered from much older rocks (Dutkiewicz et al., 2006; George et al., 2008; Wald-bauer et al., 2009). Some of this work has been challenged on the basis of C isotope data obtained by nano-SIMS (Rasmussen et al., 2008), suggesting that the recovered biomarkers instead refl ect contamination. In contrast, studies of fresh drill cores recovered from the Kaapvaal craton that were targeted to address this issue (Waldbauer et al., 2009) suggest that complex terpenoids, and particularly the steroids, could be indigenous to the samples

and represent compounds that were biosynthesized prior to the Great Oxidation Event (GOE) at ca. 2.3–2.4 Ga. The presence of steroids in older than 2.5 Ga rocks supports a raft of isotopic and trace-element proxies, some of which have been discussed already, which indicate that trace amounts of oxygen were being produced (e.g., Anbar et al., 2007; Czaja et al., 2012) and respired (Eigenbrode and Freeman, 2006) in the surface oceans well in advance of the GOE. Further drilling, using ultraclean drilling protocols and contamination tracers, is currently under way in the Pilbara craton, Australia, and this is expected to shed further light on ocean-atmosphere redox conditions in the Neoarchean.

Elucidating Paleoenvironmental Records

One of the most important discoveries in organic geochem-istry in the last 50 yr has been the recognition of some classes of organic molecules that encode signals for past climate regimes. Of key importance are the proxies for sea-surface temperature (SST) based on the degree of unsaturation in long-chain ketones (C

37 alkenones) from marine algae (the Uk37 index; Brassell

et al., 1986), and the proportions of isoprenoidal ether lipids (dubbed glycerol dibiphytanyl glycerol tetraethers or GDGT) derived from archaea (the Tex86 index; Schouten et al., 2002). In soils and paleosols, bacterial (non-isoprenoidal) ether lipids have been proposed to provide information on both temperature and pH (Weijers et al., 2006). Signifi cant effort has been expended in elucidating the specifi c organisms responsible for these proxies, as well as the physiological basis of the sedimentary temperature records that are encoded by alkenones and ether lipids. In the case of the former, prominent marine haptophyte algae such as Isochrysis galbana, Emiliania huxleyi, and Gephyrocapsa oce-anica, living primarily in the surface mixed layer and sometimes extending to the thermocline (Ohkouchi et al., 1999), are widely recognized as the primary biological sources in the oceans (Volk-man et al., 1995, 1980). Given the biogeographical trends and temperature preferences of alkenone-producing algae, it appears likely that the alkenones record lipids that were derived from a combination of cold- and warm-water adapted strains. Nutrient status, salinity, growth rate, growth stage, and temperature all infl uence the distributions of alkenones in cultured haptophytes, yet empirical calibrations of the Uk37 indices to core tops, and to cultures, appear to be very robust over a wide range of tem-peratures, environmental variables, and spatial scales (Prahl et al., 2003; Rosell-Melé et al., 1995; Sikes and Volkman, 1993). The C

37 alkenones occur in sediments that predate the appearance

of contemporary alkenone-producing haptophytes, and, because of their apparently close evolutionary relations and similarity to fossil forms, this has potentially opened application of the proxy to ancient sedimentary records on multi-million-year time scales. Paleotemperature reconstructions into deep time, however, will have to be viewed cautiously until more is known about the phys-iological function of C

37 alkenones in haptophyte algae. What is

most impressive about the alkenone SST records is the degree of concordance with independent isotopic records, including

22 Johnson et al.


O isotope variations recorded in ice cores and foraminifera (e.g., Herbert et al., 2001). It has become clear that multiple approaches are needed to understand and constrain each proxy in terms of the robustness of paleotemperature records and the effects of lateral sediment redistribution (Mollenhauer et al., 2011).

In contrast to the alkenones and their sedimentary records as expressed in Uk37 data, the temperature proxy based on archaeal ether lipids (Tex86) is less well constrained with respect to the precise sources of the isoprenoidal GDGTs and the controls on their production. The Tex86 proxy emerged from the discovery of abundant archaeal plankton communities that populate oceans, lakes, and other aquatic environments, in addition to recognition that membrane-spanning GDGTs contain variable numbers of cyclopentane rings that could be isolated from particulate organic matter and the underlying sediments (Delong et al., 1998). A particular problem, however, is that diverse Crenarchaeota and Euryarchaeota that inhabit both the water column and the sedi-ments produce some of the same compounds that comprise the proxy. In addition, other studies show that the prevalence, and production, of GDGT often takes place at great depth (Ingalls et al., 2006; Sinninghe Damsté et al., 2002), frequently near or within oxygen minimum zones. Nevertheless, evidence suggests that the distributions of GDGT in core-top sediments are similar to those of (epipelagic) marine Crenarchaeota living in the top 100 m of the water column, an observation consistent with the relatively robust core-top calibrations of SST with Tex86 (Wuch-ter et al., 2005). Numerous studies point to an origin for archaeal ether lipids in ammonia-oxidizing Marine Group 1 Crenarchaeota (Schouten et al., 2008). Evidence has been presented to show that Tex86 values agree with foraminifer proxies that indicate rapid ocean warming during the late Paleocene thermal maximum (Zachos et al., 2006). Paleotemperature records extending into Mesozoic time are considered possible (Schouten et al., 2003), but there are few avenues available for independent verifi cation. An important consequence that fl ows from the discovery of iso-prenoidal GDGT has been the realization that the Crenarchaeota are exceptionally important components of the marine carbon and nitrogen cycles (Delong, 2009; Ingalls et al., 2006).

Using Stable Isotopes to Better Understand the Origins of Biomarkers

One of the continuing challenges of organic geochemistry has been to determine the sources of individual biomarker mol-ecules, because few compounds are thought to exclusively derive from a single organism or represent a single biogeochemical pro-cess. Sterols, for example, are made or used by almost all eukary-otes, from microbes to mammals. Despite their varied chemical structures, sterols are not as diverse as the taxonomic distribu-tions that they represent. One way to constrain the origins of mol-ecules is through their stable isotope compositions. Initial work used bulk measurements (e.g., Hare et al., 1991), but a desire for improved specifi city and precision drove development of tools that allow isotopic measurements at the molecular level using

continuous-fl ow approaches that were analogous to GC-MS and LC-MS (liquid chromatography–mass spectrometry; see also previous discussion). Carbon was the fi rst and most obvious ele-ment targeted for compound-specifi c isotope analysis (CSIA), and the initial results immediately revealed the diverse origins of sedimentary organic molecules (Freeman et al., 1990; Hayes et al., 1990). Analogous methods for H (Sessions et al., 1999), N (Macko et al., 1997), and S (Amrani et al., 2009) soon followed. It is now common to measure multiple isotopic compositions of individual compounds such as chlorophylls, as well as the geo-porphyrins and maleimides that are derived from them (Chikarai-shi et al., 2005). In turn, the technical capability to make precise isotopic measurements on these elements exposed the dearth of knowledge about the ways in which these biosynthetic processes, organismic physiology, and environmental parameters infl u-ence the isotopic abundances of individual molecules (Hayes, 2001; Laws et al., 1997, 1995; Popp et al., 1998; Sessions et al., 1999; Smith and Freeman, 2006). The impact of developing and perfecting tools and techniques for compound-specifi c isotope analyses has been profound. The protocols and logic are now routinely used in archaeology, paleoclimatology, paleohydrology ecology, and forensics, and geochemists have been the instigators of many of these applications (Lichtfouse, 2000).

Radiocarbon measurements of individual organic molecules have also led to new insights into the age structure of fossil lipid assemblages, and their transport from source to sedimen-tary sinks (Pearson et al., 2001). Thus, in contemporary envi-ronmental settings, it is possible to discern the age relations of multiple paleoclimate proxies (Mollenhauer et al., 2003), thereby resolving processes such as winnowing and lateral advection of organic matter (Kusch et al., 2010; Mollenhauer et al., 2011). The robustness, scope, and fi delity of organic SST proxies based on algal alkenones and archaeal ether lipids have been signifi cantly increased through knowledge gained from compound-specifi c radiocarbon measurements (Eglinton and Eglinton, 2008).

Which Organisms Produce All These Sedimentary Lipids?

As interest turned to understanding the origins and diage-netic pathways by which organic molecules became fossilized, there was a great expansion in our knowledge of the diversity of lipids in sediments before, in fact, their biological precur-sors were identifi ed in living organisms (Ourisson et al., 1979; Rohmer, 2010; Rohmer et al., 1980). The emergence of another analytical tool, high-performance liquid chromatography cou-pled to mass spectrometry (HPLC-MS), enabled identifi cation of increasingly large and complex polar lipid molecules such that, today, geochemists have focused on tracking the distribu-tions of membrane-spanning lipids (Schouten et al., 2000a), intact polar lipids (Sturt et al., 2004), and complex biohopanoids (Talbot et al., 2008) from organisms through the water column and into sediments. Intact polar lipids can be proxies for living organisms and, often in combination with deoxyribonucleic acid (DNA), have been used to map out biogeochemical processes



in sediments (e.g., methanotrophy; Orphan et al., 2002) and the water column (e.g., Sinninghe Damsté et al., 2005), as well as to help defi ne the extent of the deep subsurface biosphere (Lipp et al., 2008; Rütters et al., 2002).

Advances in identifi cation of the diverse types of organic molecules in rocks have been concomitant with efforts to increase our understanding of the ways in which organisms pro-duce these lipids (biosynthetic pathways), the evolution of these pathways, and the array of functions performed by lipid mol-ecules in cells (physiological roles). This is one of the current frontiers of organic geochemistry, because, very often, the taxo-nomic relations are often ambiguous between molecules of inter-est in the environment or rock record and their precursor organ-isms. Based on knowledge derived from cultured taxa, we know that hopanoids, for example, are distributed very unevenly across the bacterial domain, but the reasons for this remain mysterious (Rohmer et al., 1984). Precisely how key compounds are pro-duced is only now coming to light, and efforts to deduce this have led to the discovery of a completely unknown isoprenoid biosyn-thesis pathway in bacteria (Rohmer, 2003; Rohmer et al., 1993); such fi ndings can have societal implications because they may lead to new classes of antibiotics. The revolution in genomics has, therefore, given geochemists a new tool for identifi cation of important biosynthetic genes and gene families, thereby enabling culture-independent approaches to be more nuanced in providing an understanding of the sources and function of biomarker lipids.

By way of example, we know that pentacyclic hopanoids and tetracyclic steroids are biosynthetically and evolutionarily related (Rohmer, 2010; Summons et al., 2006) because they are derived from the common acyclic precursor squalene. Genomic databases can be queried for both hopanoid and steroid cyclase genes, the sequences of which carry information on their evolu-tionary heritages (Fischer and Pearson, 2007). No longer are bio-synthetic capabilities solely based on studies of cultured organ-isms because their genomes and natural environmental samples can now be queried directly for the biosynthetic capabilities of community members (Fischer et al., 2005; Pearson et al., 2009). The unexpected occurrence of 2-methylhopanoids in cultures of a purple nonsulfur bacterium has led to a raft of new knowledge about the genes responsible for hopanoid biosynthesis (Welander et al., 2012) and the capacity to construct mutants lacking genes of interest (Welander et al., 2010), pointing to the involvement of these complex molecules in ameliorating stress (Doughty et al., 2009; Welander et al., 2009). In summary, molecular paleon-tology that is informed by individual and community genomes is providing signifi cant insights into the sources of organic mol-ecules that may be preserved in the rock record, which in turn provides fundamental constraints on the evolution of the biosyn-thetic pathways used by life.

GEOCHEMISTRY ON OTHER WORLDS

The last 50 yr have been a golden age for exploration of our solar system, and geochemistry has played a dominant role.

Meteorites were initially the only extraterrestrial samples avail-able for geochemical analysis, and these have now been supple-mented by samples returned from the Moon, a comet, an asteroid, and the solar wind. In addition, new kinds of meteorites, includ-ing samples from Mars and the Moon, have been recognized in the last fi ve decades. Signifi cant strides have also been made in geochemical analyses by remote sensing.

Extinct Radionuclides and Early Solar System Chronology

Short-lived radionuclides that were present in the solar neb-ula have decayed away, but their presence can be inferred from their radiogenic daughter isotopes. The fi rst extinct radionuclide found in meteorites, 129I, was discovered in 1961, and others followed in the next 50 yr. Additional extinct nuclides are now recognized to have been present in the solar nebula, including 10Be, 26Al, 41Ca, 53Mn, 60Fe, 107Pd, 146Sm, and 182Hf. In some cases, extinct radionuclides can provide high-resolution chronometers for early solar system events (Nyquist et al., 2009). The 146Sm-142Nd system has already been discussed here in regard to early terrestrial rocks.

One of the most useful extinct nuclides is 26Al, which decays to 26Mg with a half-life of ~730,000 yr. Its former existence was fi rst proven when Lee et al. (1977) analyzed plagioclase and other minerals in a refractory Ca-Al inclusion (CAI) from the Allende meteorite, and they determined that plagioclase had a large 26Mg excess (relative to nonradiogenic 24Mg). Because CAIs are the oldest solar system materials, based on their 207Pb-206Pb ages, their 26Al contents are taken to refl ect those at the beginning of the solar system, and other measurements are compared to the 26Al/27Al ratio and 207Pb-206Pb age of CAIs for chronology. The 26Al value refl ects production by stellar nucleosynthesis, and its occurrence in the solar nebula is commonly attributed to “seed-ing” by a nearby supernova. Refi nements in SIMS and ICP-MS techniques have permitted measurements of 26Al-26Mg systemat-ics in samples that have low Al contents, including chondrules and differentiated meteorites. An example is illustrated in Fig-ure 10, which shows different 26Al/27Al ratios for chondrules in several classes of chondrites, and this implies that chondrules formed several million years after CAIs.

The 182Hf-182W system, with a half-life of 8.9 m.y., has also proved to be very useful in dating early solar system events. Early measurements by N-TIMS (negative thermal ionization mass spectrometry) were diffi cult, but this chronometer was made truly accessible when MC-ICP-MS analysis became available (Halliday et al., 1996). Both Hf and W are refractory elements and so usually occur in chondritic proportions. Because, how-ever, Hf is lithophile and W is siderophile, the 182Hf-182W system can date metal-silicate fractionation events such as core forma-tion. Such ages, in turn, constrain the time of planet accretion. For example, the estimated age of Earth’s accretion (50% mass), based on this chronometer, is 30 to >100 m.y. after CAIs (Halli-day, 2004). More rapid accretion ages for asteroidal bodies were determined by Kleine et al. (2009).

24 Johnson et al.


Mass-Independent Isotope Fractionation and Early Solar System Conditions

One of the fi rst mass-independent fractionations (MIF) for stable isotopes was observed for O measured in CAIs in mete-orites (Clayton et al., 1973), as illustrated in Figure 11. For 17O/16O-18O/16O variations, mass-dependent fractionation defi nes a line with a slope of ~0.5. Subsequent analyses of O isotopes in bulk meteorites showed that igneous meteorites (achondrites) defi ned distinct fractionation lines parallel to that of Earth. Bulk chondrites of different classes also plot in distinct areas of the diagram. As a consequence, O isotopes have become a very use-ful criterion for meteorite classifi cation, and they place strong constraints on processes that operated early in the history of the solar system (Clayton, 2004).

The MIF trend for CAIs in Figure 11 was originally inter-preted as a mixing line between solar system gas (plotting on or above the terrestrial fractionation line) and exotic grains that contained pure 16O (plotting on an extension of the CAI line). A newer explanation is that the isotopic variations in CAIs arose from self-shielding during photodissociation of CO, a major nebular gas (Clayton, 2002). Because 16O is abundant, there were large differences between the amounts of C16O, on the one hand, and C17O and C18O on the other. The abundant C16O thus became optically thick far from the Sun; C17O and C18O remained opti-cally thin, and so were dissociated to a greater extent by radiation. This produced an inner zone in the nebula that was enriched in 17O and 18O, which reacted with H to make H

2O. The 16O-depleted

water then reacted with dust or condensates to form nebular sol-ids. In this model, the Sun, representing the nebular gas, must be 16O-rich. The fi rst measurement of O isotopes in the solar wind returned to Earth by the Genesis mission (McKeegan et al., 2011) is consistent with this model.

Figure 10. Ion microprobe measurements of initial 26Al/27Al, and corresponding ages relative to Ca-Al inclusion (CAI), in chon-drules from primitive meteorites (unequilibrated ordinary chon-drites and several classes of carbonaceous chondrites). Figure is from McSween and Huss (2010; used with permission). UOC—unequilibrated ordinary chondrite.

Nebular Condensation

Many, although not all, of the solids that now comprise solar system bodies once condensed from nebular vapor. Early attempts to determine the condensation sequence in a cooling gas of solar composition involved simple calculations, performed long before digital computers were available. The equilibrium condensation behavior of elements in a nebular gas was rigor-ously modeled by Grossman and Larimer (1974) and other work-ers since then (e.g., Petaev and Wood, 1998). The canonical con-densation sequence, showing condensing phases and the fraction of each element condensed at various temperatures, is illustrated in Figure 12. Based on experimental determinations of entropy, enthalpy, and heat capacity, equations of state describing the thermodynamic stabilities of a host of possible minerals under various conditions can be calculated. Liquids are not stable at the low pressures appropriate for the solar nebula. Some minerals in the condensation sequence do not condense directly, but form or adjust their compositions by reactions of previously condensed phases with the gas. This condensation calculation was done for the 23 elements with the highest cosmic abundances. Generally, thermodynamic data for trace elements are lacking, so chemi-cal analyses of trace elements in high-temperature minerals of chondritic meteorites guide our understanding of their condensa-tion behavior. Accordingly, in these models trace elements are allowed to condense as simple metals, oxides, or sulfi des, which are assumed to dissolve into appropriate major mineral phases. The validity of condensation calculations has been supported by studies of refractory inclusions (CAIs) in chondrites, which have bulk chemical compositions consistent with those calculated for the fi rst 5% of condensable matter (Davis and Richter, 2004). Moreover, the same minerals that comprise these inclusions are predicted to have been the earliest condensed phases (Fig. 12). There remains some controversy, however, about whether the CAIs are actually condensates or refractory residues from evapo-ration (the reverse of condensation).

Condensation calculations have also been done at different nebular pressures and using nonsolar gas compositions (Ebel and Grossman, 2000; Wood and Hashimoto, 1993; Yoneda and Gross-man, 1995). Parts of the solar nebula may have been enriched in dust, which, when vaporized, could have yielded nonsolar vapor, possibly allowing more reduced phases or even liquids to con-dense. These condensation calculations could also apply to other stars, e.g., red giants, which can have nonsolar compositions.

Organic Matter from Space

The molecular and isotopic chemistry of extraterrestrial organic matter is another area where technical innovation has spurred improved appreciation for the nature and diversity of organic compounds formed in the interstellar medium and the processes that alter them prior to their delivery to Earth in mete-orites (Derenne and Robert, 2010) and interstellar dust particles (Flynn et al., 2004). The low-molecular-weight compounds



identifi ed include many classes of biologically important mol-ecules, for example, amino acids (Kvenvolden et al., 1970), hydroxyacids and dicarboxylic acids (Cronin et al., 1993; Law-less et al., 1974; Peltzer et al., 1984), as well as nucleobases (Martins et al., 2008). Isotopic data and pyrolysis studies suggest that distinct processes are involved in formation of small mol-ecules and macromolecular material (Sephton et al., 2004; Yuen et al., 1984). Perhaps the most profound, enigmatic, and contro-versial fi nding was that some amino acids were not racemic, as had long been considered fact (Engel and Macko, 1997; Engel et al., 1990; Engel and Nagy, 1982). This discovery, originally dismissed as due to terrestrial contamination, became accepted once it was demonstrated also to be a feature of some nonpro-

Figure 11. Oxygen isotopes (relative to SMOW) measured in Ca-Al inclusion (CAI) from a carbonaceous chondrite defi ne a mass-independent fractionation trend, distinct from the terres-trial mass-dependent trend. Figure is modifi ed from Clayton et al. (1973). SMOW—standard mean ocean water.

tein amino acids (Cronin and Pizzarello, 1997). Moreover, the l- enantiomeric excess of some compounds such as isovaline in the Murchison and Orgueil meteorites appears to be related to the extent of aqueous processing, suggesting that it could refl ect amplifi cation of a small initial isovaline asymmetry. If correct, this would be inconsistent with the theory that ultraviolet (UV) circularly polarized light was the primary source of l-enrichment in amino acids. It seems possible, therefore, that early life on Earth had access to molecular building blocks with the left hand-edness that characterizes the amino acids of all life today (Engel and Macko, 2001; Glavin and Dworkin, 2009). Such results dem-onstrate the importance of “off-world” geochemistry to inform-ing us about terrestrial evolution.

Lunar Geochemistry

The Moon is a geochemical experiment conducted under vastly different conditions than Earth. The geochemical explo-ration of the Moon began with the return of samples by Apollo astronauts in 1969, which extended over an exciting 4 yr period. Depletions in volatile elements and enrichments in refractory ele-ments measured in these rocks support the giant impact hypoth-esis for the Moon’s formation (Taylor et al., 2006). Vaporization of portions of the target (Earth) and the impactor, followed by incomplete condensation in Earth orbit, can account for the frac-tionation observed in volatile and refractory elements. Measured depletions of siderophile elements in lunar rocks are consistent with models that indicate preferential incorporation of the silicate mantles of the impactor and target into the Moon, accompanied by accretion of the impactor’s core into Earth’s core.

The revelation that a magma ocean once existed on the Moon (Wood et al., 1970) has profoundly changed planetary science, and now wholesale melting has been proposed for a number of bodies, including early Earth. REE analyses came of age during

Figure 12. A model of the condensa-tion sequence for a cooling gas of solar composition at 10−4 atm pressure. Con-densed minerals are labeled in italics, and curves show the fraction of each element condensed as a function of temperature. REE—rare earth element. Figure is modifi ed from Grossman and Larimer (1974).

26 Johnson et al.


the Apollo program, as discussed earlier, and the complementary REE patterns of anorthosite from the highlands crust and basalts from the maria (Fig. 13) provide the most persuasive evidence for a lunar magma ocean (Taylor and Jakes, 1974). In this model, the thick anorthositic crust was formed by fl otation of plagioclase, which produced a positive Eu anomaly in the crust, and this was balanced by the Eu depletion that was produced by olivine and pyroxene that accumulated to form the mantle, a characteristic that was inherited by later mare basalt melts. The Eu anomaly in lunar rocks was enhanced by the Moon’s low oxidation state, which increased the proportion of Eu2+, allowing it to partition into plagioclase. The last dregs of the magma ocean were rich in Fe, Ti, and incompatible trace elements, including K, REEs, and P (the “KREEP” component), which became sandwiched between the crystallized crust and mantle. Gravitational overturn of this dense, Fe- and Ti-rich layer resulted in mixing of KREEP into magmas that subsequently intruded the anorthositic crust. The discovery of KREEP was initially made through geochemi-cal analyses of Apollo samples (Taylor et al., 2001).

Geochemistry by Remote Sensing

Geochemical analyses no longer are restricted to the labora-tory. Here, we consider two examples of geochemistry by space-craft that have altered our perceptions of what is possible using remote sensing.

The Gamma Ray Spectrometer on the Lunar Prospector orbiter analyzed the abundances of a handful of elements (Fe, Ti, K, Mg, Al, Ca, Si, Th, and H; Lawrence et al., 1998; Pretty-man et al., 2006), utilizing characteristic gamma-ray emissions produced by radioactive decay or by reactions initiated by cos-

Figure 13. Chondrite-normalized rare earth ele-ment (REE) patterns for lunar anorthosite and mare basalt. The complementary europium anom-alies for crust and mantle-derived rocks support the magma ocean hypothesis.

mic rays. Orbital measurements provide global coverage and thus are particularly useful in understanding geochemical processes at a planetary scale. Global maps of the distribution of Fe and Th, which are particularly sensitive to this technique, have been employed to distinguish lunar terranes based on their geochemi-cal characteristics (Jolliff et al., 2000). The petrology of each terrane has been interpreted unambiguously by comparison with the Fe and Th contents of returned lunar samples. Other results, made possible by measuring the neutron fl ux from the surface, included the discovery of H at the lunar poles associated with cold, permanently shadowed craters believed to contain water ice (Feldman et al., 2001).

The Mars Exploration Rovers Opportunity and Spirit car-ried Alpha Particle X-Ray Spectrometers (APXS) that analyzed the chemical compositions of hundreds of rocks and soils on the surface of Mars (Brückner et al., 2008; Gellert et al., 2006). By measuring characteristic X-rays produced by alpha particles and X-rays emitted from a radioactive source, the APXS was able to analyze all of the major elements and a few minor and trace elements. These studies were a critical part of the classifi cation of rocks encountered during rover traverses, and they provided constraints on the processes that formed these materials. Oppor-tunity analyzed sedimentary rocks that contained high contents of S, Cl, and Br, interpreted to refl ect evaporation of salt-laden brines, demonstrating that liquid water was once abundant on Mars (Squyres et al., 2004). On the other side of the planet, Spirit analyzed a variety of ancient basaltic rocks (Fig. 14) that have compositions that are distinct from those of younger Mar-tian basaltic meteorites, buttressing the argument that although heterogeneous, Mars is fundamentally a basalt-covered world (McSween et al., 2009).

Stardust in the Laboratory

Tiny motes of stardust, condensates that formed around dying stars or farther out in the interstellar medium, were dis-covered in chondritic meteorites (Lewis et al., 1987), after a long search beginning in the 1960s. These diamond nanoparti-cles were isolated by dissolving chondrites in a series of harsh acids, and at each step tracking the isotopically anomalous Xe they contained. The grains are thought to have been implanted when the outer layers of giant C-rich stars were sloughed off and condensed as diamond, and were subsequently implanted with distinctive nuclides produced when the star exploded as a super-nova. Approximately 20 types of presolar grains, including sili-con carbide (Tang and Anders, 1988) and graphite (Amari et al., 1990), as well as oxides, nitrides, and silicates, have now been found. All are distinguished by their exotic isotopic compositions (Tang and Anders, 1988).

The challenging elemental and isotopic analyses of stardust grains, which range from a few nanometers to a few micrometers in size, illustrate the great progress made in micro-analytical techniques. An example, using SIMS-analyzed C and N isotopes to fi ngerprint the sources of presolar silicon carbide grains, is



shown in Figure 15. More importantly, the isotopic compositions of presolar grains provide “ground truth” for stellar nucleosyn-thesis models, directly linking cosmochemical measurements in the laboratory to astrophysical theory. These data also (1) provide information on capture cross sections for neutrons, the capture of which makes heavier elements, (2) demonstrate where theoreti-cal models are inadequate to describe the internal structures of stars, (3) identify neutron sources for the s-process, and constrain the scale of mixing in supernovae (Nittler, 2003; Zinner, 2004). Presolar grains also fundamentally changed the way we think about the solar system’s formation. Their widespread occurrence in primitive meteorites demonstrates that a hot nebula did not vaporize all preexisting solids, overturning a view that prevailed 50 yr ago.

INTEGRATING THE PICTURE: GEOCHEMICAL CYCLES

The concept of geochemical cycles dates to the late nine-teenth century and was well established (i.e., began to appear in textbooks) by the mid-twentieth century (e.g., Goldschmidt, 1954). The modern phase of quantifying global geochemi-cal cycles arguably dates from the pioneering work of Garrels and Mackenzie (1971, 1972), who integrated major-element chemistry (including C, S, and Cl) into a steady-state recycling model for the evolution of the sedimentary mass. As described in greater detail in the following, geochemical cycles are often com-plex and inter-related, and considerable effort has been directed

Figure 14. A geochemical classifi cation diagram for volcanic rocks comparing the compositions of Martian rocks and soils derived from them in Gusev crater, analyzed by the Spirit rov-er’s Alpha Particle X-Ray Spectrometers (APXS), with labo-ratory analyses of Martian meteorites. Figure is modifi ed from McSween et al. (2009).

Figure 15. Carbon and nitrogen isotopic ratios in presolar silicon carbide grains, measured with an ion microprobe. Most grains from stars on the asymptotic giant branch (AGB) of the Hurtzsprung-Russell diagram plot above the solar composition (dashed lines). Supernovae grains have lower 14N/15N, and rare grains from no-vae, which are powered by explosive hydrogen burning, plot in the lower-left quadrant. The stellar sources of other grains are not cur-rently known. Figure is modifi ed from Zinner (2004).

toward trying to understand how cycles are linked. Many ele-mental cycles require monograph-length treatment, and the fact that there is a journal devoted solely to biogeochemical cycling (Global Biogeochemical Cycles) refl ects the sustained interest in geochemical cycles.

In geochemical cycles, the reservoirs (mass, M) and fl uxes of individual elements (or groups of closely related elements) into and out of the system (mass/unit time, F

in and F

out) are quantita-

tively accounted for over geological time (t). Accordingly, assess-ments of geochemical cycles are essentially a mass balance of the element of interest on some appropriate physical scale over some appropriate duration, and to a great extent are the natural conse-quence of quantifying the overall rock cycle (Gregor, 1992) and hydrological cycle (Fig. 16). Geochemical cycles are considered to be either “exogenic,” for those operating on or near Earth’s surface (hydrosphere, atmosphere, biosphere ± sediments), typi-cally on relatively short time scales (<~104–107 yr), or “endo-genic,” for those operating within the interior of Earth (oceanic and continental crust, mantle, core ± sediments), typically on relatively long time scales (>~104–107 yr). The physical inter-face between the endogenic and exogenic parts of an element’s geochemical cycle typically occurs in soils and the sedimentary cover. Characterization of complete global cycles requires inte-gration of both endogenic and exogenic cycles as witnessed, for example, by the relatively recent recognition that the C cycle is

28 Johnson et al.


affected by signifi cant levels of microbial activity deep within parts of Earth’s crust (e.g., Hazen et al., 2012), in addition to mantle sources.

Depending on the way in which the physical and/or tempo-ral scales are defi ned, geochemical cycles may be either open or closed, where the distinction is a function of whether or not exter-nal inputs and/or outputs take place within the defi ned reservoirs. Those concepts govern whether or not a geochemical cycle is in steady state (i.e., dM/dt = 0; F

in = F

out), leading to the further con-

cept of residence time (τ = M/F), which in turn is equivalent to the inverse of the fi rst-order rate constant (τ = 1/k) for simple lin-ear cycles, thus providing information about the response times (i.e., kinetics) of the system (Lasaga and Berner, 1998).

One problem that attracted early attention was the issue of elemental cycling through the oceans. This interest was initiated largely by the pioneering work of Sillèn (1961), who demon-strated that seawater chemistry could be in a steady-state con-dition, controlled by equilibrium reactions between atmospheric gases and marine carbonate and silicate minerals. Up until the early 1950s, the conventional wisdom was that the oceans’ salt had accumulated over the entirety of geological time. During

the 1960s–1980s, considerable effort went into determining the apparent mean oceanic residence times of the elements (τ), a con-cept fi rst introduced by Barth (1952), as well as documenting the balance between the masses of the various elements that entered the oceans from rivers, and those exiting the oceans through sedimentation (e.g., Drever et al., 1988). The ensuing research identifi ed important additional sources of elements to the oceans (e.g., basalt-seawater hydrothermal interaction) that had not been previously recognized, additional reservoirs for elements (e.g., pore waters, altered basalts, estuaries, continental shelves), and numerous mineral and biogeochemical reactions, leading to a far more complete understanding of marine chemistry (e.g., Broecker and Peng, 1982).

By evaluating geochemistry in the framework of geochemi-cal cycles, the fundamental processes that infl uence elemental distributions (geological, geophysical, chemical, biological, temporal) are better addressed (Lerman and Wu, 2007). A good example of the value of this approach in identifying previously unrecognized processes and reservoirs is the well-known “miss-ing sink” issue for carbon (Broecker, 2012). During the 1990s, the short-term (exogenic) C cycle had been quantifi ed suffi ciently

Figure 16. Representation of the pre-industrial (A.D. 1750) global carbon cycle (adapted from Sundquist and Visser, 2003). Boxes show major reservoirs, with carbon mass listed in units of Pg (1015 g), and arrows show major fl uxes of carbon (listed in units of Pg/yr). DIC—dissolved inorganic carbon.



to recognize that ~15%–20% of the CO2 delivered to the atmo-

sphere by fossil fuel combustion could not be accounted for by the recognized major C reservoirs (atmosphere, accounting for ~50%, and seawater, accounting for ~30%–35%). This led to a major research effort to identify the missing sink, which in turn resulted in recognition of the terrestrial biosphere (previously considered relatively minor) as a major reservoir for exogenic C.

The Biological Connection

The importance of biological activity in controlling the cycling of a wide range of elements has been long appreciated (e.g., Vinogradov, 1943). Attempts, however, to both defi ne and quantify global biogeochemical cycles involving elements that are essential for biological activity have only occurred somewhat recently, though they are of great geochemical interest (e.g., Gar-rels and Lerman, 1981). Biogeochemical cycles consider some or all of a wide array of biologically relevant elements in an inte-grated manner, including those that are a signifi cant part of living tissue and skeletons (C, O, H, N, P, S, Ca, Si), those that may be less involved in living matter but are important in redox processes allowing for energy transfer (Fe, Mn, in addition to N and S), and a long list of minor and trace elements that may be necessary for metabolism (e.g., biolimiting) and/or substitute for major ele-ments of biogeochemical importance (e.g., Fe, Mn, Mg, Ba, Ge, B, Mo, V, Zn, and even the REEs).

The C cycle (and its role in biogeochemical cycles) is of intrinsic interest to geochemists due to the role of C in all bio-logical activity and its importance in the geological record as a common rock-forming constituent (limestone, dolomite, carbo-naceous sediment) and as a natural resource (fossil fuels). The fact, however, that C has received by far the most attention of any of the geochemical cycles, and indeed may be the most intensely studied geochemical problem of the past 50 yr and more (Ber-ner, 2004; Broecker, 2012; Des Marais, 2001), is mainly due to the observation that atmospheric CO

2 concentrations have sys-

tematically risen since the industrial revolution at a rate that is unprecedented for the Phanerozoic, due mainly to the burning of fossil fuels (Fig. 17). It is now widely recognized that perturba-tions within the C cycle (temporal variations in atmospheric CO

2)

represent a dominant control on the changes of Earth’s current climate, as well as both short-term and long-term paleoclimate. The recent abrupt increase in atmospheric levels of this impor-tant greenhouse gas, due mainly to human activity, is also now understood to play a central role in infl uencing recent global cli-mate change and increases in mean global surface temperatures (Berner, 2003).

Redox Processes in Biogeochemical Cycles

The recent focus on element cycling, including biogeochem-ical cycles, is primarily driven by redox chemistry, both abiologic or biologic, due to its importance in controlling the composition and evolution of the ocean-atmosphere system (e.g., Raiswell

Figure 17. Monthly mean concentration of carbon dioxide, mea-sured as parts per million molecules of CO

2 in total molecules of

dry air, measured at the Mauna Loa Observatory, Hawaii (data from National Oceanic and Atmospheric Administration [NOAA] Web site (http://www.esrl.noaa.gov/gmd/ccgg/trends/), the so-called Keeling curve. The annual cycle observed in the record is a sea-sonal effect due to the much greater land mass (and vegetation) in the Northern Hemisphere, resulting in greater CO

2 uptake during

plant growth in the northern summer. These relations were fi rst rec-ognized by Charles Keeling (Keeling, 1960).

and Canfi eld, 2012). As discussed already, C is certainly the most studied redox-sensitive element that is cycled between reduced and oxidized forms today on Earth, driven largely by photosyn-thesis. This in turns drives atmospheric O

2 contents through the

extent of C burial and thus plays a central role in the long-term O cycle. This infl uences, for example, the long-term cycling of Fe, one of the most important redox elements in rocks, given its high abundance in the terrestrial planets. The redox couple with Fe was studied extensively in the 1970s, with a major focus on banded iron formations (BIFs), and later, the relation between the Fe biogeochemical and marine evolutions. Evaluation of the geochemical cycles of redox-sensitive trace elements such as Cr and Mo rose in prominence in the 1980s and 1990s.

One of the major issues addressed in studies of redox bio-geochemical cycling over the past 50 yr has been the history of atmospheric O

2 over geological time (Holland, 1962). As briefl y

noted already, a prominent approach to understanding the his-tory of atmospheric O

2 on Earth has been study of both mass-

dependent and mass-independent fractionation (MIF) of stable S isotopes, and these records are compared, along with C and Fe isotopes, in Figure 18. The range in d34S values for sulfi des in marine sedimentary rocks generally increases with decreas-ing age, an observation that has also been long recognized (e.g., Canfi eld, 2001). The general increase in the highest d34S values measured for sulfi des is broadly taken to record an increase in the d34S values of seawater sulfate, and a recent compilation of this can be found in Canfi eld and Farquhar (2009). It is almost universally accepted that the range in d34S values for sulfi des

30 Johnson et al.


Figure 18. Temporal variations in (A) banded iron formation (BIF) deposition and atmospheric oxy-gen, (B) d13C values of Ca-Mg carbonates, (C) d34S values for sulfi des, (D) mass-independent S isotope fractionation, expressed as D33S, and (E) δ56Fe values; δ56Fe values are broken out as black shales (high-S, high-C contents), Ca-Mg carbonates, and BIF sam-ples. Green band represents period of maximum BIF deposition and immediately predates the increase in atmospheric oxygen contents. BIF deposition is from Bekker et al. (2010). Atmospheric O

2 curve is from

Catling and Claire (2005), shifted to older ages based on disappearance of mass-independent S isotope frac-tionation (Farquhar et al., 2010). Early pulses of O

2

time band are based on previous studies (Anbar et al., 2007; Czaja et al., 2012; Duan et al., 2010; Kendall et al., 2010; Voegelin et al., 2010). The d13C data are from Shields and Veizer (2002). The d34S data for sulfi des and seawater sulfate are from Canfi eld and Farquhar (2009), and D33S values for sulfi des are from same source. The δ56Fe values are from sources cited in Johnson et al. (2008b), with additional data sources (Czaja et al., 2010, 2012; Heimann et al., 2010; Hof-mann et al., 2009; Planavsky et al., 2009; Steinhoefel et al., 2009; Tsikos et al., 2010; Valaas-Hyslop et al., 2008; Von Blanckenburg et al., 2008). All isotopic data refl ect bulk sample analyses.



refl ects various extents of bacterial sulfate reduction (Canfi eld, 2001, 2005). Turning to S-MIF, nonzero D33S values for marine sedimentary rocks are restricted to samples of ca. 2450 Ma age and older, and most studies interpret this to indicate very low atmospheric oxygen contents (originally discovered by Farquhar et al. [2000a], and recently reviewed by Farquhar et al. [2010]), although alternative explanations have been proposed (discussed earlier herein). The transition from large D33S values to zero D33S values at ca. 2450 Ma correlates with an increase in the range of d34S values, consistent with an increase in seawater sulfate con-tents and development of free oxygen in the atmosphere, which in turn would enhance rates of bacterial sulfate reduction (Can-fi eld, 2005; Farquhar et al., 2011).

As discussed earlier, the d13C values for Ca-Mg carbonates of Archean and Proterozoic age largely scatter closely about zero, with the exception of the 2.3–2.0 Ga Lomagundi excursion (Fig. 18). Increased organic C burial seems the most likely explana-tion for the increase in d13C values for carbonates at this time, which in turn would drive further increase in atmospheric O

2.

This would tend to increase seawater sulfate contents, providing opportunities to increase the 34S/32S fractionations produced by microbial sulfate reduction due to “excess” sulfate; this would increase the inventory of sedimentary sulfi des that have very neg-ative d34S values (Fig. 18). Accompanying the rise in atmospheric O

2 would be a loss of S-MIF, refl ected in a shift toward zero D33S

values for sulfi des younger than 2.3 or 2.4 Ga in age (Fig. 18).The largest Fe isotope excursion known in the rock record

occurs in the Neoarchean and Paleoproterozoic (Fig. 18). Because the vast majority of Fe in the crust has a δ56Fe value near zero, including low-C, low-S sedimentary rocks that have Fe contents similar to those of the average crust (e.g., Johnson et al., 2008b), deviations in the δ56Fe values from zero are signifi -cant and generally rare in terms of the Fe mass balance of Earth. The zero to positive δ56Fe values for rocks older than 3.5 Ga in age, which to date mainly include BIFs, are generally accepted to refl ect partial oxidation of marine hydrothermal Fe2+

aq, suggest-

ing that the amount of oxidant was limited (Dauphas et al., 2004), and recent Fe isotope work suggests that photic zone O

2 levels

were <0.001% of present day at this time (Czaja et al., 2013). More controversial, however, is the strong decrease in δ56Fe val-ues between ca. 3 and 2.5 Ga. Rouxel et al. (2005) and Anbar and Rouxel (2007) proposed that oxidation of marine hydrothermal Fe2+

aq during BIF genesis, or oxide precipitation on continental

shelves, produced negative δ56Fe values in seawater, which was directly incorporated into sulfi de-rich marine sedimentary rocks. As recently discussed by Czaja et al. (2012), extensive oxidation of Fe2+

aq is a likely explanation for the low δ56Fe values of Neo-

archean Ca-Mg carbonates and, in fact, provides one of several lines of evidence for intermittent oxygenation of surface environ-ments in the time leading up to the Great Oxidation Event (GOE) at ca. 2.3 or 2.4 Ga (Anbar et al., 2007; Kendall et al., 2010; Voegelin et al., 2010). Oxidation of Fe2+

aq, however, cannot easily

explain the spread in δ56Fe values for Fe-rich rocks such as BIFs, and Yamaguchi et al. (2005) and Johnson et al. (2008a, 2008b) do

not generally interpret the Fe isotope compositions of such lithol-ogies to be a direct proxy for seawater, instead favoring microbial Fe cycling as an explanation for the Fe isotope variability.

The relative paucity of highly negative δ56Fe values after the GOE (Fig. 18) has been interpreted to refl ect a decrease in the footprint of microbial Fe3+ reduction as microbial sulfate reduction increased in its impact on redox cycling. Sulfate reduc-tion would produce abundant sulfi de, which readily reacts with Fe3+ oxides, decreasing Fe3+ availability to microbial reduction. The increase in δ56Fe values that generally corresponds with a decrease in BIF deposition (Fig. 18) is consistent with a tran-sition from an Fe-dominated redox cycle before the GOE to a S-dominated redox cycle after the GOE. Collectively, the tempo-ral record of BIF deposition and C, S, and Fe isotopes displays remarkable coherence that can be well explained by expected changes in microbial redox cycling during the transition from an anoxic to oxic world.

We close by considering redox cycling on another planetary body. The recent phase of Mars exploration has led to recogni-tion of a long-lived dynamic sedimentary cycle on that planet. This has generated interest in characterizing geochemical cycles and evaluating the implications for surface processes (e.g., Grotz-inger et al., 2011), an interest that undoubtedly will increase with the successful landing of the Mars Science Laboratory. Mars may represent a valuable end member for understanding near-surface geochemical cycles in general due to the absence of plate tecton-ics. There appears to be a long-term evolution of sedimentary mineralogy, where a common occurrence of Noachian (older than 3.7 Ga) clay minerals gives way to a sulfate-rich mineralogy in the Hesperian (ca. 3.7–3.2 Ga), and fi nally to anhydrous ferric oxides in the Amazonian (younger than 3.2 Ga; Bibring et al., 2006). Recent work on the ALH84001 meteorite has provided support for clay minerals on the surface of Mars prior to 4.0 Ga (Beard et al., 2013). As on Earth, where sedimentary mineralogy (e.g., BIF) provides a proxy for geochemical evolution (Hazen et al., 2008), this secular change in sedimentary mineralogy has been interpreted to represent acidifi cation, oxidation, and desic-cation of the Martian surface over geological time. Although the current atmosphere is very thin, most models call for a more sub-stantial early Martian atmosphere that was dominated by CO

2.

Sulfur, however, is also enriched in Mars surfi cial deposits, and it now appears that some form of a S-cycle may have strongly infl u-enced surfi cial processes over much of geological time, leading to widespread low pH conditions (Halevy et al., 2007; King and McLennan, 2010; McLennan, 2012).

CONCLUDING REMARKS

We are not so bold as to predict what advances will occur in geochemistry in the next 50 yr, as we are quite sure that many of the analytical and computational gains that have been key factors to a number of geochemical advances were not envisioned by geochemists in the early 1960s. It was also probably not clear to these geochemists what the impact of plate tectonics would

32 Johnson et al.


be, nor of obtaining samples from the Moon, or of in situ analy-ses of rocks and soils on Mars. In a shorter time frame than fi ve decades, however, it seems likely that the trend toward interdis-ciplinary research, where multiple large data sets are brought to bear on geologic problems, is likely to expand. In addition, there will likely be huge gains in computational and data stor-age capacity, although analytical advances may be more muted, given the fact that many analyses are now limited by counting statistics rather than electronics, although there is certainly room for large improvements in terms of sampling effi ciency of many instruments (what gets to the detector versus what is consumed). A major unknown for the future of geochemistry, and of science in general, is the commitment to research by broader society, and such issues will undoubtedly affect what is written in a similar paper in the early 2060s.

ACKNOWLEDGMENTS

We thank the organizing committee of the 125th anniversary cel-ebration of the Geological Society of America for the opportu-nity to write this summary of progress in geochemistry. Johnson and Summons acknowledge support from the National Aeronau-tics and Space Administration (NASA) Astrobiology Institute during the time this review was prepared. McLennan acknowl-edges support from NASA. McSween acknowledges support from NASA’s Cosmochemistry Program. We thank Richard Carlson, Robert Clayton, John Hayes, Scott Samson, John Val-ley, and Richard Walker for helpful reviews of the manuscript.

REFERENCES CITED

Abouchami, W., and Goldstein, S.L., 1995, A lead isotopic study of circum-Antarctic manganese nodules: Geochimica et Cosmochimica Acta, v. 59, no. 9, p. 1809–1820, doi:10.1016/0016-7037(95)00084-D.

Adam, P., Schmid, J.C., Mycke, B., Strazielle, C., Connan, J., Huc, A., Riva, A., and Albrecht, P., 1993, Structural investigation of nonpolar sulfur cross-linked macromolecules in petroleum: Geochimica et Cosmochimica Acta, v. 57, no. 14, p. 3395–3419, doi:10.1016/0016-7037(93)90547-A.

Albarède, F., 2004, The stable isotope geochemistry of copper and zinc: Reviews in Mineralogy and Geochemistry, v. 55, p. 409–427, doi:10.2138/gsrmg.55.1.409.

Alizai, A., Clift, P.D., Giosan, L., Van Laningham, S., Hinton, R., Tabrez, A.R., and Danish, M., 2011, Pb isotopic variability in the modern– Pleistocene Indus River system measured by ion microprobe in detrital K- feldspar grains: Geochimica et Cosmochimica Acta, v. 75, p. 4771–4795, doi:10.1016/j.gca.2011.05.039.

Allègre, C.J., 1968, Comportement des systemes U-Th-Pb dans le manteau superieur et modele d’evolution de ce dernier au cours des temps geolo-giques: Earth and Planetary Science Letters, v. 5, p. 261–269, doi:10.1016/S0012-821X(68)80050-0.

Allègre, C.J., and Luck, J.-M., 1980, Osmium isotopes as petrogenetic and geo-logical tracers: Earth and Planetary Science Letters, v. 48, p. 148–154, doi:10.1016/0012-821X(80)90177-6.

Allègre, C.J., and Minster, J.F., 1978, Quantitative models of trace element behavior in magmatic processes: Earth and Planetary Science Letters, v. 38, p. 1–25, doi:10.1016/0012-821X(78)90123-1.

Allègre, C.J., and Rousseau, D., 1984, The growth of the continent through geologi-cal time studied by Nd isotope analysis of shales: Earth and Planetary Science Letters, v. 67, p. 19–34, doi:10.1016/0012-821X(84)90035-9.

Alvarez, L.W., Alvarez, W., Asaro, F., and Michel, H.V., 1980, Extraterrestrial cause for the Cretaceous-Tertiary extinction: Science, v. 208, p. 1095–1108, doi:10.1126/science.208.4448.1095.

Alves, S., Schiano, P., and Allègre, C.J., 1999, Rhenium-osmium isotopic investigation of Java subduction zone lavas: Earth and Planetary Science Letters, v. 168, p. 65–77, doi:10.1016/S0012-821X(99)00050-3.

Alves, S., Schiano, P., Capmas, F.O., and Allègre, C.J., 2002, Osmium isotope binary mixing arrays in arc volcanism: Earth and Planetary Science Let-ters, v. 198, p. 355–369, doi:10.1016/S0012-821X(02)00524-1.

Amari, S., Anders, E., Virag, A., and Zinner, E., 1990, Interstellar graphite in meteorites: Nature, v. 345, p. 238–240, doi:10.1038/345238a0.

Amelin, Y., Lee, D.C., Halliday, A.N., and Pidgeon, R.T., 1999, Nature of the Earth’s earliest crust from hafnium isotopes in single detrital zircons: Nature, v. 399, p. 252–255, doi:10.1038/20426.

Amelin, Y., Lee, D.C., and Halliday, A.N., 2000, Early-Middle Archaean crustal evolution deduced from Lu-Hf and U-Pb isotopic studies of single zir-con grains: Geochimica et Cosmochimica Acta, v. 64, p. 4205–4225, doi:10.1016/S0016-7037(00)00493-2.

Amrani, A., Sessions, A.L., and Adkins, J.F., 2009, Compound-specifi c δ34S analysis of volatile organics by coupled GC/multicollector-ICPMS: Ana-lytical Chemistry, v. 81, p. 9027–9034, doi:10.1021/ac9016538.

Anbar, A.D., 2004, Molybdenum stable isotopes: Observations, interpretations and directions: Reviews in Mineralogy and Geochemistry, v. 55, p. 429–454, doi:10.2138/gsrmg.55.1.429.

Anbar, A.D., and Rouxel, O., 2007, Metal stable isotopes in paleoceanogra-phy: Annual Review of Earth and Planetary Sciences, v. 35, p. 717–746, doi:10.1146/annurev.earth.34.031405.125029.

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, p. 1903–1906, doi:10.1126/science.1140325.

Andersen, D.J., Lindlsey, D.H., and Davidson, P.M., 1993, QUILF—A Pas-cal program to assess equilibria among Fe-Mg-Mn-Ti oxides, pyroxenes, olivine, and quartz: Computers & Geosciences, v. 19, p. 1333–1350, doi:10.1016/0098-3004(93)90033-2.

Archer, C., and Vance, D., 2006, Coupled Fe and S isotope evidence for Archean microbial Fe(III) and sulfate reduction: Geology, v. 34, p. 153–156, doi:10.1130/G22067.1.

Armstrong, R.L., 1968, A model for Sr and Pb isotopic evolution in a dynamic Earth: Reviews of Geophysics, v. 6, p. 175–199, doi:10.1029/RG006i002p00175.

Armstrong, R.L., 1981, Radiogenic isotopes: The case for crustal recycling on a steady-state no-continental-growth Earth: Philosophical Transactions of the Royal Society, ser. A, v. 301, p. 443–472, doi:10.1098/rsta.1981.0122.

Barth, T.W., 1952, Theoretical Petrology: New York, John Wiley, 387 p.Beard, B.L., and Johnson, C.M., 2004a, Fe isotope variations in the modern

and ancient Earth and other planetary bodies: Reviews in Mineralogy and Geochemistry, v. 55, p. 319–357.

Beard, B.L., and Johnson, C.M., 2004b, Inter-mineral Fe isotope variations in mantle-derived rocks and implications for the Fe geochemical cycle: Geochimica et Cosmochimica Acta, v. 68, p. 4727–4743, doi:10.1016/j.gca.2004.04.023.

Beard, B.L., Johnson, C.M., Cox, L., Sun, H., Nealson, K.H., and Aguilar, C., 1999, Iron isotope biosignatures: Science, v. 285, p. 1889–1892, doi:10.1126/science.285.5435.1889.

Beard, B.L., Johnson, C.M., Skulan, J.L., Nealson, K.H., Cox, L., and Sun, H., 2003a, Application of Fe isotopes to tracing the geochemical and bio-logical cycling of Fe: Chemical Geology, v. 195, p. 87–117, doi:10.1016/S0009-2541(02)00390-X.

Beard, B.L., Johnson, C.M., Von Damm, K.L., and Poulson, R.L., 2003b, Iron isotope constraints on Fe cycling and mass balance in oxygen-ated Earth oceans: Geology, v. 31, p. 629–632, doi:10.1130/0091-7613(2003)031<0629:IICOFC>2.0.CO;2.

Beard, B.L., Ludois, J.M., Lapen, T.J., and Johnson, C.M., 2013, Pre–4.0 billion year weathering on Mars constrained by Rb-Sr geochronology on mete-orite ALH84001: Earth and Planetary Science Letters, v. 361, p. 173–182, doi:10.1016/j.epsl.2012.10.021.

Becker, L., Poreda, R.J., Hunt, A.G., Bunch, T.E., and Rampino, M., 2001, Impact event at the Permian-Triassic boundary: Evidence from extra-terrestrial noble gases in fullerenes: Science, v. 291, p. 1530–1533, doi:10.1126/science.1057243.

Becker, R.H., and Clayton, R.N., 1972, Carbon isotopic evidence for the origin of a banded iron-formation in Western Australia: Geochimica et Cosmo-chimica Acta, v. 36, p. 577–595, doi:10.1016/0016-7037(72)90077-4.



Bekker, A., Slack, J.F., Planavsky, N., Krapez, B., Hofmann, A., Konhauser, K.O., and Rouxel, O.J., 2010, Iron formation: The sedimentary product of a complex interplay among mantle, tectonic, oceanic, and biospheric pro-cesses: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 105, no. 3, p. 467–508, doi:10.2113/gsecongeo.105.3.467.

Belousova, E.A., Kostitsyn, Y.A., Griffi n, W.L., O’Reilly, S.Y., and Pear-son, N.J., 2010, The growth of the continental crust: Constraints from zircon Hf-isotope data: Lithos, v. 119, p. 457–466, doi:10.1016/j.lithos.2010.07.024.

Bence, A.E., and Albee, A.L., 1968, Empirical correction factors for the elec-tron microanalysis of silicates and oxides: The Journal of Geology, v. 76, p. 382–403, doi:10.1086/627339.

Bennett, V.C., and DePaolo, D.J., 1987, Proterozoic crustal history of the west-ern United States as determined by neodymium isotopic mapping: Geo-logical Society of America Bulletin, v. 99, p. 674–685, doi:10.1130/0016-7606(1987)99<674:PCHOTW>2.0.CO;2.

Bennett, V.C., Brandon, A.D., and Nutman, A.P., 2007, Coupled 142Nd-143Nd isotopic evidence for Hadean mantle dynamics: Science, v. 318, p. 1907–1910, doi:10.1126/science.1145928.

Bergquist, P.R., Karuso, P., Cambie, R.C., and Smith, D.J., 1991, Sterol com-position and classifi cation of the Porifera: Biochemical Systematics and Ecology, v. 19, p. 17–24, doi:10.1016/0305-1978(91)90109-D.

Berner, R.A., 2003, The long-term carbon cycle, fossil fuels and atmospheric composition: Nature, v. 426, p. 323–326, doi:10.1038/nature02131.

Berner, R.A., 2004, The Phanerozoic Carbon Cycle: Oxford, UK, Oxford Uni-versity Press, 158 p.

Bethke, C.M., 2008, Geochemical and Biogeochemical Reaction Modeling: Cambridge, UK, Cambridge University Press, 543 p.

Beukes, N.J., and Gutzmer, J., 2008, Origin and paleoenvironmental signifi -cance of major iron formations at the Archean-Paleoproterozoic bound-ary: Society of Economic Geologists (SEG) Reviews, v. 15, p. 5–47.

Beukes, N.J., Klein, C., Kaufman, A.J., and Hayes, J.M., 1990, Carbonate petrography, kerogen distribution, and carbon and oxygen isotope varia-tions in an Early Proterozoic transition from limestone to iron-formation deposition, Transvaal Supergroup, South-Africa: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 85, p. 663–690, doi:10.2113/gsecongeo.85.4.663.

Bibring, J.-P., Langevin, Y., Mustard, J. F., Poulet, F., Arvidson, R., Gendrin, A., Gondet, B., Mangold, N., Pinet, P., Forget, F., et al., 2006, Global mineralogi-cal and aqueous Mars history derived from OMEGA/Mars Express data: Sci-ence, v. 312, no. 5772, p. 400–404, doi:10.1126/science.1122659.

Bigeleisen, J., and Mayer, M.G., 1947, Calculation of equilibrium constants for isotopic exchange reactions: The Journal of Chemical Physics, v. 15, p. 261–267, doi:10.1063/1.1746492.

Birck, J.L., 2004, An overview of isotopic anomalies in extraterrestrial materi-als and their nucleosynthetic heritage: Reviews in Mineralogy and Geo-chemistry, v. 55, p. 25–64, doi:10.2138/gsrmg.55.1.25.

Blichert-Toft, J., and Albarède, F., 2008, Hafnium isotopes in Jack Hills zircons and the formation of the Hadean crust: Earth and Planetary Science Let-ters, v. 265, p. 686–702, doi:10.1016/j.epsl.2007.10.054.

Blichert-Toft, J., Frey, F.A., and Albarède, F., 1999, Hf isotope evidence for pelagic sediments in the source of Hawaiian basalts: Science, v. 285, p. 879–882, doi:10.1126/science.285.5429.879.

Boyet, M., and Carlson, R.W., 2005, 142Nd evidence for early (>4.53 Ga) global differentiation of the silicate Earth: Science, v. 309, p. 576–581, doi:10.1126/science.1113634.

Boyet, M., and Carlson, R.W., 2006, A new geochemical model for the Earth’s mantle inferred from 146Sm-142Nd systematics: Earth and Planetary Sci-ence Letters, v. 250, p. 254–268, doi:10.1016/j.epsl.2006.07.046.

Brandon, A.D., and Walker, R.J., 2005, The debate over core-mantle interaction: Earth and Planetary Science Letters, v. 232, p. 211–225, doi:10.1016/j.epsl.2005.01.034.

Brandon, A.D., Norman, M.D., Walker, R.J., and Morgan, J.W., 1999, 186Os-187Os systematics of Hawaiian picrites: Earth and Planetary Science Let-ters, v. 174, p. 25–42, doi:10.1016/S0012-821X(99)00251-4.

Brandon, A.D., Snow, J.E., Walker, R.J., Morgan, J.W., and Mock, T.D., 2000, 190Pt-186Os and 187Re-187Os systematics of abyssal peridotites: Earth and Planetary Science Letters, v. 177, p. 319–335, doi:10.1016/S0012-821X(00)00044-3.

Brantley, S.L., and Lebedeva, M., 2011, Learning to read the chemistry of rego-lith to understand the critical zone: Annual Review of Earth and Planetary Sciences, v. 39, p. 387–416, doi:10.1146/annurev-earth-040809-152321.

Brassell, S.C., Eglinton, G., Marlowe, I.T., Pfl aumann, U., and Sarnthein, M., 1986, Molecular stratigraphy: A new tool for climatic assessment: Nature, v. 320, p. 129–133, doi:10.1038/320129a0.

Brocks, J.J., Love, G.D., Summons, R.E., Knoll, A.H., Logan, G.A., and Bowden, S.A., 2005, Biomarker evidence for green and purple sulphur bacteria in a stratifi ed Palaeoproterozoic sea: Nature, v. 437, p. 866–870, doi:10.1038/nature04068.

Broecker, W., 2012, The carbon cycle and climate change: Memoirs of my 60 years in science: Geochemical Perspectives, v. 1, p. 221–339, doi:10.7185/geochempersp.1.2.

Broecker, W.S., and Peng, T.-H., 1982, Tracers in the Sea: Palisades, New York, Lamont-Doherty Geological Observatory, 690 p.

Brückner, J., Dreibus, G., Gellert, R., Squyres, S.W., Wänke, H., Yen, A., and Zipfel, J., 2008, Mars Exploration Rovers: Chemical composition by the APXS, in Bell, J., III, ed., The Martian Surface: Composition, Miner-alogy, and Physical Properties: Cambridge, UK, Cambridge University Press, p. 58–101.

Bullen, T.D., White, A.F., Childs, C.W., Vivit, D.V., and Schulz, M.S., 2001, Demonstration of signifi cant abiotic iron isotope fractionation in nature: Geology, v. 29, p. 699–702, doi:10.1130/0091-7613(2001)029<0699:DOSAII>2.0.CO;2.

Burdett, J.W., Arthur, M.A., and Richardson, M., 1989, A Neogene seawater sulfur isotope age curve from calcareous pelagic microfossils: Earth and Planetary Science Letters, v. 94, p. 189–198, doi:10.1016/0012-821X(89)90138-6.

Burke, W.H., Denison, R.E., Hetherington, E.A., Koepnick, R.B., Nelson, H.F., and Otto, J.B., 1982, Variation of seawater 87Sr/86Sr throughout Phanerozoic time: Geology, v. 10, p. 516–519, doi:10.1130/0091-7613(1982)10<516:VOSSTP>2.0.CO;2.

Burton, K.W., Ling, H.-F., and O’Nions, R.K., 1997, Closure of the Central American isthmus and its effect on deep-water formation in the North Atlantic: Nature, v. 386, p. 382–385, doi:10.1038/386382a0.

Byrne, R.H., and Sholkovitz, E.R., 1996, Marine chemistry and geochemistry of the lanthanides, in Gschneidner, K.A., Jr., and Eyring, L., eds., Hand-book on the Physics and Chemistry of Rare Earths, Volume 23: Amster-dam, Elsevier, p. 497–593.

Canfi eld, D.E., 2001, Biogeochemistry of sulfur isotopes: Reviews in Mineral-ogy and Geochemistry, v. 43, p. 607–636, doi:10.2138/gsrmg.43.1.607.

Canfi eld, D.E., 2005, The early history of atmospheric oxygen: Homage to Robert Garrels: Annual Review of Earth and Planetary Sciences, v. 33, p. 1–36, doi:10.1146/annurev.earth.33.092203.122711.

Canfi eld, D.E., and Farquhar, J., 2009, Animal evolution, bioturbation, and the sulfate concentration of the oceans: Proceedings of the National Academy of Sciences of the United States of America, v. 106, no. 20, p. 8123–8127, doi:10.1073/pnas.0902037106.

Canfi eld, D.E., and Teske, A., 1996, Late Proterozoic rise in atmospheric oxy-gen concentration inferred from phylogenetic and sulphur-isotope stud-ies: Nature, v. 382, p. 127–132, doi:10.1038/382127a0.

Canfi eld, D.E., and Thamdrup, B., 1994, The production of 34S-depleted sul-fi de during bacterial disproportionation of elemental sulfur: Science, v. 266, p. 1973–1975, doi:10.1126/science.11540246.

Canfi eld, D.E., Poulton, S.W., and Narbonne, G.M., 2007, Late-Neoprotero-zoic deep-ocean oxygenation and the rise of animal life: Science, v. 315, p. 92–95, doi:10.1126/science.1135013.

Cao, C., Love, G.D., Hays, L.E., Wang, W., Shen, S., and Summons, R.E., 2009, Biogeochemical evidence for euxinic oceans and ecological disturbance presaging the end-Permian mass extinction event: Earth and Planetary Science Letters, v. 281, p. 188–201, doi:10.1016/j.epsl.2009.02.012.

Carlson, R.W., and Boyet, M., 2008, Composition of the Earth’s interior: The importance of early events: Philosophical Transactions of the Royal Soci-ety, ser. A, v. 366, p. 4077–4103, doi:10.1098/rsta.2008.0166.

Carlson, R.W., Pearson, D.G., and James, D.E., 2005, Physical, chemical, and chronological characteristics of continental mantle: Reviews of Geophys-ics, v. 43, RG1001, p. 1–24, doi:10.1029/2004RG000156.

Catling, D.C., and Claire, M.W., 2005, How Earth’s atmosphere evolved to an oxic state: A status report: Earth and Planetary Science Letters, v. 237, no. 1–2, p. 1–20, doi:10.1016/j.epsl.2005.06.013.

Cavosie, A.J., Valley, J.W., Wilde, S.A., and E.I.M.F, 2005, Magmatic δ18O in 4400–3900 Ma detrital zircons: A record of the alteration and recycling of crust in the Early Archean: Earth and Planetary Science Letters, v. 235, p. 663–681, doi:10.1016/j.epsl.2005.04.028.

34 Johnson et al.


Chan, L.H., 1987, Lithium isotope analysis by thermal ionization mass spec-trometry of lithium tetraborate: Analytical Chemistry, v. 59, p. 2662–2665, doi:10.1021/ac00149a007.

Chesley, J.T., and Ruiz, J., 1998, Crust-mantle interaction in large igneous prov-inces: Implications from the Re-Os isotope systematics of the Columbia River fl ood basalts: Earth and Planetary Science Letters, v. 154, p. 1–11, doi:10.1016/S0012-821X(97)00176-3.

Chikaraishi, Y., Matsumoto, K., Ogawa, N.O., Suga, H., Kitazato, H., and Ohk-ouchi, N., 2005, Hydrogen, carbon and nitrogen isotopic fractionations during chlorophyll biosynthesis in C3 higher plants: Phytochemistry, v. 66, p. 911–920, doi:10.1016/j.phytochem.2005.03.004.

Clayton, R.N., 2002, Self-shielding in the solar nebula: Nature, v. 415, p. 860–861, doi:10.1038/415860b.

Clayton, R.N., 2004, Oxygen isotopes in meteorites, in Davis, A.M., ed., Treatise on Geochemistry, Volume 1: Meteorites, Comets, and Planets: Oxford, UK, Elsevier, p. 192–142.

Clayton, R.N., Grossman, L., and Mayeda, T.K., 1973, A component of primi-tive nuclear composition in carbonaceous meteorites: Science, v. 182, p. 485–488, doi:10.1126/science.182.4111.485.

Condie, K.C., 1998, Episodic continental growth and super-continents: A man-tle avalanche connection?: Earth and Planetary Science Letters, v. 163, p. 97–108, doi:10.1016/S0012-821X(98)00178-2.

Craig, H., 1961, Isotopic variations in meteoric waters: Science, v. 133, p. 1702–1703, doi:10.1126/science.133.3465.1702.

Craig, H., 1966, Isotopic composition and origin of the Red Sea and Salton Sea geothermal brines: Science, v. 154, p. 1544–1548, doi:10.1126/science.154.3756.1544.

Craig, H., Gordon, L.I., and Horibe, Y., 1963, Isotopic exchange effects in the evaporation of water: Journal of Geophysical Research, v. 68, p. 5079–5087, doi:10.1029/JZ068i017p05079.

Creaney, S., and Passe, Q., 1993, Recurring patterns of total organic carbon and source rock quality within a sequence stratigraphic framework: American Association of Petroleum Geologists Bulletin, v. 77, p. 386–401.

Creaser, R.A., Papanastassiou, D.A., and Wasserburg, G.J., 1991, Negative thermal ion mass spectrometry of osmium, rhenium and iridium: Geo-chimica et Cosmochimica Acta, v. 55, p. 397–401, doi:10.1016/0016-7037(91)90427-7.

Criss, R.E., 1999, Principles of Stable Isotope Distribution: New York, Oxford University Press, 254 p.

Criss, R.E., and Taylor, H.P.J., 1983, An 18O/16O and D/H study of Tertiary hydrothermal systems in the southern half of the Idaho batholith: Geo-logical Society of America Bulletin, v. 94, p. 640–663, doi:10.1130/0016-7606(1983)94<640:AOADSO>2.0.CO;2.

Cronin, J.R., and Pizzarello, S., 1997, Enantiomeric excesses in meteoritic amino acids: Science, v. 275, p. 951–955, doi:10.1126/science.275.5302.951.

Cronin, J.R., Pizzarello, S., Epstein, S., and Krishnamurthy, R.V., 1993, Molecu-lar and isotopic analyses of the hydroxy acids, dicarboxylic acids, and hydroxicarboxylic acids of the Murchison meteorite: Geochimica et Cos-mochimica Acta, v. 57, p. 4745–4752, doi:10.1016/0016-7037(93)90197-5.

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.

Czaja, A.D., Johnson, C.M., Roden, E.E., Beard, B.L., Voegelin, A.R., Nag-gler, T.F., Beukes, N.J., and Wille, M., 2012, Evidence for free oxygen in the Neoarchean ocean based on coupled iron and molybdenum iso-tope fractionation: Geochimica et Cosmochimica Acta, v. 86, p. 118–137, doi:10.1016/j.gca.2012.03.007.

Czaja, A.D., Johnson, C.M., Beard, B.L., Roden, E.E., Li, W., and Moorbath, S., 2013, Biological Fe oxidation controlled deposition of banded iron formation in ca. 3770 Ma Isua Supracrustal Belt (West Greenland): Earth and Planetary Science Letters, v. 363, p. 192–203, doi:10.1016/j.epsl.2012.12.025.

Dansgaard, W., Johnsen, S.J., Muller, J., and Langway, C.C., 1969, One thou-sand centuries of climatic record from Camp Century on the Greenland Ice Sheet: Science, v. 166, no. 3903, p. 377–380, doi:10.1126/science.166.3903.377.

Dauphas, N., van Zuilen, M., Wadhwa, M., Davic, A.M., Marty, B., and Janney, P.E., 2004, Clues from Fe isotope variations on the origin of Early Archean BIFs from Greenland: Science, v. 306, p. 2077–2080, doi:10.1126/science.1104639.

Davidson, J.P., 1983, Lesser Antilles isotopic evidence for the role of sub-ducted sediment in island arc magma genesis: Nature, v. 306, p. 253–256, doi:10.1038/306253a0.

Davidson, J.P., Tepley, F., Palacz, Z., and Meffan-Main, S., 2001, Magma recharge, contamination and residence times revealed by in situ laser ablation isotopic analysis of feldspar in volcanic rocks: Earth and Plan-etary Science Letters, v. 184, p. 427–442, doi:10.1016/S0012-821X(00)00333-2.

Davidson, J.P., Morgan, D.J., Charlier, B.L.A., Harlou, R., and Hora, J.M., 2007, Microsampling and isotopic analysis of igneous rocks: Implications for the study of magmatic systems: Annual Review of Earth and Planetary Sci-ences, v. 35, p. 273–311, doi:10.1146/annurev.earth.35.031306.140211.

Davis, A.M., and Richter, F.M., 2004, Condensation and evaporation of solar system materials, in Davis, A.M., ed., Treatise on Geochemistry, Volume 1: Meteorites, Comets and Planets: Oxford, UK, Elsevier, p. 407–430.

Delong, E.F., 2009, The microbial ocean from genomes to biomes: Nature, v. 459, p. 200–206, doi:10.1038/nature08059.

Delong, E.F., King, L.L., Massana, R., Cittone, H., Murray, A., Schleper, C., and Wakeham, S.G., 1998, Dibiphytanyl ether lipids in nonthermophilic crenar-chaeotes: Applied and Environmental Microbiology, v. 64, p. 1133–1138.

DePaolo, D.J., 1981a, Neodymium isotopes in the Colorado Front Range and crust-mantle evolution in the Proterozoic: Nature, v. 291, p. 193–197, doi:10.1038/291193a0.

DePaolo, D.J., 1981b, Trace element and isotopic effects of combined wallrock assimilation and fractional crystallization: Earth and Planetary Science Letters, v. 53, p. 189–202, doi:10.1016/0012-821X(81)90153-9.

DePaolo, D.J., 1986, Detailed record of the Neogene Sr isotopic evolu-tion of seawater from DSDP Site 590B: Geology, v. 14, p. 103–106, doi:10.1130/0091-7613(1986)14<103:DROTNS>2.0.CO;2.

DePaolo, D.J., 2004, Calcium isotopic variations produced by biological, kinetic, radiogenic and nucleosynthetic processes: Reviews in Mineral-ogy and Geochemistry, v. 55, p. 255–288, doi:10.2138/gsrmg.55.1.255.

DePaolo, D.J., and Wasserburg, G.J., 1976, Nd isotopic variations and pet-rogenetic models: Geophysical Research Letters, v. 3, p. 249–252, doi:10.1029/GL003i005p00249.

Derenne, S., and Robert, F., 2010, Model of molecular structure of the insoluble organic matter isolated from Murchison meteorite: Meteoritics & Plane-tary Science, v. 45, p. 1461–1475, doi:10.1111/j.1945-5100.2010.01122.x.

Des Marais, D.J., 2001, Isotopic evolution of the biogeochemical carbon cycle during the Precambrian: Reviews in Mineralogy and Geochemistry, v. 43, p. 555–578, doi:10.2138/gsrmg.43.1.555.

Dhuime, B., Hawkesworth, C.J., Cawood, P.A., and Storey, C.D., 2012, A change in the geodynamics of continental growth 3 billion years ago: Sci-ence, v. 335, p. 1334–1336, doi:10.1126/science.1216066.

Dickin, A.P., 1995, Radiogenic Isotope Geology: Cambridge, UK, Cambridge University Press, 452 p.

Doe, B.R., and Zartman, R.E., 1979, Plumbotectonics: I. The Phanerozoic, in Barnes, H.L., ed., Geochemistry of Hydrothermal Ore Deposits: New York, Wiley Interscience, p. 22–70.

Domagal-Goldman, S., 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.

Doughty, D.M., Hunter, R.C., Summons, R.E., and Newman, D.K., 2009, 2-methylhopanoids are maximally produced in akinetes of Nostoc punctiforme: Geobiological implications: Geobiology, v. 7, p. 524–532, doi:10.1111/j.1472-4669.2009.00217.x.

Douthitt, C.B., 1982, The geochemistry of the stable isotopes of silicon: Geo-chimica et Cosmochimica Acta, v. 46, p. 1449–1458, doi:10.1016/0016-7037(82)90278-2.

Drever, J.I., Li, Y.-H., and Maynard, J.B., 1988, Geochemical cycles: The con-tinental crust and the oceans, in Gregor, C.B., Garrels, R.M., Mackenzie, F.T., and Maynard, J.B., eds., Chemical Cycles in the Evolution of the Earth: New York, John Wiley, p. 17–53.

Duan, Y., Anbar, A.D., Arnold, G.L., Lyons, T.W., Gordon, G.W., and Kendall, B., 2010, Molybdenum isotope evidence for mild environmental oxygen-ation before the Great Oxidation Event: Geochimica et Cosmochimica Acta, v. 74, no. 23, p. 6655–6668, doi:10.1016/j.gca.2010.08.035.

Dutkiewicz, A., Volk, H., George, S.C., Ridley, J., and Buick, R., 2006, Bio-markers from Huronian oil-bearing fl uid inclusions: An uncontaminated record of life before the Great Oxidation Event: Geology, v. 34, p. 437–440, doi:10.1130/G22360.1.



Dutta, S., Tripathi, S.M., Mallick, M., Mathews, R.P., Greenwood, P.F., Rao, M.R., and Summons, R.E., 2011, Eocene out-of-India dispersal of Asian dipterocarps: Review of Palaeobotany and Palynology, v. 166, p. 63–68, doi:10.1016/j.revpalbo.2011.05.002.

Ebel, D.S., and Grossman, L., 2000, Condensation in dust-enriched systems: Geochimica et Cosmochimica Acta, v. 64, p. 339–366, doi:10.1016/S0016-7037(00)00550-0.

Eglinton, G., 1970, Chemical Fossils: San Francisco, California, W.H. Freeman.Eglinton, T.I., and Eglinton, G., 2008, Molecular proxies for paleoclimatol-

ogy: Earth and Planetary Science Letters, v. 275, p. 1–16, doi:10.1016/j.epsl.2008.07.012.

Eigenbrode, J.L., and Freeman, K.H., 2006, Late Archean rise of aerobic microbial ecosystems: Proceedings of the National Academy of Sciences of the United States of America, v. 103, p. 15,759–15,764, doi:10.1073/pnas.0607540103.

Eiler, J.M., 2007, Clumped-isotope geochemistry: The study of naturally-occurring, multiply-substituted isotopologues: Earth and Planetary Sci-ence Letters, v. 262, p. 309–327, doi:10.1016/j.epsl.2007.08.020.

Elderfi eld, H., 1986, Strontium isotope stratigraphy: Palaeogeography, Pal-aeoclimatology, Palaeoecology, v. 57, p. 71–90, doi:10.1016/0031-0182(86)90007-6.

Engel, A.E.J., Engel, C.G., and Havens, R.G., 1965, Chemical characteristics of ocean basalts and the upper mantle: Geological Society of America Bulletin, v. 76, p. 719–734, doi:10.1130/0016-7606(1965)76[719:CCOOBA]2.0.CO;2.

Engel, M.H., and Macko, S.A., 1997, Isotopic evidence for extraterrestrial non-racemic amino acids in the Murchison meteorite: Nature, v. 389, p. 265–268, doi:10.1038/38460.

Engel, M.H., and Macko, S.A., 2001, The stereochemistry of amino acids in the Murchison meteorite: Precambrian Research, v. 106, p. 35–45, doi:10.1016/S0301-9268(00)00123-6.

Engel, M.H., and Nagy, B., 1982, Distribution and enantiomeric composition of amino acids in the Murchison meteorite: Nature, v. 296, p. 837–840, doi:10.1038/296837a0.

Engel, M.H., Macko, S.A., and Silfer, J.A., 1990, Carbon isotope composi-tion of individual amino acids in the Murchison meteorite: Nature, v. 348, p. 47–49, doi:10.1038/348047a0.

Epstein, S., Sharp, R.P., and Goddard, I., 1963, Oxygen isotope ratios in Ant-arctic snow, fi rn and ice: The Journal of Geology, v. 71, p. 698–720, doi:10.1086/626950.

Erwin, D.H., 2006, Extinction: How Life on Earth Nearly Ended 250 Million Years Ago: Princeton, Princeton University Press, 320 p.

Erwin, D.H., Lafl amme, M., Tweedt, S.M., Sperling, E.A., Pisani, D., and Peterson, K.J., 2011, The Cambrian conundrum: Early divergence and later ecological success in the early history of animals: Science, v. 334, p. 1091–1097, doi:10.1126/science.1206375.

Escrig, S., Doucelance, R., Moreira, M., and Allègre, C.J., 2005, Os isotope systematics in Fogo Island: Evidence for lower continental crust frag-ments under the Cape Verde southern islands: Chemical Geology, v. 219, p. 93–113, doi:10.1016/j.chemgeo.2005.02.011.

Fairbairn, H.W., Schlecht, W.G., Stevens, R.E., Dennen, W.H., Ahrens, L.H., and Chayes, F., 1951, A Co-Operative Investigation of Precision and Accuracy in Chemical, Spectrochemical and Model Analysis of Silicate Rocks: U.S. Geological Survey Bulletin 980, p. 1–71.

Farley, K.A., and Mukhopadhyay, S., 2001, An extraterrestrial impact at the Permian-Triassic boundary?: Science, v. 293, p. 2343, doi:10.1126/science.293.5539.2343a.

Farmer, G.L., and DePaolo, D.J., 1983, Origin of Mesozoic and Tertiary granite from the western United States and implications for pre-Mesozoic crustal structure: I. Nd and Sr isotopic studies in the geocline of the northern Great Basin: Journal of Geophysical Research, v. 88, p. 3379–3401, doi:10.1029/JB088iB04p03379.

Farquhar, J., Bao, H., and Thiemens, M., 2000a, Atmospheric infl uence of Earth’s earliest sulfur cycle: Science, v. 289, p. 756–758, doi:10.1126/science.289.5480.756.

Farquhar, J., Savarino, J., Jackson, T.L., and Thiemens, M.H., 2000b, Evidence of atmospheric sulphur in the Martian regolith from sulphur isotopes in meteorites: Nature, v. 404, p. 50–52, doi:10.1038/35003517.

Farquhar, J., Savarino, J., Airieau, S., and Thiemens, M.H., 2001, Observation of wavelength-sensitive mass-independent sulfur isotope effects during SO

2 photolysis: Implications for the early atmosphere: Journal of Geo-

physical Research, v. 106, p. 32,829–32,839, doi:10.1029/2000JE001437.

Farquhar, J., Wu, N.P., Canfi eld, D.E., and Oduro, H., 2010, Connections between sulfur cycle evolution, sulfur isotopes, sediments, and base metal sulfi de deposits: Economic Geology and the Bulletin of the Society of Economic Geologists, v. 105, no. 3, p. 509–533, doi:10.2113/gsecongeo.105.3.509.

Farquhar, J., Zerkle, A.L., and Bekker, A., 2011, Geological constraints on the origin of oxygenic photosynthesis: Photosynthesis Research, v. 107, no. 1, p. 11–36, doi:10.1007/s11120-010-9594-0.

Faure, G., and Hurley, P.M., 1963, The isotopic composition of strontium in oceanic and continental basalts: Application to the origin of igneous rocks: Journal of Petrology, v. 4, p. 31–50, doi:10.1093/petrology/4.1.31.

Faure, G., Hurley, P.M., and Powell, J.L., 1965, The isotopic composition of strontium in surface water from the North Atlantic Ocean: Geo-chimica et Cosmochimica Acta, v. 29, p. 209–220, doi:10.1016/0016-7037(65)90018-9.

Fedo, C.M., and Whitehouse, M.J., 2002, Metasomatic origin of quartz- pyroxene rock, Akilia, Greenland, and implications for Earth’s earliest life: Science, v. 296, p. 1448–1452, doi:10.1126/science.1070336.

Feldman, W.C., Maurice, S., Lawrence, D.J., Little, R.C., Lawson, S.L., Gas-nault, O., Wiens, R.C., Barraclough, B.L., Elphic, R.C., Prettyman, T.H., Steinberg, J.T., and Binder, A.B., 2001, Evidence for water ice near the lunar poles: Journal of Geophysical Research, v. 106, p. 23,231–23,251, doi:10.1029/2000JE001444.

Fike, D.A., Grotzinger, J.P., Pratt, L.M., and Summons, R.E., 2006, Oxida-tion of the Ediacaran ocean: Nature, v. 444, p. 744–747, doi:10.1038/nature05345.

Fischer, W.W., and Pearson, A., 2007, Hypotheses for the origin and early evo-lution of triterpenoid cyclases: Geobiology, v. 5, p. 19–34.

Fischer, W.W., Summons, R.E., and Pearson, A., 2005, Targeted genomic detec-tion of biosynthetic pathways: Anaerobic production of hopanoid bio-markers by a common sedimentary microbe: Geobiology, v. 3, p. 33–40, doi:10.1111/j.1472-4669.2005.00041.x.

Fischer, W.W., Schroeder, S., Lacassie, J.P., Beukes, N.J., Goldberg, T., Strauss, H., Horstmann, U.E., Schrag, D.P., and Knoll, A.H., 2009, Isotopic con-straints on the Late Archean carbon cycle from the Transvaal Supergroup along the western margin of the Kaapvaal craton, South Africa: Precam-brian Research, v. 169, p. 15–27, doi:10.1016/j.precamres.2008.10.010.

Fleck, R.J., and Criss, R.E., 1985, Strontium and oxygen isotopic variations in Mesozoic and Tertiary plutons of central Idaho: Contributions to Mineral-ogy and Petrology, v. 90, p. 291–308, doi:10.1007/BF00378269.

Flynn, G.J., Keller, L.P., Jacobsen, C., and Wirick, S., 2004, An assessment of the amount and types of organic matter contributed to the Earth by interplanetary dust: Advances in Space Research, v. 33, p. 57–66, doi:10.1016/j.asr.2003.09.036.

Freeman, K.H., Hayes, J.M., Trendel, J.M., and Albrecht, P., 1990, Evidence from carbon isotope measurements for diverse origins of sedimentary hydrocarbons: Nature, v. 343, p. 254–256, doi:10.1038/343254a0.

Frei, D., Liebscher, A., Franz, G., Wunder, B., Klemme, S., and Blundy, J., 2009a, Trace element partitioning between orthopyroxene and anhydrous silicate melt on the lherzolite solidus from 1.1 to 3.2 GPa and 1,230 to 1,535°C in the model system Na2O-CaO-MgO-Al2O3-SiO2: Contri-butions to Mineralogy and Petrology, v. 157, p. 473–490, doi:10.1007/s00410-008-0346-5.

Frei, R., Gaucher, C., Poulton, S.W., and Canfi eld, D.E., 2009b, Fluctuations in Precambrian atmospheric oxygenation recorded by chromium isotopes: Nature, v. 461, p. 250–253, doi:10.1038/nature08266.

French, K.L., Tosca, N.J., Cao, C., and Summons, R.E., 2012, Diagenetic and detrital origin of moretane anomalies through the Permian-Triassic bound-ary: Geochimica et Cosmochimica Acta, v. 84, p. 104–125, doi:10.1016/j.gca.2012.01.004.

Friedman, I., Redfi eld, A.C., Schoen, B., and Harris, J., 1964, The variation of the deuterium content of natural waters in the hydrologic cycle: Reviews of Geophysics, v. 2, p. 177–224, doi:10.1029/RG002i001p00177.

Friedman, I., Lipman, P.W., Obradovich, J.D., Gleason, J.D., and Christiansen, R.L., 1974, Meteoric water in magmas: Science, v. 184, p. 1069–1072, doi:10.1126/science.184.4141.1069.

Gaffney, A.M., Nelson, B.K., Reisberg, L., and Eiler, J., 2005, Oxygen-osmium isotope systematics of West Maui lavas: A record of shallow-level mag-matic processes: Earth and Planetary Science Letters, v. 239, p. 122–139, doi:10.1016/j.epsl.2005.07.027.

Gaines, S.M., Eglinton, G., and Rullkotter, J., 2009, Echoes of Life: What Fos-sil Molecules Reveal about Earth History: New York, Oxford University Press, 367 p.

36 Johnson et al.


Garrels, R.M., and Lerman, A., 1981, Phanerozoic cycles of sedimentary car-bon and sulfur: Proceedings of the National Academy of Sciences of the United States of America, v. 78, p. 4652–4656.

Garrels, R.M., and Mackenzie, F.T., 1971, Evolution of Sedimentary Rocks: New York, Norton, 397 p.

Garrels, R.M., and Mackenzie, F.T., 1972, A quantitative model for the sedi-mentary rock cycle: Marine Chemistry, v. 1, p. 27–41.

Garrels, R.M., and Thompson, M.E., 1962, A chemical model for sea water at 25°C and one atmosphere total pressure: American Journal of Science, v. 260, p. 57–66, doi:10.2475/ajs.260.1.57.

Gast, P.W., Tilton, G.R., and Hedge, C., 1964, Isotopic composition of lead and strontium from Ascension and Gough Islands: Science, v. 145, p. 1181–1185, doi:10.1126/science.145.3637.1181.

Gellert, R., Rieder, R., Brückner, J., Clark, B.C., Dreibus, G., Klingelhöfer, G., Lugmair, G., Ming, D.W., Wänke, H., Yen, A., Zipfel, J., and Squyres, S.W., 2006, The Alpha Particle X-ray Spectrometer (APXS): Results from Gusev crater and calibration report: Journal of Geophysical Research, v. 111, no. E2, doi:10.1029/2005JE002555.

George, S.C., Volk, H., Dutkiewicz, A., Ridley, J., and Buick, R., 2008, Pres-ervation of hydrocarbons and biomarkers in oil trapped inside fl uid inclusions for 2 billion years: Geochimica et Cosmochimica Acta, v. 72, p. 844–870, doi:10.1016/j.gca.2007.11.021.

Ghosh, P., Adkins, J., Affek, H., Balta, B., Guo, W., Schauble, E.A., Schrag, D., and Eiler, J.M., 2006, 13C–18O bonds in carbonate minerals: A new kind of paleothermometer: Geochimica et Cosmochimica Acta, v. 70, p. 1439–1456, doi:10.1016/j.gca.2005.11.014.

Gill, R., 1997, Modern Analytical Geochemistry: An Introduction to Quantita-tive Chemical Analysis for Earth, Environmental and Materials Scientists: Singapore, Addison Wesley Longman, 329 p.

Glavin, D.P., and Dworkin, J.P., 2009, Enrichment of the amino acid l-isovaline by aqueous alteration on CI and CM meteorite parent bodies: Proceedings of the National Academy of Sciences of the United States of America, v. 106, p. 5487–5492, doi:10.1073/pnas.0811618106.

Goldschmidt, V.M., 1954, Geochemistry: Oxford, UK, Oxford University Press, 730 p.

Goldstein, S.L., O’Nions, R.K., and Hamilton, P.J., 1984, A Sm-Nd isotopic study of atmospheric dusts and particulates from major river systems: Earth and Planetary Science Letters, v. 70, p. 221–236.

Grantham, P.J., and Wakefi eld, L.L., 1988, Variations in the sterane carbon number distributions of marine source rock derived crude oils through geological time: Organic Geochemistry, v. 12, p. 61–73, doi:10.1016/0146-6380(88)90115-5.

Green, T.H., 1994, Experimental studies of trace-element partitioning appli-cable to igneous petrogenesis—Sedona 16 years later: Chemical Geology, v. 117, p. 1–36, doi:10.1016/0009-2541(94)90119-8.

Gregor, B., 1992, Some ideas on the rock cycle: 1788–1988: Geochimica et Cos-mochimica Acta, v. 56, p. 2993–3000, doi:10.1016/0016-7037(92)90285-Q.

Gregory, R.T., and Criss, R.E., 1986, Isotopic exchange in open and closed systems: Reviews in Mineralogy, v. 16, p. 91–127.

Gregory, R.T., and Taylor, H.P.J., 1981, An oxygen isotope profi le in a section of Cretaceous oceanic crust, Samail ophiolite, Oman: Journal of Geo-physical Research, v. 86, p. 2737–2755, doi:10.1029/JB086iB04p02737.

Grice, K., Gibbison, R., Atkinson, J.E., Schwark, L., Eckardt, C.B., and Max-well, J.R., 1996, Maleimides (1H-pyrrole-2,5-diones) as molecular indicators of anoxygenic photosynthesis in ancient water columns: Geo-chimica et Cosmochimica Acta, v. 60, p. 3913–3924, doi:10.1016/0016-7037(96)00199-8.

Grice, K., Cao, C., Love, G.D., Bottcher, M.E., Twitchett, R.J., Grosjean, E., Summons, R.E., Turgeon, S.C., Dunning, W., and Jin, Y., 2005, Photic zone euxinia during the Permian-Triassic superanoxic event: Science, v. 307, p. 706–709, doi:10.1126/science.1104323.

Griffi n, W.L., Pearson, N.J., Belousova, E., Jackson, S.E., van Achterbergh, E., O’Reilly, S.Y., and Shee, S.R., 2000, The Hf isotope composition of cra-tonic mantle: LAM-MC-ICPMS analysis of zircon megacrysts in kimber-lites: Geochimica et Cosmochimica Acta, v. 64, p. 133–147, doi:10.1016/S0016-7037(99)00343-9.

Grosjean, E., Love, G.D., Stalvies, C., Fike, D.A., and Summons, R.E., 2009, Origin of petroleum in the Neoproterozoic–Cambrian South Oman Salt Basin: Organic Geochemistry, v. 40, p. 87–110, doi:10.1016/j.orggeochem.2008.09.011.

Grossman, L., and Larimer, J.W., 1974, Early chemical history of the solar system: Reviews of Geophysics and Space Physics, v. 12, p. 71–101, doi:10.1029/RG012i001p00071.

Grotzinger, J., Beaty, D., Dromart, G., Gupta, S., Harris, M., Hurowitz, J., Kocurek, G., McLennan, S., Milliken, R., Ori, G.G., and Sumner, D., 2011, Sedimentary geology: Key concepts and outstanding questions: Astrobiology, v. 11, p. 77–87, doi:10.1089/ast.2010.0571.

Halevy, I., Zuber, M.T., and Schrag, D.P., 2007, A sulfur dioxide climate feedback on early Mars: Science, v. 318, p. 1903–1907, doi:10.1126/science.1147039.

Halliday, A.N., 2004, The origin and earliest history of the Earth, in Davis, A.M., ed., Treatise on Geochemistry, Volume 1: Meteorites, Comets, and Planets: Oxford, UK, Elsevier, p. 509–557.

Halliday, A.N., Rehkemper, M., Lee, D.-C., and Yi, W., 1996, Early evolution of the Earth and Moon: New constraints from Hf-W isotope geochemistry: Earth and Planetary Science Letters, v. 142, p. 75–89, doi:10.1016/0012-821X(96)00096-9.

Halliday, A.N., Lee, D.-C., Christensen, J.N., Rehkemper, M., Yi, W., Luo, X., Hall, C.M., Ballentine, C.J., Pettke, T., and Stirling, C., 1998, Applica-tions of Multiple Collector-ICPMS to cosmochemistry, geochemistry, and paleoceanography: Geochimica et Cosmochimica Acta, v. 62, p. 919–940, doi:10.1016/S0016-7037(98)00057-X.

Hare, E.P., Fogel, M.L., Stafford, J.T.W., Mitchell, A.D., and Hoering, T.C., 1991, The isotopic composition of carbon and nitrogen in individual amino acids isolated from modern and fossil proteins: Journal of Archaeo-logical Science, v. 18, p. 277–292, doi:10.1016/0305-4403(91)90066-X.

Harper, C.L., and Jacobsen, S.B., 1992, Evidence from coupled 147Sm-143Nd and 146Sm-142Nd systematics for very early (4.5-Gyr) differentiation of the Earth’s mantle: Nature, v. 360, p. 728–732, doi:10.1038/360728a0.

Harrison, T.M., Blichert-Toft, J., Miller, W., Albarède, F., Holden, P., and Mojz-sis, S.J., 2005, Heterogeneous Hadean hafnium: Evidence of continen-tal crust at 4.4 to 4.5 Ga: Science, v. 310, p. 1947–1950, doi:10.1126/science.1117926.

Harrison, T.M., Schmitt, A.K., McCulloch, M.T., and Lovera, O.M., 2008, Early (4.5 Ga) formation of terrestrial crust: Lu-Hf, δ18O, and Ti ther-mometry results for Hadean zircons: Earth and Planetary Science Letters, v. 268, p. 476–486, doi:10.1016/j.epsl.2008.02.011.

Hart, G.L., Johnson, C.M., Hildreth, W., and Shirey, S.B., 2003, New osmium isotope evidence for intracrustal recycling of crustal domains with discrete ages: Geology, v. 31, p. 427–430, doi:10.1130/0091-7613(2003)031<0427:NOIEFI>2.0.CO;2.

Hart, S.R., 1988, Heterogeneous mantle domains: Signatures, genesis and mix-ing chronologies: Earth and Planetary Science Letters, v. 90, p. 273–296, doi:10.1016/0012-821X(88)90131-8.

Haskin, L., and Gehl, M.A., 1962, The rare-earth distribution in sediments: Journal of Geophysical Research, v. 67, p. 2537–2541, doi:10.1029/JZ067i006p02537.

Hawkesworth, C., and Schersten, A., 2007, Mantle plumes and geochem-istry: Chemical Geology, v. 241, p. 319–331, doi:10.1016/j.chemgeo.2007.01.018.

Hawkesworth, C.J., O’Nions, R.K., and Arculus, R.J., 1979, Nd and Sr iso-tope geochemistry of island arc volcanics, Grenada, Lesser Antilles: Earth and Planetary Science Letters, v. 45, p. 237–248, doi:10.1016/0012-821X(79)90126-2.

Hawkesworth, C.J., Kempton, P.D., Rogers, N.W., Ellam, R.M., and van Cal-steren, P.W., 1990, Continental mantle lithosphere, and shallow level enrichment processes in the Earth’s mantle: Earth and Planetary Science Letters, v. 96, p. 256–268, doi:10.1016/0012-821X(90)90006-J.

Hawkesworth, C., Cawood, P., Kemp, T., Storey, C., and Dhuime, B., 2009, A matter of preservation: Science, v. 323, p. 49–50, doi:10.1126/science.1168549.

Hawkesworth, C.J., Dhuime, B., Pietranik, A.B., Cawood, P.A., Kemp, A.I.S., and Storey, C.D., 2010, The generation and evolution of the continental crust: Journal of the Geological Society of London, v. 167, p. 229–248, doi:10.1144/0016-76492009-072.

Haxel, G.B., Hedrick, J.B., and Orris, G.R., 2005, Rare Earth Elements—

Critical Resources for High Technology: U. S. Geological Survey Fact Sheet 087–02, 4 p.

Hayes, J.M., 1983, Geochemical evidence bearing on the origin of aerobiosis, a speculative interpretation, in Schopf, J.W., ed., The Earth’s Earliest Bio-sphere: Its Origin and Evolution: Princeton, New Jersey, Princeton Uni-versity Press, p. 291–301.

Hayes, J.M., 2001, Fractionation of carbon and hydrogen isotopes in biosyn-thetic processes: Reviews in Mineralogy and Geochemistry, v. 43, p. 225–277, doi:10.2138/gsrmg.43.1.225.



Hayes, J.M., Freeman, K.H., Popp, B.N., and Hoham, C.H., 1990, Compound-specifi c isotopic analyses: A novel tool for reconstruction of ancient biogeochemical processes: Organic Geochemistry, v. 16, p. 1115–1128, doi:10.1016/0146-6380(90)90147-R.

Hays, L.E., Beatty, T., Henderson, C.M., Love, G.D., and Summons, R.E., 2007, Evidence for photic zone euxinia through the end-Permian mass extinction in the Panthalassic Ocean (Peace River Basin, western Can-ada): Palaeoworld, v. 16, p. 39–50, doi:10.1016/j.palwor.2007.05.008.

Hays, L.E., Grice, K., Foster, C.B., and Summons, R.E., 2012, Biomarker and isotopic trends in a Permian–Triassic sedimentary section at Kap Stosch, Greenland: Organic Geochemistry, v. 43, p. 67–82, doi:10.1016/j.orggeochem.2011.10.010.

Hazen, R.M., Papineau, D., Bleeker, W., Downs, R.T., Ferry, J.M., McCoy, T.J., Sverjensky, D.A., and Yang, H., 2008, Mineral evolution: The American Mineralogist, v. 93, p. 1693–1720, doi:10.2138/am.2008.2955.

Hazen, R.M., Hemley, R.L., and Mangum, A.J., 2012, Carbon in Earth’s inte-rior: Storage, cycling, and life: Eos (Transactions, American Geophysical Union), v. 93, p. 17–18, doi:10.1029/2012EO020001.

Hedges, J.I., and Keil, R.G., 1995, Sedimentary organic matter preservation: An assessment and speculative synthesis: Marine Chemistry, v. 49, p. 81–115, doi:10.1016/0304-4203(95)00008-F.

Heimann, A., Johnson, C.M., Beard, B.L., Valley, J.W., Roden, E.E., Spicuzza, M.J., and Beukes, N.J., 2010, Fe, C, and O isotope compositions of banded iron formation carbonates demonstrate a major role for dissimila-tory iron reduction in ~2.5 Ga marine environments: Earth and Planetary Science Letters, v. 294, p. 8–18, doi:10.1016/j.epsl.2010.02.015.

Helgeson, H.C., 1968, Evaluation of irreversible reactions in geochemi-cal processes involving minerals and aqueous solutions: I. Thermody-namic relations: Geochimica et Cosmochimica Acta, v. 32, p. 853–877, doi:10.1016/0016-7037(68)90100-2.

Herbert, T.D., Schuffert, J.D., Andreasen, D., Heusser, L., Lyle, M., Mix, A., Ravelo, A.C., Stott, L.D., and Herguera, J.C., 2001, Collapse of the Cali-fornia Current during glacial maxima linked to climate change on land: Science, v. 293, p. 71–76, doi:10.1126/science.1059209.

Hess, J., Bender, M.L., and Schilling, J.-G., 1986, Evolution of the ratio of strontium-87 to strontium-86 in seawater from Cretaceous to present: Sci-ence, v. 231, p. 979–984, doi:10.1126/science.231.4741.979.

Hildreth, W., and Moorbath, S., 1988, Crustal contributions to arc magmatism in the Andes of central Chile: Contributions to Mineralogy and Petrology, v. 98, p. 455–489, doi:10.1007/BF00372365.

Hills, I.R., and Whitehead, E.V., 1966, Triterpanes in optically active petroleum distillates: Nature, v. 209, p. 977–979, doi:10.1038/209977a0.

Hoefs, J., and Schidlowski, M., 1967, Carbon isotope composition of carbona-ceous matter from the Precambrian of the Witwatersrand system: Science, v. 155, p. 1096–1097, doi:10.1126/science.155.3766.1096.

Hofmann, A., Bekker, A., Rouxel, O., Rumble, D., and Master, S., 2009, Multiple sulphur and iron isotope compositions of detrital pyrite in Archaean sedimentary rocks: A new tool for provenance analysis: Earth and Planetary Science Letters, v. 286, no. 3–4, p. 436–445, doi:10.1016/j.epsl.2009.07.008.

Holba, A.G., Dzou, L.I.P., Masterson, W.D., Hughes, W.B., Huizinga, B.J., Singletary, M.S., Moldowan, J.M., Mello, M.R., and Tegelaar, E., 1998, Application of 24-norcholestanes for constraining source age of petro-leum: Organic Geochemistry, v. 29, p. 1269–1283, doi:10.1016/S0146-6380(98)00184-3.

Holland, H.D., 1962, Model for the evolution of the Earth’s atmosphere, in Engel, A.E.J., James, H.L., and Leonard, B.F., eds., Petrologic Studies: A Volume in Honor of A.F. Buddington: New York, Geological Society of America, p. 447–477.

Holser, W.T., 1977, Catastrophic chemical events in the history of the ocean: Nature, v. 267, p. 403–408, doi:10.1038/267403a0.

Holser, W.T., Schonlaub, H.-P., Attrep, M., Boeckelmann, K., Klein, P., Magaritz, M., Orth, C.J., Fenninger, A., Jenny, C., Kralik, M., Mauritsch, H., Pak, E., Schramm, J.-M., Stattegger, K., and Schmoller, R., 1989, A unique geochemical record at the Permian/Triassic boundary: Nature, v. 337, p. 39–44, doi:10.1038/337039a0.

House, C.H., Schopf, J.W., McKeegan, K.D., Coath, C.D., Harrison, T.M., and Stetter, K.O., 2000, Carbon isotopic composition of individual Pre-cambrian microfossils: Geology, v. 28, p. 707–710, doi:10.1130/0091-7613(2000)28<707:CICOIP>2.0.CO;2.

Hunt, J.M., 1996, Petroleum Geochemistry and Geology: New York, W.H. Freeman, 743 p.

Hunt, J.M., Philp, R.P., and Kvenvolden, K.A., 2002, Early developments in petroleum geochemistry: Organic Geochemistry, v. 33, p. 1025–1052, doi:10.1016/S0146-6380(02)00056-6.

Hurley, P.M., and Rand, J.R., 1969, Predrift continental nuclei: Science, v. 164, p. 1229–1242, doi:10.1126/science.164.3885.1229.

Hurley, P.M., Bateman, P.C., Fairbairn, H.W., and Pinson, W.H.J., 1965, Investigation of initial 87Sr/86Sr ratios in the Sierra Nevada plutonic province: Geological Society of America Bulletin, v. 76, p. 165–174, doi:10.1130/0016-7606(1965)76[165:IOISSR]2.0.CO;2.

Hurowitz, J.A., and McLennan, S.M., 2007, A 3.5 Ga record of water-limited, acidic weathering conditions on Mars: Earth and Planetary Science Let-ters, v. 260, p. 432–443, doi:10.1016/j.epsl.2007.05.043.

Ingalls, A.E., Shah, S.R., Hansman, R.L., Aluwihare, L.I., Santos, G.M., Druffel, E.R.M., and Pearson, A., 2006, Quantifying archaeal community autotrophy in the mesopelagic ocean using natural radiocarbon: Proceed-ings of the National Academy of Sciences of the United States of Amer-ica, v. 103, p. 6442–6447, doi:10.1073/pnas.0510157103.

Jackson, M.G., and Hart, S.R., 2006, Strontium isotopes in melt inclusions from Samoan basalts: Implications for heterogeneity in the Samoan plume: Earth and Planetary Science Letters, v. 245, p. 260–277, doi:10.1016/j.epsl.2006.02.040.

Jacobsen, S.B., 1988, Isotopic constraints on crustal growth and recycling: Earth and Planetary Science Letters, v. 90, p. 315–329, doi:10.1016/0012-821X(88)90133-1.

Jacobsen, S.B., and Wasserburg, G.J., 1980, Sm-Nd isotopic evolution of chondrites: Earth and Planetary Science Letters, v. 50, p. 139–155, doi:10.1016/0012-821X(80)90125-9.

Jacobsen, S.B., Ranen, M.C., Petaev, M.I., Remo, J.L., O’Connell, R.J., and Sasselov, D.D., 2008, Isotopes as clues to the origin and earliest differ-entiation history of the Earth: Philosophical Transactions of the Royal Society, ser. A, v. 366, p. 4129–4162, doi:10.1098/rsta.2008.0174.

Janik, C.J., Nehring, N.L., and Truesdell, A.H., 1983, Stable isotope geochem-istry of thermal fl uids from Lassen Volcanic National Park, California: Geothermal Resources Council Transactions, v. 7, p. 295–300.

Jansen, W., and Slaughter, M., 1982, Elemental mapping of minerals by elec-tron microprobe: The American Mineralogist, v. 67, p. 521–533.

Jerram, D.A., and Higgins, M.D., 2007, 3D analysis of rock textures: Quanti-fying igneous microstructures: Elements, v. 3, p. 239–245, doi:10.2113/gselements.3.4.239.

Jicha, B.R., Johnson, C.M., Hildreth, W., Beard, B.L., Hart, G.L., Shirey, S.B., and Singer, B.S., 2009, Discriminating assimilants and decoupling deep- vs. shallow-level crystal records at Mount Adams using 238U-230Th dis-equilibria and Os isotopes: Earth and Planetary Science Letters, v. 277, p. 38–49, doi:10.1016/j.epsl.2008.09.035.

Johnson, C.M., 1991, Large-scale crust formation and lithosphere modifi ca-tion beneath middle to late Cenozoic calderas and volcanic fi elds, western North America: Journal of Geophysical Research–Solid Earth and Plan-ets, v. 96, p. 13,485–13,507, doi:10.1029/91JB00304.

Johnson, C.M., 1993, Mesozoic and Cenozoic contributions to crustal growth in the southwestern United States: Earth and Planetary Science Letters, v. 118, p. 75–89, doi:10.1016/0012-821X(93)90160-B.

Johnson, C.M., Skulan, J.L., Beard, B.L., Sun, H., Nealson, K.H., and Brater-man, P.S., 2002, Isotopic fractionation between Fe(III) and Fe(II) in aque-ous solutions: Earth and Planetary Science Letters, v. 195, p. 141–153, doi:10.1016/S0012-821X(01)00581-7.

Johnson, C.M., Beard, B.L., and Albarède, F., 2004, Geochemistry of non-traditional stable isotopes, in Rosso, J.J., ed., Reviews in Mineralogy and Geochemistry, Volume 55: Washington, D.C., Mineralogical Society of America and Geochemical Society, 454 p..

Johnson, C.M., Beard, B.L., Klein, C., Beukes, N.J., and Roden, E.E., 2008a, Iron isotopes constrain biologic and abiologic processes in banded iron formation genesis: Geochimica et Cosmochimica Acta, v. 72, p. 151–169, doi:10.1016/j.gca.2007.10.013.

Johnson, C.M., Beard, B.L., and Roden, E.E., 2008b, The iron isotope fi nger-prints of redox and biogeochemical cycling in the modern and ancient Earth: Annual Review of Earth and Planetary Sciences, v. 36, p. 457–493, doi:10.1146/annurev.earth.36.031207.124139.

Johnson, T.M., and Bullen, T.D., 2004, Mass-dependent fractionation of sele-nium and chromium isotopes: Reviews in Mineralogy and Geochemistry, v. 55, p. 289–317, doi:10.2138/gsrmg.55.1.289.

Jolliff, B.L., Gillis, J.J., Haskin, L.A., Korotev, R.L., and Weiczorek, M.A., 2000, Major lunar crustal terranes: Surface expressions and crust- mantle

38 Johnson et al.


origins: Journal of Geophysical Research, v. 105, p. 4197–4216, doi:10.1029/1999JE001103.

Kah, L.C., Lyons, T.W., and Frank, T.D., 2004, Low marine sulphate and pro-tracted oxygenation of the Proterozoic biosphere: Nature, v. 431, p. 834–838, doi:10.1038/nature02974.

Kamo, S.L., Czamanske, G.K., Amelin, Y., Fedorenko, V.A., Davis, D.W., and Trofi mov, V.R., 2003, Rapid eruption of Siberian fl ood-volcanic rocks and evidence for coincidence with the Permian-Triassic boundary and mass extinction at 251 Ma: Earth and Planetary Science Letters, v. 214, p. 75–91, doi:10.1016/S0012-821X(03)00347-9.

Kanasewich, E.R., and Farquhar, R.M., 1965, Lead isotope ratios from the Cobalt-Noranda area, Canada: Canadian Journal of Earth Sciences, v. 2, p. 361–384, doi:10.1139/e65-029.

Kaplan, I.R., Emery, K.O., and Rittenbebg, S.C., 1963, The distribution and isotopic abundance of sulphur in recent marine sediments off south-ern California: Geochimica et Cosmochimica Acta, v. 27, p. 297–331, doi:10.1016/0016-7037(63)90074-7.

Karhu, J.A., and Holland, H.D., 1996, Carbon isotopes and the rise of atmospheric oxygen: Geology, v. 24, p. 867–870, doi:10.1130/0091-7613(1996)024<0867:CIATRO>2.3.CO;2.

Kasting, J.F., Howard, M.T., Wallman, K., Veizer, J., Shields, G.A., and Jaffres, J., 2006, Paleoclimates, ocean depth, and the oxygen isotopic composi-tion of seawater: Earth and Planetary Science Letters, v. 252, p. 82–93, doi:10.1016/j.epsl.2006.09.029.

Kaufman, A.J., and Xiao, S., 2003, High CO2 levels in the Proterozoic atmo-

sphere estimated from analyses of individual microfossils: Nature, v. 425, p. 279–282, doi:10.1038/nature01902.

Keeling, C.D., 1960, The concentration and isotopic composition of carbon dioxide in the atmosphere: Tellus, v. 12, p. 200–203, doi:10.1111/j.2153-3490.1960.tb01300.x.

Keith, M.L., and Weber, J.N., 1964, Carbon and oxygen isotopic composition of selected limestones and fossils: Geochimica et Cosmochimica Acta, v. 28, p. 1787–1816, doi:10.1016/0016-7037(64)90022-5.

Kemp, A.I.S., Hawkesworth, C.J., Paterson, B.A., and Kinny, P.D., 2006, Epi-sodic growth of the Gondwana supercontinent from hafnium and oxygen isotopes in zircon: Nature, v. 439, p. 580–583, doi:10.1038/nature04505.

Kemp, A.I.S., Wilde, S.A., Hawkesworth, C.J., Coath, C.D., Nemchin, A., Pid-geon, R.T., Vervoort, J.D., and DuFrane, S.A., 2010, Hadean crustal evo-lution revisited: New constraints from Pb-Hf isotope systematics of the Jack Hills zircons: Earth and Planetary Science Letters, v. 296, p. 45–56, doi:10.1016/j.epsl.2010.04.043.

Kendall, B., Reinhard, C.T., Lyons, T., Kaufman, A.J., Poulton, S.W., and Anbar, A.D., 2010, Pervasive oxygenation along Late Archaean ocean margins: Nature Geoscience, v. 3, no. 9, p. 647–652, doi:10.1038/ngeo942.

Keto, L.S., and Jacobsen, S.B., 1988, Nd isotopic variations of Phanerozoic paleoceans: Earth and Planetary Science Letters, v. 90, p. 395–410, doi:10.1016/0012-821X(88)90138-0.

King, P.L., and McLennan, S.M., 2010, Sulfur on Mars: Elements, v. 6, p. 107–112, doi:10.2113/gselements.6.2.107.

Kistler, R.W., and Peterman, Z.E., 1973, Variations in Sr, Rb K, Na, and ini-tial 87Sr/86Sr Mesozoic granitic rocks and intruded wall rocks in central California: Geological Society of America Bulletin, v. 84, p. 3489–3512, doi:10.1130/0016-7606(1973)84<3489:VISRKN>2.0.CO;2.

Kleine, T., Touboul, M., Bourdon, B., Nimmo, F., Mezger, K., Palme, H., Jacobsen, S.B., Yin, Q.-Z., and Halliday, A.N., 2009, Hf-W chronology of the accretion and early evolution of asteroids and terrestrial planets: Geochimica et Cosmochimica Acta, v. 73, p. 5150–5188, doi:10.1016/j.gca.2008.11.047.

Klemme, H.D., and Ulmishek, G.F., 1991, Effective petroleum source rocks of the world: Stratigraphic distribution and controlling depositional fac-tors: American Association of Petroleum Geologists Bulletin, v. 75, p. 1809–1851.

Klomp, U.C., 1986, The chemical structure of a pronounced series of iso-alkanes in South Oman crudes: Organic Geochemistry, v. 10, p. 807–814, doi:10.1016/S0146-6380(86)80017-1.

Knauth, L.P., 2005, Temperature and salinity history of the Precambrian ocean: Implications for the course of microbial evolution: Palaeogeography, Pal-aeoclimatology, Palaeoecology, v. 219, p. 53–69, doi:10.1016/j.palaeo.2004.10.014.

Knauth, L.P., and Epstein, S., 1976, Hydrogen and oxygen isotope ratios in nodular and bedded cherts: Geochimica et Cosmochimica Acta, v. 40, p. 1095–1108, doi:10.1016/0016-7037(76)90051-X.

Knesel, K.M., Davidson, J.P., and Duffi eld, W.A., 1999, Evolution of silicic magma through assimilation and subsequent recharge: Evidence from Sr isotopes in sanidine phenocrysts, Taylor Creek Rhyolite, NM: Journal of Petrology, v. 40, p. 773–786, doi:10.1093/petroj/40.5.773.

Knoll, A.H., Bambach, R.K., Canfi eld, D.E., and Grotzinger, J.P., 1996, Com-parative Earth history and Late Permian mass extinction: Science, v. 273, p. 452–457, doi:10.1126/science.273.5274.452.

Knoll, A.H., Walter, M.R., Narbonne, G.M., and Christie-Blick, N., 2004, A new period for the geologic time scale: Science, v. 305, p. 621–622, doi:10.1126/science.1098803.

Knoll, A.H., Summons, R.E., Waldbauer, J.R., and Zumberge, J., 2007, The geological succession of primary producers in the oceans, in Falkwoski, P., and Knoll, A.H., eds., The Evolution of Primary Producers in the Sea: Boston, Massachusetts, Elsevier, p. 133–163.

Kohnen, M.E.L., Sinninghe Damsté, J.S., Ten Haven, H.L., and De Leeuw, J.W., 1989, Early incorporation of polysulphides in sedimentary organic matter: Nature, v. 341, p. 640–641, doi:10.1038/341640a0.

Koopmans, M.P., Kaster, J., Van Kaam-Peters, H.M.E., Kenig, F., Schouten, S., Hartgers, W.A., de Leeuw, J.W., and Sinninghe Damsté, J.S., 1996, Diagenetic and catagenetic products of isorenieratene: Molecular indica-tors for photic zone anoxia: Geochimica et Cosmochimica Acta, v. 60, p. 4467–4496, doi:10.1016/S0016-7037(96)00238-4.

Kump, L.R., Pavlov, A., and Arthur, M.A., 2005, Massive release of hydrogen sulfi de to the surface ocean and atmosphere during intervals of oceanic anoxia: Geology, v. 33, p. 397–400, doi:10.1130/G21295.1.

Kurz, M.D., Jenkins, W.J., Schilling, J.G., and Hart, S.R., 1982, Helium iso-topic variations in the mantle beneath the central North Atlantic Ocean: Earth and Planetary Science Letters, v. 58, p. 1–14, doi:10.1016/0012-821X(82)90099-1.

Kusch, S., Eglinton, T.I., Mix, A.C., and Mollenhauer, G., 2010, Timescales of lateral sediment transport in the Panama Basin as revealed by radiocarbon ages of alkenones, total organic carbon and foraminifera: Earth and Plan-etary Science Letters, v. 290, p. 340–350, doi:10.1016/j.epsl.2009.12.030.

Kuypers, M.M.M., van Breugel, Y., Schouten, S., Erba, E., and Sinninghe Damsté, J.S., 2004, N2-fi xing cyanobacteria supplied nutrient N for Cretaceous oceanic anoxic events: Geology, v. 32, p. 853–856, doi:10.1130/G20458.1.

Kvenvolden, K., 2006, Organic geochemistry—A retrospective of its fi rst 70 years: Organic Geochemistry, v. 37, p. 1–11, doi:10.1016/j.orggeo-chem.2005.09.001.

Kvenvolden, K., Lawless, J., Pering, K., Peterson, E., Flores, J., Ponnampe-ruma, C., Kaplan, I.R., and Moore, C., 1970, Evidence for extraterres-trial amino-acids and hydrocarbons in the Murchison meteorite: Nature, v. 228, p. 923–926, doi:10.1038/228923a0.

Lasaga, A.C., and Berner, R.A., 1998, Fundamental aspects of quantitative models for geochemical cycles: Chemical Geology, v. 145, p. 161–175, doi:10.1016/S0009-2541(97)00142-3.

Lassiter, J.C., and Hauri, E.H., 1998, Osmium-isotope variations in Hawaiian lavas: Evidence for recycled oceanic lithosphere in the Hawaiian plume: Earth and Planetary Science Letters, v. 164, p. 483–496, doi:10.1016/S0012-821X(98)00240-4.

Lawless, J.G., Zeitman, B., Pereira, W.E., Summons, R.E., and Duffi eld, A.M., 1974, Dicarboxylic acids in the Murchison meteorite: Nature, v. 251, p. 40–42, doi:10.1038/251040a0.

Lawrence, D.J., Feldman, W.C., Barraclough, N.L., Binder, A.B., Elphic, R.C., Maurice, S., and Thomsen, D.R., 1998, Global elemental maps of the Moon: The Lunar Prospector gamma-ray spectrometer: Science, v. 281, p. 1484–1489, doi:10.1126/science.281.5382.1484.

Laws, E.A., Popp, B.N., Bidigare, R.R., Kennicutt, M.C., and Macko, S.A., 1995, Dependence of phytoplankton carbon isotopic composition on growth rate and [CO2]aq: Theoretical considerations and experimen-tal results: Geochimica et Cosmochimica Acta, v. 59, p. 1131–1138, doi:10.1016/0016-7037(95)00030-4.

Laws, E.A., Bidigare, R.R., and Popp, B.N., 1997, Effect of growth rate and CO2 concentration on carbon isotopic fractionation by the marine dia-tom Phaeodactylum tricornutum: Limnology and Oceanography, v. 42, p. 1552–1560, doi:10.4319/lo.1997.42.7.1552.

Lee, T., Papanastassiou, D.A., and Wasserburg, G.J., 1977, Aluminum-26 in the early solar system: Fossil or fuel?: Astrophysical Journal Letters, v. 211, p. L107–L110, doi:10.1086/182351.

Lepland, A., van Zuilen, M.A., Arrhenius, G., Whitehouse, M.J., and Fedo, C.M., 2005, Questioning the evidence for Earth’s earliest life‚ Akilia revisited: Geology, v. 33, p. 77–79, doi:10.1130/G20890.1.



Lerman, A., and Wu, L., 2007, Kinetics of global geochemical cycles, in Brant-ley, S.L., Kubicki, J.D., and White, A.F., eds., Kinetics of Water-Rock Interaction: New York, Springer, p. 655–736.

Lewis, R.S., Tang, M., Wacker, J.F., Anders, E., and Steel, E., 1987, Interstellar diamonds in meteorites: Nature, v. 326, p. 160–162, doi:10.1038/326160a0.

Lichtfouse, E., 2000, Compound-specifi c isotope analysis: Applica-tion to archaeology, biomedical sciences, biosynthesis, environ-ment, extraterrestrial chemistry, food science, forensic science, humic substances, microbiology, organic geochemistry, soil sci-ence and sport: Rapid Communications in Mass Spectrometry, v. 14, p. 1337–1344, doi:10.1002/1097-0231(20000815)14:15<1337::AID-RCM9>3.0.CO;2-B.

Lipp, J.S., Morono, Y., Inagaki, F., and Hinrichs, K.-U., 2008, Signifi cant con-tribution of Archaea to extant biomass in marine subsurface sediments: Nature, v. 454, p. 991–994, doi:10.1038/nature07174.

Long, J.V.P., 1995, Microanalysis from 1950 to the 1990s, in Potts, P.J., Bowles, J.F.W., Reed, S.J.B., and Cave, M.R., eds., Microprobe Techniques in the Earth Sciences: London, Chapman & Hall, p. 1–48.

Love, G.D., Grosjean, E., Stalvies, C., Fike, D.A., Grotzinger, J.P., Bradley, A.S., Kelly, A.E., Bhatia, M., Meredith, W., Snape, C.E., Bowring, S.A., Condon, D.J., and Summons, R.E., 2009, Fossil steroids record the appearance of Demospongiae during the Cryogenian period: Nature, v. 457, p. 718–721, doi:10.1038/nature07673.

Luck, J.M., and Allègre, C.J., 1983, 187Re-187Os systematics in meteorites and cosmological consequences: Nature, v. 302, p. 130–132, doi:10.1038/302130a0.

Luguet, A., Pearson, D.G., Nowell, G.M., Dreher, S.T., Coggon, J.A., Spetsius, Z.V., and Parman, S.W., 2008, Enriched Pt-Re-Os isotope systematics in plume lavas explained by metasomatic sulfi des: Science, v. 319, p. 453–456, doi:10.1126/science.1149868.

Lyons, J.R., 2009, Atmospherically-derived mass-independent sulfur isotope signatures and incorporation into sediments: Chemical Geology, v. 267, p. 164–174, doi:10.1016/j.chemgeo.2009.03.027.

Macko, S.A., Uhle, M.E., Engel, M.H., and Andrusevich, V., 1997, Stable nitro-gen isotope analysis of amino acid enantiomers by gas chromatography/combustion/isotope ratio mass spectrometry: Analytical Chemistry, v. 69, p. 926–929, doi:10.1021/ac960956l.

Mao, H.K., Wu, Y., Chen, L.C., Shu, J.F., and Jephcoat, A.P., 1990, Static com-pression of iron to 300 GPa and Fe0.8Ni0.2 alloy to 260 GPa: Implica-tions for composition of the core: Journal of Geophysical Research, v. 95, p. 21,737–21,742, doi:10.1029/JB095iB13p21737.

Martins, Z., Botta, O., Fogel, M.L., Sephton, M.A., Glavin, D.P., Watson, J.S., Dworkin, J.P., Schwartz, A.W., and Ehrenfreund, P., 2008, Extraterrestrial nucleobases in the Murchison meteorite: Earth and Planetary Science Let-ters, v. 270, p. 130–136, doi:10.1016/j.epsl.2008.03.026.

Mason, B., 1952, Principles of Geochemistry: New York, John Wiley, 276 p.Mason, B., 1992, Victor Moritz Goldschmidt: Father of Modern Geochemistry:

The Geochemical Society Special Publication 4, 184 p.Mason, B., and Taylor, S.R., 1982, Inclusions in the Allende Meteorite: Wash-

ington, D. C., Smithsonian Institution, Smithsonian Contributions to Earth Science, 30 p.

Masterson, A.L., Farquhar, J., and Wing, B.A., 2011, Sulfur mass-independent fractionation patterns in the broadband UV photolysis of sulfur dioxide: Pressure and third body effects: Earth and Planetary Science Letters, v. 306, p. 253–260, doi:10.1016/j.epsl.2011.04.004.

McCaffrey, M.A., Michael Moldowan, J., Lipton, P.A., Summons, R.E., Peters, K.E., Jeganathan, A., and Watt, D.S., 1994, Paleoenvironmental implica-tions of novel C30 steranes in Precambrian to Cenozoic age petroleum and bitumen: Geochimica et Cosmochimica Acta, v. 58, p. 529–532, doi:10.1016/0016-7037(94)90481-2.

McCulloch, M.T., Gregory, R.T., Wasserburg, G.J., and Taylor, H.P.J., 1981, Sm-Nd, Rb-Sr, and 18O/16O isotopic systematics in an oceanic crustal section: Evidence from the Samail ophiolite: Journal of Geophysical Research, v. 86, p. 2721–2735, doi:10.1029/JB086iB04p02721.

McKeegan, K.D., Kallio, A.P.A., Heber, V.S., Jarzebinski, G., Mao, P.H., Coath, C.D., Kunihiro, T., Wiens, R.C., Nordholt, J.E., Moses, R.W., Rei-senfeld, D.B., Jurewicz, A.J.G., and Burnett, D.S., 2011, The oxygen iso-topic composition of the Sun inferred from captured solar wind: Science, v. 332, p. 1528–1532, doi:10.1126/science.1204636.

McLennan, S.M., 2012, Geochemistry of sedimentary processes on Mars, in Grotzinger, J.P., and Milliken, R.E., eds., Sedimentary Geology on Mars: Society for Sedimentary Geology Special Publication 102, p. 119–138.

McLoughlin, N., Grosch, E.G., Kilburn, M.R., and Wacey, D., 2012, Sulfur iso-tope evidence for a Paleoarchean subseafl oor biosphere, Barberton, South Africa: Geology, v. 40, p. 1031–1034, doi:10.1130/G33313.1.

McSween, H.Y.J., and Huss, G.R., 2010, Cosmochemistry: Cambridge, UK, Cambridge University Press, 568 p.

McSween, H.Y.J., and Keil, K., 2000, Mixing relationships in the Martian regolith and the composition of globally homogeneous dust: Geochi-mica et Cosmochimica Acta, v. 64, p. 2155–2166, doi:10.1016/S0016-7037(99)00401-9.

McSween, H.Y.J., Taylor, G.J., and Wyatt, M.B., 2009, Elemental composi-tion of the Martian crust: Science, v. 324, p. 736–739, doi:10.1126/science.1165871.

Menzies, M.A., 1989, Cratonic, circumcratonic and oceanic mantle domains beneath the western United States: Journal of Geophysical Research, v. 94, p. 7899–7915, doi:10.1029/JB094iB06p07899.

Michard, A., Gurriet, P., Soudant, M., and Albarède, F., 1985, Nd isotopes in French Phanerozoic shales: External vs. internal aspects of crustal evolution: Geochimica et Cosmochimica Acta, v. 49, p. 601–610, doi:10.1016/0016-7037(85)90051-1.

Mojzsis, S.J., Arrhenius, G., McKeegan, K.D., Harrison, T.M., Nutman, A.P., and Friend, C.R., 1996, Evidence for life on Earth before 3,800 million years ago: Nature, v. 384, p. 55–59, doi:10.1038/384055a0.

Mojzsis, S.J., Harrison, T.M., and Pidgeon, R.T., 2001, Oxygen-isotope evidence from ancient zircons for liquid water at the Earth’s surface 4,300 Myr ago: Nature, v. 409, p. 178–181, doi:10.1038/35051557.

Moldowan, J.M., Dahl, J., Huizinga, B.J., Fago, F.J., Hickey, L.J., Peakman, T.M., and Taylor, D.W., 1994, The molecular fossil record of oleanane and its relation to angiosperms: Science, v. 265, p. 768–771, doi:10.1126/science.265.5173.768.

Mollenhauer, G., Eglinton, T.I., Ohkouchi, N., Schneider, R.R., Miller, P.J., Grootes, P.M., and Rullkatter, J., 2003, Asynchronous alkenone and foraminifera records from the Benguela upwelling system: Geochi-mica et Cosmochimica Acta, v. 67, p. 2157–2171, doi:10.1016/S0016-7037(03)00168-6.

Mollenhauer, G., McManus, J.F., Wagner, T., McCave, I.N., and Eglinton, T.I., 2011, Radiocarbon and 230Th data reveal rapid redistribution and temporal changes in sediment focussing at a North Atlantic drift: Earth and Planetary Science Letters, v. 301, p. 373–381, doi:10.1016/j.epsl.2010.11.022.

Muehlenbachs, K., and Clayton, R.N., 1976, Oxygen isotope composition of the oceanic crust and its bearing on seawater: Journal of Geophysical Research, v. 81, p. 4365–4369, doi:10.1029/JB081i023p04365.

Mundil, R., Pálfy, J., Renne, P.R., and Brack, P., 2010, The Triassic timescale: New constraints and a review of geochronological data, in Lucas, S.G., ed., The Triassic Timescale: Geological Society of London Special Publi-cation 334, p. 41–60, doi:10.1144/SP334.3.

Murray, A.P., Edwards, D., Hope, J.M., Boreham, C.J., Booth, W.E., Alexan-der, R.A., and Summons, R.E., 1998, Carbon isotope biogeochemistry of plant resins and derived hydrocarbons: Organic Geochemistry, v. 29, p. 1199–1214, doi:10.1016/S0146-6380(98)00126-0.

Nesbitt, H.W., 2003, Petrogenesis of siliciclastic sediments and sedimentary rocks, in Lentz, D.R., ed., Geochemistry of Sedimentary Rocks: Evolu-tionary Considerations to Mineral Deposit-Forming Environments, Vol-ume 4: Saint Johns, Canada, Geological Association of Canada GeoText, p. 39–51.

Nesbitt, H.W., and Wilson, R.E., 1992, Recent chemical weathering of basalts: American Journal of Science, v. 292, p. 740–777, doi:10.2475/ajs.292.10.740.

Nesbitt, H.W., and Young, G.M., 1984, Prediction of some weathering trends of plutonic and volcanic rocks based on thermodynamic and kinetic con-siderations: Geochimica et Cosmochimica Acta, v. 48, p. 1523–1534, doi:10.1016/0016-7037(84)90408-3.

Nier, A.O., 1940, A mass spectrometer for routine isotope abundance mea-surements: The Review of Scientifi c Instruments, v. 11, p. 212–216, doi:10.1063/1.1751688.

Nittler, L.R., 2003, Presolar stardust in meteorites: Recent advances and sci-entifi c frontiers: Earth and Planetary Science Letters, v. 209, p. 259–273, doi:10.1016/S0012-821X(02)01153-6.

Noble, R.A., Alexander, R., Kagi, R.I., and Knox, J., 1985, Tetracyclic diterpenoid hydrocarbons in some Australian coals, sediments and crude oils: Geochimica et Cosmochimica Acta, v. 49, p. 2141–2147, doi:10.1016/0016-7037(85)90072-9.

40 Johnson et al.


Norton, D., and Knight, J.E., 1977, Transport phenomena in hydrothermal sys-tems: Cooling plutons: American Journal of Science, v. 277, p. 937–981, doi:10.2475/ajs.277.8.937.

Norton, D., and Taylor, H.P.J., 1979, Quantitative simulation of the hydrother-mal systems of crystallizing magma on the basis of transport theory and oxygen isotope data: An analysis of the Skaergaard intrusion: Journal of Petrology, v. 20, p. 421–486, doi:10.1093/petrology/20.3.421.

Nyquist, L.E., Kleine, T., Shih, C.-Y., and Reese, Y.D., 2009, The distribution of short-lived radioisotopes in the early solar system and the chronol-ogy of asteroid accretion, differentiation, and secondary mineralization: Geochimica et Cosmochimica Acta, v. 73, p. 5115–5136, doi:10.1016/j.gca.2008.12.031.

Oehler, D.Z., Schopf, J.W., and Kvenvolden, K.A., 1972, Carbon isotopic stud-ies of organic matter in Precambrian rocks: Science, v. 175, p. 1246–1248, doi:10.1126/science.175.4027.1246.

Ohkouchi, N., Kawamura, K., Kawahata, H., and Okada, H., 1999, Depth ranges of alkenone production in the central Pacifi c Ocean: Global Bio-geochemical Cycles, v. 13, p. 695–704, doi:10.1029/1998GB900024.

O’Neil, J., Shaw, S.E., and Flood, R.H., 1977, Oxygen and hydrogen isotope compositions as indicators of granite genesis in the New England batho-lith, Australia: Contributions to Mineralogy and Petrology, v. 62, p. 313–328, doi:10.1007/BF00371018.

O’Neil, J., Carlson, R.W., Francis, D., and Stevens, R.K., 2008, Neodym-ium-142 evidence for Hadean mafi c crust: Science, v. 321, p. 1828–1831, doi:10.1126/science.1161925.

O’Nions, R.K., Hamilton, P.J., and Evensen, N.M., 1977, Variations in 143Nd/144Nd and 87Sr/86Sr ratios in oceanic basalts: Earth and Planetary Science Letters, v. 34, p. 13–22, doi:10.1016/0012-821X(77)90100-5.

O’Nions, R.K., Hamilton, P.J., and Hooker, P.J., 1983, A Nd isotope investi-gation of sediments related to crustal development in the British Isles: Earth and Planetary Science Letters, v. 63, p. 229–240, doi:10.1016/0012-821X(83)90039-0.

Orphan, V.J., House, C.H., Hinrichs, K.-U., McKeegan, K.D., and DeLong, E.F., 2002, Multiple archaeal groups mediate methane oxidation in anoxic cold seep sediments: Proceedings of the National Academy of Sci-ences of the United States of America, v. 99, p. 7663–7668, doi:10.1073/pnas.072210299.

Ourisson, G., and Albrecht, P., 1992, Geohopanoids—The most abundant natural-products on Earth: Accounts of Chemical Research, v. 25, p. 398–402, doi:10.1021/ar00021a003.

Ourisson, G., Albrecht, P., and Rohmer, M., 1979, The Hopanoids: Palaeo-chemistry and biochemistry of a group of natural products: Pure and Applied Chemistry, v. 51, p. 709–729.

Pancost, R.D., Crawford, N., Magness, S., Turner, A., Jenkyns, H.C., and Maxwell, J.R., 2004, Further evidence for the development of photic-zone euxinic conditions during Mesozoic oceanic anoxic events: Journal of the Geological Society of London, v. 161, p. 353–364, doi:10.1144/0016764903-059.

Park, R., and Epstein, S., 1960, Carbon isotope fractionation during pho-tosynthesis: Geochimica et Cosmochimica Acta, v. 21, p. 110–126, doi:10.1016/S0016-7037(60)80006-3.

Patchett, P.J., and Tatsumoto, M., 1980, Hafnium isotope variations in oceanic basalts: Geophysical Research Letters, v. 7, p. 1077–1080, doi:10.1029/GL007i012p01077.

Patchett, P.J., Kouvo, O., Hedge, C.E., and Tatsumoto, M., 1981, Evolution of continental crust and mantle heterogeneity: Evidence from Hf iso-topes: Contributions to Mineralogy and Petrology, v. 78, p. 279–297, doi:10.1007/BF00398923.

Patchett, P.J., White, W.M., Feldmann, H., Kielinczuk, S., and Hofmann, A.W., 1984, Hafnium/rare earth element fractionation in the sedimentary system and crustal recycling into the Earth’s mantle: Earth and Planetary Science Letters, v. 69, p. 365–378, doi:10.1016/0012-821X(84)90195-X.

Patterson, C., 1956, Age of meteorites and the Earth: Geochimica et Cosmochi-mica Acta, v. 10, p. 230–237, doi:10.1016/0016-7037(56)90036-9.

Patterson, C., and Tatsumoto, M., 1964, The signifi cance of lead isotopes in detrital feldspar with respect to chemical differentiation within the Earth’s mantle: Geochimica et Cosmochimica Acta, v. 28, p. 1–22, doi:10.1016/0016-7037(64)90052-3.

Paul, B., Wood, J.A., Hergt, J., Danyushevsky, L., Kunihiro, T., and Nakamura, E., 2011, Melt inclusion Pb-isotopic analysis by LA-MC-ICPMS: Assess-ment of analytical performance and application to OIB genesis: Chemical Geology, v. 289, p. 210–223, doi:10.1016/j.chemgeo.2011.08.005.

Pavlov, A.A., and Kasting, J.F., 2002, Mass-independent fractionation of sulfur isotopes in Archean sediments: Strong evidence for an anoxic Archean atmo-sphere: Astrobiology, v. 2, p. 27–41, doi:10.1089/153110702753621321.

Payne, J.L., and Clapham, M.E., 2012, End-Permian mass extinction in the oceans: An ancient analog for the twenty-fi rst century?: Annual Review of Earth and Planetary Sciences, v. 40, p. 89–111, doi:10.1146/annurev-earth-042711-105329.

Pearson, A., McNichol, A.P., Benitez-Nelson, B.C., Hayes, J.M., and Eglin-ton, T.I., 2001, Origins of lipid biomarkers in Santa Monica Basin surface sediment: A case study using compound-specifi c 14C analysis: Geochi-mica et Cosmochimica Acta, v. 65, p. 3123–3137, doi:10.1016/S0016-7037(01)00657-3.

Pearson, A., Leavitt, W.D., Saenz, J.P., Summons, R.E., Tam, M.C.M., and Close, H.G., 2009, Diversity of hopanoids and squalene-hopene cyclases across a tropical land-sea gradient: Environmental Microbiology, v. 11, no. 5, p. 1208–1223, doi:10.1111/j.1462-2920.2008.01817.x.

Pearson, D.G., and Wittig, N., 2008, Formation of Archean continental lithosphere and its diamonds: The root of the problem: Journal of the Geological Society of London, v. 165, p. 895–914, doi:10.1144/0016-76492008-003.

Peck, W.H., Valley, J.W., Wilde, S.A., and Graham, C.M., 2001, Oxygen isotope ratios and rare earth elements in 3.3 to 4.4 Ga zircons: Ion microprobe evi-dence for high d18O continental crust and oceans in the Early Archean: Geochimica et Cosmochimica Acta, v. 65, p. 4215–4229, doi:10.1016/S0016-7037(01)00711-6.

Peltzer, E.T., Bada, J.L., Schlesinger, G., and Miller, S.L., 1984, The chemical conditions on the parent body of the Murchison meteorite: Some conclu-sions based on amino, hydroxy and dicarboxylic acids: Advances in Space Research, v. 4, p. 69–74, doi:10.1016/0273-1177(84)90546-5.

Perry, E.C., Jr., 1967, The oxygen isotope chemistry of ancient cherts: Earth and Planetary Science Letters, v. 3, p. 62–66, doi:10.1016/0012-821X(67)90012-X.

Perry, E.C., Jr., Monster, J., and Reimer, T., 1971, Sulfur isotopes in Swazi-land system barites and the evolution of the Earth’s atmosphere: Science, v. 171, p. 1015–1016, doi:10.1126/science.171.3975.1015.

Petaev, M.I., and Wood, J.A., 1998, The condensation with partial isola-tion (CWPI) model of condensation in the solar nebula: Meteoritics & Planetary Science, v. 33, p. 1123–1137, doi:10.1111/j.1945-5100.1998.tb01717.x.

Peterman, Z.E., Hedge, C.E., and Tourtelot, H.A., 1970, Isotopic composition of strontium in sea water throughout Phanerozoic time: Geochimica et Cos-mochimica Acta, v. 34, p. 105–120, doi:10.1016/0016-7037(70)90154-7.

Peters, K.E., Walters, C.C., and Moldowan, J.M., 2005, The Biomarker Guide: Biomarkers and Isotopes in Petroleum Exploration and Earth History: Cambridge, UK, Cambridge University Press, 700 p.

Peterson, K.J., Summons, R.E., and Donoghue, P.C.J., 2007, Molecular pal-aeobiology: Palaeontology, v. 50, p. 775–809, doi:10.1111/j.1475-4983.2007.00692.x.

Pettijohn, F.J., 1975, Sedimentary Rocks: New York, Harper and Row, 628 p.Philippot, P., van Zuilen, M., Lepot, K., Thomazo, C., Farquhar, J., and Van

Kranendonk, M.J., 2007, Early Archaean microorganisms preferred ele-mental sulfur, not sulfate: Science, v. 317, p. 1534–1537, doi:10.1126/science.1145861.

Piepgras, D.J., and Wasserburg, G.J., 1980, Neodymium isotopic variations in seawater: Earth and Planetary Science Letters, v. 50, p. 128–138, doi:10.1016/0012-821X(80)90124-7.

Piepgras, D.J., Wasserburg, G.J., and Dasch, E.J., 1979, The isotopic composi-tion of Nd in different ocean masses: Earth and Planetary Science Letters, v. 45, p. 223–236, doi:10.1016/0012-821X(79)90125-0.

Planavsky, N., Rouxel, O., Bekker, A., Shapiro, R., Fralick, P., and Knudsen, A., 2009, Iron-oxidizing microbial ecosystems thrived in late Paleoprotero-zoic redox-stratifi ed oceans: Earth and Planetary Science Letters, v. 286, p. 230–242, doi:10.1016/j.epsl.2009.06.033.

Poitrasson, F., Mao, X.-L., Mao, S.S., Freydier, R., and Russo, R.E., 2003, Comparison of ultraviolet femtosecond and nanosecond laser ablation inductively coupled plasma mass spectrometry analysis in glass, mona-zite, and zircon: Analytical Chemistry, v. 75, p. 6184–6190, doi:10.1021/ac034680a.

Polyakov, V.B., and Mineev, S.D., 2000, The use of Mossbauer spectroscopy in stable isotope geochemistry: Geochimica et Cosmochimica Acta, v. 64, p. 849–865, doi:10.1016/S0016-7037(99)00329-4.

Popp, B.N., Laws, E.A., Bidigare, R.R., Dore, J.E., Hanson, K.L., and Wake-ham, S.G., 1998, Effect of phytoplankton cell geometry on carbon isotopic



fractionation: Geochimica et Cosmochimica Acta, v. 62, p. 69–77, doi:10.1016/S0016-7037(97)00333-5.

Prahl, F., Wolf, G. V., and Sparrow, M. A., 2003, Physiological impacts on alkenone paleothermometry: Paleoceanography, v. 18, 1025, doi:10.1029/2002PA000803, 2.

Prettyman, T.H., Hagerty, J.J., Elphic, R.C., Feldman, W.C., Lawrence, D.J., McKinney, G.W., and Vaniman, D.T., 2006, Elemental composition of the lunar surface: Analysis of gamma ray spectroscopy data from Lunar Prospector: Journal of Geophysical Research, v. 111, no. E12, E12007, doi:10.1029/2005JE002656.

Puchtel, I.S., Brandon, A.D., Humayun, M., and Walker, R.J., 2005, Evidence for the early differentiation of the core from Pt-Re-Os isotope system-atics of 2.8-Ga komatiites: Earth and Planetary Science Letters, v. 237, p. 118–134, doi:10.1016/j.epsl.2005.04.023.

Raiswell, R., and Canfi eld, D.E., 2012, The iron biogeochemical cycle past and present: Geochemical Perspectives, v. 1, p. 1–220, doi:10.7185/geochempersp.1.1.

Rampen, S.W., Schouten, S., Abbas, B., Elda Panoto, F., Muyzer, G., Camp-bell, C.N., Fehling, J., and Sinninghe Damsté, J.S., 2007, On the origin of 24-norcholestanes and their use as age-diagnostic biomarkers: Geology, v. 35, p. 419–422, doi:10.1130/G23358A.1.

Rasmussen, B., Fletcher, I.R., Brocks, J.J., and Kilburn, M.R., 2008, Reassess-ing the fi rst appearance of eukaryotes and cyanobacteria: Nature, v. 455, p. 1101–1104, doi:10.1038/nature07381.

Raymo, M.E., Ruddiman, W.F., and Froelich, P.N., 1988, Infl uence of late Cenozoic mountain building on ocean geochemical cycles: Geology, v. 16, p. 649–653, doi:10.1130/0091-7613(1988)016<0649:IOLCMB>2.3.CO;2.

Reed, S.J.B., 1995, Electron probe microanalysis, in Potts, P.J., Bowles, J.F.W., Reed, S.J.B., and Cave, M.R., eds., Microprobe Techniques in the Earth Sciences: London, Chapman & Hall, p. 49–89.

Richard, P., Shimizu, N., and Allègre, C.J., 1976, 143Nd/146Nd, a natural tracer: An application to oceanic basalts: Earth and Planetary Science Let-ters, v. 31, p. 269–278, doi:10.1016/0012-821X(76)90219-3.

Righter, K., 2003, Metal-silicate partitioning of siderophile elements and core formation in the early Earth: Annual Review of Earth and Planetary Sci-ences, v. 31, p. 135–174, doi:10.1146/annurev.earth.31.100901.145451.

Righter, K., and Drake, M.J., 1996, Core formation in Earth’s Moon, Mars and Vesta: Icarus, v. 124, p. 513–529, doi:10.1006/icar.1996.0227.

Rino, S., Komiya, T., Windley, B.F., Katayama, I., Motoki, A., and Hirata, T., 2004, Major episodic increases of continental crustal growth determined from zircon ages of river sands; implications for mantle overturns in the early Precambrian: Physics of the Earth and Planetary Interiors, v. 146, p. 369–394, doi:10.1016/j.pepi.2003.09.024.

Rohmer, M., 2003, Mevalonate-independent methylerythritol phosphate path-way for isoprenoid biosynthesis. Elucidation and distribution: Pure and Applied Chemistry, v. 75, p. 375–387, doi:10.1351/pac200375020375.

Rohmer, M., 2010, Hopanoids, in Timmis, K.N., ed., Handbook of Hydrocar-bon and Lipid Microbiology: Berlin, Springer, p. 133–142.

Rohmer, M., Dastillung, M., and Ourisson, G., 1980, Hopanoids from C30 to C35 in recent muds: Naturwissenschaften, v. 67, p. 456–458, doi:10.1007/BF00405643.

Rohmer, M., Bouvier-Nave, P., and Ourisson, G., 1984, Distribution of hopanoid triterpanes in prokaryotes: Journal of General Microbiology, v. 130, p. 1137–1150.

Rohmer, M., Knani, M., Simonin, P., Sutter, B., and Sahm, H., 1993, Isoprenoid biosynthesis in bacteria: A novel pathway for the early steps leading to isopentenyl diphosphate: Biochemical Journal, v. 295, part 2, p. 517–524.

Rosell-Melé, A., Eglinton, G., Pfl aumann, U., and Sarnthein, M., 1995, Atlantic core-top calibration of the U37K index as a sea-surface palaeotempera-ture indicator: Geochimica et Cosmochimica Acta, v. 59, p. 3099–3107, doi:10.1016/0016-7037(95)00199-A.

Rouxel, O.J., Bekker, A., and Edwards, K.J., 2005, Iron isotope constraints on the Archean and Paleoproterozoic ocean redox state: Science, v. 307, p. 1088–1091, doi:10.1126/science.1105692.

Rudnick, R.L., and Walker, R.J., 2009, Interpreting age from Re-Os iso-topes in peridotites: Lithos, v. 112S, p. 1083–1095, doi:10.1016/j.lithos.2009.04.042.

Russell, W.A., Papanastassiou, D.A., and Tombrello, T.A., 1978, Ca isotope fractionation on the Earth and other solar system materials: Geochi-mica et Cosmochimica Acta, v. 42, p. 1075–1090, doi:10.1016/0016-7037(78)90105-9.

Rütters, H., Sass, H., Cypionka, H., and Rullkotter, J., 2002, Phospholipid analysis as a tool to study complex microbial communities in marine sediments: Journal of Microbiological Methods, v. 48, p. 149–160, doi:10.1016/S0167-7012(01)00319-0.

Saal, A.E., Hart, S.R., Shimizu, N., Hauri, E.H., and Layne, G.D., 1998, Pb isotopic variability in melt inclusions from oceanic island basalts, Polyne-sia: Science, v. 282, p. 1481–1484, doi:10.1126/science.282.5393.1481.

Salters, V.J.M., 1996, The generation of mid-ocean ridge basalts from the Hf and Nd isotope perspective: Earth and Planetary Science Letters, v. 141, p. 109–123, doi:10.1016/0012-821X(96)00070-2.

Salters, V.J.M., and Hart, S.R., 1991, The mantle sources of ocean ridges, islands and arcs: The Hf-isotope connection: Earth and Planetary Science Letters, v. 104, p. 364–380, doi:10.1016/0012-821X(91)90216-5.

Savin, S.M., and Epstein, S., 1970, The oxygen and hydrogen isotope geo-chemistry of clay minerals: Geochimica et Cosmochimica Acta, v. 34, p. 25–42, doi:10.1016/0016-7037(70)90149-3.

Schidlowski, M., 1987, Application of stable carbon isotopes to early biochemi-cal evolution on Earth: Annual Review of Earth and Planetary Sciences, v. 15, p. 47–72, doi:10.1146/annurev.ea.15.050187.000403.

Schidlowski, M., Eichman, R., and Junge, C.E., 1975, Precambrian sedimen-tary carbonates: Carbon and oxygen isotope geochemistry and implication for the terrestrial oxygen budget: Precambrian Research, v. 2, p. 1–69, doi:10.1016/0301-9268(75)90018-2.

Schidlowski, M., Appel, P.W.U., Eichmann, R., and Junge, C.E., 1979, Car-bon isotope geochemistry of the 3.7×109-yr-old Isua sediments, West Greenland: Implications for the Archaean carbon and oxygen cycles: Geochimica et Cosmochimica Acta, v. 43, p. 189–199, doi:10.1016/0016-7037(79)90238-2.

Schouten, S., Hopmans, E.C., Pancost, R.D., and Sinninghe Damsté, J.S.,2000a, Widespread occurrence of structurally diverse tetraether membrane lipids: Evidence for the ubiquitous presence of low-temperature relatives of hyper-thermophiles: Proceedings of the National Academy of Sciences of the United States of America, v. 97, p. 14,421–14,426, doi:10.1073/pnas.97.26.14421.

Schouten, S., Van Kaam-Peters, H.M.E., Rijpstra, W.I.C., Schoell, M., and Sin-ninghe Damsté, J.S., 2000b, Effects of an oceanic anoxic event on the stable carbon isotopic composition of early Toarcian carbon: American Journal of Science, v. 300, p. 1–22, doi:10.2475/ajs.300.1.1.

Schouten, S., Hopmans, E.C., Schefuß, E., and Sinninghe Damsté, J.S., 2002, Distributional variations in marine crenarchaeotal membrane lipids: A new tool for reconstructing ancient sea water temperatures?: Earth and Planetary Science Letters, v. 204, p. 265–274, doi:10.1016/S0012-821X(02)00979-2.

Schouten, S., Hopmans, E.C., Forster, A., van Breugel, Y., Kuypers, M.M.M., and Sinninghe Damste, J.S., 2003, Extremely high sea-surface tempera-tures at low latitudes during the Middle Cretaceous as revealed by archaeal membrane lipids: Geology, v. 31, p. 1069–1072, doi:10.1130/G19876.1.

Schouten, S., Hopmans, E.C., Baas, M., Boumann, H., Standfest, S., Konneke, M., Stahl, D.A., and Sinninghe Damsté, J.S., 2008, Intact membrane lipids of “Candidatus Nitrosopumilus maritimus,” a cultivated representative of the cosmopolitan mesophilic group I Crenarchaeota: Applied and Environmental Microbiology, v. 74, p. 2433–2440, doi:10.1128/AEM.01709-07.

Seifert, W.K., and Moldowan, J.M., 1978, Applications of steranes, terpanes and monoaromatics to the maturation, migration and source of crude oils: Geochimica et Cosmochimica Acta, v. 42, p. 77–95, doi:10.1016/0016-7037(78)90219-3.

Seifert, W.K., and Moldowan, J.M., 1980, The effect of thermal stress on source-rock quality as measured by hopane stereochemistry: Physics and Chemistry of the Earth, v. 12, no. 229–237.

Sephton, M.A., Looy, C.V., Brinkhuis, H., Wignall, P.B., de Leeuw, J.W., and Visscher, H., 2005, Catastrophic soil erosion during the end-Permian biotic crisis: Geology, v. 33, p. 941–944, doi:10.1130/G21784.1.

Sephton, M.A., Love, G.D., Watson, J.S., Verchovsky, A.B., Wright, I.P., Snape, C.E., and Gilmour, I., 2004, Hydropyrolysis of insoluble carbonaceous matter in the Murchison meteorite: New insights into its macromolecu-lar structure: Geochimica et Cosmochimica Acta, v. 68, p. 1385–1393, doi:10.1016/j.gca.2003.08.019.

Sepulveda, J., Wendler, J.E., Summons, R.E., and Hinrichs, K.-U., 2009, Rapid resurgence of marine productivity after the Cretaceous-Paleogene mass extinction: Science, v. 326, p. 129–132, doi:10.1126/science.1176233.

Sessions, A.L., Burgoyne, T.W., Schimmelmann, A., and Hayes, J.M., 1999, Fractionation of hydrogen isotopes in lipid biosynthesis: Organic Geo-chemistry, v. 30, p. 1193–1200, doi:10.1016/S0146-6380(99)00094-7.

42 Johnson et al.


Sharp, Z.D., 1990, A laser-based microanalytical method for the in situ determina-tion of oxygen isotope ratios of silicates and oxides: Geochimica et Cosmochi-mica Acta, v. 54, p. 1353–1357, doi:10.1016/0016-7037(90)90160-M.

Shen, S.-z., Crowley, J.L., Wang, Y., Bowring, S.A., Erwin, D.H., Sadler, P.M., Cao, C.-q., Rothman, D.H., Henderson, C.M., Ramezani, J., Zhang, H., Shen, Y., Wang, X.-d., Wang, W., Mu, L., Li, W.-z., Tang, Y.-g., Liu, X.-l., Liu, L.-j., Zeng, Y., Jiang, Y.-f., and Jin, Y.-g., 2011, Calibrating the end-Permian mass extinction: Science, v. 334, p. 1367–1372, doi:10.1126/science.1213454.

Sheppard, S.M.F., 1986, Characterization and isotopic variations in natural waters: Reviews in Mineralogy, v. 16, p. 165–183.

Shields, G.A., and Veizer, J., 2002, The Precambrian marine carbonate isotope database: Version 1.1: Geochemistry Geophysics Geosystems, v. 3, no. 6, 12 p., doi:10.1029/2001GC000266.

Shirey, S.B., and Richardson, S.H., 2011, Start of the Wilson cycle at 3 Ga shown by diamonds from subcontinental mantle: Science, v. 333, p. 434–436, doi:10.1126/science.1206275.

Shirey, S.B., and Walker, R.J., 1998, The Re-Os isotope system in cosmo-chemistry and high-temperature geochemistry: Annual Review of Earth and Planetary Sciences, v. 26, p. 423–500, doi:10.1146/annurev.earth.26.1.423.

Sikes, E.L., and Volkman, J.K., 1993, Calibration of alkenone unsatura-tion ratios (Uk’37) for paleotemperature estimation in cold polar waters: Geochimica et Cosmochimica Acta, v. 57, p. 1883–1889, doi:10.1016/0016-7037(93)90120-L.

Sillèn, L.G., 1961, The physical chemistry of sea water, in Sears, M., ed., Ocean-ography: Washington, D.C., American Association for the Advancement of Science, p. 549–581.

Simon, J.I., Reid, M.R., and Young, E.D., 2007, Lead isotopes by LA-MC-ICP-MS: Tracking the emergence of mantle signatures in an evolving silicic magma system: Geochimica et Cosmochimica Acta, v. 71, p. 2014–2035, doi:10.1016/j.gca.2007.01.023.

Simoneit, B.R.T., Grimalt, J.O., Wang, T.G., Cox, R.E., Hatcher, P.G., and Nis-senbaum, A., 1986, Cyclic terpenoids of contemporary resinous plant detritus and of fossil woods, ambers and coals: Organic Geochemistry, v. 10, p. 877–889, doi:10.1016/S0146-6380(86)80025-0.

Sinninghe Damsté, J.S., and Koopmans, M.P., 1997, The fate of carotenoids in sediments: An overview: Pure and Applied Chemistry, v. 69, p. 2067–2074, doi:10.1351/pac199769102067.

Sinninghe Damsté, J.S., Rijpstra, W.I.C., Hopmans, E.C., Prahl, F.G., Wakeham, S.G., and Schouten, S., 2002, Distribution of membrane lipids of planktonic Crenarchaeota in the Arabian Sea: Applied and Environmental Microbiology, v. 68, p. 2997–3002, doi:10.1128/AEM.68.6.2997-3002.2002.

Sinninghe Damsté, J.S., Muyzer, G., Abbas, B., Rampen, S.W., Masse, G., Allard, W.G., Belt, S.T., Robert, J.-M., Rowland, S.J., Moldowan, J.M., Barbanti, S.M., Fago, F.J., Denisevich, P., Dahl, J., Trindade, L.F., and Schouten, S., 2004, The rise of the rhizosolenid diatoms: Science, v. 304, p. 584–587, doi:10.1126/science.1096806.

Sinninghe Damsté, J.S., Rijpstra, W.I.C., Geenevasen, J.A.J., Strous, M., and Jetten, M.S.M., 2005, Structural identifi cation of ladderane and other membrane lipids of planctomycetes capable of anaerobic ammo-nium oxidation (anammox): The FEBS Journal, v. 272, p. 4270–4283, doi:10.1111/j.1742-4658.2005.04842.x.

Smith, F.A., and Freeman, K.H., 2006, Infl uence of physiology and climate on δD of leaf wax n-alkanes from C3 and C4 grasses: Geochimica et Cosmochimica Acta, v. 70, p. 1172–1187, doi:10.1016/j.gca.2005.11.006.

Solomon, G.C., and Taylor, H.P.J., 1989, Isotopic evidence for the origin of Mesozoic and Cenozoic granitic plutons in the northern Great Basin: Geology, v. 17, p. 591–594, doi:10.1130/0091-7613(1989)017<0591:IEFTOO>2.3.CO;2.

Spooner, E.T.C., 1976, The strontium isotopic composition of seawater, and seawater-oceanic crust interaction: Earth and Planetary Science Letters, v. 31, p. 167–174, doi:10.1016/0012-821X(76)90108-4.

Squyres, S.W., Grotzinger, J.P., Arvidson, R.E., Bell, J.F., Calvin, W., Chris-tensen, P.R., Clark, B.C., Crisp, J.A., Farrand, W.H., Herkenhoff, K.E., Johnson, J.R., Klingelhaufer, G., Knoll, A.H., McLennan, S.M., McSween, H.Y., Morris, R.V., Rice, J.W., Rieder, R., and Soderblom, L.A., 2004, In situ evidence for an ancient aqueous environment at Meridiani Planum, Mars: Science, v. 306, p. 1709–1714, doi:10.1126/science.1104559.

Stacey, J.S., and Kramers, J.D., 1975, Approximation of terrestrial lead iso-tope evolution by a two-stage model: Earth and Planetary Science Letters, v. 26, p. 207–221, doi:10.1016/0012-821X(75)90088-6.

Steinhoefel, G., Horn, I., and Von Blanckenburg, F., 2009, Micro-scale tracing of Fe and Si isotope signatures in banded iron formation using femtosec-ond laser ablation: Geochimica et Cosmochimica Acta, v. 73, p. 5343–5360, doi:10.1016/j.gca.2009.05.037.

Stevenson, R.K., and Patchett, P.J., 1990, Implications for the evolution of continental crust from Hf isotope systematics of Archean detri-tal zircons: Geochimica et Cosmochimica Acta, v. 54, p. 1683–1697, doi:10.1016/0016-7037(90)90400-F.

Sturt, H.F., Summons, R.E., Smith, K., Elvert, M., and Hinrichs, K.-U., 2004, Intact polar membrane lipids in prokaryotes and sediments deciphered by high-performance liquid chromatography/electrospray ionization multistage mass spectrometry–new biomarkers for biogeochemistry and microbial ecology: Rapid Communications in Mass Spectrometry, v. 18, p. 617–628, doi:10.1002/rcm.1378.

Summons, R.E., and Powell, T.G., 1987, Identifi cation of aryl isoprenoids in source rocks and crude oils: Biological markers for the green sul-phur bacteria: Geochimica et Cosmochimica Acta, v. 51, p. 557–566, doi:10.1016/0016-7037(87)90069-X.

Summons, R.E., Brassell, S.C., Eglinton, G., Evans, E., Horodyski, R.J., Robin-son, N., and Ward, D.M., 1988, Distinctive hydrocarbon biomarkers from fossiliferous sediment of the Late Proterozoic Walcott Member, Chuar Group, Grand Canyon, Arizona: Geochimica et Cosmochimica Acta, v. 52, p. 2625–2637, doi:10.1016/0016-7037(88)90031-2.

Summons, R.E., Bradley, A.S., Jahnke, L.L., and Waldbauer, J.R., 2006, Ste-roids, triterpenoids and molecular oxygen: Philosophical Transactions of the Royal Society of London, ser. B, Biological Sciences, v. 361, p. 951–968, doi:10.1098/rstb.2006.1837.

Sundquist, E.T., and Visser, K., 2003, The geologic history of the carbon cycle, in Holland, H.D., and Turekian, K.K., eds., Treatise on Geochemistry, Volume 8, Biogeochemistry: Oxford, UK, Elsevier, p. 425–472.

Sutton, S.R., Caffee, M.W., and Dove, M.T., 2006, Synchrotron radiation, neu-tron, and mass spectrometry techniques at user facilities: Elements, v. 2, p. 15–21, doi:10.2113/gselements.2.1.15.

Swisher, C.C., III, Grajales-Nishimura, J.M., Montanari, A., Margolis, S.V., Claeys, P., Alvarez, W., Renne, P., Cedillo-Pardoa, E., Maurrasse, F.J., Curtis, G.H., Smit, J., and McWilliams, M.O., 1992, Coeval 40Ar/39Ar ages of 65.0 million years ago from Chicxulub crater melt rock and Cretaceous-Tertiary boundary tektites: Science, v. 257, p. 954–958, doi:10.1126/science.257.5072.954.

Talbot, H.M., Summons, R.E., Jahnke, L.L., Cockell, C.S., Rohmer, M., and Farrimond, P., 2008, Cyanobacterial bacteriohopanepolyol signatures from cultures and natural environmental settings: Organic Geochemistry, v. 39, p. 232–263, doi:10.1016/j.orggeochem.2007.08.006.

Tang, Ming, and Anders, E., 1988, Isotopic anomalies of Ne, Xe, and C in meteorites. II. Interstellar diamond and SiC: Carriers of exotic noble gases: Geochimica et Cosmochimica Acta, v. 52, p. 1235–1244, doi:10.1016/0016-7037(88)90277-3.

Tatsumoto, M., 1966, Genetic relations of oceanic basalts as indicated by lead isotopes: Science, v. 153, p. 1094–1101, doi:10.1126/science.153.3740.1094.

Taylor, H.P., Jr., 1980, The effects of assimilation of country rocks by mag-mas on 18O/16O and 87Sr/86Sr systematics in igneous rocks: Earth and Planetary Science Letters, v. 47, p. 243–254, doi:10.1016/0012-821X(80)90040-0.

Taylor, H.P., Jr., and Sheppard, S.M.F., 1986, Igneous rocks: I. Processes of isotopic fractionation and isotopic systematics: Reviews in Mineralogy, v. 16, p. 227–271.

Taylor, L.A., Nazarov, M.A., Demidova, S.I., and Patchen, A., 2001, A new mare basalt from Oman: Meteoritics & Planetary Science, v. 36, p. A204.

Taylor, S.R., and Jakes, P., 1974, The geochemical evolution of the Moon, in Proceedings of the Lunar and Planetary Science Conference, Volume 5: Houston, Texas, Lunar and Planetary Institute, p. 1287–1305.

Taylor, S.R., and McLennan, S.M., 1985, The Continental Crust: Its Composi-tion and Evolution: Oxford, UK, Blackwell Scientifi c Publications, 312 p.

Taylor, S.R., and McLennan, S.M., 2009, Planetary Crusts: Their Compo-sition, Origin and Evolution: Cambridge, UK, Cambridge University Press, 378 p.

Taylor, S.R., Pieters, C.M., and MacPherson, G.J., 2006, Earth-Moon system, planetary science, and lessons learned: Reviews in Mineralogy and Geo-chemistry, v. 60, p. 657–704, doi:10.2138/rmg.2006.60.7.

Tera, F., Brown, L., Morris, J., Sacks, I.S., Klein, J., and Middleton, R., 1986, Sediment incorporation in island-arc magmas: Inferences from 10Be:



Geochimica et Cosmochimica Acta, v. 50, p. 535–550, doi:10.1016/0016-7037(86)90103-1.

Thiagarajan, N., Adkins, J., and Eiler, J., 2011, Carbonate clumped isotope thermometry of deep-sea corals and implications for vital effects: Geo-chimica et Cosmochimica Acta, v. 75, p. 4416–4425, doi:10.1016/j.gca.2011.05.004.

Thirlwall, M.F., 1991, Long-term reproducibility of multicollector Sr and Nd isotope ratio analyses: Chemical Geology–Isotope Geoscience Section, v. 94, p. 85–104, doi:10.1016/0168-9622(91)90002-E.

Thode, H.G., Monster, J., and Dunford, H.B., 1961, Sulphur isotope geo-chemistry: Geochimica et Cosmochimica Acta, v. 25, p. 159–174, doi:10.1016/0016-7037(61)90074-6.

Tissot, B.P., and Welte, D.H., 1978, Petroleum Formation and Occurrence: A New Approach to Oil and Gas Exploration: New York, Springer-Verlag, 554 p.

Tomascak, P. B., 2004, Developments in the understanding and application of lithium isotopes in the earth and planetary sciences: Reviews in Mineral-ogy and Geochemistry, v. 55, p. 153–195.

Tripati, A.K., Eagle, R.A., Thiagarajan, N., Gagnon, A.C., Bauch, H., Halloran, P.R., and Eiler, J.M., 2010, 13C-18O isotope signatures and “clumped iso-tope” thermometry in foraminifera and coccoliths: Geochimica et Cosmo-chimica Acta, v. 74, p. 5697–5717, doi:10.1016/j.gca.2010.07.006.

Tsikos, H., Matthews, A., Erel, Y., and Moore, J.M., 2010, Iron isotopes con-strain biogeochemical redox cycling of iron and manganese in a Palaeo-proterozoic stratifi ed basin: Earth and Planetary Science Letters, v. 298, no. 1–2, p. 125–134, doi:10.1016/j.epsl.2010.07.032.

Twitchett, R.J., Looy, C.V., Morante, R., Visscher, H., and Wignall, P.B., 2001, Rapid and synchronous collapse of marine and terrestrial ecosys-tems during the end-Permian biotic crisis: Geology, v. 29, p. 351–354, doi:10.1130/0091-7613(2001)029<0351:RASCOM>2.0.CO;2.

Tyrrell, S., Haughton, P.D.W., and Daly, J.S., 2007, Drainage reorganization during breakup of Pangea revealed by in-situ Pb isotopic analysis of detri-tal K-feldspar: Geology, v. 35, p. 971–974, doi:10.1130/G4123A.1.

Urey, H.C., 1947, The thermodynamic properties of isotopic substances: Jour-nal of the Chemical Society, May 1947, p. 562–581.

Valaas-Hyslop, E., Valley, J.W., Johnson, C.M., and Beard, B.L., 2008, The effects of metamorphism on O and Fe isotope compositions in the Biwabik iron-formation, northern Minnesota: Contributions to Mineral-ogy and Petrology, v. 155, p. 313–328, doi:10.1007/s00410-007-0244-2.

Valley, J.W., 1986, Stable isotope geochemistry of metamorphic rocks: Reviews in Mineralogy, v. 16, p. 445–489.

Valley, J.W., and Kita, N.T., 2009, In situ oxygen isotope geochemistry by ion microprobe: Toronto, Mineralogical Associate of Canada, Short Course Volume 41, p. 19–63.

Valley, J.W., Peck, W.H., King, E.M., and Wilde, S.A., 2002, A cool early Earth: Geology, v. 30, p. 351–354, doi:10.1130/0091-7613(2002)030<0351:ACEE>2.0.CO;2.

Valley, J.W., Cavosie, A.J., Fu, B., Peck, W.H., and Wilde, S.A., 2006, Com-ment on “Heterogeneous Hadean hafnium: Evidence of continental crust at 4.4 to 4.5 Ga”: Science, v. 312, p. 1139, doi:10.1126/science.1125301.

Van Kaam-Peters, H.M.E., Rijpstra, W.I.C., De Leeuw, J.W., and Sinninghe Damsté, J.S., 1998a, A high resolution biomarker study of different lithofacies of organic sulfur-rich carbonate rocks of a Kimmeridgian lagoon (French southern Jura): Organic Geochemistry, v. 28, p. 151–177, doi:10.1016/S0146-6380(97)00132-0.

Van Kaam-Peters, H.M.E., Schouten, S., Kester, J., and Sinninghe Damsté, J.S., 1998b, Controls on the molecular and carbon isotopic composition of organic matter deposited in a Kimmeridgian euxinic shelf sea: Evi-dence for preservation of carbohydrates through sulfurisation: Geochi-mica et Cosmochimica Acta, v. 62, p. 3259–3283, doi:10.1016/S0016-7037(98)00231-2.

van Zuilen, M.A., Lepland, A., and Arrhenius, G., 2002, Reassessing the evidence for the earliest traces of life: Nature, v. 418, p. 627–630, doi:10.1038/nature00934.

Veizer, J., and Compston, W., 1976, 87Sr/86Sr in Precambrian carbonates as an index of crustal evolution: Geochimica et Cosmochimica Acta, v. 40, p. 905–914, doi:10.1016/0016-7037(76)90139-3.

Veizer, J., Compston, W., Clauer, N., and Schidlowski, M., 1983, 87Sr/86Sr in Late Proterozoic carbonates: Evidence for a mantle event at 900 Ma ago: Geochimica et Cosmochimica Acta, v. 47, p. 295–302, doi:10.1016/0016-7037(83)90142-4.

Vervoort, J.D., and Blichert-Toft, J., 1999, Evolution of the depleted man-tle: Hf isotope evidence from juvenile rocks through time: Geochi-

mica et Cosmochimica Acta, v. 63, p. 533–556, doi:10.1016/S0016-7037(98)00274-9.

Vervoort, J.D., Patchett, P.J., Blichert-Toft, J., and Albarède, F., 1999, Rela-tionships between Lu-Hf and Sm-Nd isotopic systems in the global sedi-mentary system: Earth and Planetary Science Letters, v. 168, p. 79–99, doi:10.1016/S0012-821X(99)00047-3.

Vinogradov, A.P., 1943, Biogeochemical research in the U.S.S.R.: Nature, v. 151, p. 659–661, doi:10.1038/151659a0.

Voegelin, A.R., Nagler, T.F., Beukes, N.J., and Lacassie, J.P., 2010, Molyb-denum isotopes in Late Archean carbonate rocks: Implications for early Earth oxygenation: Precambrian Research, v. 182, no. 1–2, p. 70–82, doi:10.1016/j.precamres.2010.07.001.

Voice, P.J., Kowalewski, M., and Eriksson, K.A., 2011, Quantifying the tim-ing and rate of crustal evolution: Global compilation of radiometrically dated detrital zircon grains: The Journal of Geology, v. 119, p. 109–126, doi:10.1086/658295.

Volkman, J.K., 2003, Sterols in microorganisms: Applied Microbiology and Biotechnology, v. 60, p. 495–506.

Volkman, J.K., Eglinton, G., Corner, E.D.S., and Forsberg, T.E.V., 1980, Long-chain alkenes and alkenones in the marine coccolithophorid Emili-ania huxleyi: Phytochemistry, v. 19, p. 2619–2622, doi:10.1016/S0031-9422(00)83930-8.

Volkman, J.K., Barrerr, S.M., Blackburn, S.I., and Sikes, E.L., 1995, Alke-nones in Gephyrocapsa oceanica: Implications for studies of paleo-climate: Geochimica et Cosmochimica Acta, v. 59, p. 513–520, doi:10.1016/0016-7037(95)00325-T.

Von Blanckenburg, F., Mamberti, M., Schoenberg, R., Kamber, B.S., and Webb, G.E., 2008, The iron isotope composition of microbial carbonate: Chemi-cal Geology, v. 249, p. 113–128, doi:10.1016/j.chemgeo.2007.12.001.

Wacey, D., Kilburn, M.R., Saunders, M., Cliff, J., and Brasier, M.D., 2011, Microfossils of sulphur-metabolizing cells in 3.4-billion-year-old rocks of Western Australia: Nature Geoscience, v. 4, p. 698–702, doi:10.1038/ngeo1238.

Waldbauer, J.R., Sherman, L.S., Sumner, D.Y., and Summons, R.E., 2009, Late Archean molecular fossils from the Transvaal Supergroup record the antiquity of microbial diversity and aerobiosis: Precambrian Research, v. 169, p. 28–47, doi:10.1016/j.precamres.2008.10.011.

Walker, R.J., 2009, Highly siderophile elements in the Earth, Moon and Mars: Update and implications for planetary accretion and differentiation: Che-mie der Erde, v. 69, p. 101–125, doi:10.1016/j.chemer.2008.10.001.

Walker, R.J., Carlson, R.W., Shirey, S.B., and Boyd, F.R., 1989, Os, Sr, Nd, and Pb isotope systematics of southern African peridotite xenoliths: Implications for the chemical evolution of subcontinental mantle: Geo-chimica et Cosmochimica Acta, v. 53, p. 1583–1595, doi:10.1016/0016-7037(89)90240-8.

Walker, R.J., Morgan, J.W., and Horan, M.F., 1995, Osmium-187 enrichment in some plumes: Evidence for core-mantle interaction?: Science, v. 269, p. 819–822, doi:10.1126/science.269.5225.819.

Wang, Z., Schauble, E.A., and Eiler, J.M., 2004, Equilibrium thermodynamics of multiply substituted isotopologues of molecular gases: Geochimica et Cosmochimica Acta, v. 68, p. 4779–4797, doi:10.1016/j.gca.2004.05.039.

Watanabe, Y., Farquhar, J., and Ohmoto, H., 2009, Anomalous fractionations of sulfur isotopes during thermochemical sulfate reduction: Science, v. 324, p. 370–373, doi:10.1126/science.1169289.

Weijers, J.W.H., Schouten, S., Spaargaren, O.C., and Sinninghe Damsté, J.S., 2006, Occurrence and distribution of tetraether membrane lip-ids in soils: Implications for the use of the TEX86 proxy and the BIT index: Organic Geochemistry, v. 37, p. 1680–1693, doi:10.1016/j.orggeochem.2006.07.018.

Welander, P.V., Hunter, R.C., Zhang, L., Sessions, A.L., Summons, R.E., and Newman, D.K., 2009, Hopanoids play a role in membrane integrity and pH homeostasis in Rhodopseudomonas palustris TIE-1: Journal of Bacte-riology, v. 191, p. 6145–6156, doi:10.1128/JB.00460-09.

Welander, P.V., Coleman, M.L., Sessions, A.L., Summons, R.E., and Newman, D.K., 2010, Identifi cation of a methylase required for 2- methylhopanoid production and implications for the interpretation of sedimentary hopanes: Proceedings of the National Academy of Sciences of the United States of America, v. 107, p. 8537–8542, doi:10.1073/pnas.0912949107.

Welander, P.V., Doughty, D.M., Wu, C.H., Mehay, S., Summons, R.E., and Newman, D.K., 2012, Identifi cation and characterization of Rhodopseu-domonas palustris TIE-1 hopanoid biosynthesis mutants: Geobiology, v. 10, p. 163–177, doi:10.1111/j.1472-4669.2011.00314.x.

44 Johnson et al.


Weyer, S., Anbar, A.D., Gerdes, A., Gordon, G.W., Algeo, T.J., and Boyle, E.A., 2008, Natural fractionation of 238U/235U: Geochimica et Cosmochimica Acta, v. 72, p. 345–359, doi:10.1016/j.gca.2007.11.012.

White, A.F., and Brantley, S.L., 1995, Chemical weathering rates of silicate minerals: Reviews in Mineralogy, v. 31, p. 583.

White, W.M., 2012, Geochemistry: New York, Wiley-Blackwell, 656 p.Wignall, P., and Hallam, A., 1992, Anoxia as a cause of the Permian/ Triassic

mass extinction: Facies evidence from northern Italy and the western United States: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 93, p. 21–46.

Williford, K.H., Van Kranendonk, M.J., Ushikubo, T., Kozdon, R., and Valley, J.W., 2011, Constraining atmospheric oxygen and seawater sulfate con-centrations during Paleoproterozoic glaciation: In situ sulfur three-isotope microanalysis of pyrite from the Turee Creek Group, Western Australia: Geochimica et Cosmochimica Acta, v. 75, p. 5686–5705, doi:10.1016/j.gca.2011.07.010.

Williford, K.H., Ushikubo, T., Schopf, J.W., Lepot, K., Kitajima, K., and Valley, J.W., 2013, Preservation and detection of microstructural and taxonomic correlations in the carbon isotopic compositions of individual Precam-brian microfossils: Geochimica et Cosmochimica Acta, v. 104, p. 165–182, doi:10.1016/j.gca.2012.11.005.

Winter, B.L., and Knauth, L.P., 1992, Stable isotope geochemistry of cherts and car-bonates from the 2.0 Ga Gunfl int Iron Formation: Implications for the deposi-tional setting, and the effects of diagenesis and metamorphism: Precambrian Research, v. 59, p. 283–313, doi:10.1016/0301-9268(92)90061-R.

Wood, J.A., and Hashimoto, A., 1993, Mineral equilibrium in fractionated neb-ular systems: Geochimica et Cosmochimica Acta, v. 57, p. 2377–2388, doi:10.1016/0016-7037(93)90575-H.

Wood, J.A., Dickey, J.S., Marvin, U.B., and Poweel, B.N., Lunar anorthosites and a geophysical model of the Moon, in Levinson, A.A., ed., Proceed-ings of the Apollo 11 Lunar Science Conference 1970; Volume 1: Miner-alogy and Petrology: New York, Pergamon Press, p. 965–988.

Wuchter, C., Schouten, S., Wakeham, S.G., and Sinninghe Damsté, J.S., 2005, Temporal and spatial variation in tetraether membrane lip-ids of marine Crenarchaeota in particulate organic matter: Implica-tions for TEX86 paleothermometry: Paleoceanography, v. 20, PA3013, doi:10.1029/2004PA001110.

Xie, S., Pancost, R.D., Huang, J., Wignall, P.B., Yu, J., Tang, X., Chen, L., Huang, X., and Lai, X., 2007, Changes in the global carbon cycle

occurred as two episodes during the Permian-Triassic crisis: Geology, v. 35, p. 1083–1086, doi:10.1130/G24224A.1.

Yamaguchi, K.E., Johnson, C.M., Beard, B.L., and Ohmoto, H., 2005, Bio-geochemical cycling of iron in the Archean–Paleoproterozoic Earth: Constraints from iron isotope variations in sedimentary rocks from the Kaapvaal and Pilbara cratons: Chemical Geology, v. 218, p. 135–169, doi:10.1016/j.chemgeo.2005.01.020.

Yoneda, S., and Grossman, L., 1995, Condensation of CaO-MgO-Al2O

3-SiO

2

liquids from cosmic gases: Geochimica et Cosmochimica Acta, v. 59, p. 3413–3444, doi:10.1016/0016-7037(95)00214-K.

Young, E.D., and Galy, A., 2004, The isotope geochemistry and cosmochem-istry of magnesium: Reviews in Mineralogy and Geochemistry, v. 55, p. 197–230, doi:10.2138/gsrmg.55.1.197.

Yuen, G., Blair, N., Des Marais, D.J., and Chang, S., 1984, Carbon isotope composition of low molecular weight hydrocarbons and monocar-boxylic acids from Murchison meteorite: Nature, v. 307, p. 252–254, doi:10.1038/307252a0.

Zachos, J.C., Schouten, S., Bohaty, S., Quattlebaum, T., Sluijs, A., Brinkhuis, H., Gibbs, S.J., and Bralower, T.J., 2006, Extreme warming of mid-latitude coastal ocean during the Paleocene-Eocene thermal maximum: Inferences from TEX86 and isotope data: Geology, v. 34, p. 737–740, doi:10.1130/G22522.1.

Zahnle, K., Claire, M., and Catling, D., 2006, The loss of mass-independent fractionation in sulfur due to a Palaeoproterozoic collapse of atmo-spheric methane: Geobiology, v. 4, p. 271–283, doi:10.1111/j.1472-4669.2006.00085.x.

Zartman, R.E., and Doe, B.R., 1981, Plumbotectonics—The model: Tectono-physics, v. 75, p. 135–162, doi:10.1016/0040-1951(81)90213-4.

Zindler, A., and Hart, S.R., 1986, Chemical geodynamics: Annual Review of Earth and Planetary Sciences, v. 14, p. 493–571, doi:10.1146/annurev.ea.14.050186.002425.

Zindler, A., Jagoutz, E., and Goldstein, S., 1982, Nd, Sr and Pb isotopic sys-tematics in a three-component mantle: A new perspective: Nature, v. 298, p. 519–523, doi:10.1038/298519a0.

Zinner, E., 2004, Presolar grains, in Davis, A.M., ed., Treatise on Geochemis-try; Volume 1: Meteorites, Comets, and Planets: Oxford, UK, Elsevier, p. 17–39.

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