Isotopic order, biogeochemical processes, and earth history

Geochimica et Cosmochimica Acta, Vol. 68, No. 8, pp. 1691–1700, 2004 Copyright © 2004 Elsevier Ltd Printed in the USA. All rights reserved 0016-7037/04 $30.00 ⫹ .00

Pergamon

doi:10.1016/j.gca.2003.10.023

Isotopic order, biogeochemical processes, and earth history Goldschmidt Lecture, Davos, Switzerland, August 2002 JOHN M. HAYES* Department of Geology and Geophysics, Woods Hole Oceanographic Institution Woods Hole, MA 02543, USA

Abstract—The impetus to interpret carbon isotopic signals comes from an understanding of isotopic fractionations imposed by living organisms. That understanding rests in turn on studies of enzymatic isotope effects, on fruitful concepts of isotopic order, and on studies of the distribution of 13C both between and within biosynthetic products. In sum, these studies have shown that the isotopic compositions of biological products are governed by reaction kinetics and by pathways of carbon flow. Isotopic compositions of individual compounds can indicate specific processes or environments. Examples include biomarkers which record the isotopic compositions of primary products in aquatic communities, which indicate that certain bacteria have used methane as a carbon source, and which show that some portions of marine photic zones have been anaerobic. In such studies, the combination of structural and isotopic lines of evidence reveals relationships between compounds and leads to process-related thinking. These are large steps. Reconstruction of the sources and histories of molecular fossils redeems much of the early promise of organic geochemistry by resolving and clarifying paleoenviron-mental signals. In turn, contemplation of this new information is driving geochemists to study microbial ecology and evolution, oceanography, and sedimentology. Copyright © 2004 Elsevier Ltd specific organisms, and particular sources of signals (e.g., a paleoclimatic variation). There are multiple approaches to decoding. By resolving the structures of sedimentary molecules in great detail, many workers have been able to constrain their histories. Stratigraphic variations also provide useful evidence: molecules whose abundances vary together might have a common source. The present discussion will focus on isotopic variations. An isotopic variation does not constitute an interpretable signal unless the mechanism controlling it is known. For inorganic processes, these are usually equilibrium isotope effects (Chacko et al., 2001). For organic molecules, the mechanisms prominently include kinetic as well as equilibrium isotope effects (Hayes, 2001). Large and diverse isotopic variations are common. The questions are what causes them and whether the mechanisms are consistent enough that observed fractionations can be interpreted reliably.

1. INTRODUCTION

My aim is to trace the transformation of organic geochemistry into biogeochemistry. It has been an interesting revolution with many leaders. I am honored to appear as a representative. Forty years ago, organic geochemists thought mainly about petroleum and its sources. In that milieu, Philip Abelson, Melvin Calvin, and Geoffrey Eglinton were revolutionaries. Abelson called his research “paleobiochemistry” (Abelson, 1954) and then “biogeochemistry” (Abelson and Hoering, 1960). Eglinton and Calvin (1967) took organisms and biochemistry as the starting points for their interpretations of sedimentary molecules. The impact of this new thinking was great and the future looked very promising. The development of the field was discussed using terms from information theory (Shannon, 1948; Hamming, 1986). Sedimentary organic molecules were viewed as channels linking modern observers to ancient events. The great diversity of molecular structures meant that the number of channels, and thus the flow of information, was practically unlimited. The implied system is summarized in Figure 1, which is both conceptual and cautionary. As it illustrates the hoped-for flow of information, it also reminds us that the molecular channels are coupled to remarkably diverse sources (in the information-theoretical sense) and encoders. They are also affected by an important source of noise. The tools of modern organic chemical analysis were brought to bear and data were immediately plentiful. But data can be turned into information only by accurate decoding. For organic geochemists, the essence of decoding is to know not only how the atoms within the molecule are arranged now, but also to know about their history. Decoding establishes relationships between sedimentary molecules, initial biosynthetic products,

* Author to whom ([email protected]).

correspondence

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2. INTRAMOLECULAR PATTERNS OF ISOTOPIC ORDER

A key contribution was made by Galimov (1973), who developed accurate techniques for calculating equilibrium distributions of isotopes within organic molecules. Because vibrational environments differ significantly between carbon positions, these intramolecular distributions are not expected to be uniform. In fact, Galimov’s results indicated an astonishing world of isotopic order. Figure 2 provides an arresting example. Each circle in the diagram at right represents a carbon position in a molecule of chlorophyll, with each diameter proportional to the abundance of 13C expected at isotopic equilibrium. Is natural chlorophyll really like this? Even today, no analytical technique can either confirm or refute the prediction. Moreover, that issue is secondary. The shock is that an image this exotic is required to describe the idea. Conventional depictions of molecular structure (shown at left in the figure) have no isotopic dimension. Until then, conventional thinking about molecular structures

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Fig. 1. Schematic view of the role of sedimentary organic molecules as carriers of information. The elements portrayed here (“Message source,” etc.) are those specified by information theory (Shannon, 1948). In this rigorous and formal system, the source of a signal is distinct from the chemical or biological sources of molecules themselves.

had no isotopic dimension. But suddenly, such isotopic patterns seemed likely to offer a means of constraining molecular histories. The task would be to learn what isotopic order was built into molecules during their biosynthesis and then to determine whether and how those patterns were preserved in the geosphere. The impact of Galimov’s work was amplified because Abelson and Hoering (1961) had already demonstrated that measurable isotopic order existed within amino acids. They couldn’t determine the abundance of 13C at each carbon position, but they did show that the carboxyl groups were commonly enriched relative to the rest of the molecule. In both sign and magnitude, the internal isotopic contrasts were for many amino acids similar to those predicted by Galimov’s calculations. Was this coincidental, or did it mean that thermodynamics, rather than kinetics and biosynthetic pathways, actually controlled the distribution of 13C within—and thus between— organic molecules? The clearest answer would come from examination of structurally equivalent positions in a biosynthetic product. Lipids, particularly the common fatty acids, provide the best example. Biosynthetically, these abundant products derive from the polymerization of acetate. As shown in Figure 3 all unbranched

Fig. 3. Factors affecting the distribution of 13C in n-alkyl chains. (a) Biosynthetic sources of the C atoms in a typical n-alkanoic acid (⫽ “fatty acid”). (b) The distribution of 13C expected if isotopic equilibrium prevails throughout the biosynthetic process (Galimov, 1973). (c) The distribution expected if isotopic equilibrium prevails during the biosynthesis of acetate but not during the polymerization of acetate to produce the hydrocarbon chain. (d) The distribution expected if kinetic factors govern the distribution of 13C in both acetate and the final biosynthetic product.

carbon chains in nature contain two subsets of carbon atoms. One is derived from the methyl position of the parent acetate, the other from the carboxyl position. Since all but the terminal-C positions are equivalent CH2 groups, equilibration calls for isotopic homogeneity within the chain (Fig. 3b). Several alternatives can be considered. The first (Fig. 3c) shows what would be expected if the carboxyl position in the acetate monomer were enriched in 13C and if that enrichment, rather than equilibria established during biosynthesis of the final product, controlled the distribution of isotopes. The second alternative (Fig. 3d) is interesting because it is expected if the distribution of 13C is controlled by kinetic factors. DeNiro and Epstein (1977) showed that pyruvate decarboxylase, an enzyme similar to that which produces the acetate biomonomer, has an isotope effect that discriminates against 13C at the C-2 position in pyruvate. That isotope effect would act like a dam. Carbon-13 would pile up behind it and be diverted to alternate fates (Hayes, 2001). The C2 units produced by pyruvate dehydrogenase would be depleted in 13C at the carboxyl position. As shown in Figure 4, the magnitudes of the enrichment and depletion would depend on the branching ratio. 3. DIRECT MEASUREMENTS OF INTRAMOLECULAR ISOTOPIC ORDER

As a first step in searching for such effects, we learned how to drive the Schmidt decarboxylation *O ⫹ N * O H ⫹ HN ™™™™3 RNH ⫹ C RC 2 3 2 2 2 H2SO4 Fig. 2. Two representations of the chemical structure of chlorophyll a. On the left, the conventional structural formula. On the right, a scheme that was invented by Galimov (1974) in order to describe position-specific variations in the abundance of 13C. The isotopic distribution indicated here is that expected at isotopic equilibrium.

(1)

absolutely to completion so that we could analyze carboxyl groups (Vogler and Hayes, 1979, 1980). We also synthesized carboxylic acids in which the isotopic composition of the carboxyl group was known absolutely (Vogler et al., 1978;

Isotopic order, biogeochemical processes, and earth history

Fig. 4. Branching flow of carbon at pyruvate dehydrogenase and its effect on 13C abundances (DeNiro and Epstein, 1977; Monson and Hayes, 1982). The scheme on the left summarizes the reactants and products. A carbon kinetic isotope effect is associated with the action of pyruvate dehydrogenase. As a result, the carboxyl position in acetyl coenzyme A is depleted in 13C relative to the carbonyl position in pyruvate. The graph at right depicts the relationship between the magnitude of that depletion and the branching ratio, expressed in terms of the fraction of carbon flowing to acetate.

Vogler and Hayes, 1978). Specifically, conventional Grignard syntheses (Eqns. 2 and 3) were started with a known excess of CO2, xRMgX ⫹ CO2 3 xRCO2MgBr ⫹ 共1 ⫺ x兲 CO2

(2)

3 xRCO2H xRCO2MgBr ™™™™ H⫹,H2O

(3)

the unconsumed CO2 was recovered and analyzed isotopically, a mass balance showed that all the starting CO2 was accounted for, and the ␦ value of the carboxylic acid was calculated by difference. The acids served as standards that demonstrated the adequacy of the optimized Schmidt decarboxylation (Table 1). To analyze carbon positions within hydrocarbon chains, Monson and Hayes (1982) developed methods for disassem-

Table 1. Tests of Schmidt decarboxylation.

␦13CPDB, ‰ Substrate

Predicted for carboxyl

Found for CO2

Difference

Acetic acid Na salt Hexanoic acid Na salt Me ester Octanoic acid Na salt Me ester Na salt Nonanoic acid Na salt

⫺44.4 ⫾ 1.6

⫺46.1 ⫾ 0.1 ⫺44.7 ⫾ 0.1 ⫺57.4 ⫾ 0.5 ⫺58.7 ⫾ 0.2 ⫺57.3 ⫾ 0.2 ⫺50.7 ⫾ 0.0 ⫺51.8 ⫾ 0.0 ⫺50.6 ⫾ 0.2 ⫺59.1 ⫾ 0.1 ⫺25.2 ⫾ 0.2 ⫺64.6 ⫾ 0.3

⫺1.7 ⫺0.3 0.2 ⫺1.1 0.3 ⫺0.4 ⫺1.5 ⫺0.3 ⫺0.9 ⫺0.2 0.7

a

a

⫺57.6b ⫺50.3b ⫺58.2 ⫾ 1.0c ⫺25.0b ⫺65.2 ⫾ 1.8c

Determined by pyrolysis of the sodium salt (Meinschein et al., 1974). b Independently analyzed by the method of von Unruh and Hayes (1976). c Determined by synthesis (Vogler and Hayes, 1978).

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Fig. 5. Measurement of the isotopic compositions at specific positions within hydrocarbon chains (Monson and Hayes, 1982). (a) Chemical reactions used to convert positions to carboxyl groups. (b) Measured isotopic compositions. Additionally, combustion of the parent acids yielded overall ␦ values. The average result, ⫺12.8‰, does not differ significantly from the average of the internal methyl (⫺9.5) and internal carboxyl (⫺16.0). The terminal carboxyl positions have been fractionated by an additional process (Monson and Hayes, 1980).

bling these chains and, in the process, quantitatively producing new carboxyl groups. As shown in Figure 5a, this allowed multiple, spatially resolved analyses of a parent fatty acid and its cleavage products. Carbon positions 10 and 9, highlighted in the figure and structurally equivalent in the parent acid, derive respectively from the methyl and carboxyl carbons of an acetate group. The results (Fig. 5b) showed that these structurally equivalent positions had different isotopic compositions and, therefore, that isotopic equilibrium did not prevail. As expected on the basis of the isotope effect identified by DeNiro and Epstein (1977), the carbon position derived from the carboxyl group of acetate was isotopically depleted. These and all other results (Monson and Hayes, 1982) were quantitatively consistent with control of isotopic compositions by kinetic isotope effects at two sites of fractionation. The first of these was at pyruvate dehydrogenase, where a 23‰ kinetic isotope effect accounted for the alternating pattern of isotopic order within hydrocarbon chains. The second was at fatty acid synthetase, were a kinetic isotope effect of 13‰ at the branch point between liberation of a free acid vs. continued elongation accounted for varying isotopic abundances of terminal carboxyl groups (Monson and Hayes, 1980). Particularly in retrospect, this dissection of controlling mechanisms has had multiple lines of significance. It provides direct, quantitative support for the existence of intramolecular isotopic order and for the concept that isotopic differences between molecules are in fact the attenuated reflections of isotopic differences within molecules. Together with the demonstration that kinetic isotope effects and flows of carbon within reaction networks govern isotopic fractionations, it provides a straightforward framework for the interpretation of molecular-isotopic signals.

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Fig. 6. Main processes by which phytoplankton can assimilate inorganic carbon (Wolf-Gladrow and Riebesell, 1997; Laws et al., 1997; Laws, 1998; Badger et al., 1998; Reinfelder et al., 2000; Morel et al., 2002).

4. FRACTIONATION OF CARBON ISOTOPES DURING PRODUCTION OF ORGANIC MATTER

The most sophisticated dissection of biochemical reaction networks thus far has been that dealing with the assimilation of carbon dioxide by photosynthetic or chemoautotrophic organisms. Studies of land plants, summarized by Farquhar et al. (1982), showed that isotopic fractionation was largely governed by two processes which operate in series. The first, transport of CO2 from outside the plant to the active site of a carbon-fixing enzyme, has a small isotope effect. The second, fixation (the bonding of carbon from CO2 to a larger organic molecule), usually has a much larger isotope effect. The net fractionation is either large or small, depending which of these processes controls the overall rate. More recently, the research team at Hawaii has beautifully elucidated the processes governing the fractionation of carbon isotopes by algae (Laws et al., 1995, 2002). The relevant systems are summarized in Figure 6. The prominent role of bicarbonate in some of these shows that equilibrium isotope effects can play a role in determining isotopic compositions of organic molecules. If there is a reversible reaction, if it has an isotope effect, and if there is time for equilibration to occur, it will be a controlling factor. A treatment of carbonate equilibria incorporating considerations of mass transport as well as reaction kinetics has recently been published by Zeebe and Wolf-Gladrow (2001). For algae which obtain CO2 by passive diffusion, overall fractionations observed to date can be summarized by this expression (Popp et al., 1998): ␧P ⫽ ␧f ⫺ ␤

␮ 共V/S兲 ce

would be expected to fall along a straight line with an intercept near 25‰ on the ␧P axis. On the right, the data have been transformed and a strong correlation observed. The transformation is one in which the variable, b, serves as a proxy for growth rate. For each point, b is estimated from (25 ⫺ ␧P)ce. That is, given ␧f ⬇ 25‰, b is equal to the product ␤␮(V/S) in Eqn. 4. Since ␤ and (V/S) are practically constant for populations of E. huxleyii and G. oceanica, variations in b must reflect variations in ␮, the specific growth rate. The observed correlation with concentrations of soluble reactive phosphate need not indicate that phosphate itself is the growth-limiting nutrient. Instead, the specific growth rate may be controlled by some trace micronutrient that covaries with phosphate (Morel et al., 1991). Over long periods of geologic time, where the globally averaged fractionation between dissolved inorganic carbon and algal biomass has varied widely, we have a choice between believing that all cells grew more slowly, or that the globally averaged ratio of surface area to volume was higher, or that concentrations of CO2 were higher. On those scales of time and averaging, the CO2-related interpretations, in which the increased depletion of 13C in marine organic matter prior to the Miocene is attributed mainly to higher concentrations of CO2, may still be largely correct (Arthur et al., 1985; Hayes et al., 1999). 5. ISOTOPIC SIGNALS OF BIOGEOCHEMICAL PROCESSES

Consideration of molecular-isotopic signals leads immediately to a focus on biogeochemical processes. These include not only production but also a wide variety of aerobic and anaerobic secondary processes in the water column and sediments. All of these interact to control the molecular and isotopic composition of total organic carbon (TOC) in sediments. It’s easy to focus on that end product. Throughout the Phanerozoic, isotopic mass balances allow estimation of the fraction of carbon that’s been buried in organic form (Holser et al., 1988). For example, from the isotopic compositions of carbonate carbon and organic carbon deposited 30 million years ago, we

(4)

where ␧P is the overall fractionation, ␧f is the isotope effect associated with carbon fixation, ␤ is a constant, ␮ is the specific growth rate, V and S are the volume and surface area of the algal cells, and ce is the concentration of dissolved CO2 external to the algal cell. In the modern ocean, the most important controlling factor seems to be ␮. Its power is seen in Figure 7, which pertains specifically to the alkenone-producing haptophytes, Emiliana huxleyii and Gephyrocapsa oceanica, thus eliminating size and shape as factors. On the left, the data have been plotted simply in terms of ␧P and ce. If concentrations of dissolved CO2 were the dominant controlling factor, the results

Fig. 7. Two plots summarizing the same observations of the fractionation of 13C by alkenone-producing algae (Bidigare et al., 1997; Laws et al., 2001). On the left, measured values of ␧P are plotted as a function of 1/ce, where ce is the concentration of dissolved CO2 external to the algal cells. On the right, the same data have been transformed so that b, a proxy for growth rate (defined in the figure and explained in the text) is plotted as a function of nutrient concentrations.

Isotopic order, biogeochemical processes, and earth history

can estimate that 23% of the carbon being sequestered in sediments was organic and that the rest was accounted for by carbonates. In such calculations, the surface carbon cycle is a black box that puts out carbonates and organic materials. Processes within the box are not considered. Determining why 23% of the buried carbon was organic, or understanding differences between sedimentary basins 30 million years ago, requires consideration of processes. Molecular-isotopic techniques make some key details accessible. We can, for example, estimate the isotopic difference between primary products and the TOC that was finally buried. In our first investigations of this kind, we used porphyrins— degradation products of chlorophyll—rather than algal lipids as the isotopic proxy for primary products (Hayes et al., 1989). Our aim was to partition the isotopic fractionation between carbonates and organic carbon into two components. The first, measured by the difference between carbonates and porphyrins, would reflect fractionations associated with primary production and with the formation of carbonate minerals. The second, measured by the difference between porphyrins and TOC, would reflect the sum of fractionations associated with secondary processes. The approach is summarized schematically in Figure 8. The section investigated was from the Western Interior Seaway in North America. It includes the boundary between the Cenomanian and Turonian ages and its well-known carbonisotopic excursion (Arthur et al., 1988). Depth profiles for carbonates, porphyrins, and TOC are shown in Figure 9. As in many sedimentary sequences, the 4.1‰ range of variations for the organic materials significantly exceeds the 2.3‰ range in carbonates. The differing ranges indicate that the isotopic composition of the organic material is indeed a separate signal that encodes information beyond that carried by the isotopic composition of carbonates. Values of ␦carbonate vary as a result of changes in the globally averaged inputs and outputs to the carbon cycle. They can also be affected by diagenetic processes (Veizer,

Fig. 8. Schematic representation of the partitioning of the isotopic difference between carbonate minerals and total organic carbon into two components (Hayes et al., 1989). The first, denoted by ␧1, includes ␧P and the relatively constant offsets between carbonate minerals and dissolved CO2 and between biomass and the carbon in porphyrin ring systems (the heteroaromatic portion of a chlorophyll molecule). The second, ⌬2, is the sum of fractionations associated with food webs, diagenesis, and other secondary processes.

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Fig. 9. Isotopic profiles, Greenhorn Limestone (Hayes et al., 1989). The upward triangles and solid line refer to the rising portion of the isotopic excursion as seen in carbonates. The downward triangles and dotted line refer to the falling portion.

1983). Under the guidance of Brian Popp, such influences were stringently minimized in this work by analyzing only non luminescent, micritic carbonates (Hayes et al., 1989). Values of ␦roc tend to follow those of ␦carbonate because the dissolved inorganic carbon is the reactant from which the TOC has been produced. They also include the effects of variations in ␧P, the isotopic fractionation associated with primary production, and ⌬2, the isotopic shift associated with secondary processes, namely all of those between primary products and diagenetically stabilized TOC. These include fractionations imposed by the food chain in the water column; by the benthic community; by the sedimentary microbial community; and, depending on the sediment’s history, by thermal processes. It’s useful to think in terms of relationships between precursors and products. Those associated with primary production are summarized in Figure 10. In the upper graph, the deviations from slope ⫽ 1 show that the isotopic composition of primary products did not precisely follow that of dissolved inorganic carbon. In the lower graph, the same data are recast in terms of ␧1. As values of ␦carbonate increased during the late Cenomanian, values of ␧1 and thus of ␧P dropped and then returned to (nearly) their starting value. Thanks to the Hawaiians, we know that the variations must represent some combination of changes in concentrations of CO2, rates of growth, and algal shape and size. Of course all may have varied, with partially offsetting effects. Broadly, however, the signal on the rising side of the isotopic excursion represents a perturbation in which concentrations of CO2 were drawn down and then recovered, in which rates of growth increased and then returned to their starting levels, or in which the algal community temporarily was dominated by fatter cells (larger V/S). A second precursor-product linkage exists between primary products (the biomass produced by photosynthetic organisms, here represented isotopically by the porphyrins) and sedimentary TOC. The resulting isotopic shift is symbolized by ⌬2 (Fig. 11). Remarkably, the slope relating reactants and products is not one but instead 1.5. As porphyrins are enriched in 13C, something is enriching the TOC even more strongly. This corresponds to an increase in ⌬2. In this case, the perturbation

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values of ⌬2. Increased heterotrophy could result from rising levels of O2 in the atmosphere and surface waters. But the isotopic excursion marks an oceanic anoxic event! With only slightly more confidence than perplexity, we concluded that “oxygenation of environments other than deep ocean basins must have increased” (Hayes et al., 1989). Now, stratigraphic, sedimentological, and geochemical studies by Sageman and coworkers (Meyers et al., 2001, 2004) have elegantly confirmed this interpretation. Other isotopically distinct biomarkers can serve as indicators of significant paleoenvironmental conditions or processes. The first example was provided by Summons and Powell (1986). They initially recognized that the aromatic-hydrocarbon frac-

Fig. 10. Two views of primary processes, based on the data in Figure 9. Lines and symbols as in Figure 9. (a) Isotopic composition of porphyrins, a proxy for primary biomass, plotted as a function of the isotopic composition of micritic carbonates, a proxy for dissolved inorganic carbon. The nonunit slope shows that the isotopic composition of the organic product does not simply follow that of the inorganic precursor. (b) The fractionation between porphyrins and micritic carbonates plotted as a function of the isotopic composition of the latter. Decreases in the fractionation can be attributed to the causes noted at lower left in the graph. Increases in ␦carbonate indicate increases in the fraction of carbon buried in organic form.

maximizes near the top of the isotopic excursion and returns to its preexcursion value at the end of the event. Most of the carbon loss on the pathway between primary products and TOC is due to the activities of respiring heterotrophs in the marine food chain (Longhurst and Harrison, 1989). In modern systems, each trophic step produces a 13C enrichment of 0.5–1.5‰ (Yoshii et al., 1999). The magnitude probably depends on the composition of the input (e. g., the relative abundances of lipids, carbohydrates, and proteins). Cretaceous effects may have differed. In any case, increases in the intensity of heterotrophic processes are expected to increase

Fig. 11. Two views of secondary processes, based on the data in Figure 9. Lines and symbols as in Figure 9. (a) Isotopic composition of TOC as a function of that of porphyrins, a proxy for primary biomass. As in the relationship summarized in Figure 10, the isotopic composition of the product (TOC) again does not precisely follow that of its precursor (primary biomass). (b) The same variations summarized in terms of ⌬2, which is seen to rise and fall over the course of the isotopic excursion.

Isotopic order, biogeochemical processes, and earth history

Fig. 12. (a) Isorenieratane, an aromatic polyisoprenoid and (b) isorenieratene (Schaefle´ et al., 1977), a carotenoid pigment unique to the green photosynthetic bacteria, the Chlorobiaceae.

tions of several North American Paleozoic oils contained a prominent series of compounds related to isorenieratane (Fig. 12a). In turn, it was apparently derived from isorenieratene (Fig. 12b), for which the only known source was and is the green photosynthetic bacteria, the Chlorobiaceae. These organisms are obligate anaerobes. They fix carbon and produce biomass by using energy from photons, but they rely on sulfide as an electron donor and are killed by exposure to O2. Accordingly, the depositional environments for the source rocks of these oils must have had waters that were both anaerobic (and sulfidic) and sunlit. In large bodies of water, this is a rare condition. For confirmation, Summons and Powell (1986, 1987) isolated individual compounds, converted their carbon to CO2, and measured the abundance of 13C. The aromatic polyisoprenoids were enriched in 13C relative to saturated hydrocarbons in the same oils by as much as eight permil. That isotopic signal firmly associates these compounds with the Chlorobiaceae, which are unique in their reliance on the reversed tricarboxylic acid cycle as a means of fixing carbon. Isotope effects within that cycle are much smaller than those in other pathways of carbon fixation (Hayes, 2001). Moreover, lipids can be enriched in 13C relative to biomass (van der Meer et al., 1998). In subsequent work, the group led by Jaap Sinninghe Damste´ and Stefan Schouten at the Royal Netherlands Institute for Sea Research has developed improved methods for the detection of these important indicators of water-column anoxia (Koopmans et al., 1996). They are notably abundant, for example, in North Atlantic sediments from the Cretaceous anoxic events (Sinninghe Damste´ and Ko¨ ster, 1998). The presence of products from fastidious anaerobes in open-marine sediments is notable. Extraordinary levels of stratification are required. These may have been stabilized—at least well enough to allow transient growth of the Chlorobiaceae and generation of the molecularisotopic signal— by the estimated sea-surface temperatures of 35 to 38°C (which derive from a new, molecular-organic temperature indicator; Schouten et al., 2002). Molecular-isotopic techniques have also been useful in studies of methane cycling. This is because, as a result of isotope effects associated with its production by either biological or thermal processes, methane is commonly depleted in 13C (Whiticar, 1994). The offset is so large that, for practical purposes, methane carbon is labeled almost as effectively as many artificial tracers. Values of ␦ for lipids from methanotrophs frequently drop below ⫺100‰ (Hinrichs et al., 2000).

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The isotopic depletion in methanotrophs is doubly interesting. It results in part from isotope effects associated with the assimilation of methane (Reeburgh et al., 1997). As kinetic effects, these are unsurprising: the product (biomass) is depleted in the heavy isotope. If the fractionation derived instead from equilibrium isotope effects, its sign would be reversed, with the biomass being enriched relative to the methane (Galimov, 1985). For the question of organic 13C control, kinetic or equilibrium, the relationship between methane and methanotrophs is, therefore, decisively informative: the control is kinetic (if equilibrium effects prevailed, the methanotrophic products would be enriched in 13C relative to the methane). For the other relationship between a C1 carbon source and biomass, namely that between CO2 and plants, the kinetic and equilibrium isotope effects both predict depletion, which is of course observed. Indecisive though it is, proponents of equilibrium and kinetic controls have both pointed to the “fit,” with outsiders and students being frustrated by the confusion. Isotopically depleted biomarkers signaling the recycling of methane turned up in some of the first sedimentary extracts examined using continuous-flow techniques (Freeman et al., 1990). It was no great surprise; the samples came from the Messel Shale, which formed from the mud of an Eocene swamp. The molecular structures were characteristic of aerobic methanotrophic bacteria, which are known to thrive in such settings. The challenge has been to find biomarkers characteristic of the microorganisms catalyzing the anaerobic oxidation of methane at the expense of sulfate. Depth profiles of the concentrations of methane and sulfate in sedimentary pore waters made it perfectly clear that this process was occurring (Iversen and Jørgensen, 1985). Success eventually came to Bian (1994), who found isotopically depleted crocetane (Fig. 13) in samples from transition-zone muds provided by Niels Iversen and Bo

Fig. 13. Depth profiles of porewater sulfate and methane, the rate of oxidation of methane, and the 13C content of the isoprenoid hydrocarbon, crocetane in cores raised from the seafloor of the Kattegat (Bian et al., 2001). Complementary depth profiles of sulfate, methane, and methane oxidation were first observed by Iversen and Jørgensen (1985). The role of crocetane as biomarker for methane-consuming anaerobes was discovered independently by Bian (1994) and by Thiel et al. (1999).

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carbon cycle over the course of earth history. The young scientists pursuing these quests are progressing rapidly. Some are certain to be recognized as fitting representatives of V. M. Goldschmidt’s scientific legacy. I do not claim to reach their level, still less Goldschmidt’s. I have, however, helped to build the interdisciplinary connections that now characterize biogeochemistry. Soon to be surpassed, I nevertheless have confidence in the value of that role.

Fig. 14. A gas chromatogram of lipid biomarkers extracted from sediments at gas seeps in the Santa Barbara Basin (Hinrichs et al., 2000). The peaks marked in light gray represent products derived exclusively from producers and consumers in the oceanic water column. Those marked in darker gray represent monoalkyl-(M) and dialkyl-(D) glycerol ethers produced by sulfate-reducing bacteria within the methane-consuming consortia. The peaks marked with a heavy black line are products of methanotrophic archaea.

Jørgensen. Crocetane and other products of anaerobic methanotrophy were discovered independently by Michaelis, Thiel, and coworkers (Thiel et al., 1999). Ether-linked lipids identical to those in some methanogens, but with isotopic compositions indicating derivation from methanotrophs, were discovered by Hinrichs, DeLong, and coworkers (Hinrichs et al., 1999). Much additional information about lipids from methanotrophic communities has now accumulated. The Hamburg group was the first to show that 13C-depleted products probably from sulfate-reducing bacteria were associated with those from the methane-consuming archaea (Thiel et al., 1999). A more specific connection to sulfate reducers was established by Pancost et al. (2000). Additional results from Hinrichs et al. (2000), which serve to summarize these findings, are shown in Figure 14. Important further reports have followed (Pancost et al., 2001a, b; Thiel et al., 2001a, b). At the time of its publication, the initial report from Hinrichs et al. (2000) was unique in its combination of molecularisotopic and genomic lines of evidence. The genomic approach was promptly extended by Boetius et al. (2000) at the Max Planck Institute for Marine Microbiology, who used fluorescently labeled oligonucleotide probes to show that archaea closely related to those found by Hinrichs et al. (1999) were closely associated with bacteria closely related to known sulfate reducers. This provided support for key aspects of the “consortium hypothesis” introduced by Hoehler et al. (1994). Further multidisciplinary investigations have confirmed that the aggregates observed by Boetius et al. (2000) are indeed methanotrophic (Orphan et al., 2001b) and revealed considerable diversity in anaerobic methanotrophic communities (Orphan et al., 2001a, 2002). These most recent studies illustrate the convergence of three lines of inquiry: studies of lipid biomarkers (isotopic and structural), microbial physiology and ecology, and microbial genomics. This synergy promises to revolutionize, first, our understanding of terminal processes in the modern carbon cycle and, ultimately, our understanding of the evolution of the

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