Cover Image


View/Hide Left Panel

Oxygen and Proterozoic Evolution: An Update


Harvard University

Complex events can rarely be reconstructed from single lines of evidence, even where the record is well preserved. Preston Cloud (1983)


Many authors, most notably Preston Cloud, have argued that major events in early evolution were coupled to changes in the oxygen content of the Precambrian atmosphere. Interest has focused on events close to 2000 million years ago (Ma) and 600 Ma, when increases in PO2 are thought to have stimulated the radiations of aerobically respiring eubacteria and (via endosymbiosis) protists, and macroscopic metazoans, respectively. Acceptance of these hypotheses requires (1) geochemical evidence of environmental change; (2) paleontological evidence of coeval evolutionary innovation; and (3) physiological, ecological, and phylogenetic reasons for linking the two records. New data from Paleoproterozoic weathering profiles are providing increasingly quantitative constraints on the timing and magnitude of an early Proterozoic PO 2 increase. The emerging environmental history correlates well with the known phylogeny of protists and is consistent with the fossil record. Quantitative data on possible Neoproterozoic PO2 changes remain elusive, but new geochemical data strongly indicate that the interval just prior to the Ediacaran radiation was a time of marked environmental change. An increase in PO2 may well have accompanied this event, providing oxygen levels sufficient to support the metabolism of large heterotrophs. Geochemical, paleontological, and biological data support the hypothesis that atmospheric composition significantly constrained and at times provided important opportunities for early biological evolution.


The coevolution of life and its environment has long been a principal theme in interpretations of Earth's early history. In particular, molecular oxygen has frequently been singled out as a major factor in Precambrian evolution, a view argued eloquently by Preston Cloud (1968a,b, 1972, 1983).

Why this coevolutionary view should be so popular is clear enough. Although the present day atmosphere con-

The National Academies | 500 Fifth St. N.W. | Washington, D.C. 20001
Copyright © National Academy of Sciences. All rights reserved.
Terms of Use and Privacy Statement

Below are the first 10 and last 10 pages of uncorrected machine-read text (when available) of this chapter, followed by the top 30 algorithmically extracted key phrases from the chapter as a whole.
Intended to provide our own search engines and external engines with highly rich, chapter-representative searchable text on the opening pages of each chapter. Because it is UNCORRECTED material, please consider the following text as a useful but insufficient proxy for the authoritative book pages.

Do not use for reproduction, copying, pasting, or reading; exclusively for search engines.

OCR for page 21
Effects of Past Global Change on Life 1 Oxygen and Proterozoic Evolution: An Update ANDREW H. KNOLL and HEINRICH D. HOLLAND Harvard University Complex events can rarely be reconstructed from single lines of evidence, even where the record is well preserved. Preston Cloud (1983) ABSTRACT Many authors, most notably Preston Cloud, have argued that major events in early evolution were coupled to changes in the oxygen content of the Precambrian atmosphere. Interest has focused on events close to 2000 million years ago (Ma) and 600 Ma, when increases in PO2 are thought to have stimulated the radiations of aerobically respiring eubacteria and (via endosymbiosis) protists, and macroscopic metazoans, respectively. Acceptance of these hypotheses requires (1) geochemical evidence of environmental change; (2) paleontological evidence of coeval evolutionary innovation; and (3) physiological, ecological, and phylogenetic reasons for linking the two records. New data from Paleoproterozoic weathering profiles are providing increasingly quantitative constraints on the timing and magnitude of an early Proterozoic PO 2 increase. The emerging environmental history correlates well with the known phylogeny of protists and is consistent with the fossil record. Quantitative data on possible Neoproterozoic PO2 changes remain elusive, but new geochemical data strongly indicate that the interval just prior to the Ediacaran radiation was a time of marked environmental change. An increase in PO2 may well have accompanied this event, providing oxygen levels sufficient to support the metabolism of large heterotrophs. Geochemical, paleontological, and biological data support the hypothesis that atmospheric composition significantly constrained and at times provided important opportunities for early biological evolution. INTRODUCTION The coevolution of life and its environment has long been a principal theme in interpretations of Earth's early history. In particular, molecular oxygen has frequently been singled out as a major factor in Precambrian evolution, a view argued eloquently by Preston Cloud (1968a,b, 1972, 1983). Why this coevolutionary view should be so popular is clear enough. Although the present day atmosphere con-

OCR for page 21
Effects of Past Global Change on Life tains 21% O2, there is widespread agreement that prior to the emergence of oxygenic cyanobacteria, PO2 must have been extremely low. Recent models by Canuto et al. (1983) and Kasting (1987) suggest that the prebiotic atmosphere contained no more than about 10-10 bar of molecular oxygen—enough to make hematite stable, but far too little to provide an effective ozone screen or to support aerobic metabolism. Indeed, a standard tenet of chemical evolution is that prebiotic chemistry could not have proceeded in environments containing significant amounts of O2. Increasingly well resolved phylogenetic trees (e.g., Woese, 1987) complement this perspective. These trees indicate that anaerobic organisms diverged earlier than aerobes, and that aerobes requiring high PO2(i.e., large animals) appeared later than aerobes able to function in less oxic environments. It is, therefore, attractive to link biological to environmental history; however, the entire pattern of biological evolution can potentially be explained quite differently. The facts noted in the previous paragraph really only require a specified initial condition: that is, that the Earth's prebiotic atmosphere was essentially anoxic and that the first organisms were, therefore, anaerobic. There is no a priori reason why an early radiation of cyanobacteria could not have engendered an early and rapid increase in PO2 approximating or even exceeding today's levels. Very different controls would then have to be sought for the observed evolutionary patterns. Acceptance of what might fairly be called the Cloud model requires that three criteria be satisfied: geochemical documentation of environmental change; independent paleontological evidence for coeval evolutionary innovation; and physiological, phylogenetic, and ecological reasons for linking criteria 1 and 2. In the following pages, we evaluate the rapidly accumulating data on oxygen and biological evolution during two intervals often inferred to have been critical junctures in the history of life: (1) early in the Proterozoic Eon (ca. 2000 Ma), when increases in PO2 above the Pasteur point are thought to have made possible the evolution of aerobic prokaryotes and mitochondria-bearing protists; and (2) the latest Proterozoic (ca. 600 Ma), when another substantial increase in PO2 may have made possible the initial evolution of macroscopic animals. THE EARLY PROTEROZOIC EON Geochemical Evidence for Atmospheric Change The Paleoproterozoic Era (2500 to 1600 Ma) was a time of profound environmental change (Cloud, 1968a, 1972: Holland, 1984). Two independent sedimentological observations have long been cited in support of the hypothesis that the atmosphere first accumulated significant amounts of oxygen during this interval. Banded iron formations (BIF), quintessentially Precambrian sediments composed of iron-bearing minerals and silica, are abundant in successions older than ca. 1900 Ma, but are rare in younger sequences (Figure 1.1). Continental red beds display an inverse distribution. The origin of marine iron FIGURE 1.1 A summary of geochemical and paleobiological data relevant to considerations  of Paleoproterozoic evolution and environmental change (PAL = present atmospheric level).

OCR for page 21
Effects of Past Global Change on Life formations probably requires anoxic mid- and deep oceans for the storage and transportation of ferrous iron, while it is likely that red beds can form only when terrestrial or nearshore marine sediments come in contact with atmospheric oxygen. Thus, it has been reasoned that the BIF-red bed transition marks the rise of atmospheric oxygen. Complementary information comes from detrital uraninite in Archean and earliest Proterozoic alluvial rocks. Because this uranium mineral can survive prolonged transport only in media containing little or no oxygen, the lack of detrital uraninite deposits younger than ca. 2300 Ma also points toward a significant environmental transition (Figure 1.1; Roscoe, 1969; Grandstaff, 1980; Holland, 1984). Not all scientists have accepted the validity of these observations or of their interpretation (see, for example, Dimroth and Kimberley, 1976; Clemmey and Badham, 1982; Windley et al., 1984). It has been argued repeatedly that at least some red beds antedate the end of BIF deposition, that Archean granites have paleoweathering profiles indicative of oxic environments, and that oxidized sulfur minerals (sulfates) occur in some of the oldest known sedimentary successions. All of these observations are correct, and we must ask whether they preclude the interpretation of Archean and earliest Proterozoic environments as oxygen poor. The answer appears to be no. The formation of red beds and oxidized weathering profiles on granitic substrates requires oxygen, but only in minute quantities (see, for example, Holland, 1984; Pinto and Holland, 1988)—considerably less than is needed for aerobic metabolism. Marine sulfate does not require free oxygen at all—H2S can be photooxidized anaerobically to SO42- by photosynthetic bacteria, while the photochemical oxidation of volcanogenic S and SO2 to sulfate was probably a steady source of oxidized sulfur in the Archean oceans (Walker, 1983). Towe (1990) has specifically argued for the development of aerobic respiration early in the Archean and, therefore, for the presence of 1 to 2% PAL (present atmospheric level) O2 in the atmosphere since that time. The possibility that oxygen levels reached this physiologically important threshold so early is not contradicted by the sparse geochemical data available for early Archean rocks (see below); however, Towe's model suffers from the absence of Archean O2 sinks other than Fe2+ . We believe that the neglect of volcanic gases in his model casts significant doubts on the validity of his analysis. Other arguments against the Cloud model posit that the geochemical indicators of low PO2 during the Archean and earliest Proterozoic could be the result of burial digenesis, which generally acts to reduce minerals. Equally, it has been argued that oxide facies iron formations are themselves diagenetic replacements of carbonates. Neither of these views can sustain critical scrutiny. Although both oxidation and reduction can occur during diagenesis, there is ample evidence that at least some detrital uraninite and most iron formations have a primary sedimentary origin. New data from paleosols add quantitative rigor to arguments for Paleoproterozoic environmental change (Holland and Zbinden, 1988; Holland et al., 1989; Holland and Beukes, 1990). All paleosols younger than 1900 Ma that have been studied to date are highly oxidized. Fe2+ in the parent rocks of these paleosols was oxidized quantitatively, or nearly so, to Fe3+, and was retained in the paleosols as a constituent of Fe3+ oxides or hydroxides. This is demonstrated by the near constancy of the ratio of total Fe to Al2O3 and of total Fe to TiO2 within these paleosols and their parent rocks. Paleosols older than 1900 Ma that were developed on basaltic rocks have lost nearly all iron from their upper sections. Some of this lost iron was reprecipitated in the lower sections of the paleosols. There is some evidence that iron loss from pre-1900 Ma paleosols developed on granitic rocks was much less pronounced. These observations suggest that the O2 content of the atmosphere prior to 1900 Ma was insufficient to oxidize more than a small fraction of the iron developed in soils on basaltic rocks, but was sufficient to oxidize a good deal of the much smaller amount of iron in soils developed on granitic rocks (Figure 1.2). More detailed studies of paleosols are needed to confirm the generality of these observations. If confirmed, they can be used to assign a rough value of ca. 1% PAL to the O2 content of the atmosphere between about 2700 and 2200 Ma (Pinto and Holland, 1988). Pre- FIGURE 1.2 Iron retention in Precambrian paleosols plotted in terms of R—the ratio of the oxygen to CO2 demand in the weathering of parent rocks—and geological age. Parent rocks with R < 0.025 are granitic; those with higher R are basaltic.

OCR for page 21
Effects of Past Global Change on Life FIGURE 1.3 The fraction of iron retained during the weathering of siderite of composition (Fe0.78Mg0.22)CO3 as a function of the initial concentration of O2 in groundwater and the CO2 pressure in the atmosphere with which the groundwater equilibrated before reacting with siderite. liminary data for the distribution of the rare earth elements (REEs) in paleosols suggest that Eu in paleosols has been present in the +3 valence state since at least 2750 Ma, and that the valence state of Ce in paleosols changed from +3 to +4 between 2750 and 1800 Ma. These results are consistent with those for Fe and indicate that the REEs may join iron as useful indicators of oxygen evolution in the Precambrian atmosphere. The data for paleosols developed on igneous rocks have been supplemented recently by information on a paleoweathering profile developed on carbonate facies Kuruman Iron Formation in Griqualand West, South Africa (Holland and Beukes, 1990). The profile was probably developed ca. 1900 Ma. It is highly oxidized, and the high degree of iron retention during weathering can be used to show that PO2 was probably in excess of 15% PAL (Figure 1.3). This is a higher minimum for PO2 than that set by the behavior of iron in paleosols developed in igneous rocks, and indicates that the O2 content of the atmosphere rose from about 1 to >15% PAL between 2200 and 1900 Ma (Figure 1.1). The transition inferred from paleosol data is consistent with that inferred from the time distribution of iron formations and postdates the last known occurrence of detrital uraninite ores by several hundred million years (Knoll, 1979; Walker et al., 1983). Isotopically very light organic matter in late Archean and earliest Proterozoic sedimentary rocks has also been interpreted in terms of an early appearance of environments capable of sustaining aerobic metabolism, at least locally (Hayes, 1983). These data suggest that PO2 may have increased in at least two steps: an initial rise from extremely low oxygen tensions to levels about 1 to 2% PAL, and a later increase to levels >15% PAL approximately 2100 Ma (see also Walker et al., 1983). Why oxygen levels should have increased in this manner is not clear. The origin of oxygenic cyanobacteria is poorly constrained in time, but it certainly occurred before 2100 Ma. Fossils morphologically diagnostic for the group are known only from about 2000 Ma (Golubic and Hofmann, 1976), but plausibly cyanobacterial remains have been found in early Archean cherts (e.g., Schopf and Packer, 1987). Buick (1992) has argued on sedimentological and geochemical grounds that stromatolites in lacustrine carbonates of the 2800 Ma Fortescue Group, Australia, must have been built by oxygenic photoautotrophs. If Hayes' interpretation of the carbon isotope record is correct, cyanobacteria radiated 2800 Ma or earlier. The pre-2800 Ma sedimentary record has been sampled too poorly to establish whether anomalously light carbon was widespread in early Archean lacustrine environments. Increases in atmospheric oxygen were probably occasioned by increases in primary productivity and/or decreased rates of oxygen consumption. The increase from very low O2 levels to 1 to 2% PAL may have been related to productivity increases associated with rapid continental growth and stabilization during the late Archean/earliest Proterozoic (Knoll, 1979, 1984; Cameron, 1983). In contrast, the later increase to >15% PAL does not seem to be related to a major tectonic event. The high oxygen level in today's atmosphere must be related to the role of PO2 in maintaining the redox balance of the atmosphere-biosphere-ocean-lithosphere system. However, the nature of the connection is still in dispute. Atmospheric PO2 determines the concentration of O2 in surface ocean water, but the influence of the O2 concentration in seawater on the burial efficiency of organic matter within marine sediments seems to be slight (see, for instance, Betts and Holland, 1991). Nutrients are a more likely link between PO2 and the burial rate of organic matter, and hence between PO2 and rates of long-term O2 generation. A plausible argument can be made that links the marine geochemistry of PO43- to that of iron and hence to the O2 content of the atmosphere today. If this argument turns out to be valid, then the history of atmospheric O2 may have been controlled by a complicated feedback system involving the marine geochemistry of iron and phosphorus. The rapid increase in PO2 ca. 2100 Ma may have marked the passage of the system across a threshold from one steady state to another. Paleontological Evidence for Evolutionary Innovation At first glance, the fossil record appears to provide strong support for the linkage of environmental and biological evolution. The oldest known fossils of probable eukaryotic origin are spirally coiled, megascopic remains

OCR for page 21
Effects of Past Global Change on Life in 2100 Ma shales from Michigan (Han and Runnegar, 1992). Microfossils of probable eukaryotic origin first become widespread in rocks 1800 to 1600 Ma (Figure 1.4A-D), and molecular biomarkers for eukaryotes are similarly known from rocks =1760 Ma (Summons and Walter, 1990). Unfortunately, paleobiological documentation of Paleoproterozoic evolutionary change is hampered by a serious problem. At about 1800 Ma the fossil record improves markedly (e.g., Schopf, 1983), so it is not clear that paleontological first appearances necessarily record evolutionary innovations. Paleontological evidence is certainly consistent with a model of linked early Proterozoic environmental and biological evolution, but at the present time, fossils do not provide strong, independent confirmation of such a linkage. Documentation of microfossil assemblages from a number of pre-2100 Ma localities representing diverse paleoenvironments is needed to strengthen or reject the conclusion that fossilizable protists radiated about the time the deposition of iron formations ceased. Biological Reasons for Linkage Bearing in mind the unsatisfactory state of paleontological evidence, let us ask why evolutionary change might have attended the atmospheric transitions of the Paleoproterozoic Era. At oxygen levels less than about 1% PAL (a relatively poorly defined number; see Schopf, 1983), FIGURE 1.4 A: Probable eukaryotic microfossils from the Mesoproterozoic Roper Group, northern Australia (bar = 50 µm); B: a weakly ornamented protistan cyst from the Neoproterozoic Visingsö Beds, Sweden (bar = 25 µm); C: a vase-shaped protistan microfossil from the Neoproterozoic Elbobreen Formation, Spitsbergen (bar = 50 µm); D: large process-bearing protistan microfossils from the Neoproterozoic Draken Conglomerate Formation, Spitsbergen (bar = 200 µm);, E: Ediacaran metazoan from the White Sea, USSR (bar = 1.5 mm).

OCR for page 21
Effects of Past Global Change on Life aerobic metabolism is impossible, and there is limited atmospheric protection against ultraviolet (UV) radiation that destroys DNA (Kasting, 1987). The rate of nitrate production in the atmosphere by lightning may well have been an order of magnitude lower than today (Yung and McElroy, 1979; Levine et al., 1982); however, the difference between the present-day rate of nitrate production (Borucki and Chameides, 1984) and the rate 2500 to 2000 Ma depends on the CO2 pressure in the Paleoproterozoic atmosphere. Kasting (1990) proposed that at a PCO2 of 0.2 atmosphere (atm), the NO production rate in the absence of O2 is only about a factor of two lower than at present. At a CO 2 pressure of 0.02 atm (i.e., 60 PAL), the NO fixation rate would probably be only one-tenth of the present rate. Denitrification was probably more widespread and intense in a low-O2 ocean than in the present ocean. Therefore, NO3- was almost certainly in very short supply prior to 2100 Ma, and biological N2 fixation must have been the principal source of usable nitrogen for primary producers. H2O 2 may have been an important oxidant on the early anoxic Earth (Kasting et al., 1987); it is possible that biochemical defenses against molecules generally regarded as reactive intermediates in oxygen biochemistry evolved before O2 itself became a significant constituent of the atmosphere (McKay and Hartman, 1991). The conditions described in the previous paragraph certainly apply to the biota that existed before the evolution of cyanobacterial photosynthesis. How long such conditions persisted after the advent of oxygenic photosynthesis is unclear. As noted above, the antiquity of oxygenic cyanobacteria is poorly constrained, although it could easily be as great as 3500 Ma, the age of the oldest negligibly metamorphosed sedimentary rocks (Knoll, 1979; Schopf and Packer, 1987). As noted previously, a PO2 of approximately 1 to 2% PAL appears likely for an extended period prior to 2100 Ma (see also Towe, 1990). This is an oxygen level of both biological and environmental significance. At about 1% PAL, aerobic metabolism by single-celled organisms becomes possible, while an effective ozone screen expands the ecological possibilities of life. When PO2 rose to 1 to 2% PAL, aerobic metabolism probably followed quickly in organisms already protected against oxygen toxicity. In particular, bacteria capable of aerobic respiration, with its tremendous energetic advantage over fermentation, probably radiated rapidly (and polyphyletically) from photosynthetic ancestors. Flavin-based oxygen-utilizing pathways evolved in archaebacteria and in amitochondrial eukaryotes. At this O2 level, nitrate production levels in the atmosphere may well have been significantly lower than today's (see above). Nitrogen fixers, therefore, could have retained a considerable advantage in primary production. Ancestral eukaryotes formed endosymbiotic associations with purple bacterial aerobes, gaining the benefits of aerobic respiration. Symbioses with photosynthetic prokaryotes may not have formed concurrently, however. Nitrogen fixation is unknown in plastids and appears to be prohibited (Postgate and Eady, 1988); the reasons for this are not clear, but may involve oxygen toxicity. Although some early algae might have obtained nitrogen heterotrophically, obligately photosynthetic eukaryotes (including all extant megascopic algae) are unlikely to have occurred in the absence of significant quantities of nitrate in the environment. These considerations suggest a different biological focus for the 2100 Ma oxygen event. It is not that fundamentally new metabolisms were made possible, but rather that as oxygen increased to levels above 10% PAL, nitrate availability may well have increased dramatically (see above). Obligately photosynthetic eukaryotes would then have become feasible. With their ability to avoid formation of nutrient-depleted boundary layers adjacent to cells, eukaryotic primary producers would soon have become ecologically important as primary producers. Thus, it is not surprising that 2100 Ma shales contain megascopic algae or that slightly younger rocks contain abundant acritarchs whose morphology and distribution are similar to those of younger eukaryotic phytoplankton. How does this environmental scenario compare with the known phylogeny of eukaryotes? Figure 1.5 shows evolutionary relationships among living eukaryotes as determined by Sogin et al. (1989). At the base of the tree is Giardia, a common pathogen in the digestive system of vertebrates. Biochemically, Giardia shares more features with prokaryotes than any other known eukaryote. Ultrastructurally, however, it is clearly a true eukaryote; it contains a membrane-bounded nucleus, undulipodia (9+2 flagella), and a cytoskeleton (albeit a biochemically very simple one). On the other hand, Giardia has no mitochondria and no well-developed ER or Golgi apparatus. These organisms are heterotrophic, engulfing particulate food (phagocytosis) and absorbing dissolved organic molecules. Food is metabolized by the classic Embden-Meyerhoff pathway of glycolysis. Giardia cells are not capable of classical aerobiosis, but can use oxygen as a terminal acceptor of reducing equivalents. This system uses flavins and iron-sulfur proteins, and does not include cytochromes; it appears that the cells derive little energetic benefit from this reaction (Müller, 1988). The next branches in Figure 1.5 are occupied by the microsporidia, trichomonads, and related protists. Both groups have clearly become specialized as obligate parasites (microsporidians are apparently dependent on an external source of ATP), but they retain features that complement the picture of early eukaryotes developed from Giardia.

OCR for page 21
Effects of Past Global Change on Life Figure 1.5 Summary of eukaryotic phylogeny as determined by comparisons of small subunit ribosomal RNA sequences (redrawn as a Hennigian comb from Sogin et al., 1989). Microsporidians have prokaryote-like features of ribosomal organization and lack mitochondria; unlike Giardia they have a well-developed endomembrane system. Trichomonads also lack mitochondria, and free-living species live as aerotolerant anaerobic heterotrophs (Margulis et al., 1989). Some contain small organelles called hydrogenosomes, which are thought to be anaerobic equivalents of mitochondria. Others harbor intracellular bacterial symbionts that confer specific metabolic capabilities such as cellulose catabolism (Müller, 1988). To date, RNA sequence data are available for relatively few protists, and it is entirely possible that organisms will be recognized that branch even earlier than Giardia. Nonetheless, the significant features shared by Giardia, microsporidians, and trichomonads suggest that these organisms can provide important clues to the nature of the earliest eukaryotes. They appear to have been anaerobic but aerotolerant heterotrophs, motile (using typically eukaryotic undulipodia), and endowed with a cytoskeleton and membrane system capable of endocytosis. Whether or not the first eukaryotes could have lived in the essentially O2-free environments in which life began is uncertain. Giardia is unable to synthesize most of its lipids and so must incorporate lipids from its environment (Jarroll et al., 1989). The growth environment of modern Giardia is the small intestine of vertebrate hosts, and the lipids taken up from this environment include sterols. This raises an important issue, because sterol synthesis requires molecular oxygen at concentrations of ca. =0.2% PAL (Chapman and Schopf, 1983). If sterol synthesis is a primitive feature of eukaryotes subsequently lost by Giardia, the eukaryotic cell could not have arisen until at least low levels of oxygen had accumulated in the atmosphere. This requires that the origin of cyanobacteria predate the divergence of eukaryotes, a scenario of rapid early diversification consistent with recent phylogenies that root the universal tree between the eubacteria and an archaebacterial/eukaryote clade (Iwabe et al., 1989; Woese et al., 1990), but not with those in which all three kingdoms are viewed as diverging from a simple common ancestor (Woese, 1987). Alternatively, sterol synthesis could be a later innovation of eukaryotes, with sterol incorporation by Giardia being a relatively recent phenomenon, perhaps related to its specialized habitat. The important point is that regardless of the phylogeny preferred, organisms such as Giardia and trichomonads could have existed in Archean environments containing significantly less than 1 to 2% PAL PO2. As PO2 rose to the 1 to 2% PAL level, their aerotolerance and ability to phagocytize and maintain intracellular symbionts would have positioned them well for continued evolution. All branches above the level of trichomonads are occupied principally by mitochondria-containing organisms. The kinetoplastids include both free-living bacteriophagous forms and obligate parasites (the notorious trypanosomes).

OCR for page 21
Effects of Past Global Change on Life Often coprozoic, they are in general most common in organic-rich environments of low oxygen content (Margulis et al., 1989). Kinetoplastids are closely related to a much better known group, the euglenids. Euglenids are commonly characterized by the green photosynthetic protist Euglena; however, most organisms in this group are heterotrophic, and there is reason to believe that the euglenid acquisition of plastids was a relatively recent event involving the incorporation of a green algal symbiont or chloroplast (Gibbs, 1981; Whatley, 1981). Therefore, for the purposes of this argument, undue weight should not be accorded to the euglenid plastid. Kinetoplastids and euglenids both provide a perspective on early mitochondria-bearing eukaryotes as aerobic, bacteriophagous heterotrophs capable of engulfing potential symbionts and able to thrive at relatively low PO2. Such organisms are plausible candidates for the types of eukaryotes that might have radiated during the period when PO2 stood at 1 to 2% PAL. Eukaryotes may have gained ecological prominence as micropredators and scavengers well before they became important as photoautotrophs. Other sequenced organisms that branch earlier than the main algal limbs include amebomastigotes (common soil and water organisms that live as flagellates under low-nutrient conditions but as amebas in nutrient-rich environments); entamebas (often parasitic, some without mitochondria, others with bacterial symbionts); and cellular slime molds. The "crown" of the eukaryote tree is studded with photosynthetic members (Sogin et al., 1989). As argued by Cavalier-Smith (1987), there is no ultrastructural reason why organisms capable of engulfing mitochondrial precursors could not also have incorporated cyanobacteria. The diversity of photosynthetic eukaryotes certainly indicates that once plastid acquisition became feasible, a number of protistan lineages acquired them. The barrier to protoplastid acquisition could have been environmental, and the relatively late rise of PO2 from 1 to 2% PAL to >15% PAL provides a plausible explanation for the observed phylogeny. Until this increase in PO2, the availability of NO3- may have been severely limited, giving ecological advantage to free-living, nitrogen-fixing cyanobacteria. Once oxygen levels increased, however, nitrogenase activity was inhibited (Towe, 1985) and odd nitrogen availability increased by more than an order of magnitude. Eukaryotic phytoplankton and benthos could radiate to become important parts of most surficial ecosystems, except for stressed environments such as the upper intertidal zone of restricted seaways—environments well represented in the Proterozoic fossil records. Surprisingly, increasing paleontological data suggest that the "big bang" of higher eukaryotic evolution did not occur until 1200 to 1000 Ma (Knoll, 1992b). Earlier algae, including those that formed the 2100 Ma fossils, apparently belonged to extinct lineages. Summary of the Paleoproterozoic Earth Given the requirement that three independent criteria must be satisfied, we cannot unequivocally accept the Cloud hypothesis as it relates to early Proterozoic evolution. The geochemical and paleontological record is improving rapidly. It now appears to fit well with the molecular phylogenetic record, and it is at least consistent with paleontological observation. A modified Cloud model provides the best framework for the available biological and geochemical data. THE END OF THE PROTEROZOIC EON In the Cloud model, the other principal period of linked environmental and biological evolution is the end of the Proterozoic Eon, when further increases in PO2 are thought to have allowed the evolution of macroscopic animals. Here the relative strengths of the three lines of evidence are reversed. The paleontological data are quite extensive; it is the geochemical evidence that is consistent and suggestive rather than compelling. Details of latest Proterozoic Earth history are presented in Knoll (1992a) and Derry et al. (1992); therefore, only a brief synopsis is provided here. Paleontological Data The radiation of macroscopic animals was the cardinal evolutionary event of the Neoproterozoic Era. In 1968, Cloud argued that no unequivocal animal remains are present in rocks older than the great Varangian (ca. 610 to 590 Ma) ice age, and in the ensuing 20 years, a great deal of detailed stratigraphic research has strengthened this conclusion. Hofmann et al. (1990) have reported small, simple disks of probable metazoan origin in immediately subVaranger strata from northwestern Canada, but macroscopic animal remains and traces are otherwise conspicuously absent from pre-Varanger successions. On six continents, large and diverse, but structurally simple, animals—the so-called Ediacaran fauna (Figure 1.4C)—first appear in strata that lie above Varangian glaciogenic rocks (Runnegar, 1982a; Glaessner, 1984). Metazoan trace fossils have a parallel history of appearance and diversification (Crimes, 1987). It is important to note the terms "macroscopic" and "large." Ancestral microscopic metazoans may well have evolved significantly earlier, but left no fossil record. Indeed, it is biologically appealing to posit some sort of metazoan prehistory. The significant point, however, is

OCR for page 21
Effects of Past Global Change on Life that the fossil record of macroscopic animals begins about 600 Ma. Biological Reasons for Linkage to Environmental Change Early hypotheses linking metazoan evolution to increases in atmospheric oxygen stressed the ozone screen and its consequences for UV absorption (e.g., Nursall, 1959: Berkner and Marshall, 1965; Cloud, 1968b). Such scenarios now seem unlikely in that an essentially complete ozone screen was probably in place by the time PO2 reached 1% PAL (Kasting, 1987), a threshold attained long before 600 Ma (see above). Nonetheless, oxygen is important to metazoan evolution for physiological reasons (Raff and Raff, 1970; Towe, 1970; Runnegar, 1982b). Although tiny animals, perhaps ecologically similar to extant millimeter-scale nematodes and other meiofauna, can live at a PO2 of a few percent PAL, macroscopic animals require higher oxygen concentrations to ensure the oxygenation of multiple cell layers. Large animals also require a higher Poitou support collagen manufacture, exercise metabolism, and organismic function within skeletons. Values of 6 to 10% PAL appear to be minimum values for the support of large, unskeletonized animals that have a circulatory system or, in the case of coelenterates, a system of finely divided mesentery folds (essentially thin, flat animals folded into a three-dimensional architecture; Raff and Raff, 1970; Runnegar, 1982b). Much higher oxygen concentrations, approaching present-day levels, are necessary to support macroscopic animals without circulatory systems (Runnegar, 1982b). Thus, there are two potential ways in which increasing PO2 might correlate with metazoan evolution. If tiny ur-metazoans developed circulatory systems, then a PO2 increase to more than 6 to 10% PAL would remove this environmental barrier against the evolution of macroscopic animals. On the other hand, if macroscopic size preceded the efficient internal circulation of fluids, PO2 increases to nearly modern levels would be necessary for large animal metabolism. Most discussions of oxygen and early animal evolution have tacitly or explicitly assumed the former case (e.g., Cloud, 1968b, 1976; Towe, 1970; Runnegar, 1982a); however, this is by no means demonstrated. It is not at all clear that Ediacaran animals had well-developed circulatory systems or finely divided mesenteries. Although many present-day animals live in relatively oxygen-poor waters (e.g., Thompson et al., 1986), this may tell us little about the ancestral habitat of macroscopic animals. Once circulatory and respiratory systems were invented, large animals would have been able to inhabit oxygen-poor environments previously closed to them. Geochemical Data In the preceding section it is argued that oxygen levels greater than 6 to 10% PAL may have been necessary for the evolution of large animals, but we argue earlier that PO2 probably equaled or exceeded 15% PAL as early as 2100 Ma. We, therefore, have three choices. We can reject the Cloud hypothesis; we can suggest that oxygen levels actually decreased during the Meso- or early Neoproterozoic; or we can suggest that the first macroscopic animals did not possess well developed circulatory systems and, therefore, required oxygen levels substantially greater than 15% PAL. It is certainly premature to choose the first option, and we know little about the history of atmospheric oxygen between 1900 and 600 Ma or the internal architecture of Ediacaran animals. There is no a priori reason to believe that the secular curve for atmospheric oxygen has been monotonic, although it has often been drawn that way. In the absence of convincing data, we can ask whether or not the geological record contains any evidence suggestive of immediately pre-Ediacaran environmental change. The answer is clearly yes. At least four independent lines of evidence indicate that the period immediately preceding the Ediacaran radiation was a highly distinctive epoch in Earth history, certainly involving marked environmental change and plausibly including biologically significant variations in PO2(Figure 1.6; Derry et al., 1992; Knoll, 1992a). The distinctive nature of ca. 850-600 Ma sedimentary successions is clearly seen in two lithologies. After a hiatus of more than 1000 million years (m.y.), iron formations—some of them thick and laterally extensive—reappear on five continents (Young, 1976). It is difficult to envision their formation in the absence of extensive deep ocean anoxia. In general, Neoproterozoic iron formations are associated with the other distinctive lithologies of this period: tillites and related glaciogenic rocks. At least four ice ages punctuated late Proterozoic history; the Sturtian and Varanger glaciations were arguably the most severe in our planet's history (Hambrey and Harland, 1985; Kirschvink, 1992). Independent evidence of environmental change comes from the isotopic composition of Neoproterozoic sedimentary rocks. Over the past few years, stratigraphic variations in the marine Neoproterozoic carbon, sulfur, and strontium isotopic records have been detailed. Carbon in ca. 850 to 600 Ma carbonates and organic matter is isotopically unusual in two respects—these materials are often anomalously enriched in 13C (d13C >+5%o PDB), and within this interval there are several negative d13C excursions of 6 to 8%o, at least in part associated stratigraphically with glaciogenic rocks (Knoll et al., 1986; Kaufman et al.,

OCR for page 21
Effects of Past Global Change on Life Figure 1.6 Summary of geochemical and paleobiological data relevant to considerations  of Neoproterozoic evolution and environmental change. Filled triangles indicate ice  ages; Fe indicates iron formation deposition; the cross marks a major extinction of  large, morphologically complex protistan fossils; asterisks indicate present-day values  for the carbon isotopic composition of diagenetically stabilized carbonates (left  margin) and the strontium isotopic composition of seawater (right margin). 1990). During this same interval, the 87Sr/86Sr ratio of carbonates is anomalously low (∆87Sr as low as -500; Veizer et al., 1983; Derry et al., 1989; Asmeron et al., 1991). The positive carbon isotope anomalies indicate a substantial increase in the proportional rate of organic carbon burial in the Neoproterozoic oceans. Insofar as oxygen generation is dependent on the burial of photosynthetically produced organic matter, this may signal a significant Neoproterozoic oxygen increase. This is not necessarily the case, however, because increased introduction of H2, CO, and/or other reduced materials at midocean ridges and/or from terrestrial volcanoes could have balanced the oxygen generated by the burial of excess organic carbon. Interpretation of the Sr isotopic composition of seawater remains problematic; however, the extremely low ∆87Sr values of carbonates that also exhibit anomalous 13C enrichment may well indicate a high rate of hydrothermal input into seawater during this interval. Recent models of Neoproterozoic environmental change by Knoll and Walker (1990) and Derry et al. (1992), although differing in assumptions and procedures, both suggest that during the ca. 850 to 600 Ma interval of unusual carbon and strontium isotopic signatures, PO2 remained relatively low. Both models further suggest that during latest Proterozoic time, when Sr isotopic ratios in seawater increased from their lowest to nearly their highest values in the past 1000 m.y. PO2 may have increased significantly. The relatively rapid change in the isotopic composition of Sr occurred just prior to the Ediacaran radiation. The Knoll/Walker and Derry et al. models indicate that a latest Proterozoic PO2 increase is plausible, but not that it is proven. There are as yet no direct quantitative data indicating a change in PO2. One indirect line of evidence that supports the idea of latest Proterozoic oxygen increase in the sulfur isotopic record. The sulfur isotopic composition of 850 to 600 Ma sulfates does not move antithetically to the carbon curve, as during much of the Phanerozoic. Antithetic movement was established in latest Proterozoic times—during the brief but eventful interval when the isotopic composition of Sr in seawater shifted; Ediacaran metazoans radiated; and intriguingly, most of the large, morphologically complex protists that characterize the 850 to 600 Ma microplankton record disappeared (Figure 1.4D; Knoll and Butterfield, 1989; Zang and Walter, 1989). The substantial shift in the isotopic composition of marine sulfate recorded at this time indicates a marked shift of the sedimentary sulfur reservoir toward pyrite (Claypool et al., 1980; François and Gerard, 1986)—a shift that contributed further to the production rate of O2. Summary of the Latest Proterozoic Record The Cloud model links the diversification of macroscopic animals to new evolutionary opportunities attendant on increasing PO2. Certainly, there can no longer be any doubt that the period immediately prior to the Ediacaran radiation was a time of marked environmental fluctuation. This may have included a PO2 increase, although quantification of Neoproterozoic oxygen levels and even the O2 levels at which Ediacaran-grade animals were able to function remains uncertain. As in the case of Paleoproterozoic atmospheric change and evolution, we can claim to have only two of the three required pieces of the puzzle in place; however, specific attention can now be focused on paleosols and other features of the Neoproterozoic rock record in a concentrated effort to understand the pattern of atmospheric change at or just before the first appearance of macroscopic metazoans. CONCLUSIONS In the Phanerozoic geological record, environmental change is often associated with extinction. Its link to

OCR for page 21
Effects of Past Global Change on Life evolutionary innovation is usually indirect and depends on the removal of pre-existing ecological dominants to provide evolutionary opportunity. The thesis evaluated in this chapter is that the Archean and the Proterozoic Earth were different—that major environmental changes early in Earth history directly facilitated evolutionary innovation. The ecological specificity of many eubacteria and archaebacteria indicates that at some level this must surely be true for prokaryotic organisms (e.g., Knoll and Bauld, 1989). The Cloud model suggests that environmental-biological coevolution also applies to fundamental aspects of eukaryotic evolution, specifically to the profoundly important radiations of aerobic protists and animals. Proof of the relationship remains elusive, but accumulating evidence lends new support to the model's basic tenets. Continued research is needed to strengthen the paleontological and geochemical bases on which the empirical evidence for Proterozoic evolution and environmental change rests. Specifically, paleontological and organic geochemical research on Paleoproterozoic shales and other subtidal facies is needed to document in a more satisfactory fashion the early fossil record of eukaryotic photoautotrophs. Geochemical research on Neoproterozoic paleosols and additional indicators of PO2 are required to document the possible role of changing oxygen concentrations in latest Proterozoic environmental and evolutionary events. These outstanding questions provide an agenda by means of which we may finally be able to document what many scientists have long believed—that two of the most significant radiations in the history of life are linked closely to secular variations in atmospheric oxygen. ACKNOWLEDGMENTS We acknowledge our deep debt to Preston Cloud for his articulation of fundamental problems in biological and environmental history. We thank Joseph Montoya, Mitchell Sogin, and John Postgate for useful discussions and advice. James Kasting and Kenneth Towe provided helpful reviews of the manuscript. Our research on problems of Precambrian paleontology and geochemistry is supported by NSF Grant BSR 88-17662 and NASA Grants NAGW893 (A.H.K.) and NAGW-599 (H.D.H.). REFERENCES Asmeron, Y., S. Jacobesen, and A. H. Knoll (1991). Sr isotope variations in Late Proterozoic sea water: Implications for crustal evolution, Geochimica et Cosmochimica Acta 55, 2883-2894. Berkner, L. V., and L. C. Marshall (1965). On the origin and rise of oxygen concentration in the Earth's atmosphere, Journal of Atmospheric Science 22, 225-261. Betts, J. N., and H. D. Holland (1991). The oxygen content of ocean bottom waters, the burial efficiency of organic carbon, and the regulation of atmospheric oxygen, Palaeogeography, Palaeoclimatology, Palaeoecology 97, 5-18. Borucki, W. J., and W. L. Chameides (1984). Lightning: Estimates of the rates of energy dissipation and nitrogen fixation, Reviews of Geophysics and Space Physics 22, 363-372. Buick, R. (1992). The antiquity of oxygenic photosynthesis: Evidence from stromatolites in sulphate-deficient Archaean lakes, Science 255, 74-77. Cameron, E. M. (1983). Sulphate and sulphate reduction in early Precambrian oceans, Nature 296, 145-148. Canuto, V. M., J. S. Levine, T. R. Augustsson, C. L. Imhoff, and M. S. Giampapa (1983). UV radiation from the young Sun and oxygen and ozone levels in the prebiological atmosphere, Nature 296, 816-820. Cavalier-Smith, T. (1987). The simultaneous origin of mitochondria, chloroplasts, and microtubules, Annals of the New York Academy of Sciences 503, 55-71. Chapman, D. J., and J. W. Schopf (1983). Biological and biochemical effects of the development of an aerobic environment, in Earth's Earliest Biosphere: Its Origin and Evolution, J. W. Schopf, ed., Princeton University Press, Princeton, N.J., pp. 302-320. Claypool, G. E., W. T. Holser, I. R. Kaplan, H. Sakai, and I. Zak (1980). The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation, Chemical Geology 28, 199-260. Clemmey, H., and N. Badham (1982). Oxygen in the Precambrian atmosphere: An evaluation of geological evidence, Geology 10, 141-146. Cloud, P. (1968a). Atmospheric and hydrospheric evolution on the primitive Earth, Science 160, 729-736. Cloud, P. (1968b). Pre-metazoan evolution and the origins of the metazoa, in Evolution and Environment, T. Drake, ed., Yale University Press, New Haven, Conn., pp. 1-72. Cloud, P. (1972). A working model of the primitive Earth, American Journal of Science 272, 537-548. Cloud, P. (1976). The beginnings of biospheric evolution and their biogeochemical consequences, Paleobiology 2, 351-387. Cloud, P. (1983). Banded iron-formation—A gradualist's dilemma, in Iron-Formation: Facts and Problems, A.F. Trendall and R. C. Morris, ed., Elsevier, Amsterdam, pp. 401-416. Crimes, T. P. (1987). Trace fossils and correlation of late Precambrian and early Cambrian strata, Geological Magazine 124, 97-119. Derry, L., L. S. Keto, S. Jacobsen, A. H. Knoll, and K. Swett (1989). Sr isotopic variations of Upper Proterozoic carbonates from East Greenland and Svalbard, Geochimica Cosmochimica Acta 53, 2331-2339. Derry, L., A. J. Kaufman, and S. Jacobsen (1992). Sedimentary cycling and environmental change in the Late Proterozoic: Evidence from stable and radiogenic isotopes, Geochimica et Cosmochimica Acta 56, 1317-1329. Dimroth, E., and M. M. Kimberley (1976). Precambrian atmospheric oxygen: Evidence in the sedimentary distributions of carbon, sulfur, uranium, and iron, Canadian Journal of Earth Sciences 13, 1161-1185. François, L. M., and J.-C. Gerard (1986). A numerical model of the evolution of ocean sulfate and sedimentary sulfur during

OCR for page 21
Effects of Past Global Change on Life the last 800 million years, Geochimica Cosmochimica Acta 50, 2289-2302. Gibbs, S. P. (1981). The chloroplasts of some algal groups may have evolved from endosymbiotic eukaryotic algae, Annals of the New York Academy of Sciences 361, 193-208. Glaessner, M. (1984). The Dawn of Animal Life, Cambridge University Press, Cambridge,244 pp. Golubic, S., and H. J. Hofmann (1976). Comparison of modern and mid-Precambrian Entophysalidaceae (Cyanophyta) in stromatolitic algal mats: Cell division and degradation, Journal of Paleontology 50, 1074-1082. Grandstaff, D. E. (1980). Origin of uraniferous conglomerates at Elliot Lake, Canada, and Witwatersrand, South Africa: Implications for oxygen in the Precambrian atmosphere, Precambrian Research 13, 1-26. Hambrey, M. B., and W. B. Harland (1985). The Late Proterozoic glacial era, Paleogeography, Palaeoclimatology, Palaeoecology 51, 255-272. Han, T.-M., and B. Runnegar (1992). Megascopic eukaryotic algae from the 2.1 billion-year-old Negounee iron formation, Michigan, Science 257, 232-235. Hayes, J. M. (1983). Geochemical evidence bearing on the origin of aerobiosis: A speculative hypothesis , Earth's Earliest Biosphere: Its Origin and Evolution, J. W. Schopf, ed., Princeton University Press, Princeton, N.J., pp. 291-301. Hofmann, H. J., G. M. Narbonne, and J. D. Aitken (1990). Ediacaran remains from intertillite beds in northwestern Canada, Geology 18, 1199-1202. Holland, H. D. (1984). The Chemical Evolution of the Atmosphere and Oceans, Princeton University Press, Princeton, N.J., 582 pp. Holland, H. D., and N. J. Beukes (1990). A paleoweathering profile from Griqualand West, South Africa: Evidence for a dramatic rise in atmospheric oxygen between 2.2 and 1.9 BYBP, American Journal of Science 290A, 1-34. Holland, H. D., and E. A. Zbinden (1988). Paleosols and the evolution of the atmosphere, Part I, in Physical and Chemical Weathering in Geochemical Cycles, A. Lerman and M. Meybeck, eds., Kluwer Academic Publishers, Dordrecht, pp. 61-82. Holland, H. D., C. R. Feakes, and E. A. Zbinden (1989). The Flin Flon paleosol and the composition of the atmosphere 1.8 BYBP, American Journal of Sciences 289, 362-389 Iwabe, N., K. Kuma, M. Hasegawa, S. Osawa, and T. Miyata (1989). Evolutionary relationship of archaebacteria, eubacteria, and eukaryotes inferred from phylogenetic trees of duplicated genes, Proceedings of the National Academy of Sciences USA 86, 9355-9359. Jarroll, E. L., P. Manning, A. Berrada, D. Hare, and D. G. Lindmark (1989). Biochemistry and metabolism of Giardia, Journal of Protozoology 26, 190-197. Kasting, J. F. (1987). Theoretical constraints on oxygen and carbon dioxide concentrations in the Precambrian atmosphere, Precambrian Research 34, 205-229. Kasting, J. F. (1990). Bolide impacts and the oxidation state of carbon in the Earth's early atmosphere, Origins of Life 20, 199-231. Kaufman, A. J., J. M Hayes, A. H. Knoll, and G. J. B. Germs (1990). Isotopic compositions of carbonates and organic carbon from Upper Proterozoic successions in Namibia: Stratigraphic variation and the effects of diagenesis and metamorphism, Precambrian Research 49, 301-327. Kirschvink, J. L. (1992). Late Proterozoic low latitude global glaciation: The snowball Earth, in The Proterozoic Biosphere, J. W. Schopf and C. Klein, eds., Cambridge University Press, Cambridge,pp. 51-57. Knoll, A. H. (1979). Archean photoautotrophy: Some limits and alternatives, Origins of Life 9, 313-327 Knoll, A. H. (1984). The Archean/Proterozoic transition: A sedimentary and paleobiological perspective, in Patterns of Change in Earth Evolution, H. D. Holland and A. F. Trendall, eds., Springer-Verlag, Berlin, pp. 221-242. Knoll, A. H. (1992a). Biological and biogeochemical preludes to the Ediacaran radiation, in Origins and Early Evolutionary History of the Metazoa, J. H. Lipps and P. W. Signor, eds., Plenum, New York, pp. 53-84. Knoll, A. H. (1992b). The early evolution of eukaryotes: A global perspective, Science 256, 622-627. Knoll, A. H., and J. Bauld (1989). The evolution of ecological tolerance in prokaryotes, Transactions Royal Society of Edinburgh, Earth Science 80, 209-223. Knoll, A. H., and N. J. Butterfield (1989). New window on Proterozoic life, Nature 337, 602-603. Knoll, A. H., and J. C. G. Walker (1990). The environmental context of early metazoan evolution, Geological Society of America, Abstracts with Program 22(7), A128. Knoll, A. H., J. M. Hayes, A. J. Kaufman, K. Swett, and I. M. Lambert (1986). Secular variation in carbon isotope ratios from Upper Proterozoic successions of Svalbard and East Greenland, Nature 321, 832-838. Levine, J. S., G. L. Gregory, G. A. Harvey, W. E. Howell, W. J. Borucki, and R. E. Orville (1982). Production of nitric oxide by lightning on Venus, Geophysical Research Letters 9, 893-896. Margulis, M., J. O. Corliss, M. Melkonian, and D. I. Chapman, eds. (1989). Handbook of Protoctista, Jones and Bartlett, Boston, 914 pp. McKay, C. P., and H. Hartman (1991). Hydrogen peroxide and the evolution of oxygenic photosynthesis, Origins of Life 21, 157-164. Müller, M. (1988). Energy metabolism of protozoa without mitochondria, Annual Reviews of Microbiology 42, 465-488. Nursall, J. R. (1959). Oxygen as a prerequisite for the origin of the metazoa, Nature 183, 1170-1171. Pinto, J. P., and H. D. Holland (1988). Paleosols and the evolution of the atmosphere, Part II, in Paleosols and Weathering Through Geologic Time, J. Reinhardt and W. R. Sigleo, eds., Geological Society America Special Paper 216, 21-34. Postgate, J. R., and R. R. Eady (1988). The evolution of biological nitrogen fixation, in Nitrogen Fixation: Hundred Years After, H. Bothe, M. de Bruijn, and W. E. Newton, eds., Gustav Fischer, Stuttgart, pp. 31-39. Raff, R. A., and E. C. Raff (1970). Respiratory mechanisms and the metazoan fossil record, Nature 228, 1003-1004. Roscoe, S. M. (1969). Huronian rocks and uraniferous conglomerates in the Canadian Shield, Geological Survey of Canada Paper 68-40, 1-205.

OCR for page 21
Effects of Past Global Change on Life Runnegar, B. (1982a). The Cambrian explosion: Animals or fossils? Journal of Geological Society of Australia 29, 395-411. Runnegar, B. (1982b). Oxygen requirements, biology, and phylogenetic significance of the late Proterozoic worm Dickinsonia, and the evolution of the burrowing habit, Alcheringa 6, 223-239. Schopf, J. W., ed. (1983). Earth's Earliest Biosphere: Its Origin and Evolution, Princeton University Press, Princeton, N.J., 543 pp. Schopf, J. W., and B. Packer (1987). Early Archean (3.3-billion to 3.5-billion-year-old) microfossils from the Warrawoona Group, Australia, Science 237, 70-73. Sogin, M. L., J. Gunderson, H. Elwood, R. Alonso, and D. Peattie (1989). Phylogenetic meaning of the kingdom concept: An unusual ribosomal RNA from Giardia lamblia, Science 243, 75-77. Summons, R. E., and M. R.Walter (1990). Molecular fossils and microfossils of prokaryote and protists from Proterozoic sediments, American Journal of Science 290-A, 212-244. Towe, K. M. (1970). Oxygen-collagen priority and the early metazoan fossil record, Proceedings of the National Academy of Sciences USA 65, 781-788. Towe, K. M. (1985). Habitability of the earth Earth: Clues from the physiology of nitrigen fixation and photosynthesis, Origins of Life 15, 235-250. Towe, K. M. (1990). Aerobic respiration in the Archaean? Nature 348, 54-56. Veizer, J., W. Compston, N. Clauer, and M. Schidlowski (1983). 87Sr/ 86Sr in late Proterozoic carbonates: Evidence of a ''mantle" event at 900 Ma ago, Geochimica Cosmochimica Acta 47, 295-302. Walker, J. C. G. (1983). Possible limits on the composition of the Archean crust, Nature 302, 518-520. Walker, J. C. G., C. Klein, J. W. Schopf, D. J. Stevenson, and M. R. Walter (1983). Environmental evolution of the Archean-Early Proterozoic Earth, in Earth's Earliest Biosphere: Its Origin and Evolution, J. W. Schopf, ed., Princeton University Press, Princeton, N.J., pp. 260-290. Whatley, J. M. (1981). Chloroplast evolution: Ancient and modern, Annals of the New York Academy of Science 351, 154-165. Windley, B. F., P. R. Simpson, and M. D. Muir (1984). The role of atmospheric evolution in Precambrian metallogenesis, Fortschritte der Mineralogie 62, 253-267. Woese, C. R. (1987). Bacterial evolution, Microbiology Review 51, 221-271. Woese, C. R., O. Kandler, and M. L. Wheelis (1990). Towards a natural system of organisms: Proposal for the domains Archaea, Bacteria, and Eucarya, Proceedings of the National Academy of Sciences USA 87, 4576-4579. Young, G. M. (1976). Iron-formation and glaciogenic rocks of the Rapitan Group, Northwest Territories, Precambrian Research 3, 137-158. Yung, Y., and M. B. McElroy (1979). Fixation of nitrogen in a prebiotic atmosphere, Science 203, 1002-1004. Zang, W., and M. R. Walter (1989). Latest Proterozoic plankton from the Amadeus Basin in Central Australia, Nature 337, 642-645.