6.5
Paleobiological Clues to Early Atmospheric Evolution
C Hallmann, Max-Planck-Institute for Biogeochemistry, Jena, Germany; MARUM, University of Bremen, Bremen, Germany
RE Summons, Massachusetts Institute of Technology, Cambridge, MA, USA
ã 2014 Elsevier Ltd. All rights reserved.
6.5.1
Introduction
6.5.2
Methanogenesis and the Early Atmosphere
6.5.3
Cyanobacteria and Oxygenic Photosynthesis
6.5.3.1
Stromatolites
6.5.3.2
Microfossils
6.5.3.3
Lipids of Cyanobacteria
6.5.4
Eukaryotes and Aerobiosis
6.5.4.1
Protistan Microfossils
6.5.4.2
Fossil Macroscopic Eukaryotes
6.5.4.3
Steroid Biosynthesis
6.5.5
Algal Evolution and Sulfur Gases
6.5.6
Conclusions
Acknowledgments
References
6.5.1
Introduction
Relative to its roughly 4.5-billion-year age, Earth’s accretion
has been understood as a fast process that involved contemporaneous differentiation of core and mantle (Stevenson, 1981).
Recent improvements in mass spectrometric measurements
and the application of a number of short-lived radiogenic
isotopic systems has allowed geochemists to discern a more
detailed picture of Earth’s earliest history, including evidence
for more protracted beginnings (Halliday, 2006; Chapter 2.8).
The Giant Impact hypothesis (Canup and Asphaug, 2001)
suggests that the moon-forming event occurred relatively late
in the process and may have happened as long as 110 My after
accretion began (Halliday, 2006, 2008). This event would have
removed most of the volatiles that were captured from the solar
nebulae and left Earth with magma oceans and a CO2-rich
steam atmosphere. Ultimately, when condensation of vaporized silicates was complete, the Earth would have been left with
a very hot atmosphere comprising N2 along with CO2 þ CO þ H2O þ H2 but in unknown proportions (Zahnle, 2006).
From a paleobiological perspective, the first tangible point of
reference for the early atmosphere comes from studies of detrital zircons preserved in the 3.050 Ga Jack Hills metasediments (Wilde et al., 2001). The uranium-lead isotopic ages
and geochemistry of these minerals points to the presence of
liquid oceans and at least some continental crust by about
4.2 Ga and possibly by 4.3 Ga (Mojzsis et al., 2001; Valley
et al., 2002, 2005; Watson and Harrison, 2005). From this
point forward, Earth would have been habitable with the
potential for the earliest life to add and remove atmospheric
gases. Besides the Hadean zircons, Earth’s rock record carries
little tangible information about the prebiotic state and composition of the atmosphere – apart from the absence of molecular oxygen, there is little agreement on its constituents
(Lazcano, 2001; Miller, 1993). Extraterrestrial impactors
would have contributed volatiles (H2, H2O, CO, CO2, N2)
Treatise on Geochemistry 2nd Edition
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(Kasting and Chang, 1992) and the resulting H2O ‘steam
greenhouse’ might have maintained surficial silicates in a molten state until the frequency of the impactors diminished
(Matsui and Abe, 1986; Zahnle et al., 1988).
Mantle outgassing after the stabilization of liquid water
oceans would have consisted of volatiles with an oxidation
state intermediate between that of today’s volcanic gases
(H2O, CO2, SO2, N2) and the oxidation state of volcanic gas
in equilibrium with metallic iron (Holland, 1984). Atmospheric reactions driven by solar UV radiation and electrical
discharges could have given rise to CH4, CO, NH3, and H2CO,
although there are virtually no quantitative constraints on their
relative proportions (Holland, 1984). Furthermore, the main
reductant, H2, was continuously lost by gravitational escape
(with an accelerated pace if atmospheric methane levels were
high) and this would have led to a gradual increase in the
oxidation state at Earth’s surface (Catling et al., 2001). Sedimentary rocks from this early period (i.e., 3.8 Ga) were
clearly deposited under liquid water (Appel et al., 1998; Rosing
et al., 1996). There is abundant evidence for this in the form of
pillow basalts, hydrothermal alteration profiles, and voluminous generation of granitic rocks, all of which require significant volumes of water for genesis. Given that there are
absolutely no indications for a glaciation at this early stage in
Earth history, the faint young sun hypothesis – assuming a
solar luminosity 25% lower than at present (Newman and
Rood, 1977) – requires that there be a larger greenhouse effect
(Sagan and Mullen, 1972). It has been widely hypothesized
that high concentrations of water vapor, CO2, and possibly
methane compensated for this faint young sun (Sagan and
Mullen, 1972). An alternative view (Rosing et al., 2010) posits
that a lower global albedo, attributable to lower continental
land masses and a paucity of biogenic cloud condensation
nuclei, could compensate for low solar luminosity by retaining
enough thermal energy to avoid pervasive glaciation. But the
methods of this study were criticized (Dauphas and Kasting,
http://dx.doi.org/10.1016/B978-0-08-095975-7.01305-X
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Paleobiological Clues to Early Atmospheric Evolution
2011; Reinhard and Planavsky, 2011) and it was concluded
that the Archean must have had a strong greenhouse even in
the absence of biological cloud condensation nuclei (Goldblatt
and Zahnle, 2011; Goldblatt et al., 2009) if one accepts the
absence of evidence for glaciation. It is thus likely that atmospheric CO2 levels were much higher on the early Earth than
they are today, and this accords with the elevated rates of
volcanism and the overall lower rates of continental weathering (Owen et al., 1979; Walker et al., 1981).
After the origin of life, the greenhouse gas role played by CO2
may have been supplemented by atmospheric CH4 that was
produced by methanogenic archaea (Pavlov et al., 2001b). The
10-year atmospheric residence time that methane has today
could have been up to three orders of magnitude longer during
the early Archean (Pavlov et al., 2001a). Under such conditions,
a modern methane flux that currently maintains a 1.6 ppm
atmospheric concentration would have led to concentrations
exceeding 1000 ppm (Kasting and Siefert, 2002; Kharecha
et al., 2005). Methane production could potentially lead to a
drawdown of CO2 and if the CH4/CO2 ratio exceeded a value of
0.1 (Haqq-Misra et al., 2008), it is conceivable that Earth’s
atmosphere developed an organic haze (Pavlov et al., 2001b)
akin to that currently seen in Titan’s atmosphere (Trainer et al.,
2006). Rain-down of organic aerosols, and their catabolism by
Archean biota, is one hypothesis (Pavlov et al., 2001b) that has
been put forward to account for the pronounced negative stable
carbon isotope values of sedimentary organic matter deposited
around 2.6–2.8 Ga (Figure 1). However, the prevailing explanation for this major feature of the marine carbon isotope
record involves oxidative methane cycling (Hayes, 1983,
1994). Methane, by either mechanism, is thus an atmospheric
constituent whose presence can be envisaged through the stable
carbon isotopic composition of sedimentary organic matter
remaining from the Archean Eon. However, this idea is based
on a number of assumptions including that the carbon cycle
operated in steady state mode with similar sources and sinks as
today, with comparable burial rates for organic and inorganic
carbon (Des Marais, 2001) and that biology fractionated carbon
isotopes to comparable degrees. These assumptions should not
be taken lightly and will need to be revisited as new knowledge
is gained (Zerkle et al., 2005). These ideas are elaborated in
more detail later.
The atmospheric gas with an undisputed biological involvement is molecular oxygen (O2), which is effectively supplied
only by oxygenic photosynthesis. The invention of oxygenic
photosynthesis, which employs water as an electron donor for
carbon dioxide reduction, occurred only once in a cyanobacterium, or cyanobacterial ancestor (Allen and Martin, 2007; Allen
and Williams, 2011). After the symbiotic acquisition of cyanobacteria as chloroplast organelles, the process was subsequently
distributed among algae (Gray and Doolittle, 1982; Margulis,
1970; Martin and Kowallik, 1999; Mereschkowsky, 1905;
Schimper, 1883; Timmis et al., 2004) and, much later, to their
vascular plant descendants. The presence of molecular oxygen in
Earth’s atmosphere left its marks on geological deposits and can
be traced from 2.6 Ga (Figure 1) by, amongst other features,
the disappearance of detrital uraninite and pyrite from placer
deposits, the appearance of unconformity-related uranium and
stratiform copper deposits, and the appearance of sedimentary
red beds (Lambert and Donnelly, 1991; Lambert and Groves,
1981). More recently, highly specific geochemical proxies based
on redox-sensitive trace elements (e.g., Anbar et al., 2007) and
the rare isotopes of sulfur (e.g., Farquhar et al., 2000; Ono et al.,
2006) have led to refinements of our understanding of atmospheric oxygenation. By 2.35 Ga, O2 had accumulated in the
atmosphere to levels exceeding 105 of the present atmospheric
level, or 2 ppmv as indicated, for example, by the disappearance of a mass-independent sulfur isotope signal produced by
the O2-sensitive photochemical disproportionation of atmospheric SO2 (Farquhar et al., 2007; Papineau et al., 2007;
Williford et al., 2011). In spite of this, the details of how and
when oxygenic photosynthesis rose to prominence over its
anoxygenic counterparts remains contentious (Buick, 2008)
and there is significant evidence for the rise of cyanobacteria
well in advance of an accumulation of their metabolic byproduct in the atmosphere (Altermann et al., 2006; Kazmierczak
and Altermann, 2002; Noffke, 2010). It must also be recognized
that an alternative and minority opinion exists, namely, that
Earth’s atmosphere became oxidized before 3.5 billion years ago
(e.g., Ohmoto et al., 2006).
The rise of atmospheric O2 was a revolutionary instance of
global change with profound consequences for the subsequent
evolution of Earth’s surface environment and biosphere
(Holland, 2002). Of the prevalent metabolic pathways,
oxygen-dependent respiration, or aerobiosis, provides the largest known free energy release per electron transfer. This has led
to the observation that larger and motile organisms could
probably not have evolved without adopting aerobiosis
(Berkner and Marshall, 1965; Catling et al., 2005; Cloud,
1968, 1972; Runnegar, 1991). Bioavailability of O2 is thus
widely accepted as the prime driver of evolutionary innovation
and an incessant rise in the complexity of life (Cloud, 1972).
This key role of O2 makes understanding where and when its
production started all the more important.
Although picomolar quantities of abiotic molecular oxygen
could have been produced by the disproportionation of H2O2 in
the atmosphere at high latitudes (Haqq-Misra et al., 2008) and
upon glacial melting (Liang et al., 2006), pre-photosynthetic
production rates would never suffice to oxygenate the atmosphere. Given the geographically localized and, most probably,
temporally sporadic nature of these oxidant releases, they are also
unlikely to have sustained a capacity for aerobic metabolism. The
only logical process to account for aquatic oxygen production at
rates sufficient to oxygenate the atmosphere is oxygenic photosynthesis. During the Archean, persistent O2 sinks such as a deep
marine Fe (II) reservoir and the flux of reduced volcanic gases
(Pavlov et al., 2001a) would have buffered pO2 and prevented
any rapid atmospheric accumulation (e.g., Cloud, 1972;
Waldbauer et al., 2011). It was repeatedly suggested that tectonic
evolution allowed the onset of atmospheric oxygenation around
2.5 Ga. Kump and Barley (2007) postulated that a late Archean
tectonic episode of continent stabilization (Bleeker, 2003;
Sandiford and McLaren, 2006; Sleep, 2005) led to an increase
in subaerial volcanism at the expense of more reducing submarine volcanism. This idea was refined by Gaillard et al. (2011),
who build their hypothesis on the same tectonic episode but
suggest that differences in the degassing pressure of subaerial
versus submarine erupted magmas affected the oxidation state
of volcanogenic sulfur gases, with the following consequences
for the oxidation of the atmosphere.
141
(15)
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Paleobiological Clues to Early Atmospheric Evolution
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(1)
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Figure 1 Carbon isotopes and the rise of oxygen. The columns, left to right, chart the rise of oxygen and the various geochemical and fossil indicators
that have been used as proxies. A few points on the chart of pO2 levels are constrained; (a) A whiff of oxygen above background levels is evident from a
number of geochemical indices (Anbar et al., 2007; Garvin et al., 2008; Kaufman et al., 2007; Kendall et al., 2010; Reinhard et al., 2009); (b) The
disappearance of the ‘mass-independent sulfur isotope fractionation signal’ around 2.35 Ga is indicative of atmospheric oxygen levels exceeding 105 of
the present atmospheric level (PAL) (Farquhar et al., 2007; Guo et al., 2009); (c) This rise, possibly eventually exceeding 1 PAL, characterizes the ‘Great
Oxidation Event’, as evidenced by the appearance of sedimentary red beds and a global systematic positive stable carbon isotope excursion in
carbonates termed the Lomagundi event (Karhu and Holland, 1996; Kump et al., 2011). (1) Zoned rare earth elements and oxygen isotopes in a 4.4 Ga
Zircon suggest the existence of continental crust and a liquid hydrosphere (Wilde et al., 2001). (2) The late heavy bombardment (LHB) is speculated to
have sterilized any existing surface biosphere. Full sterilization of the habitable zone is unlikely according to Abramov and Mojzsis (2009). (3) Depleted
stable carbon isotope values of graphites might indicate a biological origin for the carbon source of the graphite (Mojzsis et al., 1996; Rosing, 1999) –
but see McCollom and Seewald (2006), van Zuilen (2002, 2003), and Lepland et al. (2011) for alternative views. (4) d34S values of microscopic sulfides
were taken as evidence for microbial sulfate reduction (Shen et al., 2001). (5) Probably-biogenic stromatolites in the North Pole area, Western Australia
(Walter et al., 1980). (6) Filamentous microfossils of debated origin and nature in the Warrawoona Group (Schopf, 1993; Schopf and Packer, 1987). Also
see Section 6.5.3.2 in the main text. (7) Stromatolite reef of unquestioned biogenicity in the Strelley Pool Chert (Allwood et al., 2006, 2009). (8) Possibly
eukaryotic acritarchs from the Moodies Group (Javaux et al., 2010). Also see Figure 2(f). (9) Possibly syngenetic biomarker evidence for cyanobacteria
and eukaryotes. See Section 6.5.4.3 in main text and references therein. (10) Possibly cyanobacterial microfossils (Altermann and Schopf, 1995). (11)
Likely syngenetic sedimentary steranes co-vary with inorganic redox-sensitive proxies (unpublished). (12) Macroscopic eukaryote Grypania spiralis
(Han and Runnegar, 1992). See Figure 2(a). (13) Unambiguous fossil cyanobacteria Eoenteophysalis belcherensis (Hofmann, 1976). See Figure 2(c).
(14) Oldest absolutely unambiguously syngenetic sedimentary biomarkers from the Barney Creek Formation (Brocks et al., 2005). (15) Oldest
unambiguous eukaryotic microfossils (Javaux et al., 2001, 2004). See Figure 2(e). (16) Banded iron formations were deposited by oxidation of the deep
marine Fe (II) reservoir by yet debated mechanisms (Lambert and Donnelly, 1991; Lambert and Groves, 1981). (17) Mass-independent fractionation of
sulfur isotopes was not preserved after pO2 reached 105 of the present atmospheric level (Farquhar et al., 2007). (18) The presence of gold, uraninite
and pyrite in quartz conglomerates and placer deposits that were transported in high-energy fluvial systems argues for low atmospheric oxygen levels
(Lambert and Donnelly, 1991; Lambert and Groves, 1981). (19) Sedimentary red beds are indicative of oxidizing conditions (Lambert and Donnelly,
1991; Lambert and Groves, 1981). (20) Features interpreted as former gas bubbles in the crests of conical stromatolites have been suggested to indicate
oxygenic photosynthesis (Bosak et al., 2009). (21) After Eigenbrode and Freeman (2006). (22) After McFadden and Kelly (2011).
142
Paleobiological Clues to Early Atmospheric Evolution
The temporal offset between the invention of oxygenic photosynthesis and the point at which O2 fluxes began to overcome
sinks, allowing it to become a constant presence in the atmosphere, is subject to a continuing and vigorous debate (Anbar
et al., 2007; Ohmoto et al., 2006). The relevance of this question
lies in the fact that, during this transitional period, respiration by
marine organisms could have been locally sustained at sites of
O2 production in the surface ocean (Kasting and Chang, 1992).
Views on the lag period vary from a 400 Ma delay (Pavlov et al.,
2001a) to a very rapid process, faster than compensated for by
the carbonate–silicate weathering cycle (Kopp et al., 2005).
Further, Goldblatt et al. (2006) have suggested a nonlinear,
concentration-dependent increase in the lifetime of atmospheric
O2, which is caused by ozone UV shielding of the troposphere
once oxygen levels reach 105 PAL. This would imply that the
existence of oxygenic photosynthesis alone does not suffice to
eventually generate a high-oxygen atmosphere, and that there is
also the feasibility of a significant lag period between the origin
of oxygenic photosynthesis and the Great Oxidation Event
(GOE). Localized oxygen oases – microaerobic regions in the
surface ocean containing low O2, but in sufficient quantity to
sustain respiration (Stolper et al., 2010) and O2-dependent
biosynthesis – could thus have existed in the Archean surface
ocean (Kasting and Chang, 1992). A similar conclusion was
reached by Waldbauer et al. (2011) who found that such oxygen
oases could have persisted for a long period of time before the
geological record captured evidence for enhanced atmospheric
O2 levels. In the remainder of this chapter, the focus is on clues
to atmospheric composition that are provided by biological
remnants: macrofossils, microfossils, isotopic fossils, and
molecular fossils.
6.5.2
Methanogenesis and the Early Atmosphere
Hydrogenotrophic methanogenesis was likely one of the first
metabolisms on Earth. Aqueous alteration of olivine, during
which serpentinite is formed, would have affected komatiites
and other ultramafic rocks in the Archean as soon as oceans
had stabilized, thereby generating H2 and CH4 as byproducts.
The Lost City Hydrothermal Field, a modern ecological niche
whose biology is supported by the products of serpentinization
reactions, hosts abundant molecular evidence for the presence
and activity of methanogenic and methanotrophic archaea
(Bradley et al., 2009; Brazelton et al., 2006). Even though
methanogens appear to represent an ancient lineage based on
comparative analysis of 16S rRNA coding genes (Woese and
Fox, 1977), chemical remnants of methanogenic archaea from
the Archean have yet to be discovered: isoprenoidal lipids that
are unambiguously diagnostic for archaea have, so far, not
been detected in any analyzed Archean sediments. Isotopically
‘heavy’ carbonates, which some interpret as indicators
of methanogenesis (Dix et al., 1995), are rare. Hayes and
Waldbauer (2006) suggested that the latter can be more readily
explained by direct exchange of isotopically-heavy dissolved
inorganic carbon (DIC) with marine waters (thereby diluting
the signal) since the main sites of methanogenesis would
have been at, or above, the sediment–water interface before
the advent of oxygenic photosynthesis, which later in geological history would push methanogens into sedimentary levels,
where generated DIC would accumulate in pore waters and
have the possibility to exchange with local carbonates.
Accordingly, the main points used to argue in favor of active
and predominant methane cycling during the early Archean
are all indirect. Climate modeling has revealed the need for
additional greenhouse warming in order to compensate for the
faint young sun (Sagan and Mullen, 1972), and elevated methane levels – higher than those that can be formed abiotically
during serpentinization of ocean crust – have been invoked
as the prime candidate (Catling et al., 2001; Haqq-Misra
et al., 2008; Kasting, 2005; Kasting and Siefert, 2002; Kharecha
et al., 2005; Kirschvink et al., 2000; Pavlov et al., 2001a,b;
Roberson et al., 2011; Walker, 1977). This hypothesis is
supported by the ensuing, probably global (Kirschvink et al.,
2000) Paleoproterozoic glaciations, which are currently
thought to be a consequence of greenhouse destruction by
rising atmospheric oxygen levels and/or hydrogen escape
to space (Catling et al., 2001). In an anoxic Archean atmosphere, 1000-year methane residence times were possible
(Chapter 10.10; Kasting, 2005 and references therein). Hayes
(1983) and Schidlowski et al. (1983) were among the first to
address the origins of a pronounced anomaly in the stable
carbon isotopic composition of sedimentary organic matter
around 2.8 Ga. Kerogen d13C values that typically vary in the
range of roughly 25% to 40% VPDB drop to values as low
as 60% (Figure 1; Hayes, 1994). Methanogenesis can lead to
carbon isotope fractionations as large as 95% in the case of
CO2 reduction and 4060% for acetate fermentation. Aerobic
methanotrophy fractionates methane carbon by a further
1020% depending on the oxidation pathway and environmental variables (Jahnke et al., 1999; Summons et al., 1994;
Whiticar, 1999). Hayes (1983, 1994) suggested that such a
two-step process was likely responsible for the formation of
exceptionally light kerogens if the metabolism of dominant
primary producers is based on methane cycling. This concept
is supported by discordant changes in kerogen d13C from
shallow- and deep-water settings over a 150 My period
(Eigenbrode and Freeman, 2006): organic matter in deepwater shales remained consistently depleted in 13C with values
as low as 60%, whereas organic matter entrained in shallowwater carbonates exhibits a trend toward heavier values near
30%, indicative of oxygenic photosynthesis. Furthermore, a
number of these same rocks exhibit elevated abundances of 3bmethylhopanes, biological marker molecules for aerobic methanotrophs (Collister et al., 1992; Neunlist and Rohmer, 1985;
Summons et al., 1994), whose concentration correlates positively with the d13C of sedimentary organic matter from the
sedimentary rocks (Eigenbrode et al., 2008). While this trend
appears counterintuitive, it can be explained by the fact that
only Type I methanotrophs – inhabiting an environmental
niche defined by low CH4, elevated O2, and sufficient fixed
nitrogen – biosynthesize 3-methylhopanepolyols, the precursors to 3b-methylhopanes (Amaral and Knowles, 1995; Hanson
and Hanson, 1996). In view of the abovementioned facies
analysis (Eigenbrode et al., 2008), this finding supports the
increasing importance of oxygenic photosynthesis in late
Archean shallow-water environments, whereas deep-water facies
were likely dominated by either Type II aerobic methanotrophs
and/or ANME consortia. It should be noted, however, that
alternative explanations for the highly negative d13C anomalies
Paleobiological Clues to Early Atmospheric Evolution
at ca. 2.8 Ga exist. As already discussed, Pavlov et al. (2001a,b)
and Haqq-Misra et al. (2008) suggest that a methane-derived,
Titan-like, organic haze may have been incorporated into sediments after the settling of larger aerosols. Des Marais (2001), in
contrast, has proposed that the extremely light kerogens could
have been formed by recycling of carbon by anaerobic microbes
including chemoautotrophs. In light of the interpretation that
the organic d13C anomaly presages the rise of environments
dominated by oxygenic photosynthesis, the question remains
whether truly anaerobic methanotrophy is possible. The matter
was picked up by Hinrichs (2002), who proposed the anaerobic
option. But even the anaerobic oxidation of methane (AOM)
that is performed by consortia of methanotrophic archaea
(ANME) and sulfate reducers (Boetius et al., 2000; Hinrichs
et al., 1999; Orphan et al., 2001) relies on the availability of
suitable oxidants, SO42 or possibly NO3 (Raghoebarsing
et al., 2005), whose marine inventories were likely very low
before the rise of atmospheric O2 (Canfield, 1998; Canfield
et al., 2000; Kaufman et al., 2007). This implies that either
form of biological methane utilization is inevitably tied to
atmospheric oxygen levels sufficient to drive oxidative sulfur
and nitrogen cycles. In any variation of the scenarios described
earlier, it is currently believed that the emergence of photosynthetic organisms was preceded by a period where methane
metabolism was prominent.
6.5.3
6.5.3.1
Cyanobacteria and Oxygenic Photosynthesis
Stromatolites
Stromatolites – a term initially coined by Kalkowsky (1908) –
are defined by Walter (1976) as ‘organosedimentary structures
produced by sediment trapping, binding, and/or precipitation as a
result of the growth and metabolic activity of microorganisms, principally cyanophytes’. This definition, which specifies a biological
origin but also excludes specific mention of their characteristic
lamination, makes them a sub-classification of microbialites,
with the consequence that stromatolites that occur in the
3.45 Ga Strelley Pool Formation of the Warrawoona Group
(Allwood et al., 2006; Schopf et al., 2007; van Kranendonk
et al., 2008) (Figure 1) represent some of the oldest accepted
traces of life on Earth (van Kranendonk, 2011; Walter et al.,
1980). Other specimens from the slightly older Dresser Formation (Buick et al., 1981) are not as well preserved in outcrop,
but may, ultimately, satisfy the strict criteria for biogenicity if
additional examples can be identified in outcrop or core.
Semikhatov et al. (1979) provided a definition fundamentally different from that of Walter (‘an attached, laminated,
lithified sedimentary growth structure, accretionary away from a
point or limited surface of initiation’) and one that does not
invoke the action of biology. Thus, if we are to view stromatolites as indicators of the production or consumption of atmospheric gasses, apart from CO2, it is necessary that they are
demonstrably biogenic and carry information about organismic physiologies. Although stromatolites are most commonly
accepted as being of biological origin, a purely abiotic dynamical model of stromatolite surface growth that involves chemical precipitation (which was presumably more common
during the early Precambrian than today), sediment fallout,
and diffusive rearrangement is (at least locally) feasible
143
(Grotzinger and Rothman, 1996). While extant and geologically younger stromatolites typically form through a combination of chemical precipitation, trapping, and binding of
sedimentary particulates in sticky extracellular polymeric substances (EPS) of an actively involved microbial biofilm
(Dupraz and Visscher, 2005; Dupraz et al., 2009), some stromatolites from the earlier Precambrian are proposed to have
formed solely by the repeated precipitation of mineral laminae
(Knoll, 2003). Such mineral precipitation does not necessarily require a biological involvement and stromatolites that
are hypothesized to have formed by this mechanism are
more often those recorded from earlier in the Precambrian
(Grotzinger and Knoll, 1999). Furthermore, many Archean
stromatolite occurrences, such as those of the Strelley Pool
Formation in Western Australia (Lowe, 1980, 1983) did not
initially yield unambiguously biogenic microfossils, which led
to the questioning of their biogenic nature (Lowe, 1994). Such
doubt has since been addressed with the systematic evaluation
of stromatolite morphotypes (Hofmann et al., 1999; van
Kranendonk et al., 2003) across more than 10 km of outcropping Strelley Pool Formation (Allwood et al., 2006) that
revealed the existence of discrete stromatolitic facies. This variation is attributed to differing paleoenvironmental settings
across an isolated Archean peritidal carbonate platform. Microscale analysis of sedimentary fabrics in these stromatolites
(Allwood et al., 2009) strongly suggest an active involvement
of benthic microbes in their formation and the ability to flourish in a variety of different local settings (van Kranendonk,
2006, 2007, 2011). Lastly, recent studies have found cell-like
microstructures of different morphologies at multiple stratigraphic levels of the Strelley Pool Formation. The initial discovery was of cellular structures with diverse morphologies
reported to occur in a black chert unit of the Strelley Pool
Formation associated with stromatolites and evaporites
(Sugitani et al., 2010). Subsequently, spheroidal and ellipsoidal objects arranged in chains and clusters and tubular sheaths
were found in a basal sandstone member of the Strelley Pool
Formation and interpreted as the cellular remains of sulfurmetabolizing microbes (Wacey et al., 2011).
The biogenicity of these most ancient traces is, however,
only the first question to be addressed. It would be an even
greater accomplishment if we could deduce microbial physiologies from ancient stromatolites. The suspected involvement of
cyanobacteria in the formation of stromatolites is based on the
prominence of these organisms in microbial mats from the
extant stromatolitic reefs that are found in tropical marine
settings such as the Hamelin Pool area of Shark Bay, Western
Australia (Burns et al., 2004; Golubic, 1976; Neilan et al.,
2002) and in the Bahamas (Baumgartner et al., 2009; Reid
et al., 2000). This observation is supported by many examples
of uncontested cyanobacterial microfossils found in association with stromatolites from the Proterozoic (e.g., Awramik
and Barghoorn, 1977). In the sedimentary rocks of the 3.2 Ga
Moodies Group and the 2.9 Ga Pongola Supergroup of South
Africa that represent siliciclastic tidal flat paleoenvironments,
combinations of microbially induced sedimentary structures,
including wrinkles, desiccation cracks, and roll-up structures,
likely record the previous existence of microbial mats (Noffke
et al., 2006, 2008). Similar structures in contemporary
environments and analogues from older periods in Earth
144
Paleobiological Clues to Early Atmospheric Evolution
history, reflect stabilization by cyanobacterially-dominated
microbial communities. While such an affiliation would
imply a high likelihood of oxygenic photosynthesis (but see
Oren and Padan, 1978), morphological analogies are often
ambiguous and demand caution. Along these lines, analogies
to modern filamentous cyanobacteria that have been observed
to form small cones both in the presence and absence of
sedimentation and lithification (Bosak et al., 2009; Flannery
and Walter, 2011; Jones et al., 2002; Lau et al., 2005; Love
et al., 1983; Walter et al., 1976a) are very suggestive, but do not
categorically assure, that ancient coniform stromatolites
engaged in phototaxis and thus phototrophy.
It is not known yet whether the formation of conical morphologies in modern mats results from phototaxis or, alternatively, from enhanced growth at cone tips through some other
physiological response (Bosak et al., 2009). A relationship
between cyanobacterial cone spacing and day length is suggestive of a spatial organization dependent on the diffusion of
nutrients required for photosynthesis (Petroff et al., 2010)
and has led to the suggestion that a photosynthetic metabolism
might have been a characteristic of Precambrian coniform stromatolites. A different study using Anabaena-containing biofilms
found no dependency of light intensity on biofilm reticulation,
but rather showed that the formation of pillars was a chemotactic response (Shepard and Sumner, 2010). Thus it seems that
one of the most definitive observations regarding the potential
photosynthetic nature of ancient stromatolites, independent of
their morphology, involves changes in abundance of stromatolitic structures with paleo-water depth, which appears to be
the case in the Strelley Pool Formation (Allwood et al., 2006,
2009). A different approach to shed light on the question of
early oxygenic photosynthesis was taken by Bosak and colleagues (Bosak et al., 2009), who studied details of the apical
zones of modern cyanobacterial cones and ancient stromatolites. Donaldson (1976) was the first to notice axial porosity in a
number of Proterozoic stromatolites and attributed this to gas
bubble entrapment. Growth experiments with a cyanobacterial
mat from Yellowstone National Park revealed that enhanced
photosynthetic activity in the crestal area of cyanobacterial
cones caused the formation and entrapment of oxygen-rich
bubbles in the crest, but not flanks, of the structures. Similar
features – near circular cavities enclosed by contorted laminae
in the tips, but not sides, of cones – were observed in a set of
Proterozoic stromatolites (Bosak et al., 2009). The disrupted
crestal areas were attributed to gas release resulting from higher
metabolic activity on topographic highs and, by analogy to the
modern mats, oxygenic photosynthesis in fossil conical stromatolites up to an age of 2.7 Ga has been suggested (Figure 2(b);
Bosak et al., 2009). The need for additional research on the
associations between tufted mats and cyanobacterial oxygenic
photosynthesis takes on added importance with the recent discovery of Neoarchaean stromatolites with exquisite preservation
of cm-scale tufts (Flannery and Walter, 2011).
6.5.3.2
Several units within the 3.533.42 Ga Warrawoona Group in
Western Australia host organic microstructures that have been
(b)
(a)
18.5 mm
(c)
10 mm
5 mm
(e)
(d)
50 mm
Microfossils
(f)
35 mm
100 mm
Figure 2 Fossils indicative of aerobiosis. (a) Little doubt exists regarding a eukaryotic affinity of macrofossil Grypania Spiralis (Han and Runnegar,
1992). (b) Fossilized bubbles (see arrow) in the crestal area of conical Stromatolites were suggested as a tracer for the existence of oxygenic
photosynthesis (Bosak et al., 2009). (c) Eoenteophysalis belcherensis from the Kasegalik Formation on the Belcher Islands is the oldest unambiguous
fossil colony of coccoid cyanobacteria (Hofmann, 1976). (d) Well-preserved eukaryotic Tappania acritarch with anastomosing processes from the
Neoproterozoic Wynniatt Formation, Canada (Butterfield, 2005). (e) The process-bearing acritarch Tappania plana from the Mesoproterozoic Roper
Group, Australia (Javaux et al., 2004) is one of the oldest and widely-accepted eukaryotic microfossils. (f) The exceptionally large size of this acritarch
from the 3.2 Ga Moodies Group points toward a eukaryote (Javaux et al., 2010). In the absence of further characteristics (see main text), however, such
a classification remains ambiguous.
Paleobiological Clues to Early Atmospheric Evolution
interpreted as cellular remains. Some of the Warrawoona
microfossils have morphological resemblances to cyanobacteria, both modern and fossil (Figure 1). Sheath-enclosed
spheroidal cells and filaments were reported from the Towers
(now Dresser) Formation and from a chert unit in the Apex
basalt, respectively (Schopf, 1993; Schopf and Packer, 1987),
and were suggested to be the oldest fossil evidence for life on
Earth. The biogenicity of these structures was later called into
question (Brasier et al., 2002, 2004, 2005, 2006; Marshall
et al., 2011; Pinti et al., 2009), but also defended (de Gregorio
et al., 2009; Schopf et al., 2002, 2007). No metabolic inferences have been made regarding the microstructures subsequently discovered by Sugitani et al. (2010), although the
assemblages are morphologically diverse and considerations
of the composition, size range, abundance, taphonomic features, and spatial distributions suggest that cluster-forming
small (<15 mm) spheroids and lenticular to spindle-like
structures are probable microfossils. Thread-like structures
that are also found in the same unit of the Strelley Pool
Formation are interpreted as fossilized fibrils of biofilm,
rather than microfossils. Other chemical, isotopic, and elemental analyses of carbonaceous matter in this unit show
patterns that are consistent with, but not strictly diagnostic
of, an origin from metamorphosed microbial remains
(Glikson et al., 2008). Also, Wacey et al. (2011) found multiple and diverse cellular morphologies with size ranges typical of prokaryotic assemblages. Raman spectra from the
microfossils and d13C values in the range of 33 to 46%
VPDB are consistent with thermally mature disordered carbonaceous material of biological origin. However, given the
ubiquity of sedimentary pyrite found in close association
with sedimentary organic matter, their additional claim that
micron to sub-micron grains of pyrite in and around these
objects are a diagnostic sign of their sulfur-based metabolisms
is less assured. While metabolisms based on sulfur were very
likely extant by this time (Shen and Buick, 2004; Shen et al.,
2001), the cell morphologies and associations reported by
Wacey et al. (2011) are not fully ambiguous and it is not
until much later in Earth’s history that metabolisms can be
confidently attributed to specific kinds of microfossils. We
also note that while Raman spectra can aid in the identification of carbonaceous matter and its thermal history, this tool
does not, in its own right, provide unambiguous evidence of
biogenicity (Pasteris and Wopenka, 2003). Confocal laser
scanning microscopy coupled to Raman imagery, on the
other hand, affords a means to demonstrate, in three dimensions at high spatial resolution, both morphology and molecular composition of ancient microscopic fossils (Schopf et al.,
2005; Schopf and Kudryavtsev, 2009).
Permineralized carbonaceous microfossils, recovered from
the 2.60 Ga Campbell Group, South Africa, have been interpreted as fossil cyanobacteria based on their cellular morphologies (Altermann and Schopf, 1995). According to (Knoll,
2003), however, poor preservation limits confidence in this
interpretation and these microfossils should rather be categorized as ‘possibly cyanobacterial’. Akinetes, on the other hand,
are the resting stages of heterocystous cyanobacteria and one of
the most diagnostic of all cell types. Akinetes have been found
in rocks as old as 2.1 Ga (Tomitani et al., 2006). Confidently
assigned (Knoll, 2003) fossil cyanobacteria (Figure 2(c)) were
145
also found in cherts of the otherwise mainly dolomitic McLeary
and Kasegalik Formations in the Belcher Islands, Canada
(Hofmann, 1976), whose age of 1.8 Ga is constrained by
the overlying shales and mafic volcanics of the Flaherty Formation (Fryer, 1972). Given the compelling geological and
geochemical evidence for a GOE around 2.45 Ga (Farquhar
et al., 2007) and other geochemical data indicative of free
oxygen at 2.5–2.6 Ga (Anbar et al., 2007; Garvin et al., 2008;
Kaufman et al., 2007; Kendall et al., 2010; Reinhard et al.,
2009), we can be rather confident that cyanobacterial oxygenic
photosynthesis was well established by 2.5 Ga and thus well
before cyanobacterial affiliations can be attributed to specific
assemblages of microfossils.
6.5.3.3
Lipids of Cyanobacteria
Cyanobacteria characteristically biosynthesize mid-chain
branched mono-, di- and trimethylalkanes (Dembitsky et al.,
2001; Han and Calvin, 1970; Köster et al., 1999; Robinson and
Eglinton, 1990; Shiea et al., 1990) that can be preserved,
sometimes abundantly, in ancient sedimentary rocks
(Bauersachs et al., 2009; Kenig et al., 1995). Similar components biosynthesized by insects (Nelson and Blomquist, 1995)
can be distinguished on the basis of generally higher chain
lengths (>C24 vs. C16–22 in cyanobacteria). However, background abundances of mid-chain branched alkanes are quite
high in many sedimentary rocks and oils, even when a direct
contribution by cyanobacteria is not obvious. The presence of
mid-chain branched alkanes is therefore mostly valued as an
alert that would prompt a search for more specific indicators of
past cyanobacterial activity.
Hopanoids, on the other hand, are more definitive molecular fossils. These pentacyclic triterpenoids are biosynthesized
by a variety of bacteria and some higher plants (Ourisson et al.,
1982), and are widely assumed to play a role in mediating
membrane behavior and in assisting in stress tolerance
(Kannenberg et al., 1980; Ourisson and Rohmer, 1982; Poralla
et al., 1984; Rohmer et al., 1979; Sáenz et al., 2012; Welander
et al., 2009). The major sedimentary isomers have a 17a(H)
21b(H) stereochemistry, which differs from the predominant
biological 17b(H)21b(H) with which the precursor C35
bacteriohopanepolyols are biosynthesized. Details of the he
stereochemical conversion, which takes place during dia- and
catagenesis, are well established (Chapter 10.3; Hallmann
et al., 2011; and references therein). Hopanoid molecular
transformations give rise to several prominent homologous
series of C27–C35 hopanes that are found ubiquitously in
ancient sediments and petroleum. Apart from this most common and prominent hopane series, there was also the early
discovery that hopanes with an additional methylation at C-2
(2-methylhopanoids) were biosynthesized by cyanobacteria
and some alphaproterobacteria while C-3 methylated hopanoids (3-methylhopanoids) are found in Type 1 methanotrophs and Acetobacter species (Zundel and Rohmer, 1985
and references therein). The fully saturated catagenetic products of these 2- and 3-methylhopanes were subsequently also
characterized in sedimentary rocks and petroleum (Summons
and Jahnke, 1990), with a prevalence in carbonate and evaporite lithologies.
146
Paleobiological Clues to Early Atmospheric Evolution
A systematic evaluation found that 2-methyl bacteriohopanepolyols occur in a significant ( 30%) proportion of cultured cyanobacteria and cyanobacterial mats, as well as in
sedimentary rocks all through Earth history, with elevated
abundances in some Precambrian deposits (Summons et al.,
1999) (Figure 3). This finding led to the suggestion of a
2-methylhopane biomarker index (2-MHI), expressed as the
relative proportion of 2-methylhopanes over their desmethyl
counterparts, as an indicator of the existence and activity of
cyanobacteria and, thus, oxygenic photosynthesis during
sediment deposition.
Elevated 2-MHI values were presented as evidence for the
existence of oxygenic photosynthesis during the late Archean
(Brocks et al., 1999; Eigenbrode et al., 2008; Summons et al.,
1999). The 2-MHI enjoyed widespread acceptance until it
was discovered that common soil microbes other than
cyanobacteria – the nitrogen-fixing a-proteobacteria Bradyrhizobium japonicum (Talbot et al., 2007), Beijerinckia indica,
Beijerinckia mobilis (Vilcheze et al., 1994), and the anoxygenic
phototrophic purple non-sulfur bacterium Rhodopseudomonas
palustris (Rashby et al., 2007) – also are able to biosynthesize
these compounds de novo. Further complications followed,
such as the evidence gleaned from genomic databases that
less than 10% of extant bacteria are capable of hopanoid
biosynthesis (Pearson et al., 2007). A study of lipids and phylogenetic affiliation based on sequences of hopanoid cyclases
(sqhC) in a land-sea transect across San Salvador island revealed
low, yet evenly distributed, abundances of 2-methylhopanoids,
but no cyanobacterial sqhC genes (Pearson et al., 2009),
thus suggesting either a non-cyanobacterial source for
2-methylhopanes in the >2 mm cut of aquatic particulate
Phanerozoic
Precambrian
F/F
0.30
5
P/T
0.20
Toarcian
OAEs
0.25
2-Methylhopane index
organic matter, or a seasonally limited presence of hopaneproducing cyanobacteria. Similarly, the Venter Global Ocean
Sampling expedition did not encounter any cyanobacterial
sqhC sequences in comparable marine samples (Pearson and
Rusch, 2009). Once the genetic basis of C-2 hopanoid methylation was discovered (Welander et al., 2010), the hpnP gene
responsible for encoding the radical SAM protein that methylates hopanoids at C-2 was confirmed to be present in just
three bacterial phyla, namely, the cyanobacteria, the aproteobacteria, and the acidobacteria. Hopanoid biosynthesis, however, remains to be demonstrated in this latter group.
A physiological role for methylated hopanoids has also to be
found, although studies with cyanobacterium Nostoc punctiforme suggest a connection to the cell cycle in this organism
(Doughty et al., 2009).
Does this mean that the 2-MHI is no longer useful? Although
the biosynthetic capacities of bacterial phyla are still poorly
constrained, several lines of evidence argue in favor of
2-methylhopanes being informative as paleophylogenetic indicators. Cyanobacteria are the dominant phytoplankton in the
oligotrophic open ocean (Chisholm et al., 1988) and studies of
BHP in marine water column particulates (Saenz et al.,
2011a,b), in marine cyanobacterial mats (Allen et al., 2010),
and in recent and ancient marine sediments show that 2methylhopanoids are ubiquitous in environments where
cyanobacteria are likely to have been prevalent. Although a
definitive answer still eludes us, advances in bioinformatics
may eventually help in identifying the origin of the hpnP gene.
It is believed that cyanobacteria played a more significant
role in primary productivity during most of the Precambrian,
prior to the rise of the modern algae (Knoll et al., 2007).
0.15
12
3
9
7 8
0.10
6
4
8
0.05
00
00
28
00
26
00
24
00
22
00
20
00
18
00
16
00
14
00
12
10
600
80
0
50
0
40
0
30
0
20
0
10
0
0
0.00
Geological age (Ma)
Figure 3 The 2-methylhopane index through geological time. The relative proportion of C31 ab 2a-methylhopane over C30 ab hopane is thought to be a
possible indicator of the prominence of cyanobacteria in an ecosystem or depositional environment. A dichotomy is present at the Precambrian–
Cambrian boundary, with Phanerozoic returns to pre-Phanerozoic values during times of environmental stress. Modified from Knoll AH, Summons RE,
Waldbauer JR, and Zumberge JE (2007) The geological succession of primary producers in the oceans. In: Falkowski P and Knoll A (eds.) Evolution
of Primary Producers in the Sea, pp. 133–163. Amsterdam: Elsevier. OAE, oceanic anoxic event; P/T, Permian–Triassic boundary; F/F, Frasnian–
Famennian boundary; 1, South Oman Salt Basin (n ¼ 41), from Grosjean et al. (2008); 2, Nepa-Botuoba-Katanga oil family (n ¼ 19), from Kelly et al.
(2011); 3, Coppercap Formation (n ¼ 14), Hallmann, unpublished data; 4, Baykit High oils (n ¼ 6), from Kelly et al. (2011); 5, Shaler Supergroup (n ¼ 8),
Hallmann, unpublished data; 6, Francevillian Basin, Gabon (n ¼ 4), Hallmann, unpublished data; 7, Wittenoom Fm and Mt.McRae shale from drill hole
ABDP-9 (first series, Bitumen 1 & 2); 8, Transvaal Supergroup from drill holes GKP01 (distal facies, lower values) and GKF01 (proximal facies, higher
values), after Waldbauer et al. (2008); 9, Hamersley province cores from drill holes SV1, RHDH2a, and WRL1, after Eigenbrode and Freeman (2006).
Paleobiological Clues to Early Atmospheric Evolution
Cyanobacterially-dominated microbial communities that
probably formed vast stromatolite reefs on carbonate platforms, and were likely analogous to the modern counterparts
in Shark Bay and the Bahamas, are one likely source of ancient
sedimentary 2-methylhopanoids. However, selective preservation of continental margin sedimentary records and a paucity
of records for the deeper parts of the ocean basins preclude
a full understanding of the secular and spatial patterns
of 2-methylhopanoid production through geological time.
On the other hand, we can be confident of their preferential
occurrence in low-latitude carbonate depositional systems
and during episodes of oceanic deoxygenation (Knoll et al.,
2007). Phanerozoic sediments deposited during times of
enhanced marine euxinia and characterized by elevated 2MHI values include Cretaceous Oceanic Anoxic Events (OAE)
(Kuypers et al., 2004a,b; Dumitrescu and Brassell, 2005), the
end-Permian mass extinction event (Cao et al., 2009), and the
Frasnian–Famennian extinction event (Edwards et al., 1997;
Knoll et al., 2007). All these notable marine biodiversity crises
are characterized by lowered stable nitrogen isotopic values of
sedimentary kerogens, suggesting that a significant proportion
of sedimentary organic matter was derived from nitrogenfixing cyanobacteria (Junium and Arthur, 2007; Kuypers
et al., 2004b; Wada and Hattori, 1991). Proliferation of diazotrophic cyanobacteria was hypothesized to be a consequence
of strong euxinic conditions that altered the marine nitrogen
cycle through reduced nitrification (Kuypers et al., 2004b).
Thus, there are geological, geochemical, and biochemical
grounds for considering cyanobacteria as being responsible
for the elevated levels of 2-methylhopanes during these events
(Figure 3).
Although it is tempting to extend these observations to the
Precambrian and to suggest a cyanobacterial origin for the elevated 2-MHI values observed in Proterozoic and Archean sedimentary rocks, the current knowledge does not permit such
extrapolation. The Archean is renowned for a paucity of wellpreserved sedimentary sections amenable to organic geochemical studies. Sedimentary 15Norganic data in a stratigraphic context
are sparse (Garvin et al., 2008; Godfrey and Falkowski, 2009;
Thomazo et al., 2011). Molecular biological studies of contemporary cyanobacteria that include investigations of the
physiological roles of hopanoids (including 2-methylhopanoids)
are promising lines of research that will shed light on how
and, perhaps, when cyanobacteria first developed the capacity
to split water and released free O2 into the environment (Sessions
et al., 2009).
6.5.4
Eukaryotes and Aerobiosis
The emergence of eukaryotes was dependent on the prior, or
perhaps concurrent, establishment of oxygenic photosynthesis:
almost all eukaryotic organisms must maintain oxygen homeostasis (Semenza, 2007). Molecular oxygen is also required for
the biosynthesis of sterols – albeit at exceptionally low
concentrations (Waldbauer et al., 2011) – that are essential
components of most eukaryotic membranes. The eukaryotes
that cannot synthesize sterols de novo must acquire them
through their diet. The extremely small concentrations of dissolved molecular oxygen (<7 nM) that are required for sterol
147
biosynthesis by S. cerevisiae might be an explanation for the
observation of extant eukaryotes appearing to exist under anoxic
conditions (Danovaro et al., 2010). Ever since the symbiotic
acquisition of mitochondria (Sagan, 1967), during which
proto-eukaryotes gained the capacity for aerobic respiration,
eukaryotes have obligatorily engaged in aerobiosis, which gave
them the energetic advantage that would lead to the continuing
evolution of complexity. Eukaryotic remnants in the form of
macro-, micro-, or molecular fossils thus point to the existence
of oxygenic photosynthesis during the time of their formation
and burial, and their oldest sedimentary traces are a critical
signpost for our understanding of oceanic and atmospheric
chemistry at that time.
6.5.4.1
Protistan Microfossils
The classification of acritarchs as eukaryotic is common, but it
can also be problematic. These resistant organic microfossils
are so named because of their uncertain phylogenetic classification. Knoll et al. (2006) distinguish eukaryotic microfossils
by their size, morphology, and taphonomy: “Prokaryotes can
have processes, and they can have preservable walls. But we do
not know any prokaryote that combines these. . . characters,
nor any that exhibit such a complexity of form. . ..” When
considering size alone, there have been reports of potentially
eukaryotic acritarchs from 3.2 Ga shallow marine deposits of
the Moodies Group in South Africa (Javaux et al., 2010) and
from rocks as old as 3.4 Ga in Australia (Sugitani et al., 2007,
2009, 2010). These organic walled spheroidal microfossils
have diameters up to 300 mm but lack either wall ornamentation or complex wall ultrastructure (Figure 2(f)). Although
none of the contemporary bacterial cells that can grow to
comparably large sizes – cyanobacterial akinetes, large sulfur
bacteria, or myxobacteria – are known to occupy similar environmental habitats and to form recalcitrant biopolymers, a
eukaryotic classification remains ambiguous in the absence of
further characteristics.
The oldest certainly eukaryotic microfossils have been
recovered from Mesoproterozoic strata. The Ruyang Group
carries an age estimate of 1600–1250 Ma (Xiao et al., 1997)
and contains Shuiyousphaeridium macroreticulatum (Yan and
Zhu, 1992) and Tappania plana (Yin, 1997), both of which
meet all the three criteria of eukaryotic affinity mentioned
earlier. Tappania sp. have also been recovered from the welldated Roper Group in northern Australia (Javaux et al., 2001)
with an age just short of 1.5 Ga (Jackson et al., 1999).
Furthermore, the Roper Group acritarchs (Figure 2(e))
exhibit a taxonomically diverse assemblage with certain morphologies being limited to habitats of a specific physical environment (Javaux et al., 2001). The idea that Tappania‘s
morphology, which suggests the presence of a cytoskeleton
(Javaux et al., 2001), indicates a more evolved eukaryotic
form is, however, debatable as it was found that the eukaryotic
endocytic system (Dacks et al., 2009) as well as a cytoskeleton
(Wickstead and Gull, 2011) had probably existed already in
the last eukaryotic common ancestor (LECA). There is little
doubt that a diverse assemblage of protists inhabited the
marine realm during the Mesoproterozoic, but the inventory
of microfossils found to date cannot be used to conclusively
reconstruct earlier eukaryotic evolution.
148
6.5.4.2
Paleobiological Clues to Early Atmospheric Evolution
Fossil Macroscopic Eukaryotes
The oldest macrofossil that bears convincing characteristics of
eukaryotic algae is Grypania spiralis (Figure 2(a)). The organism
that produced these sub-centimeter-sized, spirally coiled
imprints is not known, but the relatively large size (up to
50 cm when uncoiled) and occasionally massive occurrence
of what appear to be individual bodies, as well as the morphological regularity of specimens within populations (Knoll et al.,
2006), are rather unmistakable signs of the eukaryotic nature of
these specimens. Grypania’s first appearance in the rock record
has been dated to 2.11 Ga (Han and Runnegar, 1992) or
1.87 Ga (Schneider et al., 2002) in the Negaunee iron formation, thus placing it just after the Lomagundi positive carbon
isotope anomaly that has been associated with an abnormally
large amount of carbon burial and O2 production (Karhu and
Holland, 1996; Kump et al., 2011). Grypania has been suggested
to be a sessile benthic alga (Runnegar, 1991; Walter et al.,
1976b, 1990). However, no fossil holdfast has been discovered
so far in the direct vicinity of Grypania imprints and the Negaunee Formation appears to have been deposited in deep waters
(Morey, 1983). Oxygenic photosynthesis may have been limited to the neritic zone during the Archean (Kendall et al., 2010)
and Paleoproterozoic, while current models suggest that the
ventilation of deeper oceans did not occur until the later Neoproterozoic (Canfield et al., 2007; Fike et al., 2006; Logan et al.,
1995). The existence of benthic eukaryotic organisms in such
distal deepwater facies at 2.0 Ga would thus argue for a
completely different redox structure of the water column and
cannot be reconciled with other geochemical data sets. On the
other hand, Grypania could well have been a planktonic
dweller, which is consistent with the geochemical data and
lack of holdfast. In any event, these fossils suggest the existence
of larger eukaryotes during the early Paleoproterozoic.
O
Steroid Biosynthesis
As discussed previously, all eukaryotes require sterols, or a
surrogate type of membrane component. The biosynthesis
of sterols is an oxygen-intensive process and the principal
steps are conserved across some of the deepest phylogenetic
branches in the eukaryotic domain (Summons et al., 2006).
It is also believed that the last eukaryotic common ancestor already contained the enzymes necessary for sterol
biosynthesis (Desmond and Gribaldo, 2009). A very small
number of bacteria carry the genetic capacity to biosynthesize
sterols de novo (Pearson et al., 2003), but they are not capable
of extending the steroid molecule at C-24, thereby giving
rise to distinct species of bacterial and eukaryotic sterols. The
bacteria that can make sterols have probably acquired this
capacity through secondary gene transfer (Desmond and
Gribaldo, 2009). Steroid biosynthesis is initiated by the
epoxidation of squalene to form 2,3-oxidosqualene by the
enzyme squalene monooxygenase (SQMO) in a process that
utilizes equimolar amounts of dioxygen and squalene
(Figure 4). Subsequent cyclization by oxidosqualene cyclase
(OSC) enzymes generates either lanosterol or cycloartenol
through an intermediate protosterol cation (Abe et al., 1993;
Dean et al., 1967; Tchen and Bloch, 1957a). Both lanosterol
O
NH
N
N
6.5.4.3
O
NH
R
El Albani et al. (2010) reported centimeter-sized structures,
fossilized by pyritization, from the 2.1 Ga Francevillian Basin
in Gabon, and suggested a multicellular eukaryotic affinity
based, among other points, on radial structures and inner
fold patterns that were attributed to a coordinated growth
pattern. This interpretation however, is viewed with some
skepticism and the structures might rather be representative
of diagenetic concretions or remnants of microbial mats that
experienced gas bubble buildup.
HN
O
N
R
N
N
O
18O
NH
H
NH
18OH
R
N
18O
H
O
N
18O
18O
18O
1
2
3
Figure 4 Molecular oxygen in steroid biosynthesis. The sterol biosynthetic pathway starts with the epoxidation of squalene to (3S)-2,3-oxidosqualene
by a squalene monooxygenase (SQMO). (1) Oxygen reacts with the dihydroflavin portion of a flavine-adenine-dinucleotide (FAD) to produce (2) a
hydroperoxyflavin. (3) Nucleophilic epoxidation of squalene to yield an oxidized flavin and 2,3-oxidosqualene. The oxidized FAD is regenerated by
NADPH-cytochrome C450 reductase. The oxygen of 2,3-oxidosqualene ends up as the C-3 hydroxyl oxygen in sterols. The origin of this oxygen was the
subject of much debate before Tchen and Bloch (1957a,b) used chemical reasoning followed by stable isotope labeling experiments to show that it
derives from O2 and not from water.
Paleobiological Clues to Early Atmospheric Evolution
and cycloartenol contain the C2–C3 oxygen atom from oxidosqualene bound in the form of a hydroxyl function at the C3 position. The origin of this oxygen was the subject of much
early discussion but it was shown by 18O labeling experiments
with SQMO to derive indirectly from O2, not water (Tchen and
Bloch, 1957b) (Figure 4). An additional nine moles of oxygen
are then required to remove three angular methyl groups on
the way to cholesterol.
The possibility of an initially anaerobic pathway to 2,3oxidosqualene using H2O instead of O2 (Raymond and
Blankenship, 2004) is unlikely based on chemical reasoning
and the overall number of oxygen-dependent biosynthetic
steps (Summons et al., 2006). This rationale is the basis
for using fossil steroids as proxies for oxygen availability and
thus, indirectly, the existence of oxygenic photosynthesis. Fossil
steroids have been reported from numerous Archean and
Proterozoic sedimentary rocks (Figure 1) and oil inclusions
aged up to 2.7 Ga (Brocks et al., 1999, 2003a,b; Dutkiewicz
et al., 2003, 2006, 2007; Eigenbrode, 2004; Volk et al., 2005;
Waldbauer et al., 2008), but their syngenicity has been
questioned (Brocks, 2011; Brocks et al., 2008; Rasmussen
et al., 2008). Methodological improvements (Hallmann et al.,
2011; Sherman et al., 2007), however, suggest that sequential
extractions combined with dissolution of carbonate and clay
minerals in samples can release mineral-occluded biomarkers
and allow a distinction between contaminant and syngenetic
lipids. Nevertheless, reports of Archean biomarkers should be
interpreted with caution until solid evidence from ultra-clean
scientific drilling can confirm these data.
Furthermore the presence of steroids in a sedimentary
deposit cannot be taken as a direct gauge of atmospheric
oxygen content. The dissolved oxygen threshold for the synthesis of ergosterol by yeast cells (S. cerevisiae) was determined to be 7 nM (Waldbauer et al., 2011). Even if such
microaerobic conditions were maintained in the upper water
column by oxygenic phototrophs, reduced atmospheric gases
and UV radiation would rapidly consume the outgassing flux
of O2 into an anoxic Archean atmosphere (Pavlov et al.,
2001a; Zahnle et al., 2006). Such processes would maintain
atmospheric pO2 at vanishingly low levels during the
Archean, even if oxygenic photosynthesis existed along
ocean margins by 2.7 Ga as suggested by sedimentary lipids
and inorganic proxies. It should be noted, however, that the
7 nM value represents a minimum and that dissolved oxygen
concentrations in Archean surface waters might well have
been higher wherever oxygenic photosynthesis was active.
Accordingly, direct conclusions about atmospheric pO2 during the Archean cannot be drawn from the presence of steroids; their presence, however, testifies that the saturation
of Archean oxygen sinks was proceeding, probably gradationally, changing atmospheric composition and, ultimately,
contributing to the GOE.
6.5.5
Algal Evolution and Sulfur Gases
The dominant sulfur compound in the contemporary atmosphere is SO2 which in a pre-anthropic world, was mainly
supplied by subaerial volcanism, along with typically lesser
amounts of H2S (Chapter 10.14). This total volcanic sulfur
149
flux is estimated at 60 Tg S per year (Chapter 6.7). Oxidation
of both gases ultimately generates SO42 which forms an
important contribution to cloud condensation nuclei (CCN)
and thus carries the capability to influence climate. This nonsea salt sulfate (NSS-SO42) aerosol must be distinguished
from sulfate aerosols introduced to the marine atmosphere
by sea spray and aquatic bubble bursting, which are typically
larger in size and of shorter atmospheric residence time.
Another significant contribution to CCN comes from oceanto-atmosphere fluxes of dimethylsulfide (DMS; modern flux
estimated at 25–50 Tg S per year) that forms through the bacterially mediated enzymatic conversion of dimethylsulfiopropionate (DMSP) in marine surface waters, where the latter is
produced by a variety of marine algae (Malin and Kirst, 1997;
Stefels, 2000; van Alstyne, 2008). DMS is volatile and, once
generated, transfers easily into the atmosphere, although the
latter can also be transformed in the water column, for example, to dimethylsulfoxide (DMSO), by photo-oxidation in the
presence of photosensitizers such as chlorins or organic acids.
Rates of water column oxidation of DMS are unknown, but
they are suggested to be similar to the rates of outgassing.
DMSP and DMS have received much attention for their
potential role in climate regulation (Ayers and Cainey, 2007;
Bates et al., 1987; Charlson et al., 1987). Atmospheric oxidation of DMS by NOx and HOx eventually forms NSS-SO42 and
methanesulfonic acid (MSA), which can be a source of sulfuric
acid that has the power to create new CCN aerosols. A recent
sulfur isotope study (Oduro et al., 2012) has revealed that MSA
and NSS-SO42 are indeed the principal atmospheric oxidation products of DMS and that sulfur isotope variability of
MSA and NSS-SO42 is likely not caused by contribution
of mixed DMS and non-DMS sources.
This DMS flux likely did not exist prior to the emergence of
eukaryotes. The strongest producers of DMSP are the dinoflagellates and haptophytes (Keller et al., 1989), as well as some
chrysophytes and diatoms (Stefels, 2000). While the latter
three groups have appeared only in geologically recent times,
dinoflagellates have likely existed since the Cambrian
(Moldowan and Talyzina, 1998) and possibly even much earlier (Summons et al., 1992). Furthermore, our knowledge of
the full array of biological DMSP sources remains limited.
DMS has a short estimated residence time of around 2 days
in the modern atmosphere (Chapter 10.14), but this would
have differed prior to the GOE (Waldbauer et al., 2011) due to
much faster photolysis. Rosing et al. (2010) have argued that
the absence of biologically produced CCN during the Archean
would be one factor that could have contributed to lowering
the global albedo, thereby reducing the need for an extreme
greenhouse atmosphere. While it seems that this is incorrect,
and even in the complete absence of biological CCN the
Archean must have had a significant greenhouse atmosphere
(Goldblatt and Zahnle, 2011), it is hypothesized that the CCN
argument might be used in reverse.
The emergence of eukaryotes and the radiation of algae in
the marine realm could have had a significant impact on
climate by adding a previously nonexistent flux of sulfur to
the atmosphere that might have intensely increased the
amount of CCN and thereby led to global cooling. In the preanthropogenic world, biological activity was likely the main
factor determining the concentration of CCN (Andreae, 2007).
150
Paleobiological Clues to Early Atmospheric Evolution
This has been corroborated by the observation that the air over
highly productive regions of the modern ocean contains a
much higher concentration of CCN than air above the lowproductivity gyre regions (Andreae, 2007). Furthermore, it was
modeled that CCN reduction as a consequence of global
warming and, consequently, the reduced biological productivity during the Middle Cretaceous would be a plausible means
to create the extremely high sea surface temperatures implied
by temperature proxies (Kump and Pollard, 2008). While the
Huronian glaciation ( 2.4–2.1 Ga) is widely seen as a consequence of the GOE, and was likely influenced by the destruction of an atmospheric methane greenhouse due to rising
oxygen levels, an increased abundance of CCN, provided by a
growing inventory of marine algae, may have contributed to
this episode of cooling. Although oxidative continental weathering was increasing during the GOE, the marine sulfate pool
remained small at this early stage of atmospheric oxidation
(Halverson and Hurtgen, 2007). DMSP biosynthesis starts
with the assimilation of seawater sulfate (Stefels, 2000) and,
although it is not known what effect SO42 concentrations
would have on the rates of DMSP biosynthesis, the total DMS
flux was likely dependent on the marine sulfate inventory.
While many plausible factors have been put forward as triggers
for the Huronian ‘Snowball Earth’ glaciation, a rapid increase
in the population and diversity of marine algae could have
contributed. Possibly more plausible, another algal diversification pulse much later in Earth history – which is supported by
an increase in the taxonomic richness of acritarch assemblages
around 800 Ma (Grey, 2005; Knoll et al., 2006) – could have
played a role in the onset of the Cryogenian Snowball Earth
events.
6.5.6
Conclusions
The chemical composition of the atmosphere is a key factor by
which an evolving biosphere directly influences Earth’s surface
environment. Atmospheric evolution is frequently modeled to
fulfill physical constraints, such as the need for greenhouse
warming during the Archean, but it can only be understood
from an integrated perspective that includes all aspects of
biospheric and planetary evolution. Secular variations in the
atmospheric inventories of methane and O2 have been under
debate for many years and models have been continuously
adjusted with input parameters from the fields of physics,
geology, geochemistry, and paleontology. In this brief review,
the need for additional paleobiological indicators to improve
our understanding of early atmospheric evolution is reinforced. Sedimentary organic matter and geostable lipid molecules have considerable unrealized potential in this regard
(Sessions et al., 2009) and, therefore, molecular biological
and physiological studies of the role of lipids in organismic
fitness and evolution will add to our understanding of how
Earth’s contemporary atmosphere came to be so.
Acknowledgments
The authors gratefully acknowledge the Agouron Institute, the
NASA Astrobiology Institute and the Max-Planck-Society for
support during the preparation of this review. Malcolm Walter
provided many suggestions that improved the manuscript.
Reviews by Jim Kasting, Mark van Zuilen and James Farquhar
are gratefully acknowledged.
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