organic geochemical signatures of early life on earth
TRANSCRIPT
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12.2 Organic Geochemical Signatures of Early Life on EarthRE Summons, Massachusetts Institute of Technology, Cambridge, MA, USAC Hallmann, Max-Planck-Institute for Biogeochemistry, Jena, Germany
ã 2014 Elsevier Ltd. All rights reserved.
12.2.1 Introduction 3312.2.2 Eoarchean (4.0–3.6 Ga) Biological Remnants? 3312.2.3 The Post-3.5 Ga Sedimentary Record of Stable Carbon Isotopes 3412.2.4 The Record of Organic Carbon Burial 3512.2.5 The Composition of Buried Organic Matter 3712.2.6 Visible Structures with Organic Affinities 4012.2.6.1 Organic-Walled Microfossils 4012.2.6.2 Fossil Microbial Mats, Textures, and Trace Fossils 4112.2.6.3 Stromatolites 4112.2.7 Summary and Prospects 42Acknowledgments 43References 43
GlossaryArchean Eon The geologic eon that extends from c.3.8 Ga to
the Proterozoic 2.5 Ga. The Archean Eon is in the process of
being redefined chronometrically and subdivided into the
eras of Eoarchean (4.0–3.6 Ga), Paleoarchean (3.6–3.2 Ga),
Mesoarchean (3.2–2.8 Ga), and Neoarchean (2.8–2.5 Ga).
The International Commission on Stratigraphy
currently does not recognize the lower boundary of the
Eoarchean.
Bitumen Sedimentary organic matter that is or was mobile
and soluble in organic solvents.
Fa Fraction of aromatic hydrogen in kerogen.
Hadean Eon An informal designation for the time between
the formation of the Earth c.4.5 Ga ago and the oldest
known rocks of c.3.8 Ga.
Kerogen Insoluble, macromolecular organic matter.
Ma/Ga Million/billion years before present.
Myr/Gyr Million/billion year.
Ro% Vitrinite reflectance, or vitrinite reflectance equivalent –
a proxy for degree of thermal alteration of organic matter.
VCDT Vienna Canyon Diablo Troilite, the international
standard for stable sulfur isotopic measurements.
VPDB Vienna Pee Dee Belemnite, the international
standard for stable carbon isotopic measurements.
atise on Geochemistry 2nd Edition http://dx.doi.org/10.1016/B978-0-08-095975
12.2.1 Introduction
The timing of life’s appearance on Earth is subject to excep-
tionally poor constraints. Geochemical thermometers pre-
served in 4.4–4.0-billion-year (Ga)-old zircons recovered
from a 3.5-Ga sedimentary rock attest to a watery, clement
early Hadean Eon that would have been conducive for life to
appear and proliferate (Valley et al., 2002; Watson and
Harrison, 2005). Other geochemical evidence is consistent
with the hypothesis that there were oceans, some continental
crust, and weathering processes in place by 4.3 Ga (Ushikubo
et al., 2008). However, any relict of Hadean life that may
have been present in sediments deposited in the first c.700
million years (Ma) of our planet’s history appears to have
been lost as a result of persistent impacts by asteroids, plate
subduction, weathering, or metamorphism (Schopf, 1983).
Therefore, in this brief overview we focus mainly on the
subsequent Archean Eon (3.8–2.5 Ga) for sedimentary rocks
that record clues about the nature and metabolic capacities
of Earth’s early denizens. Simultaneously, we must keep in
mind that considerable biospheric evolution likely took place
during the Hadean Eon.
12.2.2 Eoarchean (4.0–3.6 Ga) BiologicalRemnants?
Graphite occurring in the highly altered terrain of the Isua
greenstone belt in southwestern Greenland (Mojzsis et al.,
1996; Rosing, 1999) represents the oldest postulated remains
of life. The metamorphosed host rocks, which include pillow
basalts and possible turbidites, were evidently deposited in
deep water and are remnants of an early Archean (>3.75 Ga)
seafloor hydrothermal system. The origins of reduced carbon
present in apatite crystals in the Isua and Akilia metasediments
(Mojzsis et al., 1996), and putative sedimentary graphite par-
ticles with d13C values in the range�11.4% to�20.2% Vienna
Pee Dee Belemnite (VPDB) (Rosing, 1999) were originally
proposed to have a biological origin due to their depletion in13C relative to carbonates of similar age (Schidlowski, 1988).
Biology was implicated because the kinetic isotope effect asso-
ciated with the preferential uptake of 12CO2 during biological
carbon fixation (Des Marais, 2001; O’Leary, 1988) results in
organic matter being 13C-depleted compared to inorganic sub-
strates. However, a confounding complication is that CO2
reduction that takes place by abiological means can be
-7.01005-6 33
34 Organic Geochemical Signatures of Early Life on Earth
accompanied by carbon isotope fractionations of a similar
magnitude (McCollom, 2011). This result, predicted by
theory, has been demonstrated experimentally for hydrocar-
bon gases produced under simulated hydrothermal conditions
(McCollom and Seewald, 2006).
The C-isotope findings from Isua and Akilia have been
extensively debated in light of the great antiquity of this
record and because of their ambiguity. Both the proposed
sedimentary nature of the rocks and origins of the reduced
carbon have been questioned. Recent work proposes that
reduced carbon in the Isua terrain probably arose through
metasomatic decomposition of ferrous carbonates (van
Zuilen et al., 2002, 2003) and that these rocks are of little
or no biogeochemical relevance. Accordingly, additional
lines of evidence must be brought to the table before car-
bon isotopic data for organic matter can be used to infer
biological processes in the world’s oldest sediments.
12.2.3 The Post-3.5 Ga Sedimentary Recordof Stable Carbon Isotopes
It has long been held that the �26% isotopic separation
between the sedimentary inorganic (i.e., carbonate, CCARB)
and organic (reduced, CORG) carbon reservoirs provides evi-
dence of biological carbon fixation (Schidlowski, 1983, 2001).
This hypothesis follows from the observation of 27–31% range
of carbon isotopic fractionations associated with the RuBisCO
proteins at the heart of the Calvin–Benson–Bassham (CBB)
cycle of autotrophic carbon fixation (O’Leary, 1988). This
makes empirical sense in the context of our simplified view
of the carbon-bearing compartments within the global carbon
cycle, where these reservoirs are characterized by distinct car-
bon isotopic compositions conforming to known equilibrium
and kinetic fractionation factors (Figure 1).
Fresh organic matter
Sedimentary organic matter
Metamorphic and ignreduced carbon
-40 -30 -20
d 13C (‰
0–103 years
107–109 years
106–109 years
103–108 years
Biosyn
Decomposition and burial
Pressure and heat
Subd
Figure 1 Biogeochemical carbon cycle (reproduced from Des Marais DJ (2Precambrian. Reviews in Mineralogy and Geochemistry 43: 555–578, with perare shown in correspondence to time spans (left y-axis) needed to traverse eacobserved d13C values.
However, as discussed above, nonbiological processes
could generate reduced carbon with a stable carbon isotopic
signature that is indistinguishable from that formed during
biological carbon fixation. Hydrothermal systems, in particu-
lar, are environments of the early Earth that may have wit-
nessed production of organic matter without biological
intervention (McCollom and Seewald, 2007; Proskurowski
et al., 2008). Therefore, claims of biogenesis based on isotopic
data alone cannot be taken at face value. They must be accom-
panied, for example, by a rigorous evaluation of geological and
petrographic data that specify the stratigraphic context of sam-
ples, metamorphic grades, and the potential for diagenetic and
metasomatic overprinting. Sedimentological, geochemical,
and other features that inform us about the paleoenvironmen-
tal setting are also paramount considerations. The sediments of
the c.3.35 Ga Strelley Pool Formation are a case in point. They
are generally well preserved for rocks of Paleoarchean age,
contain pristine carbonate in places, and include diverse stro-
matolites as well as a suite of geological features that are
consistent with deposition in a shallow coastal marine envi-
ronment (Allwood et al., 2006, 2009, 2010) – an interpreta-
tion supported by the distribution and abundances of rare
earth elements (Allwood et al., 2010). In other words, the
sedimentary setting is one in which we would expect to
encounter photosynthetic carbon fixation if this metabolic
mechanism already existed. Organic matter is preserved in
the laminations of the stromatolites and in stratigraphically
equivalent black chert deposits that represent silicified sedi-
ments. The kerogen occurs as clasts and globules deposited
together with other detrital materials that are finely dissemi-
nated throughout the chert matrix. Bulk d13C values for
kerogen from stromatolites range from �28.3% to �35.8%(Marshall et al., 2007) consistent with carbon fixation via
the CBB cycle. A number of localities within the same forma-
tion host preserved carbonaceous objects with diverse mor-
phologies that are interpreted as microfossils (Sugitani et al.,
Mantlecarbon
Marbleeous
Carbonates
Marine HCO3
−
-10 0 +10
VPDB)
HydrosphereAtmosphereBiosphere
Sedimentary
Metamorphic
Mantle–crust
thesis
uction
Decomposition
Outgassing
Weathering
Cycle
CO2
(sea, atm.)
001) Isotopic evolution of the biogeochemical carbon cycle during themission from Mineralogical Society of America). Subcycles (right y-axis)h of the subcycles. The x-axis boxes roughly correspond to the ranges of
Organic Geochemical Signatures of Early Life on Earth 35
2010). Recent data reporting microscopic and geochemical
evidence for sulfur-metabolizing microbes in the same forma-
tion (Wacey et al., 2011a,b) add to the accumulating evidence
for a range of biological processes taking place during the
deposition of the Strelley Pool Formation. Wacey et al.
(2011a,b) identified chains and clusters of organic microstruc-
tures which they identified as microfossils based on features
that included hollow cell lumens, nitrogen-containing cell
walls, evidence of taphonomic degradation, and d13C values
in the range of �33% to �46% VPDB. Associated pyrite crys-
tals had D33S values between �1.65% and þ1.43% and d34Svalues from �12% to þ6% Vienna Canyon Diablo Troilite
(VCDT). These observations are consistent with evidence for
the biogenicity of the carbon and sulfur isotopic signals in
other sections of the Strelley Pool Formation (Bontognali
et al., 2012).
The long-term constancy of the average isotopic composi-
tions of inorganic (da) and organic carbon (do), in post-3.5 Ga
sediments, together with their �26% offset, has been cited as
evidence of a continuously active biogeochemical carbon cycle
(Des Marais, 2001; Des Marais et al., 1992; Hayes, 1993; Hayes
et al., 2002; Schidlowski, 1983, 2001; Schidlowski et al., 1983).
A long-term average value of da near 0%, when the crustal
average is ��6%, implies the existence of a 13C-depleted,
crustal organic carbon reservoir, while excursions of da from
the long-term average value are viewed andmodeled as intervals
of enhanced organic carbon burial or weathering (Des Marais,
1997; Des Marais et al., 1992; Hayes, 1993; Hayes
and Waldbauer, 2006; Hayes et al., 1983). This evidence is
largely based on globally distributed sample sets rather than
discrete samples, formations, or sedimentary basins. In addi-
tion, basin-scale C-isotopic data from the Late Neoarchean to
Early Paleoproterozoic also record large positive and negative
shifts in carbonate d13C attributable to perturbations in the
global carbon cycle involving events of enhanced carbon burial
(Aharon, 2005; Baker and Fallick, 1989; Bekker et al., 2008;
Melezhik et al., 1999, 2007) or weathering (Kump, 1991;
Kump and Arthur, 1999). In a recent example, Kump and
colleagues studied the isotopic compositions of the oxidized
and reduced carbon phases in carbonate rocks and organic
carbon-bearing shales from the Zaonega Formation in the
Paleoproterozoic Onega Basin on the southeastern margin of
the Fennoscandian Shield. Here they identified a strong negative
d13C excursion, which they correlated with a probably synchro-
nous anomaly in the Francevillian Basin of Gabon. This
‘Shunga-Francevillian anomaly’ was then attributed it to intense
oxidative weathering of rocks in the aftermath of the ‘Great
Oxidation Event’ (Kump et al., 2011).
In evaluating these concepts, it must be recognized that some
of the C-isotopic data, especially for organic carbon in highly
mature terrains, may be compromised by elevatedmetamorphic
grades and/or metasomatism (Hayes et al., 1983; Schidlowski,
2001). Thermal alteration changes the d13C values (do) of sed-imentary organic matter (kerogen) and in general it can be
assumed that most Archean kerogens have experienced lower
greenschist metamorphism (�300 �C) with concomitant shifts
in d13C values by asmuch as 3% (DesMarais, 1997; Hayes et al.,
1983). Additional processes involving hydrothermal alteration
can influence the d13C value and these include isotope exchange
with CO2-rich fluids (Kitchen and Valley, 1995; Robert, 1988);
exchange of crustal carbon with carbon from the mantle is a
further factor for consideration (Hayes and Waldbauer, 2006).
These processes can shift the d13C of sedimentary organic matter
to significantly higher values and potentially lower the d13C of
carbonate (da) and, as a consequence, complicate the interpre-
tation of the isotopic data since the primary biological and the
carbonate reference signal tend to converge. However, in spite of
the fact that the crustal records of da and do can be affected by
diagenesis, metamorphism, and exchange with the mantle, the
offset of their average values appears to be one of the most
robust proxies for the existence of biological processes on Earth.
In light of the relatively stable long-term apparent isotopic
fractionation (D13C) between da and do, it is logical to infer
that marine primary productivity in the surface waters of the
early ocean represents the prime input of organic matter in
(reasonably) well-preserved Archean and Proterozoic sedimen-
tary rocks. However, we can make no conclusion about the
biota responsible for ancient primary productivity based on
the carbon isotopic record alone. Since no distinct difference
exists in the organic carbon isotopic composition of biomass
produced by oxygenic versus the different modes of anoxygenic
photosynthesis (for a different view, see Nisbet et al., 2007),
the carbon isotopic record obtained from Precambrian sedi-
ments provides no direct evidence for the onset of oxygenic
photosynthesis. Notably, however, in the 2.8–2.5 Ga Fortescue
and Hamersley sedimentary sequences of the Pilbara Craton,
where there are extensive and paleoenvironmentally con-
strained records of organic and inorganic carbon isotopes,
persistent trends are observed (e.g., Eigenbrode and Freeman,
2006; Hayes et al., 1983). In recent work, Eigenbrode and
Freeman (2006) observed a 13C-enrichment of �10% in
organic carbon from shallow-water carbonate rocks relative
to coeval deep-water sediments. In addition, organic carbon
from shallow-water environments has a very wide (�29%)
range in values ranging from �57% to �28%, which is in
marked contrast to the 13C-depleted and more narrow range
of �40% to �45% for organic carbon from deep-
water sediments. Eigenbrode and Freeman (2006) posit that
the deep-water signals reflect assimilation of methane or other13C-depleted substrates like it has been hypothesized earlier
(Hayes, 1983, 1994; Hinrichs, 2002; Schoell and Wellmer,
1981; Watanabe et al., 1997). They also propose that the
progressive 13C-enrichment in organic matter from shallow
settings from 2.8 to 2.5 Ga reflects the expansion of aerobic
ecosystems and oxygen-respiring communities as a conse-
quence of the early advent of oxygenic photosynthesis, which
is discussed in detail below (Hayes, 1993).
12.2.4 The Record of Organic Carbon Burial
Hydrogen escape, either directly as a volcanic output (Kasting,
1993) or after photolysis of methane or water in the upper
atmosphere (Catling et al., 2001), probably played a key role
in changing the oxidation state of the earliest Earth. After the
appearance of oxygenic photosynthesis – the only mechanism
capable of producing O2 in appreciable amounts – O2 sinks that
include respiration, reduced volcanic gases together with ferrous
iron, manganese, sulfide, and hydrogen from subsea weathering
of fresh oceanic crust, first had to be satisfied before O2 could
begin to accumulate in the atmosphere. Burial of a fraction of
the organic carbon that was formed with the O2 – that is, its
36 Organic Geochemical Signatures of Early Life on Earth
sequestration in the crust – is required to prevent consumption
of all oxygen and allows its progressive accumulation in the
atmosphere and ocean system (Broecker, 1970; Holland, 1978,
1984; van Valen, 1971). During most of Earth’s history autotro-
phy most likely was the dominant source of the organic carbon
entering sediments and, therefore, the crustal inventory. Assim-
ilation of CO2 and its reduction into ‘fixed’ organic compounds
requires parallel oxidation of an electron donor. Volcanogenic
H2, Fe2þ, S2�, and simple reduced carbon compounds were
probably the initial redox partners associated with CO2 reduc-
tion by methanogens and acetogens and this would have led to
some accumulation of the corresponding oxidized species
(Hayes and Waldbauer, 2006; Sleep and Bird, 2007, 2008).
However, this process is inefficient. It has been estimated that
nonphotosynthetic ecologies would be hampered by levels of
primary productivity as low as 10�4 of those of a photosynthetic
world (Sleep and Bird, 2008), thereby imposing strict limits to
the rates at which reduced carbon could be buried. Energy
harvested from sunlight, therefore, would have enhanced the
rates of autotrophic carbon fixation, carbon burial, and release
of oxidizing power to surface environments. Still, the rates
would have been significantly lower than today (Figure 2) and
inherently limited by the fluxes of electron donors provided by
surface and subsea volcanism to microbial communities living
at or near the sea surface. It is thought that black shales in
themselves are a biosignature for a photosynthetic biosphere.
Black shales also have special taphonomic significance in that
they can survive deep burial and high-grade metamorphism
(Sleep and Bird, 2007).
The initiation of oxygenic photosynthesis, where water
assumed the role of the electron donor for CO2 reduction,
required the development of complex light-harvesting systems
and appears to have resulted from the combination of two
preexisting anoxygenic photosynthesis pathways via interme-
diate steps (Blankenship, 1992, 2010; Blankenship and
Hartman, 1998). Molecular evidence supports this metabolic
merge and provides evidence identifying the green and purple
sulfur bacteria as the sources of the original biochemical
machinery that now resides as photosystems I and II in the
Mantlecarbon
Fresh organic matter
Marble
Sedimentary organic matter
Metamorphic and igneousreduced carbon
Carbonates
Marine HCO3
-
9000
8990
2
1.60.4
2
9
60
45
10
6
50
10
PModern
CO2
(sea, atm.)
Figure 2 Fluxes in the biogeochemical carbon cycle (reproduced from Desduring the Precambrian. Reviews in Mineralogy and Geochemistry 43: 555–5correspond to those illustrated in Figure 1 and the two schemes are models thphotosynthesis.
thylakoids of cyanobacteria and in their descendants, the chlo-
roplasts of green algae and vascular plants (Xiong et al., 2000).
Once unconstrained by fluxes of electron donors from volca-
nogenic sources, carbon fixation through oxygenic photosyn-
thesis would lead to vastly enhanced rates of primary
productivity. It has been estimated that the onset of oxygenic
photosynthesis would have increased global organic produc-
tivity and carbon burial by at least one (Canfield et al., 2006)
and possibly two to three (Des Marais, 2000) orders of magni-
tude. Resultant environmental oxidation would, however, only
be transient. Stoichiometric reversal readily takes place
through respiration. In the modern ocean typically less than
0.5% of primary organic matter survives transit through the
water column to be buried and preserved (Hedges and Keil,
1995; Wakeham and Canuel, 2006). The proportion of organic
matter that does escape remineralization, however, breaks the
redox balance: burial of organic matter leads to an excess of
oxidized species in the surface environment. Any portion of the
organic matter that has passed through the geological carbon
subcycle and becomes exposed at the surface is again suscepti-
ble to weathering with further consumption of oxygen. Many
factors have ultimately contributed to the present-day oxidized
state of Earth’s surface environment but two of them are para-
mount. Firstly, the initial state of the crust was purely volcanic
and the relative proportion of sedimentary rocks increased
through the Hadean and Early Archean, thereby also increasing
the potential size of the crustal reduced carbon reservoir
(Taylor and McLennan, 1985). Secondly, this process is largely
irreversible. With the growth of continental crust and its sedi-
mentary carbon reservoir, the burial flux of organic matter and
other reduced species can be used as a proxy to partially recon-
struct past atmospheric oxygenation although the details of its
inception and how it progressed remain subject of intense
debate (Bjerrum and Canfield, 2004; Canfield et al., 2000;
Des Marais, 2001; Des Marais et al., 1992; Hayes and
Waldbauer, 2006; Sessions et al., 2009).
The fraction of buried organic carbon can be reconstructed
from the stable carbon isotopic compositions of co-occurring
organic and inorganic carbon so long as these rocks have
Mantlecarbon
Fresh organic matter
Marble
Sedimentary organic matter
Metamorphic and igneousreduced carbon
Carbonates
~20
~13
?
164
20
?
~40
45
~30
?>20
~7
reoxygenic photosynthesis
CO2
(sea, atm.)
Marine HCO3
-
Marais DJ (2001) Isotopic evolution of the biogeochemical carbon cycle78, with permission from Mineralogical Society of America). Subcyclesat illustrate the comparative fluxes before and after the advent of oxygenic
Organic Geochemical Signatures of Early Life on Earth 37
not been too severely altered. The process of thermal matura-
tion preferentially cleaves 12Cd12C over 12Cd13C bonds in
kerogen, which leads to false, heavier residual organic d13Cvalues – although assessment of thermal maturity and recon-
struction of primary values is possible in certain cases (Des
Marais, 1997; Hayes et al., 1983). Taking this into consider-
ation, simplified models constructed using assumptions of a
carbon cycle operating in steady state, together with da and dodata from the extensive compilations that now exist, have
enabled accounts of carbon burial during the Precambrian
(Des Marais, 1997, 2001; Des Marais et al., 1992; Hayes and
Waldbauer, 2006) and through Phanerozoic time (Berner,
2003; Berner and Raiswell, 1983) that are consistent with
geological observations of progressive accumulation of oxi-
dants at the surface including the atmospheric inventory of
O2. At the same time, there are numerous complexities and
uncertainties that preclude deeper understanding of the incep-
tion and progress of the Archean carbon cycle on Earth
(Bjerrum and Canfield, 2004; Fischer et al., 2009; Hayes and
Waldbauer, 2006; Kump, 2008).
Early Archean isotope data from the 3.2–3.5 Ga volcano-
sedimentary sequences of the Pilbara and Kaapvaal cratons
display average da values of 0�2% (Veizer et al., 1989) and
do values between ��25% and �42%. These values encom-
pass the wide range of discrimination exhibited during
carbon assimilation by CBB autotrophs and would be consis-
tent with – but not compelling evidence for – the existence of
oxygenic photosynthesis, since chemoautotrophic microor-
ganisms such as methanogens and anoxygenic phototrophic
bacteria (Schidlowski et al., 1983) can produce similar frac-
tionations. A significant change in the range of do values
becomes apparent around 2.9 Ga. Values as low as �65%cannot be explained by autotrophy alone, even under high-
CO2 atmospheric concentrations. Extremely depleted biomass
can be produced by the recycling of fermentation-derived CO2
or acetate (Eigenbrode and Freeman, 2006; House et al., 2003)
or by the assimilation of acetate from acetogens that compete
for H2 with organisms such as methanogens and sulfate-
reducing bacteria (Gelwicks et al., 1989; Whiticar, 1999). It is
however unlikely that a microbiosphere engaging in these
metabolisms alone is capable of generating the large abun-
dances of very light organic matter found during this time
period. Another plausible scenario involves an active cycle of
methane oxidation and assimilation (Coleman et al., 1981;
Hayes, 1983, 1994; Hinrichs, 2002; Schoell and Wellmer,
1981), with the sharp drop in sedimentary do values possibly
representing an increase in the availability of oxidized electron
acceptors as a consequence of the advent of oxygenic photo-
synthesis. Molecular support for this hypothesis comes from a
positive correlation between abundances of 3b-methylhopanes
(3-MH), biomarkers for methanotrophic bacteria, and kerogen13C in the same samples (Eigenbrode et al., 2008). While one
would intuitively expect a negative correlation, the observation
makes sense in the way that these 3-MH are biosynthesized
only by type-I methanotrophs, which occupy a specific niche
with methane availability but higher O2 levels (Amaral and
Knowles, 1995; Hanson and Hanson, 1996) than present in
the deep basinal areas that are dominated by type-II methano-
trophs. The 3-MH-producing methanotrophs thus thrived
alongside photoautotrophs, which would have produced bio-
mass in much greater abundances. This explains why a light
isotopic signature of methane assimilation is not evident in the13C-enriched 3-MH samples. Further evidence for this sugges-
tion might be revealed when in situ multielement isotope
analyses of microscopic fossils and kerogen fragments at
small spatial scales become better calibrated, understood, and
more widely applied (e.g., Williford et al., 2011).
12.2.5 The Composition of Buried Organic Matter
All documented occurrences of Archean sedimentary organic
matter take place in terrains that have experienced alteration
through tectonism, hydrothermal activity, and ionizing radiation.
Accordingly, the kerogens that remainhave beenoverprintedwith
considerable loss of the primary characteristics. Preserved bitu-
mens largely consist of carbonaceous globules and seams of
highly reflective (�4.0% Ro) pyrobitumen (Buick et al., 1998;
Gray et al., 1998; Rasmussen, 2005; Rasmussen and Buick, 2000)
together with minute traces of hydrocarbons preserved in shales,
carbonates (Brocks et al., 1999, 2003a; Eigenbrode et al., 2008),
and in the fluid inclusions of psammitic quartz crystals
(Dutkiewicz et al., 1998, 2006; George et al., 2008). While the
C-isotopic data for this material can be useful (see above), it has
proven largely intractable for the application of molecular tech-
niques to trace its primary nature.
Kerogen represents the only solid phase in the sedimentary
organic carbon reservoir that has remained in situ and immobile
since it was first deposited. As mentioned above, all known
Archean rocks and, therefore, kerogens have seen metamorphic
grades to at least lower greenschist (Prehnite–Pumpellyite) facies.
Most kerogens are notoriously difficult to characterize
because of the heterogeneous nature of the components and
their polymeric nature. A primary measure of composition
comprises the elemental abundances, especially in respect to
C, H, O, N, and S. Thermal maturation leads to progressive loss
of H, N, O, and S relative to carbon, which is the end-stage
product in metamorphosed sediments. The ratio of H over C
(H/C) is a measure of the relative proportion of all hydrogen
and carbon remaining in the macromolecular carbonaceous
network at each stage of the thermal trajectory. Fresh biomass
is characterized by H/C values of �1.0–2.0, where algal
and bacterial biomass dominated by lipids have higher values
(H/C �2.0), and organic matter (e.g., plants) that characteris-
tically contains a higher abundance of incorporated oxygen has
lower values (H/C �1.0). Heat-driven release of hydrocarbons,
nitrogen, and CO2 leads to residual kerogens with progres-
sively lower H/C values; this is the same process that involves
generation and migration of petroleum phases from rocks with
high total organic carbon contents. Although the final H/C
value at any given stage of thermal maturity will be a function
of the original value of the carbonaceous matter as well as with
the molecular composition, values below 0.5 are generally
regarded as representing mature organic matter that has lost
most, if not all, of its capacity to generate hydrocarbons. Early
studies using hydrous pyrolysis (e.g., Lewan et al., 1985),
where water is used as a source of hydrogen to enhance kero-
gen cracking and product yields, failed to detect any hydrocar-
bons where kerogens older than 1.6 Ga were heated under
closed system conditions (Hoering and Navale, 1987) and,
for a long time, this result discouraged exploration for molec-
ular biosignatures in the record of Archean rocks.
38 Organic Geochemical Signatures of Early Life on Earth
Solid-state 13C NMR spectroscopy (Smernik et al., 2006),
laser Raman and Fourier transform infrared (FTIR) spectroscopy
(Marshall et al., 2005) reveals that Early Archean kerogens from
a stromatolite in the Strelley Pool Formation are highly aromatic
(fa varying from 0.90 to 0.92) and contain only minor aliphatic
carbon or carbon-oxygenated (C–O) functionalities (Marshall
et al., 2007). The Raman carbon first-order spectra for the iso-
lated kerogens are typical of spectra obtained from disordered
sp2 carbons with low two-dimensional (2D) ordering (biperio-
dic structure). The implications of the Raman data are low 2D
ordering throughout the carbonaceous network, which indicates
the incorrect usage of the term graphite in the literature to
describe the kerogen or carbonaceous material in the Warra-
woona cherts. Hydrocarbons produced during high-temperature
experiments, where these kerogens were pyrolyzed in a streamof
high-pressure hydrogen (hydropyrolysis or HyPy; Love et al.,
1995), contain one-ring to seven-ring polycyclic aromatic hydro-
carbons that were covalently bound to the kerogen, as well as
some alkanes (linear, branched, and cyclic), which were most
probably trapped in the microporous network of the kerogen.
The polycyclic aromatic hydrocarbons have mainly C1- and C2-
alkylation while C3þ-substituted aromatics are low in abun-
dance. This study showed for the first time that correlations
exist between elemental H/C ratios, Raman spectroscopic
parameters (ID1/IG, ID1/(ID1þ IG), and La), and the degree
of alkylation of bound polyaromatic molecular constituents
generated by HyPy (Figure 3). Molecular profiles of the HyPy
products of Strelley Pool Chert kerogens and mature Mesopro-
terozoic kerogen from Roper Group (c.1.45 Ga), which is
undoubtedly microbial in origin, were very similar providing
one line of evidence for biogenicity even though no specific
biomarker structures could be identified. A combination of
Raman spectroscopy, for identifying the best-preserved kero-
gens, used together with HyPy for liberating chemically bound
molecules from these kerogens offers a sound and potentially
productive strategy for evaluating the biological origins of
Earth’s oldest preserved organic matter.
A similar approach, combining chemical analysis,
spectroscopy, and pyrolysis, was used by Derenne and col-
leagues to study kerogen from the chert facies of the Strelley
Pool Formation (Derenne et al., 2008). A measured elemental
H/C of 0.62, together with solid-state 13C nuclear magnetic
resonance spectroscopic signals for a significant fraction of ali-
phatic carbon, suggested a lower level of thermal metamor-
phism compared to the kerogens studied by Marshall et al.
(2007). A curie point pyrolysate was dominated, as would
have been expected, by aromatic hydrocarbons but also con-
tained suites of long-chain hydrocarbons comprising alkanes,
alkenes, and alkyl benzenes, all with pronounced odd-over-even
carbon number preferences. This compound distribution is a
diagnostic feature for organic matter of biological origins
(Summons et al., 2008).
Bitumens preserved in Archean rocks comprise extractable
hydrocarbons as well as hydrocarbons preserved in fluid inclu-
sions, in crystalline minerals and pyrobitumens. There are many
reported instances of pyrobitumens, preserved as globules or
nodules (thucolites) and carbon seams, which likely record
former episodes of petroleum generation and migration
(Rasmussen, 2005; Rasmussen and Buick, 2000). These deposits
very often occur in association with gold or uranium
mineralization and the bitumen can be found coating grains
of detrital uraninite, monazite, xenotime, zircon, and thorite.
Multiple processes can be responsible for trapping this once-
mobile organic matter and would include thermal metamor-
phism of hydrocarbon-bearing porous reservoirs. The bitumi-
nous grain coatings likely result from in situ irradiation from
solid particles containing uranium and thorium (Nagy et al.,
1991; Parnell, 1988). As with younger radiation-altered
organic materials, these bitumens are resistant to molecular
characterization because of their high degree of cross-linking
(Dahl et al., 1988). Migration of bitumen from its source
requires that the original kerogen was concentrated and had
an H/C ratio sufficient to allow a fluid phase to form and
overcome the adsorptive capacity of the rock’s mineral matrix.
It is rather unlikely that petroleum could ever be generated
from sedimentary rocks that formed before the advent of
photoautotrophy.
Solvent-soluble organic compounds (bitumens), mostly
hydrocarbons, have been obtained from rocks of all ages. How-
ever, there are very few examples where their occurrence in
Archean sediments is supported by robust evidence for their
syngenicity. Water-soluble organics, including amino acids, car-
bohydrates, nucleic acids, and other directly biological products,
would never have survived unaltered in the thermal regime of
the Archean greenstone belts. Hydrocarbons are considerably
more stable but it would still be difficult to envisage ways in
which they could be preserved in the oldest (>3 Ga) terrains of
northern Australia and southern Africa where total organic car-
bon contents are low and the surviving kerogens haveH/C ratios
approaching zero. The 2.8–2.4 Ga Fortescue and Hamersley
sequences of the Pilbara Craton, and the Griqualand rocks of
the Kaapvaal Craton however contain an abundance of organic
carbon-rich sediments that would be classed as potential petro-
leum source rocks if they were younger than 500 My old. Black
shales and carbonates from these successions have been studied
intensively for almost two decades. The initial studies by Brocks
and coworkers (Brocks, 2001; Brocks et al., 1999, 2003a,b)
reported traces of hydrocarbons, including triterpenoids diag-
nostic for bacteria and eukaryotes in close association with
organic carbon rich black shales but not in the interbedded
low-total organic carbon content sediments and volcanics. Thor-
ough, for the time, analyses and arguments posited that the
hydrocarbons were ‘likely syngenetic’ but the possibility of con-
tamination from younger sediments could not be completely
ruled out (Brocks et al., 2003a).HyPy experiments conducted on
kerogens isolated from some of these rocks yielded predomi-
nantly aromatic assemblages of hydrocarbons and failed to pro-
duce the saturated steroids and hopanoids that were present in
the solvent extracts. In another study, the distributions of
methylated hopanoids in similarly aged and preserved sedi-
ments could be correlated with the isotopic compositions of
associated kerogens and, for the first time, provided data that
related a mobile organic component to one that was in situ
(Eigenbrode et al., 2008). Subsequent studies by Brocks and
others have cast doubt on the validity of this work. Firstly,
hydrocarbons found in the Pilbara cores include molecules
that can be traced to contamination from plastic and are of
undoubted anthropogenic origin (Brocks et al., 2008). Secondly,
the spatial patterns of hydrocarbons and, especially, their con-
centrations near to the external surfaces of cores are proposed as
1.0
0.8
0.6
0.4
0.2
051 52 53 54 55 56 57 58
1.0
0.8
0.6
0.4
0.2
00 0.1 0.2 0.3 0.4 0.5
1.0
0.8
0.6
0.4
0.2
031 33 35 37 39 41
R2= 0.919
R2= 0.984
R2= 0.967
Met
hylp
hena
nthr
enes
/Phe
nant
hren
e
H/C atomic ratio
ID1/(ID1+ IG) (%)
La (nm)
Figure 3 Parallel changes in the composition and maturation of Archean organic matter as shown by comparison to a molecular maturation parameter(S(methylphenanthrenes)/phenanthrene). Lowering of the H/C atomic ratio is due to progressive cracking of high-H/C molecules, while it appearsthat parameters measured by Raman spectroscopy, which reflect the crystallinity of organic matter, are also robust indicators of thermal overprinting ofkerogen. After from Marshall CP, Love GD, Snape CE, et al. (2007) Structural characterization of kerogen in 3.4 Ga Archaean cherts from the PilbaraCraton, Western Australia. Precambrian Research 155: 1–23.
Organic Geochemical Signatures of Early Life on Earth 39
evidence for petroleum-derived contamination ofmuch younger
age (Brocks, 2011). Thirdly, in situ carbon isotopic analyses of
kerogens and solid bitumen phases in core material from the
Hamersley are proposed to exclude any genetic relationship
between the kerogen and soluble bitumen components of the
organic matter (Rasmussen et al., 2008). Although the approach
is a sound one, the samples utilized for in situ isotope analyses
by Rasmussen et al. (2008) were different (J. Brocks, personal
communication, 2011) from those studied for extractable
hydrocarbons by Brocks et al. (1999, 2003a,b). Accordingly, in
this case, the comparisons are invalid and do not constitute a
robust test of biomarker syngenicity.
40 Organic Geochemical Signatures of Early Life on Earth
In more recent work, improved methods were applied to
the analysis of hydrocarbons present in cores recovered during
the Agouron Griqualand Drilling Project, where over 2500 m
of well-preserved Late Archean to earliest Proterozoic Transvaal
Supergroup sediments, dating from c.2.67 to 2.46 Ga were
recovered (Sherman et al., 2007a; Waldbauer et al., 2009).
New approaches which have been implemented include a
conventional extraction (denoted bitumen 1) after which the
rock was demineralized with hydrochloric and hydrofluoric
acids to afford a mineral-occluded component (denoted bitu-
men II). This fraction has been shown to have a lower apparent
maturity than the freely extractable organics in different aged
sediments as well as subtle differences in biomarker parameters
that are responsive to lithology (Sherman et al., 2007b;
Nabbefeld et al., 2010). These differences suggest that the
mineral-occluded hydrocarbon fraction is distinct and less
prone to external contamination. In another approach, bio-
marker profiles of stratigraphically correlated intervals from
diverse lithofacies in two boreholes, separated by 24 km as
well as across a c.2 Gy unconformity, provided support for
the syngeneity of the extractable hydrocarbons. These analyses
were accompanied by a raft of other geological and isotopic
studies that provide a sound paleoenvironmental context in
which to interpret the biomarker data (Fischer et al., 2009;
Knoll and Beukes, 2009; Ono et al., 2009). Further work on
these cores and a recently completed drilling campaign in the
Pilbara Craton, where samples from three holes were recovered
using clean procedures and only water as lubricant, should
provide additional evidence with which to evaluate the synge-
nicity of the biomarkers isolated from these Neoarchean basins.
Fluid inclusion hydrocarbons comprise a special class of
bitumens in that they are encased in crystalline minerals such
as calcite, quartz, and feldspar (Bhullar et al., 1999; Jensenius
and Burruss, 1990; Munz, 2001). They are visible under the
microscope and form a preserved record of fluid migration th-
rough the sedimentary system in which they are found (George
et al., 2008). In Phanerozoic sediments, hydrocarbon-filled fluid
inclusions in sandstones can be geochemically mapped to epi-
sodes of petroleum migration and entrapment (Dutkiewicz
et al., 2004; George et al., 2004) and have been used to recon-
struct the filling histories of petroleum reservoirs (Dutkiewicz
et al., 1998, 2006). Oil-bearing fluid inclusions have been dis-
covered in Archean successions from the Pilbara and Kaapvaal
cratons (Buick et al., 1998) and in the fluvial-deltaic to marine
Paleoproterozoic Huronian Supergroup in Canada (Dutkiewicz
et al., 2006). In the case of the 2.45 Ga sediments of the
Matinenda Formation at Elliot Lake, Canada, oil – possibly
migrated from the conformably overlying McKim Formation –
was trapped in inclusions within quartz and feldspar crystals
before c.2.2 Ga and was present in quantities sufficient to allow
detailed characterization. The range of compounds detected
included n-alkanes, acyclic isoprenoids, monomethylalkanes,
aromatic hydrocarbons, low-molecular-weight cyclic hydrocar-
bons, and traces of complex polycyclic biomarkers including
steranes and triterpanes. In other words, the hydrocarbons com-
prised a similar distribution to those detected in earlier studies
(e.g., Brocks et al., 1999; Waldbauer et al., 2009). Molecular
maturity parameters showed that the oil was generated in the
oil window; there was no evidence of cracking, an observation
attributed to the fact that such inclusions are closed systems with
high fluid pressures with an absence of minerals that might
catalyze decomposition. The biomarker geochemistry of Mati-
nenda Formation fluid inclusion oils suggests that oxygenic pho-
tosynthesis was extant at the time of source rock deposition at
c.2.2 Ga. The methodology developed in this study, with its low
detection limits and low system blanks, could help to resolve the
controversies surrounding Archean shale-hosted biomarkers
(Dutkiewicz et al., 2006; George et al., 2008).
12.2.6 Visible Structures with Organic Affinities
12.2.6.1 Organic-Walled Microfossils
Microfossils can yield unambiguous insight into the existence of
early life on Earth but their interpretation can be complicated by
a multitude of factors. The classification of organic particles or
inorganic coatings on mineral grains as microfossils attempts to
ascribe taxonomic or metabolic affinity to these objects, and
even their original nature are among the most common topics
of debate and argument.
The oldest putative microfossils are Paleoarchean in age and
were discovered in the Barberton Greenstone Belt of South
Africa (Walsh and Lowe, 1985) and the in Pilbara Craton of
northwestern Australia (Awramik et al., 1983; Schopf and
Packer, 1987). The latter were found in the Towers Formation,
an assemblage of thick chert units alternating with basaltic,
felsic volcanic and clastic sedimentary units, and in a chert
unit of the Apex Basalt, both of which are stratigraphically
located within the c.3.5 Ga Warrawoona group. Some of the
spheroidal and filamentous morphologies of the structures
found at the Chinaman Creek locality (Schopf and Packer,
1987) resemble modern and fossil cyanobacteria, which led
to an interpretation that these might be the earliest biological
remnants and the suggestion that oxygen-producing photoau-
totrophy might have already had developed at that point in
geological history (Schopf, 1993; Schopf and Packer, 1987).
However, the host rock that was initially interpreted as a
shallow marine siliceous deposit has been alternatively inter-
preted as a hydrothermal vein chert (Brasier et al., 2002;
Van Kranendonk, 2006). In the Brasier et al. (2002) study,
Raman data on carbonaceous particles were used to suggest
that these represent amorphous graphite, formed by Fischer–
Tropsch-type abiotic syntheses. Ever since a keen debate has
been waged between critics (Brasier et al., 2002, 2004, 2005,
2006; Marshall et al., 2011; Pinti et al., 2009) and adherents
(De Gregorio et al., 2009; Schopf et al., 2002, 2007) of the
Warrawoona microfossil theory. While some points of criti-
cism, such as issues over the nature of branching, were validly
refuted as artifacts of the automontaging software used to create
the photomicrographs (Schopf, 2004), criticism regarding
the biogenicity of these microfossils continues unabated. Most
recently, Marshall et al. (2011) studied objects in the Apex chert
that superficially resemble those reported earlier by Schopf and
Packer (1987) and claimed erroneously that they were fractures
filled with quartz and hematite. Notably, they studied material
that was unrelated to the original discoveries. However, Schopf
and colleagues have shown through confocal laser microscopy
and Raman imagery that the Apex carbonaceousmatter is struc-
turally and chemically complex and that the Apex microbe-like
features represent ‘authentic biogenic organic matter’ (Schopf
Organic Geochemical Signatures of Early Life on Earth 41
andKudryavtsev, 2011). Although someApex chert objectsmay
be pseudo-fossils, there is sound evidence for biology in this
and other units of the Warrawoona and overlying Sulfur
Springs successions. Filamentous microfossils were found in
c.3.2 Ga deep-sea massive volcanogenic sulfide deposits
(Rasmussen, 2000) and amore recent finding of 3.4 Ga organic
microfossils has been heralded even by former skeptics (Wacey
et al., 2011a,b). In the latter study, organicmicrostructures were
associated with pyrite crystals that were interpreted as primary
metabolic by-products of the microbes. The stable carbon and
sulfur isotopic signature of the fossil cell walls and pyrite crys-
tals were taken as an indicator of a sulfur-based metabolism.
This suggestion is in line with previous studies that presented
Paleoarchean sulfur isotopic data and documented the antiq-
uity of bacterial sulfur metabolism (Bontognali et al., 2012;
Philippot et al., 2007; Shen et al., 2001, 2009; Ueno et al.,
2008; Wacey et al., 2011a,b).
The assignment of a metabolic or even taxonomic affinity to
microfossils of Archean age is laden with complications
because of poor preservation, the prevalence of ambiguous
characteristics, and lines of evidence that frequently are circum-
stantial. Photosynthetic metabolism was ascribed to organic
particles interpreted to represent microfossils in the c.3.4 Ga
Buck Reef chert, but this interpretation was primarily based on
a stratigraphic distribution that is limited to shallow marine
sedimentary settings (Tice and Lowe, 2004). Similarly, the
existence of oxygenic photosynthesis has been ascribed to fos-
sil microorganisms whose morphologies resembled those of
modern cyanobacteria (Altermann and Schopf, 1995; Schopf,
1993; Schopf and Packer, 1987). Such identifications are facil-
itated in younger deposits: the oldest microfossils that are
classified with a certain degree of confidence as cyanobacterial
on the basis of high morphological similarities to modern
cyanobacteria (Hofmann, 1976; Knoll, 2002) are of Paleopro-
terozoic age. Similarly, the oldest certainly eukaryotic micro-
fossils are found in Mesoproterozoic strata (Han and
Runnegar, 1992; Javaux et al., 2001, 2003; Knoll et al., 2006;
Walter et al., 1990; Yan and Zhu, 1992; Yin, 1997). The ques-
tion of eukaryotic life during the Archean is debated (see
Chapter 6.5) and microfossils are inconclusive in providing
an answer. Javaux et al. (2010) reported structures that would
by all means deserve a eukaryotic classification by their large
size, but in the absence of further distinguishing criteria (Knoll
et al., 2006) such an interpretation remains inconclusive.
12.2.6.2 Fossil Microbial Mats, Textures, andTrace Fossils
Regions dominated by siliciclastic sedimentation are typically
not prime localities in the search for Archean fossil life due to a
very low level of in situ mineral formation and a generally poor
preservation potential for biomass – particulate organic matter
but also organic microfossils (but see Javaux et al. (2010) for
an interesting exception). However, benthic microbiota may
still influence sedimentary structures, even if none of the
organic matter of the mat is preserved over time. Extracellular
polymeric substances aid in a surficial consolidation of both
clastic and carbonate sediment piles (Decho et al., 2005;
Dupraz et al., 2009), which leads not only to an increased
erosional resistance but also to the formation of characteristic
textures upon further sedimentary burial or desiccation
(Noffke et al., 2001). Of such microbially induced sedimentary
structures (MISSs), the most prominent are wrinkle structures,
also termed elephant-skin texture (Gehling, 1999; Runnegar
and Fedonkin, 1992), desiccation cracks, and roll-up struc-
tures. While they can be a life-marker, information on taxon-
omy or metabolism is absent unless specific microfossils have a
taphonomic niche provided by the mat. Even the biological
source of perceived MISSs cannot always be certain as it can be
hard to distinguish true MISSs from irregularities on bedding
surfaces that arise from purely physical processes such as
impressions from moving foam, or small-scale load structures
among many others (Porada and Bouougri, 2007). Several
MISSs in Archean sedimentary rocks have however been criti-
cally evaluated and thought to be remnants of microbial mat
growth. Sandstones of the 2.9 Ga Mozaan Group contain wrin-
kle structures that host filamentous textural features on a
microscale (Noffke et al., 2003) and similar textural remnants
of presumably bacterial mats are found in the 300-My-older
Moodies Group (Noffke et al., 2006). Although incapable of
pinpointing taxonomy or metabolism with certainty, MISSs in
Archean rocks provide supportive evidence for the existence of
life during the Paleoarchean (Noffke, 2008).
12.2.6.3 Stromatolites
Stromatolites are generally accepted to be organosedimentary
structures produced by sediment trapping, binding, and/or pre-
cipitation as a result of the growth and metabolic activity of
microorganisms (Walter et al., 1976). Several details of their
formation are however debated. Foremost, the aforementioned
definition places them into the realm of biogenic structure,
which might not always be the case. Semikhatov et al. (1979)
provided an alternative definition of Stromatolites that does not
involve the action of biology: “. . . attached, laminated, lithified
sedimentary growth structures, accretionary away from a point
or limited surface of initiation.” Abiological formation of stro-
matolitic structures is indeed possible by chemical precipitation
(Grotzinger and Rothman, 1996). The absence of microstruc-
tures indicative of detrital trapping and binding – promoted by
bacterial extracellular polymeric substance, or EPS – in many
Precambrian stromatolites (Knoll, 2002) has led to the idea that
some of these structures could, in theory, have formed abioti-
cally. A second point of debate in the formation of stromatolites
involves the nature of the biological component. While
assumed principally cyanobacterial in an early definition by
Walter (1976), this is not necessarily the case as a variety of
mat-building microbes could engage in the formation of
stromatiform-lithified mats. Attempts to prove a cyanobacterial
involvement, which would lend credibility to fossil stromato-
lites as indicators for photosynthetic oxygen production, have
been pursued on the basis of modern observations (Burns et al.,
2004; Golubic, 1976; Neilan et al., 2002; Reid et al., 2000;
Walter et al., 1972, 1976), as well as analogies in cone spacing
(Petroff et al., 2010) and trapped crestal bubbles (Bosak et al.,
2009) betweenmodern and ancient coniform stromatolites. For
a more detailed analysis, the reader is referred to Chapter 6.5.
The oldest known stromatolites have been reported from
the 3.49 Ga Dresser Formation in the North Pole area of
the Pilbara Craton in Western Australia (Walter, 1983; Walter
42 Organic Geochemical Signatures of Early Life on Earth
et al., 1980). Here, a bed of laminated domical stromatolites
(Buick et al., 1981, 1995; Groves et al., 1981; Walter, 1983;
Walter et al., 1980) has been argued to represent the oldest
morphological trace of life on Earth. Somewhat younger rocks
at 3.35 Ga from the same region host the next oldest diverse
assembly of stromatolitic buildups. Carbonate units intercalated
in the Strelley Pool Formation contain stromatolites that have
been first reported and discussed by Lowe (1980, 1983, 1994).
It was, however, the reexamination of a locality that was first
discovered by Alec Trendall in 1984 (the ‘Trendall locality’) –
exhibiting exceptional morphological preservation over only a
few square meters – that revived the study of early Archean
stromatolites (Hofmann et al., 1999). Following that, a system-
atic study of diverse morphotypes occurring across >10 km of
the outcropping Strelley Pool Formation revealed the existence
of multiple discrete stromatolitic facies (Allwood et al., 2006)
that appear to occupy different paleoenvironmental settings
across an Archean peritidal platform. Apart from these morpho-
logical and contextual clues, a further argument for biogenicity
was based on observed differences between the laminae situated
on stromatolitic cones and those between them. The observa-
tions suggest a mainly mechanical deposition of grains in the
cone interspaces, whereas different processes – most plausibly
explained by microbial influence – must have acted on the cone
crests (Allwood et al., 2006). The question of biogenicity was
studied in greater detail by additional microscale analyses of
sedimentary fabrics (Allwood et al., 2009). An additional sign
of biological origins comes from the observation that cohesive
layers of organic material formed at regular intervals at the
surface of domal stromatolites and that those laminae adhered
to the steep stromatolite margins without exhibiting a prefe-
rential thickening in topographic lows. Furthermore, matches
between changing depositional modes of laminae and their
” C. Hallmann
” C. Hallmann
1.
3. 4.
2.
Figure 4 The earliest remnants of life on Earth. Stromatolites from the TrenWarrawoona Group in Northwestern Australia. Stratigraphy modified from Makerogen in 3.4 Ga Archaean cherts from the Pilbara Craton, Western Australia.Hallmann and Roger Summons.
thickness suggest a transition from microbial accretion by
trapping and binding toward accretion by precipitation, indicat-
ing the adaptation of stromatolitic systems to changing envir-
onments. In combination, all these observations make a strong
case for the existence of microbial mat communities by
3.45 Ga – a hypothesis that seems to have now found general
acceptance in the scientific community (Figure 4).
12.2.7 Summary and Prospects
Early life studies will always be subject to debates about what
constitutes a genuine fossil and what does not. Early life was
entirely microbial and comprised of interdependent commu-
nities of organisms that fed on each other and, thereby,
recycled most of the material that they processed. The isotopic
records of carbon, sulfur, and, potentially, other elements are
our best clues to the fact that life was present and driving
biogeochemical cycles at a global scale. Inevitably, however,
the oldest visible objects and chemical remnants of life in the
sedimentary record constitute an imperfect record (Knoll,
2012). They tend to be corrupted by the ravages of time and
temperature and reflect only those biological processes that
have preservable remains. Evidence of biogenicity of any puta-
tive fossil must include establishing a robust environmental
context based on sound geological and geochemical under-
standing as well as a preservation mechanism that is consistent
with that environmental setting (Summons et al., 2011).
Despite the problems of metamorphism and contamination
inherent in the Archean sedimentary record, there is still a
substantial legacy of past biological activity to explore, dissect,
and catalog, especially with the number of recent and pro-
posed boreholes being drilled into unweathered and,
Wyman Fm
Euro Basalt
Strelley Pool Chert
Panorama Fm
Apex Basalt
Mt Ada Basalt / Duffer Fm
Dresser Fm
North Star Basalt
V V V
V V V
V V V
V V V
V V V
V V V
V V V
V V V
V V V
V V V
V V V
V V V
V V V
V V V
V V V
Trendall locality
3.33 Ga
3.35 Ga
3.47 Ga
3.47 Ga
3.49 Ga
3.52 Ga
3.43−3.46 Ga
Schopf locality
Awramik locality
North pole stromatolites
(Antarctic Chert Member)
Basalt
Felsic volcanics
Chert
Conical stromatolites
Domal stromatolites
V VV
*
*
* *
dall locality (1 and 3) and the Dresser Formation (2 and 4) of thershall CP, Love GD, Snape CE, et al. (2007) Structural characterization ofPrecambrian Research 155: 1–23. Copyright of photographs by Christian
Organic Geochemical Signatures of Early Life on Earth 43
therefore, better-preserved sedimentary sequences (Garvin
et al., 2009; Kaufman et al., 2007; Knoll and Beukes, 2009).
A new paradigm to search for early life should be based
on the application of sound sedimentological principles and
combinations of emergent instrumental techniques. In situ
screening of organic matter using laser Raman imagery will
help identify the best-preserved materials for further study
while, at the same time, providing nonintrusive visible and
chemical data on microscopic object of interest. Systematic
evaluation of morphologies and multielement isotopic data
for the preserved organic matter at small spatial scales (House
et al., 2000; Rasmussen et al., 2008; Williford et al., 2011) can
enable the recognition of heterogeneities, which typically
characterize biological systems. Such information is largely
invisible in bulk sample analyses. Hydrocarbon analyses on
individual preserved fluid inclusions promise to reveal new
insights into molecular fossil distributions that carry signals
diagnostic for specific biological processes, including oxygenic
and anoxygenic photosynthesis, respiration, and methane
cycling (e.g., Dutkiewicz et al., 2006).
There are very likely archives of Earth’s early life that remain
to be discovered. Remote as they may be, meteorites on the
Moon and Mars could record early terrestrial crust that was not
destroyed by subsequent resurfacing. It is also possible that
some mantle rocks preserve isotopic records of organic carbon
that was once processed by living organisms. Our most acces-
sible prospects, however, are the vast expanses of Earth’s
Archean greenstone belts. Outcrops of, and cores drilled
into, these rocks may yet reveal zones of exceptional preserva-
tion of organic matter that contain valuable chemical and
microscopic fossils. Recent discoveries of large and complex
microfossils suggest that there is much undiscovered material
ripe for detailed microchemical and isotopic analyses (e.g.,
Sugitani et al., 2010). Claims of remnant Hadean crust
(Adam et al., 2012; O’Neil et al., 2008), although controver-
sial, indicate that much remains to be learned about early
Earth’s rock record.
Acknowledgments
The authors gratefully acknowledge the Agouron Institute and
the NASA Astrobiology Institute for support during the prepa-
ration of this review. Christian Hallmann thanks the Max-
Planck-Society for support. Malcolm Walter provided many
invaluable suggestions that improved the manuscript and we
thank Kliti Grice for her review of the submitted version.
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