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DOI 10.1007/s12210-009-0061-9
REVIEW
The coevolution of life and environments
Andrew H. Knoll
Published online: 22 October 2009
Ó Springer-Verlag 2009
James Hutton, the late eighteenth century father of Geology, recognized the goodness of fit
between organisms and the environments in which they live. To Hutton, this presented a
conundrum, because simple observation showed that the Earth is constantly changing.
Trees and shrubs grow on mountainsides, but erosion strips sediments from highlands and
carries them to the sea. Clams and seaweeds live in bays and estuaries, but the sediments
eroded from mountains are continually delivered to the ocean, where they fill in those
embayments. How can species be maintained in the face of such flux? Hutton’s solution,
presented in his seminal Theory of the Earth, was elegant in its simplicity: environmental
constancy is maintained dynamically (Hutton 1788). Through time, the uplift of sedimentary basins approximately balances erosion, perpetuating the mountains and bays for
which species are manifestly well designed. Half a century later, Darwin (1859) stood this
logic on its head. Species are no more constant than the Earth, reasoned Darwin; populations adapt via natural selection to the changing biological and physical circumstances of
their environment.
The idea that planetary surfaces not only change through time, but also change directionally first appeared in the writings of an astronomer, Percival Lowell, in a description of
Mars. Lowell (1908) famously, or infamously, claimed that he could discern linear features
on the Martian surface, and he interpreted these as canals built by technologically
advanced Martians to carry seasonal melt water from the glaciated poles to parched
populations at low latitudes. From this, Lowell concluded that Mars was once much wetter
than it is today, a view widely accepted by present day planetary scientists although
decidedly not for the reasons advanced by Lowell. A century after its publication, however,
Lowell’s general point seems strikingly modern: planetary surfaces provide potentially
habitable worlds that change systematically through time. Now that we have the
This article belongs to a special issue dedicated to the Meeting ‘‘Il mondo dopo Darwin’’, Accademia
Nazionale dei Lincei, Rome, 11 February–12 February 2009.
A. H. Knoll (&)
Department of Organismic and Evolutionary Biology, Harvard University,
Cambridge, MA 02138, USA
e-mail: [email protected]
URL: http://www.fas.harvard.edu/*knollgrp/people.htm
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geochemical tools to reconstruct our own planet’s past, we have come to appreciate that
Earth, as well, has experienced both long-term state changes and transient perturbations in
an environmental history more than 4 billion years long. This is the context in which
biological evolution has proceeded.
The conventional fossil record of plant remains and animal shells, bones and burrows
extends backward in time for about 575 million years. A much older biological record,
however, is written in different form—in tiny microfossils of bacteria and eukaryotic
microbes preserved in rocks that antedate the oldest animals, in fossil molecules that record
the presence of organisms not easily preserved (or at least recognized) on the basis of
conventional microfossils, in stromatolites that accreted through interactions between
microbial communities and the physical environment, and in the isotopic composition of
the carbon and sulfur in precipitated minerals and organic matter. Collectively, these
records enable paleontologists to record a history of life more than 3 billion years in
duration (Knoll 2003).
What about Earth’s environmental history? That, too, is encrypted in sedimentary rocks,
specifically in their chemistry. At the broadest level, Earth’s environmental history is
written in three great chapters. On the Archean ([2,500 million years old) Earth, little or
no free oxygen was present in the atmosphere or ocean. In contrast, during the Phanerozoic
Eon (literally, the age of visible animals: 542 million years ago to the present), the oceans
have, in most times and places, been oxygenated from top to bottom. In between lies the
Proterozoic Eon, a long lasting intermediate state of Earth’s biosphere characterized for
much of its extent by limited oxygen in the atmosphere and surface ocean, but sulfidic
waters at depth.
What evidence supports such a tripartite view of planetary history? First, as recognized
decades ago, the widespread distribution of iron formation in Archean sedimentary basins
records early oceans quite different from those of the present day. The transport of iron
through the ocean requires anoxic water masses capable of carrying Fe(II) in solution;
hence, the first-order conclusion that anoxic, ferruginous water masses characterized much
of the oceans’ volume before 2,450–2,320 million years ago. More recent evidence from
the distributions of oxygen-sensitive minerals, ancient soil horizons, and mass independent
sulfur isotopic fractionation make it clear that, at best, oxygen was a rare and transient
feature of air and water in Earth’s youth. First-order changes occur in all lines of evidence
around 2,400 Ma, recording the initial rise of dioxygen in the atmosphere and surface
ocean (sometimes called the Great Oxidation Event; Holland 2006).
For many years, majority opinion held that earliest Proterozoic redox change resulted in
oceans somewhat like today’s, with oxygenated deep waters removing iron by the precipitation of iron oxides. A decade ago, however, Canfield (1998) proposed a dramatically
different alternative. Earliest Proterozoic oxidation, he hypothesized, generated only a
modest amount of O2, perhaps a few percent of modern levels. This would be enough to
oxygenate the atmosphere and surface ocean, but too little to spread oxygen throughout
subsurface water masses of the ocean. However, there was a second prominent product of
Earth’s initial oxygenation, sulfate ion (SO42-), and this, suggested Canfield, was key to
iron removal from the oceans.
In Canfield’s view, export of organic matter from the photic zone would have supported
anaerobic respiration in anoxic water masses below the surficial mixed layer, with
microbial sulfate reduction providing a principal means of recycling organic carbon. The
metabolic by-product of sulfate reduction is H2S, and it was this sulfide, not oxygen, that
reacted with ferrous iron, removing it from the oceans as pyrite (Fe2S). In consequence, by
ca. 1,800 million years ago the Proterozoic Eon came to have sulfidic waters beneath the
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surface mixed layer, introducing a longlasting intermediate state of the biosphere distinct
from both the early Earth in which life first took root and our familiar Phanerozoic world.
How do we test such hypotheses? The answer is to search in ancient sedimentary rocks
for geochemical signatures of environmental state. The chemistry of iron in fine-grained
sediments provides insights into the redox state of water masses directly above the
accumulating sediments. Specifically, both the proportion of iron sequestered in precipitated minerals such as oxides, carbonates or sulfides and the proportion of those precipitated minerals present as pyrite differ systematically between muds (or their lithified
equivalents, shales) bathed by oxic waters and those deposited beneath deeper, anoxic
waters. Careful analyses of the iron in mid-Proterozoic basins show that sediments accumulating near the margins of marine basins formed beneath oxygen-bearing seawater,
whereas sediments depositing in deeper environments within basin interiors, accumulated
beneath anoxic water masses—just what Canfield predicted. In complementary research,
Brocks et al. (2005) identified fossil pigments synthesized by purple and green sulfur
photosynthetic bacteria in basinal shales from a 1,640 million year old succession in
northern Australia. Such molecules require that sulfide existed within the photic zone at the
time the sediments were deposited; that is, the oxygen minimum zone of this basin contained sulfide and extended to within a few tens of meters of the sea surface.
In summary, then, a number of independent lines of evidence confirm Canfield’s
hypothesis that during Earth’s middle age the oxygen minimum zone tended toward euxinia (anoxic and sulfidic conditions). The state of deeper water masses is less clear.
Canfield originally proposed that all water masses beneath the mixed layer should have
been sulfidic, but such a state is not easily reconciled with hypotheses of low O2 in the
atmosphere and surface ocean—in total, the supplies of oxidants and reductants have to be
in balance. More likely, deep oceans were dysoxic or anoxic, but not sulfidic, with euxinic
water mixed in dynamically like the swirls of a marble cake. As we shall see, however, it
was the state of the OMZ that most affected mid-Proterozoic life.
What was the nature of biological diversity in the three stages of Earth’s environmental
history? Most paleontological research focuses on the Phanerozoic Eon, and it is clear that
since the Cambrian diversification of marine animals and subsequent radiations of plants
and animals on land, Earth’s biota has maintained the broad phylogenetic pattern observable today. We know relatively little about Archean life, but have confidence that it was
microscopic—the major microbial metabolisms that sustain biogeochemical cycles
evolved during this interval. In anoxic early oceans, the iron and carbon cycles were
probably much more closely linked than they are today.
If the basic biology of bacterial and archaeal metabolism took shape in the first great
chapter of Earth history and the macroscopic world of complex multicellularity characterizes its third chapter, what was life like in chapter two, the long Proterozoic Eon?
Exceptionally, well-preserved microfossils show that cyanobacteria, the sole bacteria to
evolve oxygen-producing photosynthesis, were widespread and diverse in middle age
oceans (Knoll 2003). This does not mean that cyanobacterial diversification is completely
frozen in the past, but it does suggest cyanobacteria with essentially modern morphologies
(including limited cell differentiation) and physiological capabilities thrived in the Proterozoic biosphere.
Stromatolites provide independent testimony to the ecological importance of cyanobacteria in middle age oceans. Laminated sedimentary structures formed by the interaction
of microbial mats and physical processes, stromatolites occur in Proterozoic carbonates
deposited from tidal flats to the base of the photic zone. Microbial mats must, thus, have
been nearly ubiquitous on ancient seafloors, a distribution that would later diminish as
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seaweeds and sessile invertebrates evolved to compete for space on the seafloor and
animals came to graze on benthic microbial communities. Other microbial groups left their
calling card in a different way, imparting chemical signatures to Proterozoic rocks. Biochemical fractionations of both C and S isotopes tell us that the metabolisms that cycle
these elements through modern oceans operated, as well, in Proterozoic seas. It is the redox
state of the oceans and not the repertoire of metabolic processes that separates Earth’s
middle age from more recent times.
Much uncertainty attends the origins of eukaryotic cells, but the last common ancestor
of living eukaryotes, a cell with mitochondria capable of aerobic respiration, likely postdates the Great Oxidation Event. Certainly, microfossils as old as ca. 1,800 million years
old can be interpreted as eukaryotic on the basis of both morphology and wall ultrastructure. A modest diversity of eukaryotic microfossils occurs in shallow marine rocks
deposited between then and 800 Ma. Most cannot be assigned to a specific taxonomic
group, but beautifully preserved red algae in silicified carbonates from Arctic Canada
indicate that crown group divergence, photosynthesis and simple multicellularity had all
come to characterize the domain by 1,200 Ma. As discussed more fully below, increasing
eukaryotic diversity characterizes rocks ca. 800 Ma and younger.
The late Proterozoic transition to more fully oxic oceans appears to have been complex
(e.g., Scott et al. 2008). Rapidly increasing geochemical data suggest that oxic conditions
became widespread in the oxygen minimum zone 580–550 Ma. Recently, however,
Canfield et al. (2008) reported an earlier loss of sulfidic OMZs, replaced not by oxic
waters, but by (one last time) ferruginous subsurface water masses. The proposed
decoupling of sulfide loss and oxygen gain allows paleontological and biogeochemical
tests of biological and environmental coevolution.
Before evaluating the evolutionary consequences of environmental transition, we must
ask how the intermediate state of Earth’s biosphere could have been maintained for so
long. One possibility is that primary producers populated a feedback system that kept O2
levels moderate and the oxygen minimum zone sulfidic. Johnston et al. (2009) have proposed that sulfide in the oxygen minimum zone supported significant primary production
by anoxygenic phototrophs (recorded by the biomarker molecules noted above), partially
decoupling carbon fixation from oxygen generation. In consequence, fluxes of organic
matter into the OMZ exceeded that of oxygen, although for energetic reasons O2 remained
the oxidant of choice in respiration. This imbalance perpetuated oxygen depletion and, in
consequence, bacterial sulfate reduction in the OMZ, regenerating the H2S required to
sustain anoxygenic photosynthesis (Johnston et al. 2009).
Such biogeochemical feedbacks could, and apparently did, sustain Earth’s middle age
for more than a billion years. What, then, initiated Neoproterozoic transition to a more
modern biosphere? In combination, erosion of the surface sulfur reservoir and increased
iron fluxes associated with supercontinental break up apparently tipped the geochemical
balance in subsurface water masses back to iron 150 million years or more before latest
Proterozoic oxygen increase. As alternative electron donors declined, oxygenic photosynthesis came to dominate primary production throughout the photic zone, and for the first
time in Earth history, organic carbon burial came to be balanced quantitatively by oxygen
production. Thus, enhanced burial of organic matter associated with high latest Proterozoic
rates of sediment accumulation resulted in higher levels of O2 in the atmosphere and
oceans.
The foregoing suggests that biological processes played a critical role in sustaining
Earth’s long middle age, but that the physical Earth facilitated environmental transition.
Setting the fossil record within the framework of late Proterozoic environmental change,
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what might we conclude about the evolutionary consequences of environmental transition?
First, sulfide is generally toxic to eukaryotes—among other effects, it interferes with
cytochrome oxidase activity in mitochondria. For this reason, the widespread presence of
sulfidic water masses just a few tens of meters below the sea surface should have dampened eukaryotic evolution in all but the shallowest reaches of the ocean. Existing paleontological and biomarker geochemical data do, in fact, suggest that Neoproterozoic
collapse of OMZ euxinia coincided in time with a major increase in eukaryotic diversity in
the oceans. Perhaps, 90% of the microscopic eukaryote species described from welldocumented ca. 800–720 million-year-old fossil assemblages have no representation in
older rocks. Steranes, molecular fossils sourced mostly by eukaryotes, become abundant
features of the organic geochemical record at this time, as well. Besides sulfide loss,
increased nitrate and Mo availability in the photic zone may have facilitated algal
expansion (Anbar and Knoll 2002).
Of course the defining biological feature of Earth’s third biospheric state is complex
multicellularity. Three-dimensional multicellularity in which only a subset of cells are in
direct contact with the environment must have been limited to small sizes by environmental oxygen levels in organisms dependent on the diffusion to oxygenate interior cells
(Knoll and Hewitt 2009). With rising oxygen, organism size would increased, leading to
increasingly strong oxygen, nutrient and molecular signal gradients within tissues. In a
functional sense, complex multicellularity can be understood as the circumvention of
diffusion by active transport of nutrients and signaling molecules, and this is the breakthrough that permitted the radiations of bilaterian animals; complex red, green and brown
algae; and fungi. Consistent with this scenario, biomarker molecules and microfossils
record nascent animals in rocks 632–650 million years old, or older, but macroscopic
metazoans and motile bilaterian animals with high rates of exercise metabolism enter only
575–555 million years ago, at the time of Ediacaran oxygen transition. Equally, whereas
simple multicellular red algae evolved 1,200 million years ago, three-dimensionally
complex florideophyte reds first occur in ca. 551 million-year-old rocks from China, and,
while simple coenocytic and multicellular green algae occur in earlier deposits, macroscopic seaweeds likely to record complex multicellularity in this group appear only near
the Proterozoic–Cambrian boundary.
Nothing in the foregoing downplays the importance of organism–organism interactions in large scale evolutionary pattern. Indeed, Butterfield (2007) has argued compellingly that the introduction of animals permanently changed ecosystem structure and
the nature of selective pressures in the oceans. To paraphrase G. E. Hutchinson, however, the evolutionary play takes place in an environmental as well as an ecological
theater, and one with continually shifting sets and props. It is difficult to account for the
deep history of life without placing it firmly in the context of Earth’s environmental
evolution.
Paleontology’s chief contribution to the Neodarwinian Synthesis, George Gaylord
Simpson’s (1944) Tempo and Mode in Evolution, hardly mentioned environmental history
and devoted less than a page to mass extinction; Simpson’s objective was to show how
evolutionary processes illuminated by population biologists could account for evolutionary
pattern in the geologic record. By the 1970s, however, it had become clear that paleontological pattern is not simply the product of population genetics played out on million
year time scales, an intellectual shift most famously summarized by Steven Stanley’s
epigram: ‘‘macroevolution is decoupled from microevolution’’. Today, this ‘‘decoupling’’
is discussed in terms of evolutionary hierarchies and, key in my estimation, Earth’s
unpredictable environmental history. Indeed, to be first approximation, macroevolutionary
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pattern in the fossil record may largely reflect microevolutionary processes played out on
an environmentally dynamic planet.
This fundamental dynamic, in which populations continually track moving environmental targets, provides a required framework for assessing the evolutionary consequences
of current global change. It helps us to understand the rich evolutionary history recorded in
Phanerozoic rocks, with plankton to pachyderms responding to shifting continents, oscillating climate, large, but transient environmental perturbation and, of course, changes in
biological components of perceived environment. On the time scale of Phanerozoic history, such environmental interactions are as important as population genetics to an
understanding of how the biological diversity of our present moment came to be.
In the twenty-first century, the tools are finally in place to understand the coevolution of
Earth and life across the entirety of Earth history, documenting environmental events that
Hutton could scarcely have imagined, a pre-Cambrian evolutionary history that Darwin
predicted, but never expected to establish, and a long-term directionality in Earth’s
planetary history that Lowell would have applauded. Indeed, in the new millennium,
outcrop scale investigation of environmental history has been extended to Mars confirming
that directional change is the rule there, too. The big picture underscores the view that life
has truly been a planetary phenomenon, shaping and being shaped by Earth’s dynamic
surface, from the moment it began.
References
Anbar AD, Knoll AH (2002) Proterozoic ocean chemistry and evolution: a bioinorganic bridge? Science
297:1137–1142
Brocks JJ, Love GD, Summons RE, Knoll AH, Logan GA, Bowden S (2005) Biomarker evidence for green
and purple sulfur bacteria in an intensely stratified Paleoproterozoic ocean. Nature 437:866–870
Butterfield NJ (2007) Macroevolution and macroecology through deep time. Palaeontology 50:41–55
Canfield DE (1998) A new model for Proterozoic ocean chemistry. Nature 396:450–453
Canfield DE, Poulton SW, Knoll AH, Narbonne GM, Ross G, Goldberg T, Strauss H (2008) Ferruginous
conditions dominated later Neoproterozoic deep-water chemistry. Science 321:949–952
Darwin C (1859) On the origin of species by means of natural selection. J. Murray, London
Holland HD (2006) The oxygenation of the atmosphere and oceans. Philos Trans R Soc B Biol Sci 361:903–
915
Hutton J (1788) Theory of the Earth; or an investigation of the laws observable in the composition,
dissolution, and restoration of land upon the globe. Trans R Soc Edinb 1(2):209–304
Johnston DT, Wolfe-Simon F, Pearson A, Knoll AH (2009) Anoxygenic photosynthesis modulated Proterozoic oxygen and sustained Earth’s middle age. Proc Natl Acad Sci USA 106:16925–16929
Knoll AH (2003) Life on a young planet: the first three billion years of evolution on Earth. Princeton
University Press, Princeton
Knoll AH, Hewitt D (2009) Complex multicellularity: phylogenetic, functional and geological perspectives.
In: Sterelny K, Calcott B (eds) The major transitions revisited. Vienna series in theoretical biology,
MIT Press
Lowell P (1908) Mars as the abode of life. New York, The McMillan Company
Scott C, Lyons TW, Bekker A, Shen Y, Poulton SW, Chu X, Anbar AD (2008) Tracing the stepwise
oxygenation of the Proterozoic ocean. Nature 452:456–458
Simpson GG (1944) Tempo and mode in evolution. Columbia University Press, New York
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