Scripta Varia

Biodiversity and Global Change: From Creator to Victim to a new Revolution

Tim Lenton, University of Exeter, UK 

[Note to the reader: This is a preliminary work-in-progress, put together in haste, with incomplete and uneven referencing. I am unsure whether the scope is too broad or the level too elementary and whether large sections are superfluous. The last section on the projected consequences of climate change for biodiversity is at a very early stage of writing. Hence I welcome any frank feedback.]


The premise of this paper is that we owe our very existence to the activities of past and present life forms, which have created a world that we could inhabit.1 This is true not just in the evolutionary sense that we are descended from earlier life forms, but in the Earth system sense that the atmosphere would be unbreathable and the climate intolerable were it not for the accumulated actions of other members of the biosphere, past and present.1 This fundamental role of biodiversity in maintaining our life support system is strangely under-recognised by utilitarian arguments for preserving nature. Hence the first section briefly reviews which life forms and which functions they perform have been particularly important in creating a world we could inhabit in the first place.

We are part of biodiversity and have been born out of this world, only to be transforming it now in ways that are bad for us and bad for much of the rest of life. As the first section reviews, this is not the first time that life has radically transformed the planet with damaging consequences.1 However, it is the first time that a single animal species has wielded such world changing power. Past revolutionary transformations of the biosphere took life to the brink of total extinction in events such as ‘snowball Earth’. Our very existence dictates that life survived such past scrapes with disaster1 – but the flipside is that there is no guarantee that the biosphere will necessarily survive what is unfolding now. After past close shaves, it typically took millions of years for the slow workings of Earth system dynamics and the blind watchmaker of natural selection to restore a well-functioning, self-regulating biosphere.1 We don’t have the luxury of waiting that long.

We are meant to be Homo sapiens – wise man (sic) – instead we are practicing an act of cosmic stupidity: In extinguishing biodiversity we are eroding our own life-support system. The irony of this would surely not be lost on any watching deity. The second section reviews how we came to be planet changers and how climate change interacting with other global changes is predicted to impact biodiversity. This turns to some critical reflection on just how big a threat the unfolding extinction could be to the core functioning of the biosphere, or conversely whether the organisms that really run the planet can acclimate and adapt to the global changes we are causing.

Finally, rather than just chart the biodiversity dimensions of our demise, the aim here is to leave the reader with a chink of light to cling to – to consider how if we really are wise, and we learned to properly value biodiversity, could we become good citizens of a flourishing future biosphere? If so, what might that revolution look like?

How life created a world we could inhabit

The aim of this section is to sketch in broad terms how the past evolution of life has created a world that we could inhabit in the first place.1 Along the way we identify some crucial roles of different types of life – from prokaryotes to fellow multicellular eukaryotes – in creating and maintaining a habitable biosphere for advanced multicellular life forms including ourselves. The perspective here is that life and the Earth have co-evolved in the sense that the evolution of life has shaped the planet, changes in the planetary environment have shaped life, and together this can be viewed as one process.2 When we look at this co-evolution over Earth history, a relatively few revolutionary changes leap out, in which the Earth system was radically transformed.1 Each of these revolutionary changes depended on the previous one, and without them we would not be here – and nor would much of the biodiversity that is currently described and projected to be under threat (i.e. complex eukaryote life forms). To help orientate the reader, Figure 1 provides a timeline of major events in Earth history.

Origins of the biosphere

The most fundamental change in the history of the Earth system started with the origin of life, which appeared on Earth remarkably soon after our planet became continuously habitable.

The formation of the solar system is dated from the oldest meteoritic material at 4.567 billion years ago. The Earth and the other planets are younger than this, because they had to form from the gravitational collisions and accumulation of material spinning around the early Sun – in a process called accretion. During the accretion of the Earth there were some truly massive collisions, one (or several) of which formed the Moon 4.470 billion years ago. Whilst the Earth was still forming, the gas giants Jupiter and Saturn had finished accreting. Their gravitational pull disrupted the band of asteroids between Mars and Jupiter, sending some of them off on elliptical orbits that crossed the inner Solar System. Crucially, this brought water and other volatile substances, including nitrogen and carbon dioxide, to the early Earth. Remarkably, some tiny bits of the Earth’s crust from this time are still present at the surface today, in the form of grains of the mineral zircon that can be precisely dated. The oldest is 4.374 billion years old and was originally part of a granite rock, indicating that the continental crust had started to form in the first 100 million years of the planet’s history. Even more remarkably the isotopic composition of oxygen in the zircon suggests that oceans of liquid water were present on the Earth at that early time. However, the onslaught from outer space was not over. The Earth and all the inner Solar System suffered a ‘Late Heavy Bombardment’ by asteroids between about 4.1 and 3.9 billion years ago. Some of these impacts were large enough to have evaporated the early oceans and thus temporarily rendered the planet uninhabitable. Thus the Earth did not become ‘continuously habitable’ until the end of this bombardment around 3.85 billion years ago.

Remarkably the first tentative evidence for life on Earth comes within ~50 million years of the end of the Late Heavy Bombardment, 3.8 billion years ago. As soon as there are sedimentary rocks that could record the presence of life they suggest it is there. Some of the first evidence takes the form of small particles of graphite – organic carbon – with an isotopic composition (d13C) consistent with carbon fixation by some form of chemosynthesis or photosynthesis.3,4 Putative microbial structures have also been described from the same rock sequence5. By 3.5 billion years ago the first putative microscopic fossils of life appear,6,7 although not everyone is convinced the fossil structures were made by biology.8,9 There are also more convincingly biogenic sedimentary structures formed by microbial mats.10 By 3.26 billion years ago there are microfossils of cells caught in the act of division.11

The earliest biosphere was made up exclusively of prokaryotes – bacteria and archaea – two kingdoms of life which divided very early. Metabolically the earliest life forms may have consumed compounds in their environment that could be reacted to release chemical energy, but a shortage of chemical energy on a global scale would have severely restricted the productivity of such a biosphere. One possibility is that early archaea consumed hydrogen from the atmosphere and carbon dioxide to make methane, but such a methanogen-based biosphere would have been restricted to around a thousandth of the productivity of the modern marine biosphere.12,13

A more productive global biosphere would have arisen when early life began to harness the most abundant energy source on the planet – sunlight. Photosynthesis fixing carbon dioxide from the atmosphere appears to have evolved very early in the history of life. Whilst the ~3.8 billion year old graphite has a carbon isotopic composition consistent with photosynthesis, some scientists argue that there are non-biological ways to make graphite with this isotopic signature. However, 3.5 billion years ago the earliest carbonate sediments have an isotopic signature that suggests significant organic carbon burial globally, which must have been supported by photosynthesis.

The first photosynthesis was not the kind we are familiar with, which splits water and spits out oxygen as a waste product. Instead, early photosynthesis was ‘anoxygenic’ – meaning it didn’t produce oxygen. It could have used a range of compounds, in place of water, as a source of electrons with which to fix carbon from carbon dioxide and reduce it to sugars, including hydrogen (H2) or hydrogen sulphide (H2S) in the atmosphere, or ferrous iron (Fe2+) dissolved in the ancient oceans. The early biosphere fuelled by anoxygenic photosynthesis would have been limited by the supplies of these electron donors, all of which are a lot less abundant than water.

In fact, shortage of materials would have posed a more general problem for life within the early Earth system: The fluxes of materials coming into the surface Earth system from volcanic and metamorphic processes today are many orders of magnitude less than the fluxes due to life at the surface of the Earth today, indicating that today’s biosphere is a phenomenal recycling system (Figure 2).

Early life thus had to evolve the means of recycling the materials it needed to metabolize – in other words, to establish global biogeochemical cycles. We have a very scant record of how and when this happened, but a few clues suggest it was very early in the history of life. For example, the phylogenetic tree of prokaryote life suggests that many recycling metabolisms, such as methane production, evolved early on. The ease or difficulty of evolving recycling has also been explored by seeding computer models with ‘artificial life’ forms and leaving them to evolve. In these ‘virtual worlds’ the closing of material recycling loops emerges as a very robust result.14,15

If the early biosphere was fuelled by anoxygenic photosynthesis, plausibly based on hydrogen gas, then calculations suggest that once the biological recycling of this gas had evolved, the early biosphere might have achieved a global productivity up to 1% of the modern marine biosphere.12,13 If early anoxygenic photosynthesis used the supply of reduced iron upwelling in the ocean then its productivity would have been controlled by ocean circulation and might have reached up to 10% of the modern marine biosphere.12,13

All of this early evolution played out in a world essentially devoid of oxygen (Figure 3).

The Oxygen Revolution

The innovation that super-charged the early biosphere was the origin of oxygenic photosynthesis, which uses abundant water as an electron donor. This was not an easy process to evolve.16 To split water requires more energy – i.e. more high energy photons of sunlight – than any of the earlier anoxygenic forms of photosynthesis. Evolution’s solution was to wire together two existing ‘photosystems’ in one cell and bolt on the front of them a remarkable piece of biochemical machinery that can rip apart water molecules. The result was the first cyanobacterial cell – the ancestor of all organisms performing oxygenic photosynthesis on the planet today.

Current evidence suggests oxygenic photosynthesis took up to a billion years to evolve, with the first compelling evidence appearing around 3-2.7 billion years ago.17 The smoking gun is chemical evidence for oxygen leaking into the environment and reacting with metals that are highly sensitive to the presence of oxygen. For example, molybdenum is mobilized from continental rocks by reacting with oxygen and appears for the first time in ocean sediments around 2.7 billion years ago. Once oxygenic photosynthesis had evolved, the productivity of the biosphere would no longer have been restricted by the supply of substrates for photosynthesis, as water and carbon dioxide were abundant. Instead, the availability of nutrients, notably nitrogen and phosphorus would have become the major limiting factors on the productivity of the biosphere – as they still are today.

Once there was a source of oxygen on the planet it is tempting to assume that the concentration of oxygen in the atmosphere would have steadily risen – a bit like filling a bath with the plug in. But oxygen did not rise in the atmosphere immediately or steadily. Instead it remained a trace gas for hundreds of millions of years. We know this because a very peculiar ‘mass independent fractionation’ (MIF) of sulphur isotopes is preserved in sediments more than 2.45 billion years old.18 This MIF signature can still be produced by photochemistry of sulphur gases in the atmosphere today, but it cannot be preserved in today’s sediments because all the sulphur first goes through a homogenising reservoir of sulphate in the ocean. Prior to 2.45 billion years ago that sulphate reservoir must have been absent due to a lack of oxygen to produce it. The MIF signature indicates that high energy ultraviolet radiation streamed through the lower atmosphere and therefore the ozone layer was absent, requiring that oxygen (from which ozone is made) was at a concentration of less than 2 parts per million in the atmosphere (Figure 3).

Oxygen could remain at such a low concentration for hundreds of millions of years because there was an excess input flux of reduced materials hungry to react with it, including reduced iron injected into the ocean through mid-ocean ridges, and reduced gases such as hydrogen and hydrogen sulphide entering the atmosphere via volcanoes. The rate of their reaction with oxygen increases with oxygen concentration, thus producing a negative feedback system that stabilized the oxygen concentration at trace levels where the sink of oxygen matched the source.19 To use the bath metaphor for oxygen in the atmosphere; the plug was out and the plug hole was large, creating a low, stable level for oxygen.

This stability broke down after hundreds of millions of years, when atmospheric oxygen jumped up in concentration in an event known as ‘The Great Oxidation’ around 2.4 billion years ago (Figure 3). The mass independent fractionation of sulphur isotopes stops, indicating that oxygen had risen sufficiently to convert all sulphur to sulphate before it was deposited in marine sediments. The fact that the MIF signature has never returned suggests the permanent formation of an ozone layer. Massive deposits of oxidized iron appeared in the form of the first sedimentary ‘red beds’. Rusted (oxidized) iron also appeared in ancient soils for the first time. These indicators reveal that oxygen concentration increased by several orders of magnitude from less than 100,000th to circa 1-10% of its present atmospheric level.20 The indicators of oxygen rise all remain with us to the present day, indicating that the Great Oxidation was never reversed.

Whilst the origin of cyanobacteria performing oxygenic photosynthesis was ultimately responsible for the Great Oxidation, there must have been other long-term changes in the Earth system that slowly oxidized the surface Earth system. Fundamental among these was the loss of hydrogen atoms to space. This is a tiny flux on today’s Earth, because water is frozen out of the atmosphere at the ‘cold trap’ between the troposphere and stratosphere. Hence hardly any hydrogen-containing gases make it to the top of the atmosphere. However, much of the organic carbon created by early oxygenic photosynthesis would have been recycled to the atmosphere as methane by archaea. In the resulting methane-rich early atmosphere, much more hydrogen could escape to space and this had the effect of oxidising the surface of the Earth.21 This drove the Earth system towards a tipping point where the balance of inputs to the atmosphere shifted from an excess of reduced material to an excess of oxygen.

The abruptness of the Great Oxidation suggests that at this point a strong positive feedback process kicked in and propelled the rise of oxygen.19,22 The formation of the ozone layer was crucial to this transition, because it temporarily suppressed the consumption of oxygen. Ultraviolet radiation catalyses a series of reactions in which oxygen combines with methane to produce carbon dioxide and water (thus reversing the production of oxygen and methane by the biosphere). Without an ozone layer, this process for removing oxygen was rapid and efficient. But once enough oxygen built up for the ozone layer to start to form, this would have shielded the atmosphere below from UV and temporarily slowed down the removal of oxygen. More oxygen would produce more ozone, letting through less UV and further suppressing oxygen consumption in a positive feedback process. Models suggest that this positive feedback was strong enough to temporarily go into ‘runaway’ producing a rapid oxygen rise.19,22,23 However, the Earth system would soon have stabilized again at a higher oxygen level with the oxygen sinks again matching the sources.

When oxygen jumped up at the Great Oxidation this caused a decline in atmospheric methane concentration,23 slowing the further oxidation of the Earth. This decline in methane could help explain why when oxygen rose there were a series of glaciations. One of these glaciations, 2.2 billion years ago, reached to low latitudes near the equator and was probably the first ‘snowball Earth’ event. The Great Oxidation was also followed by a large pulse of organic carbon burial recorded in carbon isotopes. This may have been caused by increased oxygen reacting with sulphide in rocks on the continents producing sulphuric acid that dissolved phosphorus out of rocks and fuelled productivity in the oceans.24 If so, it reinforced the transition to an oxidising world. By 1.85 billion years ago this instability in the carbon cycle and climate had settled down and the Earth entered a long period of stability known rather unflatteringly as ‘the boring billion’.

The Complexity Revolution

The turmoil of the Great Oxidation created a world much more conducive to aerobic (oxygen-utilising) life forms. There was a lot more energy to go around in the post-oxidation world, because respiration of organic matter with oxygen yields an order of magnitude more energy than breaking food down anaerobically. Among the organisms to take advantage of this energy source were the first eukaryotes – complex cells with a nucleus and many other distinct components.

Eukaryotes are profoundly different from the prokaryotes that preceded them, yet they were created from the fusion of once free living prokaryotes. The mitochondria – the energy factory – in eukaryote cells were once free living aerobic bacteria, and the plastids in plant and algal cells – where photosynthesis occurs – were once free living cyanobacteria. These cellular components were acquired in ancient symbiotic mergers. The symbiotic merger that gave rise to mitochondria provided an abundant energy source to the ancestral eukaryote cell. Eukaryotes also rearranged how they copy genetic information – copying many chromosomes in parallel – whereas prokaryotes copy their DNA in one long loop. These innovations enabled eukaryotes to express many more genes than prokaryotes, and this ultimately gave them the capacity to create more complex life forms with multiple cell types.

The origin of eukaryotes is shrouded in mystery and controversy, as biologists do not agree on what marks the start of the lineage or what constitutes fossil evidence for eukaryotes. The earliest claims for biomarker evidence of eukaryotes 2.7 billion years ago are now thought to represent contamination with younger material.25 A couple of cryptic 2.5 billion year old ‘acritarch’ fossils might be the resting stages of early eukaryotes, but the name itself means they are of ‘confused origin’. Some 1.9 billion year old spiral fossils that are visible to the naked eye might be eukaryotic algae26 (called Grypania) but could also be colonial cyanobacteria. Molecular clocks suggest the last common ancestor of all eukaryotes lived roughly 1.8-1.7 billion years ago.

Eukaryotes only slowly realized their ability to build more complex life forms with differentiated cell types. Most of the fossils from Earth’s middle age – the Proterozoic Eon – are the rather cryptic ‘acritarchs’. Much rarer eukaryote body fossils include 1.5 billion year old Tappania, which might be a fungus,27 and 1.2 billion year old Bangiomorpha pubescens, which is a multicellular red alga (seaweed) assigned to a modern order.28

Researchers are still puzzling over what held back the evolution of complex life during ‘the boring billion’, but many see environmental constraints playing a key role. For most of the Proterozoic Eon the surface ocean remained dominated by prokaryotes and the deep oceans remained anoxic – i.e. devoid of oxygen.17 At intermediate depths some of these anoxic waters became ‘euxinic’, meaning that sulphate in the water was reduced to hydrogen sulphide, which is toxic to many eukaryotes.

Eventually the deadlock was broken in the latter parts of the Neoproterozoic Era (1000-542 million years ago), which witnessed a spell of climatic turbulence, the partial oxygenation of the deep oceans, and the rise of the first animals. The first signs of change began around 740 million years ago, when biomarkers of algae become more prevalent in ocean sediments and the diversity of eukaryote fossils starts to increase. This would have made the biological pump of carbon from the surface to the deep ocean more efficient.29 There were also productive microbial ecosystems on the land at the time30 and conceivably eukaryotic fungi, algae and lichens (a symbiotic merger of the two) could have been part of those early land ecosystems – although we have no fossil evidence either way.

Meanwhile, plate tectonics was breaking up the supercontinent Rodinia and scattering the resulting land masses in an unusual configuration, with much of the land in the tropics. This would have produced very efficient silicate weathering of the continents,31 potentially enhanced by biology.32 That in turn would have drawn down atmospheric carbon dioxide and cooled the planet. Somehow the climate got so cold that an extreme ‘snowball Earth’ glaciation – the Sturtian – was triggered around 715 million years ago. Glaciation reached equatorial latitudes and lasted tens of millions of years, consistent with the time it would take to build up enough carbon dioxide to melt the ice.

The climate turmoil did not end there. A second extreme glaciation – the Marinoan – was triggered, ending 635 million years ago. It was followed by a massive deposit of carbonate rock called a ‘cap carbonate’ – consistent with the snowball Earth hypothesis.33 In the super-hot and wet aftermath of snowball Earth, weathering would have occurred incredibly rapidly, supplying calcium and magnesium ions to the ocean that would combine with the excess carbon dioxide in the atmosphere and ocean to produce a massive deposit of carbonate sediments.

Perhaps the greatest puzzle about these extreme glaciations is how the ancestors of complex life survived them. Biomarker and molecular clock evidence suggests that simple animals in the form of sponges had already evolved,34 along with multicellular algae and fungi. Yet complex life did not flourish until after the glaciations. First there are fossils of what are thought to be animal embryos, alongside algae and fungi. Then the first large fossil organisms, the ‘Ediacaran biota’ appear around 575 million years ago. Whilst their biological affinity is debated, at least some were probably animals. They were followed tens of millions of years later by mobile grazing animals – both on the sediments and as zooplankton in the water column.

What triggered this burst of animal evolution? Relatively large, mobile animals need more oxygen than the sedentary creatures including sponges that came before them.29 Intriguingly, the first evidence for oxygenation of parts of the deep oceans appears 580 million years ago, shortly before the appearance of Ediacaran fossils at depth in the ocean. However, there had been oxygen in the shallow waters of the ocean for more than a billion years before this. It may be that evolution caused oxygenation rather than vice versa.29 By increasing the efficiency of carbon and phosphorus removal into sediments, the rise of sponges and algae may have oxygenated the ocean, improving conditions for ongoing animal evolution. The revolution in biological complexity culminated in the “Cambrian Explosion” of animal diversity 540 to 515 million years ago, in which modern food webs were established in the ocean. However, as animals began to burrow and bioturbate the ocean sediments this trapped phosphorus in them and ultimately lowered marine productivity and atmospheric oxygen levels.35 As a result anoxia became more widespread in the early Paleozoic oceans and animals may have limited their own expansion.

The thing that finally created a modern world was the rise of plants on land, beginning around 470 million years ago and culminating in the first global forests by 370 million years ago. Plants accelerated chemical weathering of the land surface, lowering atmospheric carbon dioxide levels and potentially cooling the planet into the Late Ordovician glaciations,36 as well as the later Carboniferous-Permian glaciations.37 The rise of plants doubled global photosynthesis, and with it the long-term source of oxygen from the burial of organic carbon. This increased atmospheric oxygen levels,38 finally fully oxygenating the deep ocean (Figure 3). The record of fossil charcoal starting 420 million years ago tells us that oxygen has remained above 15-17% of the atmosphere since then. This in turn provided one of the necessary conditions for our kind of intelligent life to ultimately evolve: Our brains are especially energy-hungry and if the partial pressure of oxygen in the air drops by about a third, brain function really starts to suffer.

Importantly plants and their associated mycorrhizal fungi also appear to have played a key role in regulating both atmospheric oxygen and carbon dioxide levels. Atmospheric oxygen is stabilised by negative feedbacks involve fires suppressing vegetation and thus the long-term source of oxygen.39 Atmospheric carbon dioxide is limiting to vegetation productivity which in turn is the key control on the long-term sink of CO2 from the weathering of silicate rocks, providing a potent negative feedback.40 A case can be made that oxygen, carbon dioxide and climate are now more tightly regulated than they have been previously in Earth history, thanks to the involvement of complex life in the feedbacks, and that the resulting stability has provided an important pre-condition for the further evolution of complex life forms1. Certainly the biodiversity of complex life has been on an impressive up-curve over Phanerozoic time. There have been setbacks of varying sizes – including ‘mass’ extinctions. However, the background rate at which species are lost has been declining over Phanerozoic time,41 which can be taken as one measure of increasing Earth system stability on the longest timescales.

On the somewhat shorter timescale of mass extinctions, however, Earth system processes can be part of the problem. Whilst scientists tend to look for external causes to mass extinctions – usually in the form of a massive meteorite strike and/or an episode of intense terrestrial volcanism creating a ‘large igneous province’ of lava flows – such external triggers have clearly been amplified by processes internal to the Earth system, which has at least one Achilles heel: There have been repeated episodes of expanded ocean anoxia that have played a key role in mass extinction events. The largest mass extinction of all at the End Permian, 252 million years ago, saw widespread ocean anoxia. Indeed it saw the Earth system revert back towards earlier states, with depletion of the ozone layer as well. The spread of ocean anoxia can be propelled by a potent positive feedback process whereby it promotes the recycling of phosphorus from ocean sediments (by biological and chemical means), which when upwelled fuels more productivity, which as it sinks creates more respiratory demand for oxygen consumption.42 The interaction of this relatively fast positive feedback with the slower negative feedbacks that stabilise atmospheric oxygen43 may have caused the Earth system to oscillate between anoxic and oxygenated oceans for ten million years after the End Permian, and again later on in the middle of the Cretaceous period.44

The End Cretaceous extinction, 65 million years ago, finally cleared dinosaurs out of a niche space they had dominated for over 150 million years, allowing mammals to diversify into ecological prominence. In an unusual warm event 55 million years ago – the Paleocene-Eocene Thermal Maximum – three major mammal orders diversified,45 including our own – the Primates – and the odd-toed and even-toed ungulates, many species of which we would ultimately domesticate.

Over the past 40 million years, atmospheric carbon dioxide levels have fallen and the climate has got cooler and drier. The cause of the long-term cooling is still being argued over – with a partial role for the uplift of mountains driving faster weathering rates and associated carbon dioxide uptake,46 and also a role for declining long-term inputs of carbon dioxide from volcanic degassing.47 The correspondence between CO2 and climate is not a perfect one on this long timescale as the latest proxies show stable, low CO2 for the last ~25 million years, whilst the climate fluctuated and underwent overall cooling.48 One hypothesis argues that the stable low CO2 was thanks to very strong negative feedback from the role of plants in weathering, because the majority of (C3) plants were near CO2 starvation.48

In this stable low CO2 world of the Miocene Epoch (23-5.3 Ma), grasslands expanded abruptly in two phases around 17 and 6 million years ago, such that they now cover about a third of the Earth’s productive land surface. These phases of grassland expansion may have been propelled by co-evolution with mammal grazers and also by a potent positive feedback involving fires: Grasslands encourage fires which encourage grasslands, because frequent fires prevent forests regenerating. In their second phase of expansion, grasslands colonized large parts of Africa including the Great Rift Valley – the place where our evolutionary lineage diverged from chimpanzees, around 6 million years ago (Figure 4). Around 4 million years ago in the Pliocene Epoch (5.3-2.6 Ma) our hominine ancestors began to walk upright – conceivably as an adaptation to moving through the newly created savannah between clumps of woodland.

Just as our ancestors began to develop stone tool use, the Earth’s overall cooling trend culminated in a series of Northern Hemisphere ice age cycles of increasing severity and decreasing frequency (Figure 5). This change in climate dynamics marks the onset of the Quaternary Period (2.6-0 Ma). It is important not least because of what it can tell us about the stability of the ‘background’ climate system at present. One can read the transition of the Earth system from a single stable climate state in the preceding Pliocene Epoch to the progressively deeper and stronger oscillations of the Quaternary glaciations, as a system that is not only getting cooler but is also destabilising.1 The pattern of ‘saw tooth’ global oscillations (Figure 5) whereby the climate cools progressively into an ice age then snaps rapidly out of it, only for the cycle to repeat soon after, is a classic example of a system that, although it is bounded by negative feedback, contains a strong amplifier (positive feedback) – as should be familiar to students of electrical engineering.

Currently our understanding of the key feedbacks involved in the glacial-interglacial cycles is incomplete. Textbooks still emphasise the role of the Milankovitch cycles of changing solar insolation as the cause of the ice ages – but in no way can they explain the amplitude or the now dominant 100,000 year period of the glaciations – because the orbital signal is particularly weak at that frequency. Instead, variations in the Earth’s orbit are at best an irregular ‘pacemaker’ of the ice ages, and there must be an inherent, roughly 100,000 year ‘relaxation oscillator’ in the Earth system’s own dynamics.49 Current theories for that oscillator range invoke either the natural timescale of growth and decay of ice sheets or a long inherent timescale in the carbon cycle. What we can say with more confidence is that at the ‘termination’ of an ice age the Earth system goes into near runaway positive feedback, with release of carbon from the deep ocean playing a key role in amplifying global climate change.

The Quaternary climate changes (Figure 5) in turn provoked widespread speciation in mammals, including our hominine lineage. The resulting environmental instability50 may have played a role in our evolution as unusually intelligent, highly social primates. One idea is that when the environment is changing – but not too frequently or unpredictably – it pays to be smart and to cooperate in social groups to help adapt to the changing conditions.

Biodiversity as victim: How we are causing global change now

There are several key points to take away from the preceding brief history of the biosphere. Firstly, microbes (prokaryotes and unicellular eukaryotes) still play a central role in running all the major biogeochemical cycles that sustain life. Secondly, multicellular eukaryotes – plants, fungi, and animals – have also played an important role in creating and maintaining an oxygen-rich, CO2-poor, cool and stable world in which we could ultimately evolve. Thirdly, this is a bad time to be perturbing the Earth system, because it is unusually unstable.

Armed with this information, we now turn from biodiversity as creator to biodiversity as victim – taking an Earth system view of the global change that humans are causing. The underlying premise is that anthropogenic global change is a systemic phenomenon – climate change and other drivers of biodiversity loss, including land-use change, invasive species, over exploitation, and nutrient loading are all intertwined manifestations of expanding human activities.

A systemic view of anthropogenic global change

The term ‘Anthropocene’ has been coined to describe a new geological epoch in which human activities are transforming the Earth system at a global scale. Whilst there is much academic debate about whether this really is a new epoch, and if so, when it began, the salient point is that humans are now a planetary force. Here we briefly trace how humans have gone from being just another great ape to planet changers (Figure 4), and how today’s global changes have their roots in much earlier and more localised innovations.

The intentional use of fire set our ancestors apart from all other species, because it was the first innovation that extended energy use beyond the human body. Controlled use of fire may have started 1.5 million years ago,51 and was certainly occurring by 800,000 years ago in Africa52 and by 400,000 years ago in Europe.53 The use of fire for cooking gave Homo erectus more energy in their diet from the cooking of meat and a more diverse diet (by detoxifying foods).54 The shift to hunting energy-rich meat in turn triggered the formation of social groups that settled around campsites and divided labour, causing an escalation in human social evolution.55

Between about 400,000 and 250,000 years ago, stone tool technologies became more elaborate and brain size increased rapidly. Anatomically modern humans (Homo sapiens) first appeared in East Africa around 200,000 years ago. Sometime after that, our ancestors experienced a bottleneck in population of 10,000 or fewer breeding pairs. Descendants of this founding group emerged out of Africa and began to spread around the world roughly 65,000 years ago. Their migration was facilitated by one of a series of periodic wet phases of the Sahara after a mega-drought in Africa from 135,000 to 90,000 years ago.

As modern humans arrived in new continents, they triggered the extinction of other large mammals or ‘megafauna’.56,57 This began in Australia before 44,000 years ago, in Europe over 30,000 years ago, in North America 11,500 years ago, and in South America 10,000 years ago. Extinction was less severe in Africa, perhaps because existing species were already habituated to and wary of hunting humans. Fire was the first ‘tool’ that enabled early humans to start changing their environment on a large scale. Human use of fire in hunting shifted ecosystems toward grasslands. This helps explain why herbivores that browse on trees (rather than eating grasses) suffered most in the megafauna extinctions. Our ancestors may also have hunted some large herbivores to extinction, thus leaving carnivores and scavengers to suffer from a lack of food. Human use of fire in Australia helped maintain desert scrubland over large areas of the continent. This in turn may have inhibited the return of the monsoon into the continental interior when the Earth system entered the present Holocene interglacial epoch.58,59 If so, it may represent the first large-scale impact of humans on the climate system. Other human fire driven ecosystem tipping events, include rapid landscape transformations in the mesic environments of New Zealand,60 the wet tropical forests of the pre-Columbian Amazon,61 and across the savannas and woodlands of Africa.62

As the Earth system exited the last ice age there was a major fluctuation in the climate of the Northern Hemisphere. An abrupt warming around 14,700 years ago was followed by a marked cooling 12,700 years ago and a further abrupt warming 11,500 years ago. During the cool period known as the ‘Younger Dryas’, people in the Eastern Mediterranean region who had been collecting abundant wild cereals for food began domesticating the first cereal crops, perhaps in response to the regional drying effects of climate change. As the Earth system settled into the stable Holocene interglacial state, around 10,500 years ago, the Sahara re-entered one of its wet and green phases, turning the region encompassing the Nile, Euphrates and Tigris rivers into the fabled Fertile Crescent. Farming began there with the domestication of wheat, barley, peas, sheep, goats, cows and pigs. Farming also arose independently elsewhere in the world,63 around 8,500 years ago in South China, 7,800 years ago in North China, 4,800 years ago in Mexico, and 4,500 years ago in Peru and Eastern North America.

The relatively abrupt and independent occurrence of farming all over the world suggests it may have been held back by environmental conditions before the Holocene. Low ice age levels of carbon dioxide and the volatile glacial climate would certainly not have helped establish agriculture. Once established, farming increased the energy input to human societies. This ‘Neolithic revolution’ caused an increase in human fertility (soon followed by an increase in mortality), which increased the population from 6 million to over 30 million between 6,000 and 4,000 years ago and perhaps as high as 100 million by 2,000 years ago. However, one of the downsides of farming was that sedentary, high-density agricultural civilizations were more sensitive to climate change than mobile foraging societies – with abrupt shifts in tropical climate during the Holocene linked to the collapse of several ancient societies.

The increased population and energy flows due to farming increased the material inputs to, and waste products from, societies. The resulting environmental effects began early in the Holocene, but their scale is much debated.64,65 Irrigation began around 8000 BP in Egypt and Mesopotamia, leading to some localized salination and siltation of the land, reducing crop yields and encouraging a shift in agricultural crop from wheat to more salt-tolerant barley.66 The use of manure as fertilizer may have begun as early as 9000 BP in SW Asia and 7000 BP in Europe.64,67 The clearance of forests to create agricultural land and supply biomass energy and wood from 8000 BP onwards, reduced the carbon storage capacity of the land, transferring CO2 to the atmosphere.68 Cumulative carbon emissions may have approached 300 PgC by 500 BP contributing ~20 ppm to atmospheric CO2 levels. The biogeophysical effects of forest clearance also affected the climate, regionally and remotely.69 Anthropogenic sources of methane started around 5000 BP with the irrigation of rice paddies and have contributed to changes in atmospheric CH4 concentration over the past ~3000 years.70

The ‘Early Anthropocene’ hypothesis71,72 argues that the new epoch began thousands of years ago, with changes in atmospheric CO2 and CH4 caused by the Neolithic revolution. However, others argue that natural changes in the climate and carbon cycle can explain most of the changes in atmospheric CO2 and CH4 during the Holocene. For example, variations in the Earth’s orbit meant that 6,000 years ago the Northern Hemisphere was warmer than today and therefore supported more vegetation, both in the boreal regions and across much of North Africa – creating a ‘green Sahara’. This helps explain somewhat lower atmospheric CO2 levels in the early Holocene. As the orbital forcing steadily declined there was a relatively abrupt drying and expansion of the Sahara desert, around 5,000 years ago. Models predict this was due to a shift between alternative steady states of the vegetation-climate system in North Africa. This ‘browning of the Sahara’ together with a retreat of boreal forests from the highest Northern latitudes added CO2 to the atmosphere.

Over the past two millennia, our records of past climate change improve, with multiple proxies for climate variability including tree ring and ice core records and temperatures from boreholes. These records reveal slow fluctuations between somewhat warmer and cooler intervals on Northern Hemisphere land surfaces, including the Medieval Warm Period (circa 950-1250AD) and the Little Ice Age (circa 1550-1850AD). Intervals of cooler climate correlate with poor agricultural production, war, and population decline, but any causal links are argued over. Ice core records reveal some variations in atmospheric composition including a 10 ppm decline in CO2 500 years ago, which was also a time when human biomass burning declined, and has been argued to be due to plague-induced human population decline which allowed large areas to reforest and take up carbon. However, the ‘Early Anthropocene’ hypothesis71,72 continues to be controversial, partly because pre-industrial societies were limited in the energy supplies with which they could transform their environment.

Many researchers link the start of the Anthropocene with the Industrial Revolution, because the accessing of fossil fuel energy greatly increased the impact of humanity on the Earth system. The Industrial Revolution marks the transition from societies fuelled largely by recent solar energy (via biomass, water and wind) to ones fuelled by concentrated ‘ancient sunlight’. Although coal had been used in small amounts for millennia, for example for iron making in ancient China, fossil fuel use only took off with the invention and refinement of the steam engine. Thomas Newcomen’s demonstration of a working steam engine in 1712, followed by James Watt’s improvements to it in 1769, gave a great boost to coal extraction, by draining mines of water. The steam engine was also used to convert fossil fuel energy into mechanical power in manufacturing and transport. This created a potent positive feedback loop that propelled the Industrial Revolution.

The exploitation of concentrated fossil fuel energy triggered a massive expansion of population, food production, material consumption, and associated waste products. Human population doubled between 1825 and 1927 from 1 to 2 billion, doubled again by 1975 to 4 billion, and is on course to double again by 2030 to 8 billion. With the industrial revolution, food and biomass have ceased to be the main source of energy for human societies. Within 150 years, from 1850 to 2000, global human energy use increased tenfold73 from 56 to 600 EJ yr-1, such that less than a tenth of total energy input to human societies is now contained in annual food production. By year 2000 the annual global energy flux through human societies was one third of the global terrestrial NPP74 and greater than the total global energy flux through all non-human heterotrophic biomass.73

The step increase in energy capture with industrialization is associated with fundamental changes in global material cycles. In industrial economies ~80% by weight of the total annual outflow of materials is CO2, making the atmosphere the largest waste reservoir of the industrial metabolism.75 Between 1850 and 2000 global CO2 emissions from combustion of fossil fuels and materials processing increased 125-fold from 54 to 6750 TgC yr-1 and reached 9140 TgC yr-1 in 2010.76 For some elemental cycles our collective activities now exceed the activities of the rest of the biosphere combined. Between 1860 and 2005 anthropogenic creation of reactive nitrogen grew more than tenfold77, from ~15 to 187 TgN yr-1. The excess reactive nitrogen was transferred to other environmental pools, partly denitrifying to atmospheric N2 and N2O, but also contributing to eutrophication and acidification of terrestrial and coastal marine ecosystems, to global warming and to tropospheric ozone pollution. Analogous human-induced acceleration affected the P-cycle. Much of this escalation of human impact on the Earth system has occurred since the end of the Second World War – in a transition dubbed the ‘Great Acceleration’.

Basics of climate change

The theory of the greenhouse effect is Victorian physics. As early as 1896, Svante Arrhenius calculated that a doubling of atmospheric CO2 from its preindustrial concentration would warm the world by around 5°C. This laborious calculation by hand, which took him two years, remains within the range of estimates of ‘climate sensitivity’ from the latest Earth system models. The current best estimate is around 3°C.

By the end of the 19th century, ship-borne temperature measurements were being regularly made. These, together with thermometer readings from land-based weather stations, have enabled climate scientists to piece together what is called the ‘instrumental temperature record’. It shows global warming of around 0.85°C from 1880 to 2012, around 0.5°C of which has occurred since 1980. The rise in temperature is global in extent (whereas the Medieval Warm Period and Little Ice Age were only regional phenomena). The global temperature has not risen at a steady rate – there are some periods of stable temperature (e.g. the 1940s and 1950s) and some spells of more rapid rise (e.g. the 1980s and 1990s). This is to be expected because even in the absence of human activities there is natural variability in the climate, producing warmer and colder spells – which when superimposed on a rising trend gives periods of no warming and periods of rapid warming.

One factor that can cool the climate is the injection of tiny reflective sulphate aerosol particles into the atmosphere, which scatter sunlight (sending some of it back to space). Sulphate aerosols can come from volcanic eruptions (such as Mt Pinatubo in 1991), or from fossil fuel burning, especially the combustion of sulphurous (brown) coals. However, when it enters into solution, sulphate forms sulphuric acid and hence acid rain. In order to curb acid rain, successful efforts have been made to scrub sulphur dioxide out of power station flue gases. This in turn reduced its cooling effect on the climate, unmasking the increasing greenhouse effect, and contributing to global warming.

Projections of global temperature change depend fairly linearly on the cumulative emissions of CO2 up to a given time, i.e. how much fossil fuel we burn without capturing and storing the CO2 given off (Figure 6). Roughly speaking, each 500 billion tonnes of carbon emitted will give 1°C of global warming. Thus, we have burned around 400 billion tonnes of fossil fuel carbon already and have experienced 0.8°C of warming. If we want to stay under 2°C of warming we need to limit our emissions to a trillion (1000 billion) tonnes of carbon. Whereas if we burn all 5000 billion tonnes of known fossil fuels we can eventually expect around 10°C of warming. Whether this thought experiment can ever be realized is highly questionable because <10°C warming could be so damaging as to prevent us from burning all the fossil fuel.

On the short timescale of the next few decades, temperature projections are not very dependent on emissions pathway, because the climate system is still responding to the energy imbalance caused by the accumulation of past greenhouse gas emissions. Also, natural variability in heat uptake and storage by the ocean can affect surface temperatures considerably on decadal timescales.

On the long timescale of a millennium, temperature change still depends on the total cumulative emissions of carbon.78 However, by then the Earth system will have apportioned the CO2 we have added between the atmosphere, ocean and land surface. The fraction remaining in the atmosphere – known as the ‘airborne fraction’ – will depend on the total amount of carbon we emit. At a minimum it will be about 20%. But simple and intermediate complexity models tell us that the fraction increases exponentially with the amount of carbon added. As temperature depends on the natural logarithm of atmospheric carbon these two effects combine to give a linear relationship between carbon emitted and global warming.

The relationship between carbon in the atmosphere and global temperature change is captured in a concept called ‘climate sensitivity’. This is defined as the global warming caused by doubling the CO2 content of the atmosphere, once the heat content of the ocean has adjusted and various ‘fast’ feedbacks have operated. Our best estimate is that it is near 3°C but it could be in the range 1.5-5°C. It is uncertain because our models differ over the strength of different feedbacks, and over the long-term heat uptake by the deep ocean, and observations cannot completely constrain these properties. To the ‘climate sensitivity’ we can add the sensitivity of atmospheric CO2 concentration to a given CO2 emission, which depends on feedbacks between the climate and the carbon cycle. On longer timescales there are further ‘slow’ feedbacks, for example involving the melt of ice sheets, which add to warming. The resulting ‘Earth system sensitivity’ to CO2 could be as much as twice the climate sensitivity.

Whilst much of the behaviour of the Earth system can be described as ‘linear’ and predictable with our current models, there is a class of ‘non-linear’ change that is much harder to predict and potentially much more dangerous. It involves ‘tipping points’ – where a small perturbation triggers a large response from a part of the Earth system – leading to abrupt and often irreversible changes.79 Tipping points can occur when there is strong positive feedback within a system, which creates alternative stable states for a range of boundary conditions. When changes in the boundary conditions cause the current state of a system to lose its stability, a tipping point occurs, triggering a transition into the alternative stable state. Thankfully it is very difficult to pass a tipping point at the planetary scale – rare examples from Earth history are the switches into (and out of) ‘snowball Earth’ events. However, several sub-systems of the Earth system are thought to exhibit alternative stable states and tipping points – and I have dubbed those parts of the Earth system that can exhibit tipping points ‘tipping elements’.79 Amongst them are several candidates that could be tipped by human-induced global change (Figure 7). They can be divided into those involving abrupt shifts in modes of circulation of the ocean or atmosphere (or the two of them coupled together), those involving abrupt loss of parts of the cryosphere, and those involving abrupt shifts in the biosphere – which are of particular interest here.

Consequences for biodiversity

Climate change is already affecting species distributions and phenology80 and despite ongoing arguments over methodology, is generally agreed to pose a major future threat to biodiversity of complex life forms, especially terrestrial ones.81 It is also widely acknowledged that climate change interacts with other global changes, notably land-use change,82 in a systemic way to threaten biodiversity.

Early projections of the effects of climate change focused on the loss of terrestrial plant and animal (mostly mammal) species, based on the projected movement – and often contraction – of their ‘bioclimatic envelope’ (i.e. the area defined by the correlation between climate variables and observed distribution) under a given climate change scenario.83 This was combined with the well-known species-area relationship, to project ‘committed extinction’.83 Limits of maximum dispersal rates and no dispersal were considered, giving truly alarming lower-upper estimates of 15-37% committed extinction under a mid-range 2050 climate warming scenario.83 If even remotely accurate this would put climate change on a par with habitat loss from land use change,84 as the greatest threat to terrestrial biodiversity. No timescale was assigned to how long it would take the ‘committed’ extinctions to unfold, thus opening the possibility that if global temperature could be brought down in the future – either by natural or artificial (geoengineering) means – this could save species,83 essentially by bringing their bioclimatic envelope back to where they are.

This prominent early study83 provoked a range of responses, and there have followed many studies attempting to project extinction risk from climate change – reviewed elsewhere.81 The great majority of studies still use bioclimatic envelope models to project changes in range. A few studies have used more aggregated dynamic global vegetation models to project shifts in biomes and then downscaled from that to assess extinction risk. To assess the actual loss of species, as well as species-area relationships, IUCN status methods and dose-response relationship models have also been used.

Early on it was noted that the lack of consideration of distinct ecotypes within a species, each having a more restricted bioclimatic envelope than the species as a whole, could lead to underestimation of species loss.85 Conversely the lack of consideration of acclimation (phenotypic plasticity) and, more pertinently, genetic adaptation could lead to an over-estimation of committed extinction.

Those versed in recent Earth history have noted that the abrupt climate changes of the Quaternary period provoked nothing like the level of species extinction being projected under smaller levels of future climate change. Instead the mega-fauna extinctions that did occur are mostly (if not universally) attributed directly to human activities, occurring as they do synchronously with the spread of Homo sapiens across the planet.57 This begs the question of how species kept up with the movement of their bioclimatic envelope in the past. One line of reasoning is that past climate changes were less rapid giving more time for migration/dispersal or genetic adaptation of populations. However, whilst this is true for the overall pace of deglaciation, some past abrupt warming events (e.g. at the start of the Bolling-Allerod or the end of the Younger Dryas), whilst not fully global in scale, were more rapid than anything considered in the future projections. Despite this, several relatively sedentary (e.g. tree) species in the most affected regions were able to move at an impressive rate. This suggests a critical role for rare long-range dispersal events86 (and/or for nascent islands of cryptic diversity outside recognised climate envelopes).

Whilst the focus of existing studies on the fate of terrestrial plant and animal species is undoubtedly important, the types of biodiversity being considered in the assessments is extremely restrictive and hard to connect to fundamental questions about the future functioning of ecosystems and the biosphere. There is relatively little consideration of marine life or indeed of whole kingdoms of life – the archaea and bacteria (prokaryotes), protists (unicellular eukaryotes), and fungi. The focus on diversity at the species level is problematic for prokaryotes (where the notion of species is somewhat meaningless thanks to rampant horizontal gene transfer) and may be missing the possibility that the greatest threats to the biosphere from climate change could come at much larger levels of biological organisation. For the ecosystem or Earth system thinker there is surprisingly little consideration of effects of global change on functional diversity. In particular, would such relatively high levels of projected species loss compromise the functioning of ecosystems and the biosphere, and to what degree?

Taking a more top-down approach to the biodiversity threats posed by climate change returns us to the possibility of wholesale loss of even large-scale biomes due to the passing of tipping points.

Starting on the land surface, some biomes (particularly in the tropics) are strongly coupled to the atmosphere through positive feedbacks. For example, the ‘green Sahara’ state that was present 6000 years ago supported an atmospheric circulation that brought moisture into what is now a desert. Today the Amazon rainforest recycles water to the atmosphere thus helping maintain the rainfall that supports the forest. It also suppresses fires. However, if the climate dries regionally – as has been seen in recent Amazon drought years (2005, 2010) – this can lead to dieback of trees and a shift to a more devastating fire regime. If grasses begin to encroach into the forest these encourage fires which destroy tree saplings and support an alternative grassland or savannah state (a positive feedback). Grasslands are already thought to be an alternative stable vegetation state for large parts of the Amazon basin under present rainfall.87,88 In the future, if the region dries out, widespread dieback of the Amazon rainforest has been projected in some models.89,90

Biomes may also pass climate thresholds to their viability, due to the dominant organisms reaching physiological limits or due to abrupt increases in disturbance factors linked to climate change. Currently several regions of temperate and boreal forests are experiencing widespread tree dieback due to drought and heat-stress91 and to attack by bark beetles that are thriving in a warmer climate.92,93 In some future projections, large areas at the southern boundary of the boreal forests are abruptly lost and replaced with steppe grasslands,94,95 as the summers become too hot for the dominant trees and disturbance by fires and bark beetle attacks increases.

Abrupt dieback of large areas of forests would clearly negatively impact the biodiversity they contain and could happen faster than the movement of species’ bioclimatic envelopes. Loss of large areas of tropical or boreal forests would also feedback CO2 to the atmosphere, making a significant contribution to cumulative CO2 emissions and temperature rise.

Turning to ocean biomes, coral reefs are under threat of large-scale loss due to a combination of drivers, including ocean warming and ocean acidification.96,97 Coral bleaching events, linked to ocean warming have become much more widespread and detrimental in recent decades, and marine biologists are already talking about tropical coral reefs being at a ‘point of no return’.97 Cold-water corals that grow down to 3000m depth will be the first to be affected by ocean acidification as the saturation horizon of aragonite (the crystalline form of calcium carbonate out of which they are constructed) shallows. Once bathed in corrosive waters, under saturated in aragonite, the skeletons and shells may dissolve and the reefs collapse. With unabated CO2 emissions an estimated 70% of the presently known deep-sea coral reef locations will be in corrosive waters by the end of this century98.

An even larger scale tipping point threat to ocean ecosystems comes from the increased input of phosphorus and nitrogen to the land, which is already fuelling anoxic conditions in freshwaters and some shelf seas. If this continues for centuries and humans refine all the known phosphate rock reserves and turn them into fertiliser then this will significantly increase the phosphorus content of the ocean.99 This risks triggering a global oceanic anoxic event, because anoxia enhances phosphorus recycling from shelf-sea sediments fuelling more productivity and anoxia – a potent positive feedback.44,99 Ocean anoxia is further encouraged by warming of the ocean, which reduces the solubility of oxygen in the water and tends to stratify the ocean, isolating deeper deoxygenating waters from the atmosphere. As the fossil record attests, widespread ocean anoxia is a potent cause of mass extinction.

Taking a more bottom-up approach to extinction threat, the fundamental question of ‘can life adapt?’ to global change has become the focus of a rapidly expanding recent body of work.100,101 Here adapt is taken in the broad sense to cover both physiological acclimation and genetic adaptation. The question is particularly pertinent for the marine ecosystem as it is based fundamentally on microbial primary producers with considerable phenotypic plasticity, and (thanks to short generation times) considerable potential for rapid evolutionary adaptation. Thus, climate change will have consequences for the phytoplankton, but changes in function may be more pertinent than levels of extinction.

As an example, we have developed an evolutionary ecosystem model (‘EVE’) that captures the physiological acclimation of individual phytoplankton cells and the genetic adaptation of populations.102-104 Work so far has focused on acclimation responses to ocean warming, based on molecular genetic insights into the temperature sensitivity of core biochemistry.105 In particular having shown that ribosomal RNA functions more efficiently at higher temperatures, we predict that the ribosomal content of cells will decline in a warmer ocean and with that their N:P ratio will go up. The resulting predictions of organism N:P composition across temperature and resource gradients in the contemporary ocean compares well to data.104 Such a fundamental future change in the composition of organisms at the base of the food chain will have a range of biogeochemical and ecological consequences. For example, we expect it to trigger greater nitrogen depletion of ocean waters and a resulting selective pressure for increased nitrogen fixation.

To wrap-up, as it stands we are not in a great scientific position to answer whether the threat to biodiversity is a threat to the functioning of the biosphere (i.e. our life-support system), because that is a long way from the questions most scientists in this field have been trying to answer. I argue it could be such an existential threat if we seriously erode the role of life in global biogeochemical cycling – for example if we seriously impair the effect of plants and soil communities on rock weathering then we could begin the transition to a much hotter world, unable to support complex eukaryote life forms.106 Such threats, like oceanic anoxic events, may seem distant by human timescales, but from the perspective of geologic time they could be imminent.

Conclusion: A new revolution?

Whilst human transformation of the planet was initially unwitting, now we are increasingly collectively aware of it. We humans have conscious foresight and a sense of purpose that (as far as we know) has never been part of the Earth system before. This changes the Earth system fundamentally, because it means that one species can (in principle at least) consciously, collectively shape the future trajectory of our planet. We know our current way of living is both unsustainable and is eroding biodiversity (in many senses of the word), and this might threaten our very life-support system. This should be more than enough existential threat to be triggering a rear guard action, yet little real change is happening. As the Pope’s encyclical letter on care for our common home makes clear, this must be due to a failure in our value system.107

Specifically we fail to recognise and embody in our social and economic systems the fundamental existential value of the biosphere. However, it doesn’t take much, even within the current narrow paradigm of economic valuation to get a very different result. For example, if we simply recognise that the biosphere provides essential services to us that cannot be substituted by something in the market place,108 and furthermore that a part of those services is under threat of irreversible loss,79 then even if we barely value those services in market terms, they should still be exerting a huge influence on our ‘social planning’ – including triggering policy decisions to work much harder and faster to mitigate climate change.109 This is because the irreplaceable and un-substitutable biosphere does not grow like the economy – hence even in mainstream economics it effectively becomes scarcer and therefore more precious the longer we continue in a growth paradigm.110

Whether the current growth paradigm thus contains the intellectual seeds of its own demise is a moot point given the current political economies of knowledge. Nevertheless, it seems worth reflecting if we could change our value system to properly embrace our reliance on the biosphere, and see ourselves as an integral part of it, could we chart a very different way forward? In particular, rather than heading (now with our eyes open) toward an impending apocalypse, or frantically trying to retreat into lower energy and material consumption, could we help create a positive revolutionary change of the Earth system? And could such a ‘good Anthropocene’ be truly good for all and not just some rich elite? I won’t pretend to have all the answers, but a future world powered by sustainable energy with near-closed recycling of materials could conceivably support billions of people and allow the rest of the biosphere to flourish – as outlined elsewhere.1 Yes it will require technology as well as some reorganisation of the economic system (at the very least), and no, we don’t have a tremendous track record in the wise use of either. Still I submit that when staring at an unfolding ecocide, with the stakes so high, rather than just bash the rest of society over the head with what we are doing wrong, our duty as scientists is to help articulate a vision of a better future that we can work towards together, and provide a map for how to start navigating there from here.


1               Lenton, T.M. & Watson, A.J., Revolutions that made the Earth. (Oxford University Press, Oxford, 2011).

2               Lenton, T.M., Schellnhuber, H.J., & Szathmáry, E., Climbing the co-evolution ladder. Nature 431, 913 (2004).

3               Rosing, M.T., 13C-Depleted Carbon Microparticles in >3700-Ma Sea-Floor Sedimentary Rocks from West Greenland. Science 283, 674-676 (1999).

4               Ohtomo, Y., Kakegawa, T., Ishida, A., Nagase, T., & Rosing, M.T., Evidence for biogenic graphite in early Archaean Isua metasedimentary rocks. Nature Geosci 7 (1), 25-28 (2014).

5               Nutman, A.P., Bennett, V.C., Friend, C.R.L., Van Kranendonk, M.J., & Chivas, A.R., Rapid emergence of life shown by discovery of 3,700-million-year-old microbial structures. Nature 537 (7621), 535-538 (2016).

6               Schopf, J.W., Microfossils of the Early Archean Apex Chert: New Evidence of the Antiquity of Life. Science 260, 640-646 (1993).

7               Schopf, J.W., Fossil evidence of Archaean life. Philosophical Transactions of the Royal Society B 361, 869-885 (2006).

8               Brasier, M.D., Green, O.R., Jephcoat, A.P., Kleppe, A.K., Van Kranendonk, M.J., Lindsay, J.F., Steele, A., & Grassineau, N.V., Questioning the evidence for Earth's oldest fossils. Nature 416, 76-81 (2002).

9               Brasier, M.D., McLoughlin, N., Green, O., & Wacey, D., A fresh look at the fossil evidence for early Archaean cellular life. Philosophical Transactions of the Royal Society of London Series B - Biological Sciences 361 (887-902) (2006).

10            Noffke, N., Christian, D., Wacey, D., & Hazen, R.M., Microbially Induced Sedimentary Structures Recording an Ancient Ecosystem in the ca. 3.48 Billion-Year-Old Dresser Formation, Pilbara, Western Australia. Astrobiology 13 (12), 1103-1124 (2013).

11            Knoll, A.H. & Barghoorn, E.S., Archean Microfossils Showing Cell Division from the Swaziland System of South Africa. Science 198 (4315), 396-398 (1977).

12            Kharecha, P., Kasting, J., & Siefert, J., A coupled atmosphere–ecosystem model of the early Archean Earth. Geobiology 3 (2), 53-76 (2005).

13            Canfield, D.E., Rosing, M.T., & Bjerrum, C., Early anaerobic metabolisms. Philosophical Transactions of the Royal Society B: Biological Sciences 361 (1474), 1819-1836 (2006).

14            Williams, H.T.P. & Lenton, T.M., The Flask model: Emergence of nutrient-recycling microbial ecosystems and their disruption by environment-altering 'rebel' organisms. Oikos 116 (7), 1087-1105 (2007).

15            Boyle, R.A., Williams, H.T.P., & Lenton, T.M., Natural selection for costly nutrient recycling in simulated microbial metacommunities. Journal of Theoretical Biology 312, 1-12 (2012).

16            Allen, J.F. & Martin, W., Evolutionary biology: Out of thin air. Nature 445 (7128), 610-612 (2007).

17            Lenton, T.M. & Daines, S.J., Biogeochemical Transformations in the History of the Ocean. Annual Review of Marine Science 9 (1), 31-58 (2017).

18            Farquhar, J., Bao, H., & Thiemens, M., Atmospheric Influence of Earth's Earliest Sulfur Cycle. Science 289 (5480), 756-758 (2000).

19            Goldblatt, C., Lenton, T.M., & Watson, A.J., Bistability of atmospheric oxygen and the great oxidation. Nature 443, 683-686 (2006).

20            Daines, S.J., Mills, B., & Lenton, T.M., Atmospheric oxygen regulation at low Proterozoic levels by incomplete oxidative weathering of sedimentary organic carbon. Nature Communications 8, 14379 (2017).

21            Catling, D.C., McKay, C.P., & Zahnle, K.J., Biogenic Methane, Hydrogen Escape, and the Irreversible Oxidation of Early Earth. Science 293, 839-843 (2001).

22            Claire, M.W., Catling, D.C., & Zahnle, K.J., Biogeochemical modelling of the rise in atmospheric oxygen. Geobiology 4 (4), 239-269 (2006).

23            Daines, S.J. & Lenton, T.M., The effect of widespread early aerobic marine ecosystems on methane cycling and the Great Oxidation. Earth and Planetary Science Letters 434, 42-51 (2016).

24            Bekker, A. & Holland, H.D., Oxygen overshoot and recovery during the early Paleoproterozoic. Earth and Planetary Science Letters 317-318, 295-304 (2012).

25            Rasmussen, B., Fletcher, I.R., Brocks, J.J., & Kilburn, M.R., Reassessing the first appearance of eukaryotes and cyanobacteria. Nature 455, 1101-1104 (2008).

26            Han, T.M. & Runnegar, B., Megascopic eukaryotic algae from the 2.1-billion-year-old negaunee iron-formation, Michigan. Science 257 (1992).

27            Butterfield, N.J., Probable Proterozoic fungi. Paleobiology 31 (1), 165-182 (2005).

28            Butterfield, N.J., Bangiomorpha pubescens n. gen., n. sp.: implications for the evolution of sex, multicellularity, and the Mesoproterozoic/Neoproterozoic radiation of eukaryotes. Paleobiology 26 (3), 386-404 (2000).

29            Lenton, T.M., Boyle, R.A., Poulton, S.W., Shields, G.A., & Butterfield, N.J., Co-evolution of eukaryotes and ocean oxygenation in the Neoproterozoic era. Nature Geoscience 7 (4), 257-265 (2014).

30            Lenton, T.M. & Daines, S.J., Matworld – the biogeochemical effects of early life on land. New Phytologist, doi:10.1111/nph.14338 (2016).

31            Donnadieu, Y., Godderis, Y., Ramstein, G., Nedelec, A., & Meert, J., A `snowball Earth' climate triggered by continental break-up through changes in runoff. Nature 428 (6980), 303-306 (2004).

32            Lenton, T.M. & Watson, A.J., Biotic enhancement of weathering, atmospheric oxygen and carbon dioxide in the Neoproterozoic. Geophysical Research Letters 31 (5), L05202 (2004).

33            Hoffman, P.F., Kaufman, A.J., Halverson, G.P., & Schrag, D.P., A Neoproterozoic snowball earth. Science 281, 1342-1346 (1998).

34            Love, G.D. et al., Fossil steroids record the appearance of Demospongiae during the Cryogenian period. Nature 457 (7230), 718-721 (2009).

35            Boyle, R.A., Dahl, T.W., Dale, A.W., Shields-Zhou, G.A., Zhu, M., Brasier, M.D., Canfield, D.E., & Lenton, T.M., Stabilization of the coupled oxygen and phosphorus cycles by the evolution of bioturbation. Nature Geoscience 7 (9), 671-676 (2014).

36            Lenton, T.M., Crouch, M., Johnson, M., Pires, N., & Dolan, L., First plants cooled the Ordovician. Nature Geoscience 5 (2), 86-89 (2012).

37            Berner, R.A., The Rise of Plants and Their Effect on Weathering and Atmospheric CO2. Science 276, 544-546 (1997).

38            Lenton, T.M., Dahl, T.W., Daines, S.J., Mills, B.J.W., Ozaki, K., Saltzman, M.R., & Porada, P., Earliest land plants created modern levels of atmospheric oxygen. Proceedings of the National Academy of Sciences 113 (35), 9704-9709 (2016).

39            Lenton, T.M. & Watson, A.J., Redfield revisited: 2. What regulates the oxygen content of the atmosphere? Global Biogeochemical Cycles 14 (1), 249-268 (2000b).

40            Schwartzman, D.W. & Volk, T., Biotic enhancement of weathering and the habitability of Earth. Nature 340, 457-460 (1989).

41            Harnik, P.G. et al., Extinctions in ancient and modern seas. Trends in Ecology & Evolution 27 (11), 608-617.

42            Van Cappellen, P. & Ingall, E.D., Benthic phosphorus regeneration, net primary production, and ocean anoxia: A model of the coupled marine biogeochemical cycles of carbon and phosphorus. Paleoceanography 9, 677-692 (1994).

43            Van Cappellen, P. & Ingall, E.D., Redox stabilisation of the Atmosphere and Oceans by Phosphorus-Limited Marine Productivity. Science 271, 493-496 (1996).

44            Handoh, I.C. & Lenton, T.M., Periodic mid-Cretaceous Oceanic Anoxic Events linked by oscillations of the phosphorus and oxygen biogeochemical cycles. Global Biogeochemical Cycles 17 (4), 1092 1010.1029/2003GB002039 (2003).

45            Gingerich, P.D., Environment and evolution through the Paleocene-Eocene thermal maximum. Trends in Ecology & Evolution 21 (5), 246-253 (2006).

46            Raymo, M.E. & Ruddiman, W.F., Tectonic forcing of late Cenozoic climate. Nature 359, 117-122 (1992).

47            Mills, B., Daines, S.J., & Lenton, T.M., Changing tectonic controls on the long-term carbon cycle from Mesozoic to present. Geochem. Geophys. Geosyst. 15, 4866-4884 (2014).

48            Pagani, M., Caldeira, K., Berner, R., & Beerling, D.J., The role of terrestrial plants in limiting atmospheric CO2 decline over the past 24 million years. Nature 460 (7251), 85-88 (2009).

49            Crucifix, M., Oscillators and relaxation phenomena in Pleistocene climate theory. Philosophical Transactions of the Royal Society A: Mathematical, Physical and Engineering Sciences 370 (1962), 1140-1165 (2012).

50            Magill, C.R., Ashley, G.M., & Freeman, K.H., Ecosystem variability and early human habitats in eastern Africa. Proceedings of the National Academy of Sciences 110 (4), 1167-1174 (2013).

51            Wrangham, R.W., Jones, J.H., Laden, G., Pilbeam, D., & Conklin-Brittain, N.L., The raw and the stolen: Cooking and the ecology of human origins. Current Anthropology 40, 567-590 (1999).

52            Pausas, J.G. & Keeley, J.E., A Burning Story: The Role of Fire in the History of Life. BioScience 59 (7), 593-601 (2009).

53            Roebroeks, W. & Villa, P., On the earliest evidence for habitual use of fire in Europe. Proceedings of the National Academy of Sciences 108 (13), 5209-5214 (2011).

54            Boyden, S.V., Biohistory: The Interplay Between Human Society and the Biosphere, Past and Present. (Parthenon Publishing Group, 1992).

55            Brown, K.S., Marean, C.W., Herries, A.I.R., Jacobs, Z., Tribolo, C., Braun, D., Roberts, D.L., Meyer, M.C., & Bernatchez, J., Fire As an Engineering Tool of Early Modern Humans. Science 325 (5942), 859-862 (2009).

56            Barnosky, A.D., Koch, P.L., Feranec, R.S., Wing, S.L., & Shabel, A.B., Assessing the Causes of Late Pleistocene Extinctions on the Continents. Science 306 (5693), 70-75 (2004).

57            Sandom, C., Faurby, S., Sandel, B., & Svenning, J.-C., Global late Quaternary megafauna extinctions linked to humans, not climate change. Proceedings of the Royal Society B: Biological Sciences 281 (1787) (2014).

58            Lynch, A.H., Abramson, D., Görgen, K., Beringer, J., & Uotila, P., Influence of savanna fire on Australian monsoon season precipitation and circulation as simulated using a distributed computing environment. Geophysical Research Letters 34 (20), L20801 (2007).

59            Miller, G., Mangan, J., Pollard, D., Thompson, S., Felzer, B., & Magee, J., Sensitivity of the Australian Monsoon to insolation and vegetation: Implications for human impact on continental moisture balance. Geology 33 (1), 65-68 (2005).

60            McWethy, D.B., Whitlock, C., Wilmshurst, J.M., McGlone, M.S., Fromont, M., Li, X., Dieffenbacher-Krall, A., Hobbs, W.O., Fritz, S.C., & Cook, E.R., Rapid landscape transformation in South Island, New Zealand, following initial Polynesian settlement. Proceedings of the National Academy of Sciences 107 (50), 21343-21348 (2010).

61            Nevle, R.J., Bird, D.K., Ruddiman, W.F., & Dull, R.A., Neotropical human-landscape interactions, fire, and atmospheric CO2 during European conquest. The Holocene 21 (5), 853-864 (2011).

62            Archibald, S., Staver, A.C., & Levin, S.A., Evolution of human-driven fire regimes in Africa. Proceedings of the National Academy of Sciences 109 (3), 847-852 (2012).

63            Diamond, J. & Bellwood, P., Farmers and Their Languages: The First Expansions. Science 300 (5619), 597-603 (2003).

64            Ellis, E.C., Kaplan, J.O., Fuller, D.Q., Vavrus, S., Klein Goldewijk, K., & Verburg, P.H., Used planet: A global history. Proceedings of the National Academy of Sciences 110 (20), 7978-7985 (2013).

65            Ruddiman, W.F., The Anthropocene. Annual Review of Earth and Planetary Sciences 41 (1), 45-68 (2013).

66            Jacobsen, T. & Adams, R.M., Salt and Silt in Ancient Mesopotamian Agriculture: Progressive changes in soil salinity and sedimentation contributed to the breakup of past civilizations. Science 128 (3334), 1251-1258 (1958).

67            Bogaard, A., Heaton, T.H.E., Poulton, P., & Merbach, I., The impact of manuring on nitrogen isotope ratios in cereals: archaeological implications for reconstruction of diet and crop management practices. Journal of Archaeological Science 34 (3), 335-343 (2007).

68            Kaplan, J.O., Krumhardt, K.M., Ellis, E.C., Ruddiman, W.F., Lemmen, C., & Klein Goldewijk, K., Holocene carbon emissions as a result of anthropogenic land cover change. The Holocene 21 (5), 775-791 (2011).

69            Devaraju, N., Bala, G., & Modak, A., Effects of large-scale deforestation on precipitation in the monsoon regions: Remote versus local effects. Proceedings of the National Academy of Sciences 112 (11), 3257-3262 (2015).

70            Mitchell, L., Brook, E., Lee, J.E., Buizert, C., & Sowers, T., Constraints on the Late Holocene Anthropogenic Contribution to the Atmospheric Methane Budget. Science 342 (6161), 964-966 (2013).

71            Ruddiman, W.F., The Anthropogenic Greenhouse Era Began Thousands of Years Ago. Climatic Change 61, 261-293 (2003).

72            Ruddiman, W.F., The Early Anthropogenic Hypothesis: Challenges and Responses. Reviews of Geophysics 45, RG4001 (2007).

73            Lenton, T.M., Pichler, P.P., & Weisz, H., Revolutions in energy input and material cycling in Earth history and human history. Earth Syst. Dynam. 7 (2), 353-370 (2016).

74            Haberl, H., Erb, K.H., Krausmann, F., Gaube, V., Bondeau, A., Plutzar, C., Gingrich, S., Lucht, W., & Fischer-Kowalski, M., Quantifying and mapping the human appropriation of net primary production in earth's terrestrial ecosystems. Proceedings of the National Academy of Sciences 104 (31), 12942-12947 (2007).

75            Matthews, E. et al., The Weight of Nations. Material Outflows from Industrial Economies. World Resources Institute, 2000.

76            Marland, G., Andres, R.J., & Boden, T.A. Global CO2 Emissions from Fossil-Fuel Burning, Cement Manufacture, and Gas Flaring: 1751-2005 in Trends: A Compendium of Data on Global Change (Carbon Dioxide Information Analysis Center, Oak Ridge National Laboratory, Oak Ridge, Tenn., U.S.A., 2008).

77            Galloway, J.N., Townsend, A.R., Erisman, J.W., Bekunda, M., Cai, Z., Freney, J.R., Martinelli, L.A., Seitzinger, S.P., & Sutton, M.A., Transformation of the Nitrogen Cycle: Recent Trends, Questions, and Potential Solutions. Science 320 (5878), 889-892 (2008).

78            Lenton, T.M., Land and ocean carbon cycle feedback effects on global warming in a simple Earth system model. Tellus 52B (5), 1159-1188 (2000).

79            Lenton, T.M., Held, H., Kriegler, E., Hall, J., Lucht, W., Rahmstorf, S., & Schellnhuber, H.J., Tipping Elements in the Earth's Climate System. PNAS 105 (6), 1786-1793 (2008).

80            Parmesan, C. & Yohe, G., A globally coherent fingerprint of climate change impacts across natural systems. Nature 421 (6918), 37-42 (2003).

81            Bellard, C., Bertelsmeier, C., Leadley, P., Thuiller, W., & Courchamp, F., Impacts of climate change on the future of biodiversity. Ecology Letters 15 (4), 365-377 (2012).

82            Oliver, T.H. & Morecroft, M.D., Interactions between climate change and land use change on biodiversity: attribution problems, risks, and opportunities. Wiley Interdisciplinary Reviews: Climate Change 5 (3), 317-335 (2014).

83            Thomas, C.D. et al., Extinction risk from climate change. Nature 427 (6970), 145-148 (2004).

84            Sala, O.E. et al., Global Biodiversity Scenarios for the Year 2100. Science 287 (5459), 1770-1774 (2000).

85            Harte, J., Ostling, A., Green, J.L., & Kinzig, A., Biodiversity conservation: Climate change and extinction risk. Nature 430 (6995) (2004).

86            Nathan, R., Long-Distance Dispersal of Plants. Science 313 (5788), 786-788 (2006).

87            Hirota, M., Holmgren, M., Van Nes, E.H., & Scheffer, M., Global Resilience of Tropical Forest and Savanna to Critical Transitions. Science 334 (6053), 232-235 (2011).

88            Staver, A.C., Archibald, S., & Levin, S.A., The Global Extent and Determinants of Savanna and Forest as Alternative Biome States. Science 334 (6053), 230-232 (2011).

89            Cox, P.M., Betts, R.A., Jones, C.D., Spall, S.A., & Totterdell, I.J., Acceleration of global warming due to carbon-cycle feedbacks in a coupled climate model. Nature 408, 184-187 (2000).

90            Jones, C., Lowe, J., Liddicoat, S., & Betts, R., Committed ecosystem change due to climate change. Nature Geoscience 2, 484-487 (2009).

91            Allen, C.D. et al., A global overview of drought and heat-induced tree mortality reveals emerging climate change risks for forests. Forest Ecology and Management 259 (4), 660-684 (2010).

92            Kurz, W.A., Dymond, C.C., Stinson, G., Rampley, G.J., Neilson, E.T., Carroll, A.L., Ebata, T., & Safranyik, L., Mountain pine beetle and forest carbon feedback to climate change. Nature 452, 987-990 (2008).

93            Kurz, W.A., Stinson, G., Rampley, G.J., Dymond, C.C., & Neilson, E.T., Risk of natural disturbances makes future contribution of Canada's forests to the global carbon cycle highly uncertain. Proceedings of the National Academy of Sciences 105 (5), 1551-1555 (2008).

94            Joos, F., Prentice, I.C., Sitch, S., Meyer, R., Hooss, G., Plattner, G.-K., Gerber, S., & Hasselmann, K., Global warming feedbacks on terrestrial carbon uptake under the Intergovernmental Panel on Climate Change (IPCC) emissions scenarios. Global Biogeochemical Cycles 15 (4), 891-907 (2001).

95            Lucht, W., Schaphoff, S., Erbrecht, T., Heyder, U., & Cramer, W., Terrestrial vegetation redistribution and carbon balance under climate change. Carbon Balance and Management 1, 6 (2006).

96            Hoegh-Guldberg, O. et al., Coral Reefs Under Rapid Climate Change and Ocean Acidification. Science 318 (5857), 1737-1742 (2007).

97            Veron, J.E.N., Hoegh-Guldberg, O., Lenton, T.M., Lough, J.M., Obura, D.O., Pearce-Kelly, P., Sheppard, C.R.C., Spalding, M., Stafford-Smith, M.G., & Rogers, A.D., The coral reef crisis: the critical importance of <350ppm CO2. Marine Pollution Bulletin 58, 1428-1437 (2009).

98            Guinotte, J.M., Orr, J., Cairns, S., Freiwald, A., Morgan, L., & George, R., Will human-induced changes in seawater chemistry alter the distribution of deep-sea scleractinian corals? Frontiers in Ecology and the Environment 4, 141-146 (2006).

99            Watson, A.J., Lenton, T.M., & Mills, B.J.W., Ocean de-oxygenation as a destabilizing feedback in the Earth system. Philosophical Transactions of the Royal Society (submitted).

100         Parmesan, C., Ecological and Evolutionary Responses to Recent Climate Change. Annual Review of Ecology, Evolution, and Systematics 37, 637-669 (2006).

101         Hoffmann, A.A. & Sgro, C.M., Climate change and evolutionary adaptation. Nature 470 (7335), 479-485 (2011).

102         Clark, J.R., Daines, S.J., Lenton, T.M., Watson, A.J., & Williams, H.T.P., Individual-based modelling of adaptation in marine microbial populations using genetically defined physiological parameters. Ecological Modelling 222 (23-24), 3823-3837 (2011).

103         Clark, J.R., Lenton, T.M., Williams, H.T.P., & Daines, S.J., Environmental selection and resource allocation determine global patterns in picophytoplankton cell size. Limnology and Oceanography 58 (3), 1008-1022 (2013).

104         Daines, S.J., Clark, J.R., & Lenton, T.M., Multiple environmental controls on phytoplankton growth strategies determine adaptive responses of the N:P ratio. Ecology Letters 17 (4), 414-425 (2014).

105         Toseland, A., Daines, S., Clark, J.R., Kirkham, A., Strauss, J., Uhlig, C., Lenton, T.M., Valentin, K., Pearson, G., Moulton, V., & Mock, T., The impact of temperature on marine phytoplankton resource allocation and metabolism Nature Climate Change 3, 979-984 (2013).

106         Lenton, T.M. & von Bloh, W., Biotic feedback extends the life span of the biosphere. Geophysical Research Letters 28 (9), 1715-1718 (2001).

107         Francis, P., Laudato Si’: On Care for Our Common Home. (Libreria Editrice Vaticana, Vatican City, Italy, 2015).

108         Sterner, T. & Persson, U.M., An Even Sterner Review: Introducing Relative Prices into the Discounting Debate. Review of Environmental Economics and Policy 2 (1), 61-76 (2008).

109         Cai, Y., Judd, K.L., Lenton, T.M., Lontzek, T.S., & Narita, D., Environmental tipping points significantly affect the cost−benefit assessment of climate policies. Proceedings of the National Academy of Sciences 112 (15), 4606-4611 (2015).

110         Hoel, M. & Sterner, T., Discounting and relative prices. Climatic Change 84 (3-4), 265-280 (2007).


Estinzione biologica

Come salvare l'ambiente naturale da cui dipendiamo Workshop PAS-PASS, Casina Pio IV, 27 febbraio-1... Continua