S269 Earth and Life
"The Dynamic Earth"
SUMMARY
In the final analysis, the Earth is hospitable to life because of its particular orbit around the Sun, which determines the amount of solar radiation that reaches it. The amount of solar radiation actually available to warm the Earth's surface is determined by how much is reflected rather than absorbed, ie by its albedo. AT the termperatures that obtain at the Earth's surface, water can exist as solid, liquid or gas: were it not for the presence of liquid water, life could not exist on Earth.
The fixation of carbon by primary producers (mainly plants) is the basis of all life on Earth.
Seasonal changes in incoming solar raidation are a result of the tilt of the Earth's axis in relation to the orbital plane (ie if the Earth's axis were at right- angles to the orbital plane, there would be no seasons). Over time-scales of tens of thousands to hundreds of thousands of years, there are cyclical variations in incoming solar raidation caused by changes in the degree of eccentricity of the Earth's orbit, in the angle of tilt of the Earth's axis, and in the direction in which the axis points. These cycles have periodicities of ~110 000 years, ~40 000 years and ~22 000 years, respectively, and are known as the Milankovich cycles or the Milankovich-Croll cycles. (See Hilma's page).
Most of the radiation energy emitted by the Sun is in the wavelength range of 0.15-5μ, which includes ultraviolet radiation, visible radiation and infrared radiation. Solar radiation is often referred to as short-wave radiation, to distinguish it from longer-wavelength (thermal) radiation emitted by the Earth (and any other body) as a consequence of having been warmed.
The luminosity of the Sun itself varies with time. It is thought that the amount of radiation emitted by the early Sun was only 70-75% of what it is now. There is evidence for variation in solar luminosity over an ~11-year cycle and over an ~80-90-year cycle (and there may well be longer-term variations, so far undetected).
Although the Earth's surface is heated unevenly by the Sun, the redistribution of heat by winds and currents ensures that low latitudes do not continually heat up and high latitudes continually cool down.
The continents have a much lower thermal capacity than the oceans, and so heat up and cool down much faster. This strongly affects the Earth's surface temperature distribution, particularly in the land-dominated Northern Hemisphere; in the oceanic areas of the Southern Hemisphere, temperature decreases more smoothly from Equator to poles, the temperature zones simply shifting northwards and southwards with the seasons.
Heat is redistributed over the surface of the Earth by winds, surface currents and the thermohaline curculation. Convection in the atmosphere is driven by warming from below; that in the deep ocean is driven by cooling (plus evaporation and/or ice formation) at the surface. The evaporation, transport and condensation of freshwater is intimately linked to the redistribution of heat. Rising air generally has a high moisture content (as at the Intertropical Convergence Zone); sinking air (as at the subtropical highs) is usually dry.
The Coriolis force acts to deflect flows to the right in the Northern Hemisphere and to the left in the Southern Hemisphere. The Coriolis effect increase with increasing latitude, with the result that atmospheric flow in low latitudes is largely vertical, and takes the form of Hadley cells, while that at higher latitudes is more slantwise or horizontal, forming anticyclonic and cyclonic weather systems. The formation of mid-latitude weather systems is determined by the behaviour of the polar jet stream, whose undulations at the top of the troposphere lead to the transport of enormous amounts of warm air polewards and cold air equatorwards. Extreme undulations of the polar jet stream may lead to severe droughts and unusual weather patterns. The ocean also has 'weather' in the form of mesoscale eddies, which cause problems for those modelling climate because they are an order of magnitude smaller than atmospheric cyclones and anticyclones.
Tropical regions where the ITCZ moves seasonally over land are subject to seasonally reversing winds (monsoons) and hence dry seasons alternating with seasons of heavy rainfall.
The atmosphere and ocean are tightly coupled by positive feedback loops, especially in tropical areas. Tropical cyclones are a small-scale manifestation of ths coupling; El Niño-Southern Oscillation (ENSO) events are Pacific-wide and have effects in the other oceans and at higher latitudes. Connections between widely separated phenomena, linked by the same climatic event, are known as teleconnections. Positive feedback loops cause instability in systems; negative feedback loops are stabilizing.
While light, suitable temperatures and availability of water are the factors limiting primary production (and hence most other life) on land, in the oceans the limiting factors are light and the availability of nutrients. Nutrients depleted from the sunlit surface layer may be returned to it via upwelling. Upwelling occurs along the Equator, along western coastlines under the Trade Winds, and as a result of cyclonic winds.
The formation of North Atlantic Deep Water results in the release to the atmosphere of large amounts of heat, reponsible for warming north-western Europe.
Gases in the atmosphere, notably carbon dioxide and water vapour, absorb outgoing long-wave radiation and re-radiate energy, much of it back to the Earth's surface. This trapping of radiant energy (which also involves clouds) is known as the greenhouse effect.
In the hydrological cycle, water moves through various reservoirs, of which the largest by far is the ocean. Residence time in a given reservoir is defined as the mass in the reservoir at any one time divided by the rate of transport into (or out of) it.
Rainwater is weakly acidic and interacts with minerals in rocks and soil causing chemical weathering. This results in constituents of the minerals (notably Ca2+, HCO3- and SiO2) going into solution, and eventually entering the ocean. The ocean plays a major role in the cycling of elements through the Earth system.
While phytoplankton live in near-surface waters, in the photic zone, animals and bacteria live throughout the ocean, at all depths. This is possible because the cold water masses sinking at high latitudes carry dissolved oxygen down into the deep ocean.
Clouds both reflect incoming solar radiation and absorb long-wave radiation. Although they are a very important component of the climate system, and potentially the main driving mechanism for changes in the Earth's radiation budget, their role is by no means fully understood. Cloud condensation nuclei include sulfate aerosols, produced by volcanoes and industry, as well as forming from DMS, a waste product of phytoplankton populations.
Climatic change on a particular time-scale is likely to be driven by forcing factors acting on a similar time-scale. For example, the 11-year sunspot cycle could drive decadal climatic change through the types of mechanisms that come into play during El Niño events, involving strong coupling between atmosphere and ocean in the tropics.
The chemistry of the carbon atom means that it has a unique role in the living world. Its ability to share electrons with other carbon atoms allows the construction of large complex molecules of carbon, hydrogen and oxygen, of which organic material is built.
Carbon fulfils two essential roles in the biosphere: it is the primary component of living tissue (see 1 above) and, in it gaseous forms (CO2 and CH4), it allows the surface of the Earth to be warm enough to support life.
The main biogeochemical connection between Earth and life is the global carbon cycle: the movement of carbon through the atmosphere, biosphere, lithosphere and ocean. The global carbon cycle involves interlinking cycles over three major time-scales: (i) the terrestrial carbon cycle, driven by biological processes and acting over time-scales of months/years to decades; (ii) the marine carbon cycle involving chemical, biological and physical components, and acting over an intermediate time-scale of up to hundreds of thousands of years; (iii) the geological carbon cycle, involving rocks and sediemtns, and acting on time-scales of up to hundreds of millions of years.
The terrestrial carbon cycle is driven by the fixation of atmospheric carbon into organic matter via photsynthesis, and involves large fluxes of carbon between the atmosphere and fairly small organic reservoirs (living organic matter and organic debris). The average residence time of carbon in plant biomass is about nine years. Estimates of annual net primary production (flux of carbon into the terrestrial biomass reservoir) and of standing stock are problematic, and involve determining the areas occupied by different biomes.
The marine carbon cycle is linked to the terrestrial carbon cycle through the carbon carried to the sea in rivers in organic and inorganic form (both dissolved and particulate). The surface ocean is, on average, in equilibrium with the overlying atmosphere and so the fluxes between the two more or less balance. Fluxes of carbon (CO2) across the air-sea interface may be driven by physical processes (particularly sinking of cold water masses in high latutides, and upwelling), and by biological activity (ie algal photosynthesis which, when combined with rapid sinking of organic matter into the deep sea, constitutes the biological pump). In both cases, however, the mechanism whereby CO2 is 'pushed out' or 'drawn into' the ocean is chemical. The linked equilibria determine the relative proportions of CO2 gas and the various forms of dissolved inorganic carbon: H2CO3 (carbonic acid), HCO3- (bicarbonate ion, the main constituent in seawater) and CO3- (carbonate ion); these quilibria are together known as the carbonate system. Biological activity produces inorganic particulate carbon in the form of shells and skeletons which eventually sink to the sea-bed; whether they are dissolved or accumulate depends on the chemistry of the deep ocean water and the rate of supply of remains. The depth at which the proportion of calcium carbonate remains falls to less than 20% of the total sediment is known as the carbonate compensation depth.
In the geological carbon cycle, carbon in the organic and inorganic products of weathering on land is carried to the sea in rivers, takes part in the marine carbon cycle, and - in the case of a very small proportion is preserved and buried, eventually to be returned to the atmosphere or ocean via volcanism or via weathering and/or oxidation. The geological carbon cycle acts on a time-scale of up to hundreds of millions of years, and invloves small fluxes between large reservoirs; carbonate and organic sedimentary rocks together store more than 99.9% of the carbon on Earth.
Extracting and burning fossil fuels shortcuts the geological-scale return flux of carbon and has currently brought the global carbon cycle into disequlibrium. As a result,, the concentration of CO2 in the atmosphere is rising (as are those of other greenhouse gases), but it is not rising as much as expected. Some of the 'extra' CO2 has almost certainly been taken up by the ocean; the rest (or some of it) is being taken up by increased net primary productivity of land plants, notably in forests.
Long-term stability of the carbon cycle is controlled by negative feedbacks acting on the atmospheric carbon reservoir via rates of silicate weathering on land followed by deposition and preservation of carbonates in marine sediments, and (possibly) the deposition and preservation of organic carbon, mainly in marine sediments.
Volcanoes are conduits linking deep mantle processes to atmospheric composition and hence climate; they exhale large amounts of water, CO2 and SO2. Over geological time-scales, volcanoes have an important role in the carbon cycle, but in the short term their contribution to atmospheric CO2 has been overwhelmed by the anthropogenic flux. Major eruptions, like that of Mount Pinatubo in 1991, pump large amounts of SO2 into the stratosphere. Over time, the SO2 combines with water to form sulfuric acid aerosols (H2SO4). These change the Earth's radiative balance, reflecting a proportion of incoming solar radiation back into space.
The effects of an eruption on climate are dictated by:
the magnitude of the eruption;
the sulfur content of the magma;
whether or not the volcanic gases get into the stratosphere;
the location of the eruption site on the globe;
the type of eruption.
Eruptions taking place in the topics are most likely to have global effects. Explosive eruptions are more likely than effusions of lavas to inject gases into the stratosphere.
The radiative effects of major explosive eruptions can be large, reducing the amount of solar radiation reaching the Earth's surface by more than 10%. The effects of most explosive eruptions, however, are short-lived (2-3 years), as aerosols fall out of the stratosphere. Because of the long response time of the Earth's climate system as a whole, brief volcanic events do not have a significant effect on climate. While the effects of some large historic eruptions such as that of Tambora in 1815 (followed by the 'Year without a Summer' in 1816) are well documented, it is not possible to scale directly up from these to deduce the effects of much larger prehistoric eruptions. For instance, there may be a limit to the aerosol loading the stratosphere can sustain.
Eruptions of flood basalts such as those of the Columbia River Province may involve effusion of more than 1000 km3 of sulfur-rich basalt lava over periods of 10-100 years. Because these eruptions are sustained over longer periods than great explosive events, their radiative and environmental effects may be more profound. The 1783 Laki Fissure eruption in Iceland, which yielded 14 km3 of lave, was followed by an exceptionally severe winter in North America and Europe, and widespread starvation of humans in Iceland and elsewhere. The close coincidence in timing between great episodes of flood basalts in Earth history with major mass extinctions - such as those at the Cretaceous-Tertiary and Paleozoic-Mesozoic boundaries - suggests that mass extinctions may be linked with the environmental and climatic effects of the flood basalts.
Major flood basalt provinces are themselves expressions of plumes rising through the Earth from the core-mantle boundary. Occasional massive convective overturning events in the mantle may lead to 'superplumes', expressed at the surface in the generation of large igneous provinces such as the Ontong-Java Plateau. The additional CO2 released into the atmosphere by superplumes may be linked with the climatic warming which occurred during the Cretaceous Period.
Climatic conditions for a particular continental area change sumply as a result of that landmass moving over the globe through different climatic belts. In the case of the landmass that is now southern Britain, paleomagnetic measurements and paleoclimatic indicators in the rock record show that over the past 450 million years or so, this particular area of continental crust moved northwards from southern mid-latitudes to northern mid-latitudes, crossing a number of global climatic belts (dry subtropical, wet tropical/equatorial, dry subtropical again) in the process. The global distribution of paleoclimatic indicators for the Carboniferous and Permian provide evidence consistent with the existence of a supercontinent (Pangea), which had an ice-cap around the South Pole and tropical forests in the vicinity of the Equator. There may have been high-latitude forests later in the Permian. Reconstructions of continental configurations using paleoclimatic indicators must be apporached with caution because: (1) paleoclimatic indicators may be ambiguous; (2) climatic belts do not, in reality, run simply east-west; (3) we cannot know for sure whether the tilt of the Earth's axis, and hence the seasonally varying distribution of solar energy over the Earth's surface, was similar in the past to what it is today.
Climate models using highly simplified continental configurations suggest that a tropical 'ring world' would b esignificantly warmer than polar 'cap worlds'. It is thought that the geography of the Earth may have approximated to the former 700-600 million years agao, and to the latter (with one polar ice-cap) during the late Carboniferous.
It is thought that the changing distribution of continents in response to plate-tectonic processes affects global climate on the million-year time-scale through its effect on the radiation budget, and hence indirectly on the hydrological cycle and weathering (which affects the CO2 concentration of the atmosphere). Furthermore, the closing and opening of oceanic gateways as continents change their relative positions has a significant effect on shallow and deep oceanic circulation, and can contribute to global warming and cooling on significantly shorter time-scales.
The break-up of Pangea began at about 200 Ma. The opening of a low-latitude seaway may have contributed to global warming in the Cretaceous, at about 100 Ma. Around this time, deep water masses were probably warm and very saline, having formed at low latitudes. It is though that thermal isolation resulting from the initiation of the Antarctic Circumpolar Current at 25 Ma accelerated cooling of Antarctica and the growth of the south polar ice-cap.
Ice Ages are periods when the Earth has large areas of ice-cover. Within Ice AGes, there are glacials and interglacials when ice-sheets alternately grow and retreat (and sea-levels fall and rise). Between 1.5 and 1 million years ago, these alternations occurred with a cyclicity of ~40 000 years (corresponding to the tilt component of the Milankovich cycles) but by 600 000 years ago the 110 000-year eccentricity cycle had become the most dominant. Our knowledge of changes in ice volume (and hence sea-level), and of global temperature, largely derives from δ18) values obtained from the remains of foraminiferans (notably planktonic species).
Relative sea-level change results form a combination of eustatic and isostatic (epeirogenic) sea-level change. Eustatic changes are global in extent, whereas isostatic changes result from local or regional uplift or subsidence of the lithosphere. Eustatic sea-level changes are due either to changes in the volume of ocean waters (resulting from the formation and melting of ice-sheets), or to changes in the size and shape of the ocean basins (resulting either from the formation of new oceanic crust, notably as submarine plateaux, or from the replacement of a few large ocean basins by a number of smaller ones). During periods of global warming, some sea-level rise is attribuatable to expansion of the ocean water; however, an increas of 10 deg C throughout the water column in all oceans would lead to a sea-level rise of only ~ 10m.
Over the past 540 million years, there have been two periods of high global sea-level - in the Ordovician-Silurian and the Cretaceous - while sea-levels were generally low in the permo-Carboniferous and from the later Tertiary to the present time. The distribution of shallow-water sediments, and paleogeographic maps of the proportion of continents flooded at different periods in the Earth's history, can be used to determine past changes in sea-level. During glacial periods over the past 2.5 million years, sea-level has been as much as 150 m below its present-day level, and in the Cretaceous (about 100 million years ago), it was probably 250 m higher than it is today (and some researchers believe that it was even more than this). Because of the extensive areas of flat coastal plains and continental shelves, fairly small changes in sea-level can dramatically alter the relative areas of land and shallow coastal seas, with important implications for the operation of the carbon cycle, and hence for climate.
Global warming during the Cretaceous was probably the result of the addition to the atmosphere of huge amounts of CO2 as a result of volcanism (as flood basalts and as a consequence of increased rates of sea-floor spreading). However, the warming was counteracted by the removal of atmospheric CO2, through deposition and preservation of carbon in the ocean. Large amounts of organic carbon were preserved and buried in the warm anoxic oceans, and large amounts of inorganic carbon were precipitated in shallow coastal seas (and, to some extent, via the newly established deep-water 'carbonate factory').
Snow-covered mountains (particularly those at low latitudes) affect the global climate through their effect on the Earth's albedo. Over geological time-scales, the uplift of mountains as a result of continental collision can affect the operation of the global carbon cycle, and hence the amount of CO2 in the atmosphere, in various ways. Subduction of ocean-floor sediemtns and collision of continental masses bothe result in decarbonation, whereby silica and calcium carbonate react together to produce CO2, which eventually escapes to the atmosphere. However, mountains are sites of vigorous weathering, physical at high altitudes and chemical lower down; chemical weathering (followed by accumulation and burial of carbon in the ocean) results in a loss of CO2 from the atmosphere. (In addition, continental collision may result in large volumes of carbon in rocks being removed from contact with the atmosphere for many millions of years).
Mountain ranges and plateaux affect the climate physically by redirecting air masses around and/or over them. Moisture-laden winds release their precipitation on the windward side while the leeward side (in the case of the HImalayas, the Tibetan Plateau) is dry. The rise of the Himalayas and Tibet is believed to have intensified the strength of the South-West Monsoon by providing a source of heat (in particular, latent hear), at a critical position in the atmospheric circulation.
The collision between India and Eurasia at about 50 Ma was reponsible for uplift of the Tibetan Plateay and teh Himalayas. Continued thickening of the lithosphere beneath Tibet led to convective thinning followed by rapid uplift.
The rate of uplift of the Tibetan Plateau is not well established. Attempts to measure it are being made by use of fossil plants and dating of normal faults. There is some evidence that the plateau reached its present elevation between 14 and 7 million years ago.
The history of climate change in southern Asia is not well established. Attempts to determine it include studies of foraminiferans in the Indian Ocean, studies of δ13C (to throw light on plant growth) and analysis of the types of sediments eroded off the Himalayas. There is some evidence to suggest strengthening of the monsoon between 9 and 6 million years ago.
The Earth's climate has cooled markedly over the past 50 million years. The Raymo-Ruddiman hypothesis postulates that uplift of teh Himalayas and Tibet imposed cooling on the global climate, by strengthening the South-West Monsoon and hence increasing rates of chemical weathering: increased weathering of silicates followed by accumulation and burial of carbonates and/or organic carbons in the ocean is assumed to result in the long-term removal of CO2 from the atmosphere. To prevent 'runaway cooling', the rapid removal of CO2 by weathering would need to be partially compensated for. Possible mechanisms for replenishing atmospheric CO2 include decarbonation at subduction zones and the oxidation of organic carbon from exhumed sediments.
The steady-state, carbon cycle model (ie the BLAG model) interprets the role of mountains as providing a negative feedback loop that stabilizes fluctuations in temperature resulting from variations in volcanism associated with the production of new sea-floor.
The isotope ratio 87Sr/86Sr of the world's oceans has increased over the past 50 million years. This has been interpreted as indicating increased weathering rates due to the uplift of Tibet and more especially the Himalayas. This interpretation is rupported by the Sr fluxes and isotope ratios of rivers currently eroding the Himalayas.
Sediments laid down in the Arabian Sea over the past 400 000 years show cyclical variations in the strength of the South-West Monsoon winds, in primary productivity and in aridity of continental landmasses. These variations are related to glacial/interglacial cycles, i.e. to the Milankovich cycles. The challenge for those wishing to study climatic change is to disentagle the various types of change occurring in response to different forcing factors over a number of different time-scales.
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