56 —■—■—■—1—■—■—■—1—■—■—■—1—■—■—■—1—■—■—■— 1984 1988 1992 1996 2000 2004


FlG. 1.2. The variation of the global cloud cover based on ISCCP data for the period 1984-2004.

cloud cover and the radiative properties of the many types of cloud. Cloud cover is highly variable both temporally and spatially. Cloud overlap can modify rapidly the resultant scattering and absorbing properties of the cloudy sky. The radiative properties of droplets are different from ice crystals, and different condensation processes can occur rapidly.

In addition to the airborne droplets or ice crystals that make up clouds, lower concentrations of much smaller solid and liquid particles (aerosols) exist everywhere and can affect climate both directly and indirectly. Some have a complex chemical composition that results in a variety of radiation absorbing and scattering properties. Those generated naturally by volcanism and sea-spray evaporation, consisting mainly of sulphate and chlorides, are scatterers that tend to produce a cooling of the planet, while those that arise from biomass and fossil-fuel burning also contain carbonaceous materials that are strongly absorbing and thus warming agents. Other types of aerosols include irregularly shaped particles of various types of minerals, produced by aeolian processes, especially dust storms. The water content of aerosols, which depends on the ambient relative humidity, plays an important role in their radiative properties. Their atmospheric residence time is of the order of days to weeks, and they have a highly variable global distribution.

Tropospheric aerosol loads are thought to have increased over recent years due to increased anthropogenic emissions of particles and their precursor gases. The inclusion of aerosols is now necessary in climatic change studies that examine changes in surface and atmospheric temperatures, snow and ice-cover extent, sea-level rise, precipitation changes, frequency and intensity of extreme weather events, and desertification, since aerosols crucially affect the radiation budget at the top of the atmosphere, in the atmosphere, and at the surface, and hence atmospheric dynamics as well as evaporation and surface energy balance. Aerosol particles affect the Earth's climate by influencing its radiation energy budget in two ways; i) directly by scattering and absorbing solar radiation, and ii) indirectly, by modifying the cloud microphysical properties. Aerosols can act as cloud condensation nuclei, affecting cloud optical and radiative properties (cloud albedo and optical thickness) or cloud amount, lifetime and precipitation efficiency. In climate research, these are known as the first and second indirect aerosol effects, respectively.

Modern climate models must attempt to incorporate both the direct and indirect radiative forcing effects of aerosols. However, despite significant progress in understanding the role of aerosols in climate, it still remains a large uncertainty. Aerosols exhibit large spatial-temporal variability and heterogeneity associated with their short atmospheric lifetime and complex interactions with clouds in terms of both their physical and chemical properties. In order to improve the aerosol parameterization schemes in climate models, and the accuracy of estimated aerosol radiative forcings, monitoring of the spatial and temporal distributions of aerosol physical, chemical, optical and radiative properties is required both on global and local scales. To this aim, a worldwide effort has been undertaken since the early 1990s, in order to produce a global aerosol climatology by combining satellite-based observations (such as Total Ozone Mapping Spectrometer, TOMS, Moderate Resolution Imaging Spectro-Radiometer, MODIS, Multiangle Imaging Spectroradiometer, MISR, Polarization and Directionality of the Earth's Reflectance, POLDER, with measurements from ground-based monitoring networks such as the Aerosol Robotic Network (AERONET). To retrieve aerosol properties accurately, these data sets are complemented by field experimental campaigns, both ground based and airborne, conducted in climatically important regions. These serve at the same time as calibration and validation tools for satellite data.

1.5 Radiative equilibrium and radiative forcing

The dominant factor determining the surface temperature of a planet is the balance between net incoming solar radiation and outgoing thermal infra-red radiation. That is, radiative equilibrium can be assumed for most practical purposes, although strictly speaking it is never achieved, as time-dependent and non-radiative processes are always present. The processes that control the net solar radiation absorbed by the planet are complex and, as we have seen, involve the multiple scattering and absorption of solar radiation by atmospheric molecules, clouds, aerosols and the surface. The outgoing thermal infra-red radiation is controlled by the temperature of the surface and by the gases, clouds, and other particles in the atmosphere. The temperature is determined in turn by the local energy balance.

In the absence of an atmosphere and without significant internal heat sources, it is straightforward to calculate the solar radiation absorbed by the surface of a planet and its re-emission as infra-red radiation to space, and hence the surface temperature. If an atmosphere is present that contains infra-red active molecules (which means, in general, those composed of different types of atoms, such as H2O and CO2) these control the amount of infra-red radiation absorbed by the atmosphere, preventing it from escaping directly to space. Radiation from the atmosphere back towards the surface raises its temperature above that of an airless body with the same albedo, giving rise to what we (somewhat inaccurately) call the greenhouse effect. The atmosphere thus plays a crucial role in determining the climate of a planet, particularly since the greenhouse gases are present on the Earth in quite small proportions that are prone to change due to natural and man-made phenomena (e.g. large volcanic eruptions, and extensive deforestation, respectively).

The response to any change affecting the radiative properties of the atmosphere or the surface is ultimately a change in the surface temperature of the planet and hence its climate. The term radiative forcing is used to denote an externally imposed perturbation in any component of the radiative energy budget of the climate system. Radiative forcing is a measure of the change in radiative fluxes within the atmosphere-surface system, especially those of the net solar and terrestrial fluxes at the top of the atmosphere and at the Earth's surface. In making assessments of climate change, bodies such as the IPCC have focused on the changes in radiative forcing between pre-industrial times and the present, the recent past, and the future. The ability to monitor the TOA fluxes and climatological parameters over the globe by satellites, e.g. ERBE (Earth Radiation Budget Experiment) for radiative fluxes and ISCCP (International Satellite Cloud Climatology Project) for cloud cover and albedo, has provided climate modellers the opportunity to validate their predictions and improve their understanding of radiative process against measurements on a planetary scale. The radiative forcing due to increases of the well-mixed greenhouse gases from 1750 to 2000 is estimated by IPCC to be 2.43 W m~2, broken down by species in Table 1.2. The net forcing can be compared to a value of 1366 W m~2 for the total solar irradiance (TSI) (also known as the solar constant), defined as the total radiative energy flux from the Sun at the mean distance of the Earth's orbit.

The tropospheric increase in ozone corresponds to a positive radiative forcing of 0.35 W m2. Ozone forcing varies considerably by region and responds much more quickly to changes in emissions than the long-lived greenhouse gases, such as CO2. The observed depletion of the stratospheric ozone layer from 1979 to 2000 is estimated to have caused a negative radiative forcing (-0.15 W m2 ). Assuming

Table 1.2 Pre-industrial (1750) and present (1998) abundances of well-mixed greenhouse gases and the radiative forcing due to the change in abundance. Volume mixing ratios for CO2 are in ppm, for CH4 and N2 O in ppb, and for the rest in ppt. (Source: IPCC 2001)


Abundance 1750

Abundance 1998

Radiative forcing (W m 2)























full compliance with current halocarbon regulations, the positive forcing of the halocarbons will be reduced, as will the magnitude of the negative forcing from stratospheric ozone depletion as the ozone layer recovers over the twenty-first century.

1.6 Atmospheric circulation

Much of the radiative behaviour of the atmosphere is a response to the transfer of energy within the climate system by dynamics. Of particular importance for the global radiative energy balance is the poleward transfer of energy, to balance the decrease with latitude of the input from the Sun. About half of this transfer, amounting to several petawatts (1 PW = 1015 W) occurs in the atmosphere, and the rest in the oceans. The ocean circulation is coupled to that of the atmosphere by momentum and heat transfer between near-surface winds and the water surface.

The circulation of the atmosphere transports momentum, cloud, water, ozone and pollution, as well as heat, around the globe. The mean or 'general' circulation consists primarily of rising air in the strongly heated low-latitude regions, with descent over the poles where radiative cooling exceeds heating. The warm air cools during its movement polewards, descends and moves at low levels back towards the equator, completing the large circulation pattern known as a Hadley cell. The rapid rotation of the Earth means the Coriolis effect adds a component parallel to the equator to the poleward flow in the Hadley cell, which closes the cell, by forcing the north-south flow into an east-west direction, before it can reach high latitudes. Further cells develop polewards of this and the mean circulation consists of three circulation cells, the Hadley, Ferrel, and Polar cells, in each hemisphere (Fig. 1.3).

The rising air at the equator creates a band called the intertropical convergence zone, which draws in low-level air from the subtropics. In this region, also known as the doldrums, horizontal pressure gradients are weak and winds are light. From there air rises towards the tropopause and flows polewards, gradually deflecting

FlG. 1.3. A schematic diagram of the general circulation of the atmosphere of the Earth, as described in the text. On a fast-rotating planet like the Earth, the thermally driven equator-to-pole flow is deflected by the Coriolis acceleration, towards the west for equatorwards motion, and vice versa, in both hemispheres. This renders a single equator-to-pole cell unstable, and instead three cells, known as the Hadley, Ferrel and Polar cells, are found to dominate the general circulation when wave and chaotic motions are averaged out.

FlG. 1.3. A schematic diagram of the general circulation of the atmosphere of the Earth, as described in the text. On a fast-rotating planet like the Earth, the thermally driven equator-to-pole flow is deflected by the Coriolis acceleration, towards the west for equatorwards motion, and vice versa, in both hemispheres. This renders a single equator-to-pole cell unstable, and instead three cells, known as the Hadley, Ferrel and Polar cells, are found to dominate the general circulation when wave and chaotic motions are averaged out.

under the influence of the Coriolis force, towards the east, until by about 30° latitude the wind is predominantly zonal, forming the subtropical jet stream. The upper air then sinks back to the surface creating a high-pressure zone with low surface winds known as the horse latitudes. From there the air moving back towards the equator to complete the Hadley cell is deflected west, creating the trade winds.

The surface air moving polewards from the horse latitudes is deflected westwards by the Coriolis force, producing the westerlies. In these latitudes, roughly 30 to 60°, upper air winds tend to blow towards the poles and are deflected eastwards to the polar jet stream at the high-latitude limit of the Ferrel cell. Poleward of that, the rapidly decreasing distance of the surface from the axis of rotation means the atmosphere in the polar cell tends to spin up to conserve its angular momentum, forming the polar vortex. Inside this, air descends rapidly before commencing its journey equatorwards. Where it meets the mild air moving polewards in the Ferrel cell, it forms the subpolar low, a region noted for storm formation.

The long-term mean atmospheric circulation accounts for only about 25% of the polewards transport of heat required to balance the solar input and the global radiative cooling of the Earth. Another ^25% is by eddies, including midlatitude storm systems and waves and turbulence on a wide range of temporal and spatial scales, and the rest in the ocean. The rate at which this transfer takes place seems to be such as to maximize the rate of entropy production, dS/dt. This would be zero if the transport stopped, and also if the motions were very rapid. The actual situation is somewhere in between and is postulated, with some support from studies using general circulation models, to be that for which dS/dt finds a maximum value.

Thus it appears that dynamical energy transfer within the climate system not only tends towards a state of maximum entropy, as required by the Second Law of Thermodynamics, but also that it moves in this direction at the fastest possible rate. The explanation, broadly speaking, is that the processes involved are complex and chaotic, and so offer an essentially infinite number of multiple pathways towards the state of maximum disorder. The absorption, transfer and emission of radiation by the climate system, which involves a larger rate of entropy production than the dynamical transport of heat, does not have this property and does not behave in this way.

1.7 The hydrosphere and the cryosphere

The hydrosphere and cryosphere are the liquid and solid water components of the climate system. The hydrosphere comprises the surface-water bodies such as rivers, lakes, and oceans that cover about 70% of the Earth's surface, and groundwater in aquifers. Most of the precipitation on land returns to the atmosphere via evaporation, a smaller part runs into the sea, and some soaks into underground aquifers. Runoff to the sea and oceans delivers nutrients and salts that affect their heat capacity and hence modify the oceanic circulation. The cryosphere comprises the ice sheets of Greenland and Antarctica, continental glaciers, snow cover, sea ice in the Arctic and Southern oceans and permafrost. Its high reflectivity to solar radiation provides a cooling mechanism that can change relatively rapidly as condensation and melting processes have the potential to alter the planetary albedo from 5% to 80%. Ice sheets store a large amount of water and so are a potential source of sea-level variations.

1.7.1 Oceanic circulation

Any understanding of the role of the atmospheric circulation in global energy balance is not complete without a corresponding appreciation of the energy budget of the ocean. However, a crucial difference between the two, apart from the obvious differences in mass, density and heat capacity, is the almost completely negligible role played in ocean energetics by radiative transfer. Liquid water is effectively opaque to electromagnetic radiation at virtually all wavelengths, a further consequence of which is the difficulty of gathering data by remote sensing.

The ocean circulation has two fairly distinct components. The top few hundred metres, containing the mixed and thermocline layers, is driven primarily by drag between the water surface and the prevailing winds. The deep-ocean circulation, however, is dominated by differences in density resulting from temperature and salinity gradients. In both cases, the long-term average motions (that is, with waves and turbulence averaged out) are determined mainly by the balance between the main driving force (wind drag and density gradients, respectively) and the Coriolis force that enters as a result of the Earth's rotation.

The deep-ocean layer constitutes more than 90% of the mass of the ocean, i.e. all but the top few hundred metres of depth on average. It extends all the way to the surface near the poles, where the water is particularly cold and dense, resulting in strong downwelling towards the ocean floor. This rapid sinking motion is focused in relatively small areas near Greenland in the North Atlantic, and in the Weddell Sea in the Antarctic. Here, particularly cold and dense surface waters are found due to low solar heating, strong evaporative cooling aided by the prevailing high winds, and brine enrichment by freezing. The last process occurs in reverse when polar icecaps melt, and become a source of relatively fresh, low-density water, like melting snowfields on land or heavy rainfall. One of the concerns in a world undergoing global warming is that the fresh water released by melting ice caps will reduce the salinity, the density, and hence the downwelling in the polar seas, and so slow down or even reverse the deep-ocean circulation.

Over the rest of the globe, the vertical motions in the ocean are generally much weaker, and generally upwards to balance the strong downward flow in the Arctic and Antarctic. The resulting slow, global, overturning motion known as the thermohaline circulation transfers heat, salt and tracers around the globe. The motions near the surface are mainly the result of wind drag, transferring momentum from the atmosphere to the surface and near-surface layers of the water. The wind also generates surface waves that further affect the coupling between wind and water, and produce vertical as well as horizontal motions. These cause turbulent mixing of the top ~ 10-100 m, producing the nearly isothermal mixed layer.

Met Office Hadiey Centre

FiG. 1.4. The North Atlantic Drift, or Gulf Stream, showing the flow of warm tropical surface currents northwards to the North Sea, and the flow of deep cold currents southwards. Two convection areas convert warm surface water to cold deep water that drives the Gulf Stream circulation. This circulation is responsible for the milder climate in the Eastern North Atlantic region compared to the Western. ©Crown copyright (UK Met Office/Hadley Centre 2005)

Met Office Hadiey Centre

FiG. 1.4. The North Atlantic Drift, or Gulf Stream, showing the flow of warm tropical surface currents northwards to the North Sea, and the flow of deep cold currents southwards. Two convection areas convert warm surface water to cold deep water that drives the Gulf Stream circulation. This circulation is responsible for the milder climate in the Eastern North Atlantic region compared to the Western. ©Crown copyright (UK Met Office/Hadley Centre 2005)

The general pattern of surface currents in the ocean is well known from centuries of navigation and measurement. In the Northern Hemisphere, each of the ocean basins contains two counter-rotating gyres (oceanic current systems of planetary scale driven by global wind systems) in equilibrium with the wind forcing. The Southern Hemisphere circulation has a single, large gyre, because of the existence of the strong Antarctic circumpolar current, where the water can travel right around the globe without meeting a continental barrier. The gyres are characterized by currents that are much stronger on the western boundary than on the eastern. The western boundary current in the North Atlantic is the North Atlantic Drift or Gulf Stream, which transports 100 million cubic metres of warm water per second in a stream about 100 km wide, off the east coast of North America (Fig. 1.4).

Solar energy may be stored locally by the sea or ocean during summer and released back to the atmospheric environment during the winter months (§8.9). Absorbed solar energy can also be transferred regionally, affecting climate on a subglobal scale, as done by the Gulf Stream. As shown in Fig. 1.4, heat is transported by surface waters flowing north and eastward, with cold saline waters from the North returning at depth, and so conditions at the same latitude are very different either side of the Atlantic basin.

The oceanic large heat capacity provides a buffer against rapid climatic changes. However, oceans control the climate system by redistributing heat from the warm equatorial regions to the cold polar regions. Ocean circulation can be modified by long-term changes in precipitation, continental runoff, evaporation, sea-ice formation, cloud cover and prevailing winds. Variations in oceanic heat transport can alter global distributions of heat and sea-surface temperatures. There is thus a feedback since sea-surface temperatures control evaporation, atmospheric water vapour, clouds, precipitation and surface albedo (snow cover), and hence global climate.

1.7.2 El Niño southern oscillation (ENSO)

The important effect that ocean circulation changes can have on the global climate is illustrated by the role of certain natural circulation patterns in the interannual and longer-term variability of the climate. The strongest natural fluctuation of climate on interannual timescales is the ENSO phenomenon, a large, quasiperiodic climatic variation that arises from an interaction between the tropical Pacific Ocean and the overlying atmosphere.

ENSO is a natural cycle that couples the ocean-atmosphere system over the tropical Pacific and operates on a timescale of 2-7 years. Once developed, it causes a shift in the seasonal temperature and precipitation patterns in many different regions of the world, since heating of the tropical atmosphere creates changes in the global atmospheric circulation. Thus, ENSO is a dominant source of inter-annual climate variability around the world. Normally, the equatorial Pacific Ocean is characterized by warm waters in the west and cold waters in the east. The ENSO warm phase (El Niño) is associated with an unusual warming of the eastern and central equatorial Pacific. La Niña, the ENSO cold phase, is the counterpart to El Nino, often following it. It is characterized by cooler than normal sea surface temperatures across the equatorial eastern Pacific. Thus, ENSO is an oscillation between warm and cold events with a peak that typically occurs late in the calendar year (late December to early January). Both El Nino and La Niña events last for about a year, but they can last for as long as 18 months. During ENSO, a feedback between atmospheric and ocean properties is observed. Sea-surface temperature (SST) anomalies induce wind-stress anomaly patterns that in turn produce a positive feedback on the SST. Variations of the above properties cause significant changes in other oceanic and atmospheric variables, e.g. the mean depth of the thermocline, the water-vapour content of the atmosphere and the relative distributions of low, middle and high clouds.

Water vapour and clouds are the main regulators of the radiative heating of the planet since changes in these parameters modulate the variability in the radiation fluxes that control the heating or cooling of the Earth's surface and atmosphere. The radiation field, in turn, influences SST and atmospheric water vapour. Thus ENSO involves complex climatic processes and feedbacks that make its onset time, duration, strength and spatial structure difficult to predict. International monitoring programmes of the coupled atmosphere-ocean system started in the Pacific around 1985 and led to the TAO-TRITON (Tropical Atmosphere Ocean project/Triangle Trans Ocean Buoy Network) array of moored buoys. The aim of this programme is to provide real-time measurements of winds, sea-surface temperature, subsurface temperature, sea level and ocean flow that help in the understanding of the physical processes responsible for ENSO. The variability and the spatial distribution of the ocean and atmospheric variables are not the same for all ENSO events. Thus, a definition of ENSO is necessary for the study of this phenomenon. The phase and strength of ENSO events are defined by an index. Several different indices have been used in the literature, mostly based on SST, although there is one index, the Southern Oscillation index, which is related to air-pressure differences at sea level, between Darwin (Australia) in the west and Tahiti in the east. The SST-based indices are obtained from the SST anomalies with respect to average values over some specified region of the ocean. There has also been an effort to combine several atmospheric-oceanic variables into a single index like the multivariate ENSO index. Averages of 850 mb wind, outgoing longwave radiation (OLR) at the top of the atmosphere as well as precipitation over specific regions are also used, although not often, to monitor ENSO.

The Earth's climate system is driven by the radiative energy balance between the solar shortwave radiation (SW) absorbed by the atmosphere and the surface of the Earth and the thermal longwave radiation (LW) emitted by the Earth to space. In this respect, ENSO events are expected to be associated with the spatial and temporal variability of the radiative energy balance over the tropical and subtropical Pacific. The net heat flux into the ocean plays a key role in ENSO evolution and it is a significant variable in the models that have been developed to make ENSO predictions. The variation of the net heat flux during ENSO events is of paramount importance to the dynamics of the system. The net heat flux into the ocean is a small residual of four terms, the downward shortwave radiation at the surface (DSR), the latent heat loss, the sensible heat transfer and the net longwave radiation to the Earth's surface (NSL). The NSL is the difference between the downward longwave radiation (DLR) at the Earth's surface and the Earth's surface thermal emission. The DLR at the Earth's surface is a very important component of the surface radiation budget with variations arising from increases in greenhouse gases or from changes in other atmospheric properties that occur during ENSO events (Intergovernmental Panel on Climate Change, IPCC 2001).

The resulting extreme phases of the cycle, the warm El Nino and cold La Niña events, manifest themselves primarily in variations (anomaly) around the average values of sea-surface temperature, as shown in Fig. 1.5 for the central Pacific region (5N-5S, 170-120 W). Also shown is the time series of the computed down-welling longwave radiation (DLR) anomaly using a radiation-transfer model and


FlG. 1.5. El Niño index based on the sea-surface temperature anomaly in the central Pacific region 5N-5S, 170-120 W (data from NOAA). Negative indices correspond to La Niña events. Also shown is the computed downwelling longwave radiation (DLR) anomaly for the region. (Pavlakis et al. 2007)

ISCCP cloud data and NCEP/NCAR water vapour and atmospheric temperature data. The DLR anomaly follows very closely the SST anomaly. During the El Niño there are large enhancements in precipitation in the central and eastern Pacific region, and reductions in the western Pacific. The SST is warmer over most of the Pacific region and pressure gradients are smaller, so the trade winds weaken. During the La Niña the signs of the anomalies are reversed, the water in the eastern Pacific region is colder than normal, the surface pressure and easterly trade winds increase, and rainfall in the central and eastern Pacific decreases as a result of the colder SST conditions. Indices of El Niño evolution have been used to monitor the strength of the events. El Niño has been quantified in terms of simple indices corresponding to times when the sea-surface temperature anomalies in central Pacific regions exceed some value. For example, El Nino events are correlated with SST anomalies in the (5N-5S, 170-120W) region when they exceed 0.4 °C, and are enough to produce perceptible impacts in Pacific Rim countries. However, such a definition does not discriminate between major, moderate, and minor El Nino events.

1.7.3 North Atlantic oscillation (NAO)

The North Atlantic Oscillation is a well-defined pressure pattern that consists of opposing variations of barometric pressure near Iceland and near the Azores. It has a strong influence via so-called 'teleconnections' on the climate of Europe and parts of Asia. Its pronounced influence on the climate of the Atlantic basin has been known for more than two centuries. Its positive phase is associated with a westerly current, between the Icelandic low-pressure area and the Azores high-pressure area, which carries cyclones and their associated frontal

NAO Index (December-March) 1864-2000

NAO Index (December-March) 1864-2000

1870 1880 1890 1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000

FlG. 1.6. Winter (December-March) index of the NAO based on the difference of normalized sea-level pressure (SLP) between Lisbon, Portugal, and Stykkishol-mur/Reykjavik, Iceland from 1864 through 2000. The indicated year corresponds to January (e.g. December 1949-March 1950). The average winter SLP data at each station were normalized by division of each seasonal pressure by the long-term mean (1864-1983) standard deviation. The heavy solid line represents the index smoothed to remove fluctuations with periods less than 4 years. (Source: Hurrell and Dickson 2005)

1870 1880 1890 1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000

FlG. 1.6. Winter (December-March) index of the NAO based on the difference of normalized sea-level pressure (SLP) between Lisbon, Portugal, and Stykkishol-mur/Reykjavik, Iceland from 1864 through 2000. The indicated year corresponds to January (e.g. December 1949-March 1950). The average winter SLP data at each station were normalized by division of each seasonal pressure by the long-term mean (1864-1983) standard deviation. The heavy solid line represents the index smoothed to remove fluctuations with periods less than 4 years. (Source: Hurrell and Dickson 2005)

systems towards Europe. However, the pressure difference between Iceland and the Azores fluctuates on timescales of days to decades, and can be reversed at times. The variability of NAO has considerable influence on regional climate variability in Europe, in particular in wintertime (Hurrell and Dickson 2005). Surface air temperature and sea-surface temperature across wide regions of the North Atlantic Ocean, North America, the Arctic, Eurasia, and the Mediterranean are significantly correlated with NAO variability (Fig. 1.6).

When the NAO index is positive, there are enhanced winter westerlies across the North Atlantic carrying relatively warm moist maritime air over Europe and across Asia. Also, stronger northerlies over Greenland and northeastern Canada carry cold air southward and decrease land temperatures and SST over the northwest Atlantic. Over North Africa and the Middle East there is cooling and warming of North America, associated with the stronger clockwise flow around the subtropical Atlantic high-pressure centre. Anomalously low precipitation rates occur over much of Greenland and the Canadian Arctic during high NAO index winters, as well as over much of central and southern Europe, the Mediterranean, and parts of the Middle East. In contrast, more precipitation than normal falls on Iceland through to Scandinavia.

FlG. 1.7. The Greenland Ice Sheet contains just over 7 m of sea-level equivalent. The contours represent 250 m differences in ice-sheet depth, with the highest contour at 3000 m. The ice sheet covers about 1.7 million km2 and represents about 10% of the world's total fresh water reserves. Greenland extends between 60-85 N and 70-20 W. ©Crown copyright (UK Met Office/Hadley Centre 2005)

FlG. 1.7. The Greenland Ice Sheet contains just over 7 m of sea-level equivalent. The contours represent 250 m differences in ice-sheet depth, with the highest contour at 3000 m. The ice sheet covers about 1.7 million km2 and represents about 10% of the world's total fresh water reserves. Greenland extends between 60-85 N and 70-20 W. ©Crown copyright (UK Met Office/Hadley Centre 2005)

1.7.4 Ice sheets

Global warming due to a strengthened greenhouse effect has the potential of melting important large ice sheets, such as the West Antarctic and Greenland Ice Sheets. The former contains ice that would, if melted, be equivalent to 6 m of global sea-level rise, while the latter contains 7 m sea-level equivalent. Much larger amounts of water are locked up as ice in Antarctica, but for this to become free would require a warming of more than 20 °C. The Greenland Ice Sheet, on the other hand, is expected to start to decrease in size as summer temperatures rise (Fig. 1.7). This was until recently thought to be a slow process, taking about 1000 years to decrease the ice by 50%, but evidence is accumulating that suggests it may be considerably faster, and already underway. A reduction in ice sheets, in snow cover, and in sea ice would alter the Earth's surface albedo, leading to enhanced global warming.

1.8 The land surface and biosphere

1.8.1 Land-surface albedo

The plant life covering the land surface affects the moisture content of the soil and aquifers and the surface albedo. Deforestation and desertification lead to a lower water-holding capacity of the surface, to spectral reflectance increases, to increases in surface temperature and hence increased thermal emission to space. Large-scale deforestation in the humid tropics, e.g. South America (particularly Amazonia), Africa, and Southeast Asia has been identified as the most important ongoing land-surface process.

Desertification in subtropical regions has led to land degradation and to the reduction of groundwater levels in aquifers that in turn have led to a cessation of river discharge to the sea, thus affecting marine ecosystems. If the depletion of soil moisture is severe, plant root zones can separate from the deeper water sources, such as aquifers, thus initiating processes of desertification. This is especially acute in regions with distinct wet and dry seasons (e.g. Mediterranean, tropical Australia). The driving force of land-surface processes is the radiative energy input. Temperature gradients that are created in the soil and between the surface and the atmosphere control water vapour fluxes. Photosynthesis is controlled by the visible radiation and through it biological activity and the carbon cycle. The albedo of bare soils is relatively low for moist fertile soils rich in organic matter, while degraded soils become brighter as in deserts. The albedo of desert soils then depends on the mineral components, the brightest being quartz, gypsum and salt deposits. Contamination with iron oxides makes the surface appear yellowish or red (e.g. Australian deserts). Eventually, winds can expose the underlying rock further modifying the land-surface albedo (e.g. Sahara). The magnitude of the land-surface albedo from the ultraviolet to the infra-red depends also on the vegetative canopy structure. Broad-leaf plants and low plants have high reflectance in the near infra-red, while tall and dense canopies are able to scatter light more efficiently, increasing solar-radiation absorption by the land surface. Reflectance by plants is low (below 10%) in the visible and higher in the near infra-red (as high as 60%).

1.8.2 Carbon dioxide sequestering

Carbon dioxide is sequestered from the atmosphere both by the oceans, the land surface and the biosphere. Carbon dioxide is very soluble in the oceans due to carbonate chemistry. The total dissolved CO2 in the oceans is about fifty times that in the atmosphere and the rate of uptake is limited by vertical mixing. However, in the upper oceanic layers where mixing occurs down to about 100 m due to winds, the total mass of inorganic carbon dissolved as CO2 is about equal to that in the atmosphere. This mixed layer plays a crucial role in the exchange of CO2 between the atmosphere and oceans on timescales of about 10

years. In the terrestrial biosphere, photosynthesis and the uptake of CO2 from the atmosphere provide the bulk of the organic carbon that is stored. On land, carbon dioxide is taken up by ecosystems, increasing with biological activity. This uptake is limited by the relatively small fraction of plant carbon that can enter long-term storage as wood and humus. On the other hand, according to the IPCC, the CO2 released by deforestation in the tropics between 1980 and 1990 was more than compensated by other terrestrial sinks. Such storage can range between 10 and 200 years. Geochemical cycles that lead to sedimentation of carbonates from weathering of volcanic rocks by reactions of CO2 dissolved in surface waters have characteristic timescales of thousands of years.

1.9 The climate record

Table 1.3 is a summary of the changes to the climate system that have occurred in the twentieth century, as recorded by the IPCC. Most striking are the dramatic increases in the greenhouse gases CO2, CH4 and N2O that have occurred since pre-industrial times. While tropospheric ozone has also increased, the abundance of the same gas in the stratosphere has decreased, most significantly between 1970 and 2000. There has been a rise in global mean surface temperature of 0.6±0.2 °C over the twentieth century, accompanied by a mean sea level rise of about 1-2 mm. The Arctic sea-ice extent and thickness have declined, while El Niño events have become more frequent, persistent and intense, particularly during the last 20-30 years.

1.9.1 Temperature trends

Figure 1.8 shows the current best estimate of the change in global average surface temperature from 1861 to 2004, both in terms of interannual variability and the overall trend. The data show an upward trend from 1920 to 1940, followed by a cooling until 1975, and a marked rise since then, with the warmest years in the period since 1981. The measurements are taken from thousands of weather stations over the globe, on land, ships, buoys, and satellites. Clearly, the data tends to be less reliable the earlier it gets; particular care is needed in using the data before 1920.

1.9.2 Sea-ice extent

Figure 1.9 shows the decrease in Arctic sea-ice extent since 1970, amounting to about 1 million km2 or about 2.5% per decade. Trends in sea-ice thickness are more difficult to measure but are estimated at as much as 40% in the Arctic. In Antarctica, there appears to have been no significant long-term trend in sea-ice extent.

Table 1.3 Changes in the twentieth century to the climate system. (Source: IPCC 2001)


Observed Changes

Atmospheric CO2

Terrestrial biospheric CO2 exchange

Atmospheric CH4

Atmospheric N2O

Tropospheric O3

Stratospheric O3

Global mean temperature

Northern Hemisphere surface temperature

Continental precipitation

Heavy precipitation events Frequency and severity of drought

Global mean sea level

Duration of river and lake ice cover

Arctic sea-ice extent and thickness

Non-polar glaciers Snow cover


El Nino events

280 ppm between 1000-1750 AD, 368 ppm in 2000 AD, 31 ±4% increase

A source of 30 GtC between 1800 and 2000; in 1990s a net sink of about 14 ±7 GtC 700 ppbv from 1000 to 1750; 1750 ppbv in 2000 AD, 151 ±25% increase

270 ppbv from 1000 to 1750; 316 ppbv in 2000 AD, 17 ±5% increase

Increase of 35 ±15% between 1750 and 2000; regional variations decreased between 1970 and 2000; varies with altitude and latitude

Increased by 0.6 ±0.2 °C over the twentieth century

Increase over the twentieth century greater than any other century in last 1000 years Increased by 5-10% over the twentieth century in Northern Hemisphere; decreased in North Africa and Mediterranean

Increased in mid and high northern latitudes Increase in parts of Asia and Africa in recent decades

Increased at a rate of 1-2 mm/year during twentieth century

Decrease by about 2 weeks in the twentieth century in mid to high northern latitudes Thinned by 40% in recent decades late summer-early autumn; decrease in extent by 10-15% since 1950s spring-summer

Widespread retreat during the twentieth century Decreased in extent by 10% as measured since 1960 by satellites

Thawed and degraded in parts of polar, subpolar regions

More frequent, persistent and intense the last 2030 years

1.9.3 Extreme events

'Extreme' weather events are those that deviate significantly from the mean, like heat waves, droughts and floods, particularly those with impacts that can be catastrophic to ecosystems, agriculture and society. According to the IPCC, extreme events (defined to be within the upper or lower ten percentiles) have

Global average near-surface temperatures 1861-2004

Global average near-surface temperatures 1861-2004

FlG. 1.8. Global average near-surface temperatures from 1861 to 2004, shown relative to the last decade of the nineteenth century. The observations are a combination of near-surface air and sea-surface temperatures, corrected to minimize errors in measurement practices and artifacts. Annual averages are shown as bars and the solid line shows the smoothed trend. ©Crown copyright (UK Met Office/Hadley Centre 2005)

Met Office Hadley Centre

FlG. 1.8. Global average near-surface temperatures from 1861 to 2004, shown relative to the last decade of the nineteenth century. The observations are a combination of near-surface air and sea-surface temperatures, corrected to minimize errors in measurement practices and artifacts. Annual averages are shown as bars and the solid line shows the smoothed trend. ©Crown copyright (UK Met Office/Hadley Centre 2005)

increased despite the fact that total precipitation has decreased or remained constant. Since most extreme weather involves precipitation, the implication is that precipitation events are less frequent but are heavier on the average. As with temperature rise, there is emerging evidence that the incidence of both severe drought and severe wetness was roughly constant during the first part of the twentieth century but has increased in recent decades, correlated with a shift towards more warm ENSO events. In many mid and high latitude regions, the day-to-day temperature variability has decreased, the daily minimum temperature has increased, and the freeze-free period is longer. Since 1950 there has been a significant reduction in unseasonably low temperatures across much of the globe, with a smaller increase in the frequency of abnormally high temperatures.

Figure 1.10 shows the change in the rate of occurrence of 3-day rainfall intensity events, which have been shown to be important precursors to flooding. Hadley Centre predictions suggest that man-made climate change will tend to intensify the water cycle, increasing the rainfall at higher latitudes, a trend that is consistent with recent observations of increases in the rate of discharge from Eurasian rivers into the Arctic.

FlG. 1.9. Change in Arctic sea-ice extent since 1970. In deriving the trend, the extent of sea ice was defined as the area within which the concentration of sea ice is greater than 15%. ©Crown copyright (UK Met Office/Hadley Centre 2005)

1.9.4 Sea-level trends

The longest record of direct measurements of sea level comes from tide gauges. The IPCC has estimated that the rate of global mean sea-level rise during the twentieth century was in the range 10-20 cm, and that this rise was greater than that during the nineteenth century. However, no significant acceleration in the rate of rise was detected during the course of the twentieth century. Satellite measurements over the past decade indicate a rise of 2.5 mm/year in the global mean, but with large regional variations.

Over the last ice age cycle, sea level ranged from 5 m higher than today's, at the time of the last interglacial about 120 thousand years ago, to 120 m below today's, at the depth of the last ice age 21000 years ago when glaciers were at their maximum extent. Figure 1.11 shows relative sea level from measurements dating back to 1700.

1.10 Projections of future climate

In its 2001 report, the IPCC assessed the predictions of global climate models for the period 1990-2100 to illustrate the possible changes to the climate expected under future greenhouse-gas concentrations and sulphate aerosol loadings. The calculations were made by a number of different models and research groups, and

FlG. 1.10. The change in maximum annual 3-day rainfall events in northern midlat-itudes from the average, over the period 1961 to 1990. (Source: IPCC 2001)


1700 1720 1740 1760 1780 1800 1820 1840 1860 1880 1900 1920 1940 1960 1980 2000


1700 1720 1740 1760 1780 1800 1820 1840 1860 1880 1900 1920 1940 1960 1980 2000

FlG. 1.11. Relative sea level since 1970, from stations in Northern Europe. The scale bar on the left indicates a change of ±100 mm. (Source: IPCC 2001)

based on a range of scenarios for projected changes in atmospheric composition. All are based on runs or integrations of atmosphere-ocean general circulation models (AOGCMs), the results of which apply to spatial scales of hundreds of kilometres and larger (see Chapter 11). Some of the simulations also included the effects of ozone and/or the indirect effects of aerosols, but most did not attempt to include the less well understood effects arising from land-use changes, mineral dust, black carbon, etc. None of the simulations included estimates of changes in the radiation from the Sun, or of the effect of future volcanic eruptions on atmospheric aerosol concentrations.

1.10.1 Emission scenarios and global warming

As part of its work the IPCC produced a Special Report on Emissions Scenarios, known as SRES, which made a range of projections of the greenhouse-gas emissions expected from human activity in the next 100 years. These ranged from optimistic (controlled population levels, use of new, cleaner technology and so forth) to pessimistic ('business as usual'). The following is an account of their main conclusions and recommendations. Greenhouse gases Fossil-fuel burning is expected to remain the dominant influence on trends in atmospheric CO2 concentration during the twenty-first century. As the amount in the atmosphere increases, the ocean and the land will take up a decreasing fraction of anthropogenic CO2 emissions and the net effect will be to further increase the atmospheric CO2 concentration. By 2100, carbon-cycle models project a CO2 mixing ratio of 540 to 970 ppm, 90 to 250% above the value of 280 ppm in the year 1750. Since uncertainties, especially about the feedback from the biosphere, introduce a variation of about -10 to +30% around each scenario, the total predicted range is 490 to 1260 ppm (75 to 350% above the 1750 concentration).

Changing land use could influence the atmospheric CO2 concentration. If all of the carbon released by historical land-use changes could be restored to the terrestrial biosphere over the course of the century (e.g. by reforestation), the CO2 concentration would be reduced by 40 to 70 ppm. Without this, stabilization of atmospheric CO2 concentrations at 450, 650 or 1000 ppm would require global anthropogenic CO2 emissions to drop below 1990 levels, within a few decades, about a century, or about two centuries, respectively, and continue to decrease steadily thereafter. Eventually CO2 emissions would need to decline to a very small fraction of current emissions.

Estimates of the future concentrations of the non-CO2 greenhouse gases vary considerably across the SRES illustrative scenarios, with CH4 changing by +190 to +1970 ppb from a present concentration of 1760 ppb, N2O changing by +38 to +144 ppb (present concentration 316 ppb), total tropospheric O3 changing by -12 to +62%, and a wide range of changes in concentrations of HFCs, PFCs and SF6. In some scenarios, total tropospheric O3 would become as important a radiative forcing agent as CH4. Aerosols The SRES scenarios include the possibility of either increases or decreases in anthropogenic aerosols (e.g. sulphate aerosols, biomass aerosols, black and organic carbon aerosols) depending on the extent of fossil-fuel use and the success or failure of policies to abate polluting emissions. In addition, natural aerosols (e.g. sea salt, dust and emissions leading to the production of sulphate and carbon aerosols) are projected to increase as a result of changes in climate. Marker scenarios The IPCC selected four main 'marker' emission scenarios, illustrated in Fig. 1.12. Scenario A1 (high emissions) describes a future world of very rapid economic growth, a global population that peaks in midcentury and declines thereafter, and the rapid introduction of new and more efficient technologies. A1FI is a variant of this which remains fossil-fuel intensive. A2 (medium-high emissions) describes a very heterogeneous world, with an underlying theme of self-reliance and preservation of local economic and technological identities. In B2 (medium-low emission) the emphasis is on local solutions to economic, social and environmental sustainability, a continuously increasing global population, intermediate levels of economic development, and less rapid and more diverse technological change than in the B1 and A1 storylines. B1 (low emissions) describes a convergent world with the same global population, that peaks in midcentury and declines thereafter, as in the A1 storyline, but with rapid change in economic structures toward a service and information economy, with reductions in material intensity and the introduction of clean and resource-efficient technologies. IPCC stresses that it is not possible to attach probabilities to each scenario.

1.10.2 Climate projections for the twenty-first century

The IPCC climate projections are meant to give the breadth of the impacts of increasing greenhouse emissions over the next century, over a realistic range of emission scenarios and models. Although models continue to improve, at present the choice of model is found to make a difference comparable to the choice of scenario. Of course, many uncertainties remain; our understanding of the physical processes that govern climate and our ability to simulate these highly non-linear processes are imperfect and some of the most important determinants of climate, such as clouds and aerosols and their radiative properties, are certainly inadequately simulated. The fact that the models do not show a consensus over future changes in interannual variability as a result of the El Nino-Southern Oscillation (ENSO) points to other limitations in the current generation of models. Nevertheless, a number of key conclusions emerged from the IPCC study that, if not totally robust, certainly require attention and precautionary or preventative action. Projections of temperature rise In general, the troposphere and surface warms, while the stratosphere cools. The land warms faster than the ocean, and there is greater relative warming at high latitudes. The increase in mean

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