This page presents a brief introduction to the Earth's climate. It begins with the global energy budget, and then describes the Earth's circulation as a response to this. It will treat the climate system as a heat pump that takes heat from the tropics and pumps it towards the poles. The introduction is divided into the following sections:

1. Energy Budget
2. Global Atmospheric Circulation
3. Oceanic Circulation
4. Atmosphere Ocean Interaction
5. Greenhouse Effect

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For any balanced budget, what comes in must equal what goes out. In the case of planets orbiting the Sun, this means that the incoming solar radiation must equal the outgoing radiation. Otherwise, the planet will either get hotter or cooler. Balancing the global energy budget is a fundamental aspect of the climate system.

Consider a beam of sunlight hitting the Earth at the equator as shown in Figure 1. The beam is approximately at right angles to the Earth's surface, so the amount of Earth it is spread over (area a), is the same as the width of the beam. Closer to the poles, a beam of the same width covers a much bigger amount of the Earth (area b), because it arrives at a different angle to the Earth's surface. This means that the surface of the Earth receives more energy in the tropics per unit area than it does at the poles. In a similar way, at midday, when the sun is highest in the sky, a sunbeam of a certain width illuminates an area smaller than the same sunbeam would illuminate at dawn or dusk, when the sun is lower in the sky. The sun therefore feels hottest at midday.

Figure 1. Solar radiation arrives at a different angle to the Earth's surface at the poles than at the equator. In consequence, the area that the beam covers is smallest at the equator and gets larger towards the poles (area b is bigger than area a), so the Earth receives more heat per unit area at the equator than it does at the poles.

Some of the incoming solar radiation (which is mainly ultraviolet, visible light and short wavelength infrared) is reflected or scattered directly back into space by the atmosphere and some is absorbed by the Earth (see Figure 2). Once the radiation is absorbed, the Earth's surface re-emits this energy at a longer wavelength in the form of thermal radiation - heat.

Figure 2. The Earth's annual radiation budget. The numbers are all in W/m² (Watts per square meter), a measure of energy. Of the incoming radiation, 49% (168÷342) is absorbed by the Earth's surface. That heat is returned to the atmosphere in a variety of forms (evaporation processes and thermal radiation, for example). Most of this back-scattered heat is absorbed by the atmosphere, which then re-emits it both up and down. Some is lost to space, and some stays in the Earth's climate system. This is what drives the Greenhouse Effect [Figure adapted from Kiehl & Trenberth, 1997].

Figure 3 shows how the distribution of incoming solar ( shortwave) and outgoing (longwave) terrestrial radiation varies with latitude (distance from the equator). The tropics are net absorbers of energy as the amount of absorbed solar energy is greater than the amount of outgoing terrestrial radiation. Conversely, the poles are net emitters. This should mean that the tropics are constantly heating up and the poles cooling down, but they're not. The Earth must therefore continuously pump heat from the tropics to the poles: it is a heat engine. The circulation of the Earth's atmosphere and oceans is the dominant pumping mechanism. They carry approximately equal amounts of energy from the equator to the poles.

Figure 3. Shortwave radiation (from the Sun) and longwave radiation (heat emitted by the Earth) vary with latitude. The difference between the two shows that the Earth is a net absorber of energy (i.e. absorbed energy > outgoing energy) in the tropics, and a net emitter (outgoing energy > absorbed energy) in the polar regions. This is a plot of zonal mean radiation; that is, it shows how the radiation varies with latitude but not longitude. If you imagine a circle around the globe at each latitude, the radiation has been averaged around the circle, because in this case the variation with longitude is less interesting than the variation with latitude.

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The circulation of the atmosphere is responsible for about 50% of the transport of energy from the tropics to the poles. The basic mechanism is very simple: hot air rises in the tropics (convection), reducing the pressure at the surface and increasing it higher up. This forces the air to spread away polewards at high levels, and to be drawn in at low levels. As the warm, polewards moving air comes into regions with less incoming solar radiation, it cools and sinks, thus completing the circulation.

If the Earth were not rotating, we would see this very simple pattern: hot air would rise in the tropics, move away from the equator, gradually cool, sink at high latitudes near the poles, and finally re-circulate near the surface towards the equator (see Figure 4).

Figure 4. If the Earth were not rotating tropical air would rise, travel towards the pole, cool and sink before returning to the equator. The dominant flow at all heights would be along lines of longitude.

However, the Earth's rotation complicates things. For a given point on the Earth's surface to do a full rotation, it has to travel a lot further at the equator (2 π times the radius of the Earth; i.e. 6,371 km), than at mid-latitudes, and at the poles it doesn't have to travel at all, just rotate. Speed is defined as a distance divided by time, so for a full rotation in 24 hours, this means that the speed of the surface of the Earth is greatest at the equator, and falls with increasing latitude.

Now, imagine a cannonball fired towards North from the Equator. In addition to its northward speed, the cannonball also has the same easterly speed as the Earth from which it was launched. But, as it travels further North, the Earth underneath it is moving slower than the Earth at the equator was. So the cannonball appears to drift to the east in flight (as shown in Figure 5 below). This is called the Coriolis effect or force after the French engineer Gustave-Gaspard Coriolis who discovered it (cannonballs and all) in 1835. The Coriolis force is the reason why the upper level, polewards travelling air (wind!) is westerly (west to east); whereas the equatorwards travelling, surface winds are easterly (east to west). In the tropics, these easterlies are known as the trade winds.

Figure 5. Animation showing the Coriolis effect. Keep your eye on the white spot (Santa, perhaps), which moves in a straight line up the page according to a stationary viewer above the North Pole (you, in this case). But from the perspective of a viewer on the (rotating) Earth, the spot appears to deflect to the right.

In the case of the atmosphere, this means that winds travelling polewards get a bigger and bigger westerly speed. This peaks in the sub-tropical jet streams where air speeds are typically 40m/s in the upper troposphere. With such large vertical velocity gradients, the air becomes unstable, and waves develop in the westerly flow. We experience these as the low pressure systems which regularly pass over the North of Europe. These systems mix the air, which results in the transport of cold air equatorwards and of warm air polewards. Their net effect is the transport of heat polewards and they set up the so-called Ferrell cell, which rotates in the opposite sense to the Hadley Cell.

In the Polar regions, the circulation pattern is very similar to the Hadley Cell and is called the Polar Cell. It is driven by the ascent of warmer air and the descent of colder air. The mid-latitude jet stream is found in the upper troposphere where the Ferrell and Polar cells meet.

Figure 6. Idealisation version of the Earth's atmospheric circulation. Air is heated and rises in the tropics, drifting north before sinking around 30° North and South in the Hadley Cell circulation. At the surface, air flow is easterly and is known as the trade winds. In the mid-latitudes (about 30°-60°) the circulation is dominated by large-scale wave activity and extra-tropical storms (the reason European weather is so unsettled!), which manifests itself as the Ferrell Cell. At high latitudes the simple, convection driven circulation returns and is called the Polar Cell. Regions of high and low surface pressure are marked ‘H’ and ‘L’, respectively.

You can read more about the circulation of the atmosphere in:

• Clouds and Rain
• The Diurnal Cycle
• The Seasonal Cycle
• Land and Sea
• The Tropics
• The Mid-Latitudes
• Polar Regions

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The circulation of the oceans is responsible for about 50% of the transport of energy from the tropics to the poles. As in the atmosphere, the circulation is driven by the heating of surface waters in the tropics, and cooling at the poles. Cold surface currents travel equatorwards and warm surface currents travel polewards. The full, large scale circulation pattern of the ocean is called the thermohaline circulation (see Figure 15), because it is driven by differences both in temperature and salt concentrations. When water evaporates or freezes, it leaves behind its salt, making the remaining water more saline and therefore more dense. The North Atlantic Deep Water, for example, is formed by water in the Greenland Sea, which is both very cold and very salty, and therefore sinks and spreads equatorwards.

This all means that the three-dimensional structure of the ocean is very complicated, and, as yet, relatively little is known about it. Figure 7 shows a highly simplified view of how the world's oceans circulate. It would take any individual water molecule about a thousand years to do a complete circuit!

Figure 7. The global oceanic circulation (highly simplified). Also known as the oceanic conveyor. Source: www.CLIVAR.org. The red part of the conveyor represents the net transport of warm water in the top 1000m or so of the oceans, and the blue part the net transport of cold water below the thermocline.

The ocean has a greater capacity for storing heat than the atmosphere, which means it reacts slower than the atmosphere to changes in the balance in incoming/ outgoing radiation. This means that ocean temperatures change more slowly that atmospheric ones, whether this is on a diurnal, seasonal or climate time scale.

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The oceans and atmosphere interact in many different ways. There can be a net exchange of heat, salt, water and momentum between them.

When wind blows over the ocean, energy is transferred from the wind (slowing it down) to the surface layers, some of which then drives ocean currents. Water can evaporate more easily into warm air, especially if it is windy. As it evaporates, it removes heat from the ocean. If it then condenses to form a cloud droplet, it releases the heat into the air. This is one of the main ways the hurricanes get their energy.

Salt is continuously brought into the oceans by the rivers draining off the continents, which carry minerals dissolved from the rocks they run over, and deposited as sediment on the ocean floor. Water evaporating or freezing at the oceans' surface leaves the remaining water saltier, but rain, which is not salty, dilutes the salt concentration of the surface ocean. In addition, when it's really windy, salty droplets of ocean water can be blown into the air, and these can form the basis of cloud droplets.

The air and the ocean are continuously exchanging heat. As the ocean has a higher heat capacity, it takes longer to adjust to changes in incoming radiation, and therefore tends to change temperature slower. This means that the surface of the sea is usually a different temperature to the air immediately above it, and heat is transfered between the ocean and the atmosphere.

There are many feedback mechanisms between the oceans and the atmospheres. For example, evaporating water can condense in the atmosphere to form clouds. These reflect both incoming and outgoing radiation (which is why cloudy nights feel warmer than clear ones) and so determine the temperature of the ocean surface.

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In the nineteenth century various scientists (such as Joseph Fourier) explained that the atmosphere can, like an ordinary greenhouse, retain energy radiated into it from outside. The greenhouse analogy isn't very exact, but the name certainly stuck.

In the 1860s John Tyndall explained that certain gases, including water vapour and carbon dioxide (CO₂), don't affect visible light but absorb longer wavelength radiation (infrared, heat). He suggested that these gases insulate the Earth.

The actual process works like this (see figure 2): visible incoming sunlight either gets reflected (for example by clouds, or aeroplanes), or passes unhindered through the atmosphere, and gets absorbed by the surface of the Earth, thus heating it. The Earth radiates heat from the surface back into the atmosphere, where it can pass back into space, or get reflected again, or, because it has now got a longer wavelength than before, it can get absorbed by the water vapour, carbon dioxide, methane and other greenhouse gases which are present in the atmosphere. As the water vapour/ methane/ carbon dioxide molecules absorb the longwave radiation, they heat up, and in turn re-radiate long wave radiation in all directions. Some is lost to space, but some of it also gets radiated back to the surface, again warming it.

This naturally occurring process helps keep the Earth warm enough for liquid water to exist. Without greenhouse gases, the average temperature at the Earth's surface would reach only -17ºC, approximately 33ºC colder than it actually is!

Now, what if the concentrations of these insulating gases increase? We might expect the process described above to intensify. In fact, this is just what the Nobel Prize-winning Swedish chemist Svante Arrhenius did in 1896. By knowing how CO₂ absorbs heat radiation from the surface of the Earth, he calculated what would happen if the amount of CO₂ in the atmosphere were doubled. He estimated that a doubling of CO₂ would lead to an average global surface temperature increase of 2 °C. This is consistent with modern predictions.

This approach, while still a handy first guess, considers the climate system in the absence of any feedback processes. Feedbacks are processes in which outputs from the process have an effect on the inputs to that same process. Sometimes feedback processes act to offset or inhibit a change (negative feedback), and sometimes they act to amplify a change (positive feedback). Examples of negative feedback include the maintenance of your body temperature: when you get too warm, you trigger various mechanisms (e.g. perspiration) to cool you down and vice versa. A common example of positive feedback is often associated with amplified music or speech, when the microphone is placed too close to a loudspeaker... someone speaks/ sings/ plays into the microphone, the noise is amplified, and comes out of the speaker. If some of this amplified noise goes back into the microphone, it gets amplified again etc. etc. and the end result is a deafening whine.

There are many examples of feedbacks in the climate system. If the atmosphere gets warmer, ice melts. Ice reflects a lot of incoming solar radiation, so if it melts, less gets reflected, more gets absorbed by the Earth and the atmosphere gets warmer; a positive feedback. On the other hand, if there is more carbon dioxide in the atmosphere, some plants grow faster, absorbing more carbon dioxide and eventually reducing its amount in the atmosphere; a negative feedback.

Because of the complexity of the climate system, due to the presence of feedbacks within it, we need to try to represent the whole system as thoroughly as possible in order to simulate the likely changes. We need to be able to understand how and where feedbacks act, and how large they are.

You can read more about the possible feedbacks from increasing carbon dioxide in:

• Carbon Dioxide and Climate Change

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Clouds and precipitation (rain or snow, depending on where you are) form mainly where warm, humid air is forced to rise. As it rises, it expands and cools and the amount of water vapour that can be held is reduced. Any surplus water vapour condenses and forms droplets, which we see as clouds and may become large enough to fall to the Earth's surface as rain or snow. Figure 8 shows that the main regions of ascent are in the tropics and in mid-latitudes (around 50°-60°), while the main regions of descent, the sub-tropics and poles, are dry, arid regions.

Figure 8. Satellite picture of clouds at 1830 GMT on 22/4/2003. Copyright 2003 EUMETSAT.

In the tropics, the ascent is vigorous and huge cumulonimbus clouds (thunder clouds) form, reaching over 10km high and frequently grouping into clusters so that, in Figure 8, the tropics are clearly marked out by a long line of clouds. They form preferentially over the oceans, where there is a source of warm water to evaporate. However, in the mid-latitudes, ascent is more localised and less deep, resulting in more shallow, individual cumulus.

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If you watch Figure 9, you can see the daily pattern of heating and cooling. The hottest part of the land surface moves west during the day. You can also see that the ocean surface temperature doesn't vary anywhere near as much during the day as the land (especially in the centre of a large continent, such as Africa), as it takes longer to heat up/ cool down.

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The seasonal cycle in the atmosphere is driven by the fact that the Earth's axis is not at right angles to the sun (it is actually 23° away from perpendicular ). This means that, at different times of year, different latitudes get the most incoming solar radiation. At the equinoxes, the sun is overhead at the equator, at the June solstice, the sun is over the Tropic of Cancer and at the December solstice, it is over the Tropic of Capricorn. This means that, in June, July and August (northern hemisphere summer), the northern hemisphere is warmer than the southern hemisphere. Similarly in December, January and February, the southern hemisphere is warmer. These months are not symmetrical about the solstice (for example, we do not talk about the November, December, January season) because the climate system tends to lag the sun: it takes a while to heat up or cool down.

The seasonal cycle has many effects on the climate. In Figure 10, you should be able to see the ITCZ shift northwards and southwards with the seasons. The whole associated general atmospheric circulation pattern shifts with it. There are smaller scale effects as well, for example the monsoons. Tropical cyclones can only be found when the oceans are warmest, at the end of the summer season.

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The distribution of land and sea distorts the simple picture of global circulation; land heats up and cools down faster than water, leading on a large scale to the Asian monsoon but on a smaller scale to sea breezes, a common phenomenon at the coast, where winds blow from the sea during the day, but from the land during the night. The presence of continents which break up the ocean obviously disrupts the ocean circulation. The presence of mountain ranges deflects the atmospheric flow (for example, the Himalayas affect the monsoon pattern), while patterns of precipitation are determined to a large extent by land-sea contrasts, continental land masses, mountain ranges and so on.

There is a lot less land in the southern hemisphere than in the northern hemisphere, so the atmospheric circulation is a lot simpler. For example, the storm tracks are more continuous around the Earth.

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The Tropics, defined as the region between the tropics of Cancer and Capricorn, have a climate dominated by the large scale convection associated with the Inter-Tropical Convergence Zone (the ITCZ). This moves with the seasons, according to which latitude is closest to the sun. At the equinoxes, the sun is closest to the equator, at the December solstice, the sun is over the Tropic of Capricorn and at the June solstice, it is over the Tropic of Cancer. The most rapid ascent of hot air, associated with the formation of towering cumulonimbus clouds, is found at the ITCZ. These are often the source of the heavy rains and violent thunderstorms of the tropics.

Tropical climates are normally hot and humid and usually show far less seasonality in temperature than extra-tropical climates. On the other hand, other climate features, such as rainfall and wind patterns, can shown pronounced regularity, such as the monsoons.

The most dramatic weather systems found in the tropics are tropical cyclones: called hurricanes in the Atlantic, Caribbean and eastern Pacific, cyclones in the Indian Ocean and typhoons in the western North Pacific. They are low pressure systems, typically 200-2,000 km across, with wind speeds greater than 120 km/hour. They consist of deep cumulonimbus clouds, up to 12 km high, spiralling around a central, clear eye where air is descending. They form over warm tropical oceans, but cannot form equatorwards of 5°, as the Coriolis force is too weak. They rapidly decay when they move over land and are cut off from their source of warm water.

Figure 11. Hurricane Lili, 2/10/2002 in the Gulf of Mexico.

The model we are using isn't great at producing hurricanes, mostly because the grid is too coarse for the relevant processes to operate.

The monsoon is another important feature of the tropical climate, and is a result of land/sea differences and the seasons. Continental land masses cool down and heat up faster than oceans because their thermal heat capacity is lower. This means that, in winter, the air above the continents is colder than the air above the oceans. The same processes which cause the large scale atmospheric circulation then operate, and there is ascent over the oceans, descent over the continents and surface level flow from the continents to the oceans. In the summer the reverse happens. The seasonal reversing winds are called the monsoon (derived from the Arabic word for season, mausim), and most affect the Indian Ocean and western tropical Pacific. The monsoon pattern interacts with the large scale atmospheric circulation and is affected by the orography (the shape of the land surface, for example the Himalayas), which together produces a complicated weather pattern in south-west Asia.

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The mid-latitudes (roughly 30°-60°) are dominated by the weather systems that form when the Hadley Cell becomes unstable and breaks down into a series of alternating low and high pressure systems. The regions where these systems are concentrated are known as storm tracks. In the Southern Hemisphere, where there is little land, the storm track is fairly continuous around the Earth. However, in the Northern Hemisphere, storm tracks are only seen over the oceans. This is because friction is much greater over the uneven surface of the land, and slows down any winds blowing over it. Figure 12 shows the Northern Hemisphere storm tracks.

Figure 12. The storm tracks in the Northern Hemisphere in winter (December, January & February average between 1979-1997). The quantity shown is a measure of the kinetic energy in the air associated with the storms. Data from the ERA15 observational data set.

The climate in mid-latitudes is highly seasonal, being warmest when the Sun is highest in the sky (see also Figure 1) at the summer solstice. It is also governed by patterns in land and sea. Britain is warmer than most places at a comparable latitude thanks to the energy transported polewards by the North Atlantic Current and the westerly winds. It also has a much smaller seasonal cycle than, say, Siberia, because it is surrounded by water which reacts slowly to changes in the incoming solar radiation.

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Polar regions get the least incoming solar radiation. Within the Arctic/ Antarctic circles, there is complete darkness for some of the year. The ice caps themselves have important feedbacks on the climate system as they are highly reflective; incoming solar radiation is reflected back into space before it is absorbed (see Figure 2). The freezing of water around the polar ice caps is a very important mechanism driving the thermo-haline circulation.

There is very little rain or snow in polar regions, due to the predominant descent of air (see Figure 6). In the winter, it is permanently dark and very cold. In the summer, it is permanently light and not quite so cold!

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The climate of the Earth is constantly changing, in response to changes in the incoming solar radiation, the patterns of the continents, the amount of dust in the atmosphere, the chemical composition of the atmosphere and many other factors.

One of the factors which is thought to affect surface temperatures is the atmospheric concentration of carbon dioxide. Carbon dioxide is a ‘greenhouse gas’. This means that it does not reflect much incoming solar radiation, but it does strongly absorb outgoing, long wave, thermal radiation, re-emitting it back towards the surface and warming the atmosphere.

Atmospheric concentrations of carbon dioxide have been increasing in the past 200 years or so since the Industrial Revolution began. The source is mainly the burning of fossil fuels (coal, oil and gas) - for transport, industry, electricity or heat. The rest is due to land use change, such as deforestation. Figure 13 shows the atmospheric concentration of carbon dioxide in the past 1000 years (data have come from ice cores, direct measurements in recent years etc., if you're interested in this, read ‘The two-mile time machine’ by Richard B. Alley) and various estimates of how carbon dioxide concentrations will behave in the next 100 years, depending on how we react to legislation on carbon emissions. The concentrations used in the standard and doubled CO₂ experiments of climateprediction.net are marked.

Figure 13. The global atmospheric concentration of CO₂ in parts per million (ppm) (left) measured over the past 1000 years and estimated for the next 100 years (right). Source: IPCC Third Assessment Report. The CO₂ concentrations used in climateprediction.net experiments; 282ppm and 564ppm, are marked.

Scientists are still uncertain exactly how the Earth-climate system will respond to such changes in carbon dioxide and other changes to the composition of the atmosphere.

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Figure 14. The Structure of the Atmosphere. Air rises in the Inter Tropical Convergence Zone (ITCZ) to the top of troposphere, then spreads polewards. Jet streams are found just below the tropopause (the interface between the troposphere and stratosphere) where the Hadley and Ferrell, or Ferrell and Polar cells converge. The regions of surface high and low pressure are shown. In the stratosphere, the air is much more stable and less well mixed than the troposphere.

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Figure 15. Simplified Model of the Thermohaline circulation in the Oceans. Below the well-mixed, warm surface layer there is a theromocline; a zone in which temperature falls rapidly with depth. Below that is the cold, stable, deep ocean.

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