The atmosphere

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The atmosphere is a largely homogeneous mixture of gases, both horizontally and vertically, over the height range important for climate: namely the troposphere and stratosphere (Fig. 1.3). The composition of this apparently stable mixture, air, is shown in Table 1.1. The balance of the dominant constituents of air is thought to have evolved considerably over the lifetime of the planet; for instance, oxygen is likely to have been a product, rather than a necessity, of life (the Gaia hypothesis). The climate system's immense natural variability will be a recurring theme of our discussion throughout the book.

The temperature of the atmosphere varies strongly both in the vertical and with latitude. The latter is due to an imbalance in the radiation received over the Earth's surface throughout the year due to the planet's orbit and obliquity (Fig. 1.4 and §1.8). The circulation of both the atmosphere and ocean are ultimately derived from this energy imbalance; they act to counter it, in the ratio of about 3:2 respectively.

The vertical temperature distribution shown in Fig. 1.3 comes about because the atmosphere is basically heated from two sources: the ground and the upper

2 Note that this average planetary albedo assumes that the hypothetical, sterile, Earth has the same net reflectivity as the real Earth-atmosphere system. Thus this albedo is not the surface reflectivity (see §1.4).

Table 1.1. The major constituents of the atmosphere

Gaseous constituent Molecular form Proportion (%)

Nitrogen N2 78.1

Oxygen O2 20.9

Argon Ar 0.93

Water vapour H2O variable: 0.1-1

Carbon dioxide CO2 0.037

Methane CH4 0.000175

Nitric oxide N2O 0.000032

Ozone O3 variable c. 0.000005

Fig. 1.3. Zonal mean vertical profile of temperature during June at 45°N.

Fig. 1.3. Zonal mean vertical profile of temperature during June at 45°N.

stratosphere (although we will see in the next sub-section that this is a significant simplification). The ground (or ocean surface) is a heat source since some 49% of the incoming solar radiation is absorbed there. There is also an important heat source between 30 and 50 km above the ground, in the ozone layer. When ozone, O3, absorbs a photon of ultra-violet light - denoted h v because this is the energy of a photon of frequency v - the molecule dissociates in the process to form an oxygen molecule and an energetic oxygen atom, O1D, where one electron is displaced into a higher energy state than in the ordinary oxygen atom (see Appendix B for the Periodic Table of the Elements and a discussion of electron orbitals). This can then react with an oxygen molecule to reform ozone as part of the Chapman cycle:

The air molecule, M (that is, predominantly N2 or O2), is necessary in the second reaction in (1.4) as the reaction produces excess energy. This is carried

Fig. 1.4. Contour plot of daily average insolation at the top of the atmosphere as a function of season and latitude. The contour interval is 50 Wm-2. The heavy broken line indicates the latitude of the sub-solar point at noon. [Fig. 2.6 of Hartmann (1994), Global Physical Climatology. Reprinted with permission from Academic Press.]

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec away by M thereby stabilizing O3, which would otherwise dissociate. Such a reaction is called exothermic. Hence the air becomes warmer, as the temperature of a medium is merely a reflection of the average kinetic energy of its molecules.

The reactions in equation (1.4) are only part of the full Chapman cycle, which also contains reactions involving photo-dissociation of O2, and reactions between the excited oxygen atoms themselves and O2. Some of these are likewise exothermic, adding to the energy which is transferred, via chemical reaction, from solar radiation to the middle atmosphere. There are many other reactions involving ozone, some of which will be discussed in §7.2.1.

The lower atmosphere is therefore heated both from above and from below. Between these regions is a zone, in the lower stratosphere, where the energy from these heated regions only weakly penetrates. This is strongly stratified, which means that there are large vertical gradients in the concentrations of trace constituents of the air and the potential temperature (see Appendix C). The tropopause, at the bottom of the stratosphere where the gradients are greatest, resists penetration by cloud convection, or even diffusion. The well-mixed region below this, the troposphere, is the part of the atmosphere that we will be largely concerned with, because of its direct interaction with the oceans.

The strong heating of the surface at the equator (Fig. 1.4) makes the air less dense, forcing it to rise. Air flows towards this region of rising air, which tends to be concentrated in a narrow band around the globe known as the InterTropical Convergence Zone, or ITCZ. Aloft, the rising air moves polewards to compensate for the surface flow. In the late seventeenth century, when Halley first proposed this mechanism for driving the atmospheric circulation (modified 50 years later by Hadley) it was believed that this Hadley cell extended to the polar regions. This seemed logical, as polar air is cold, and so relatively dense, and should therefore flow towards the low pressure regions of the tropics in order to transfer heat from the equator to the poles and so maintain the Earth's thermal equilibrium. By the nineteenth century this idea was seen to be too simplistic. The Coriolis force, due to the solid Earth and moving atmosphere revolving at slightly different rates (see §2.5.2), gives this converging near-surface wind a

Fig. 1.4. Contour plot of daily average insolation at the top of the atmosphere as a function of season and latitude. The contour interval is 50 Wm-2. The heavy broken line indicates the latitude of the sub-solar point at noon. [Fig. 2.6 of Hartmann (1994), Global Physical Climatology. Reprinted with permission from Academic Press.]

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov Dec

Fig. 1.5. Schematic

Hadley cell cross-section of the zonal mean circulation in the

Hadley cell troposphere. The dotted upper constant feature. The crosses (and E) show where the surface flow has an easterly component, while W shows

troposphere. The dotted upper region of the Ferrel cell indicates that it is a less constant feature. The crosses (and E) show where the surface flow has an easterly component, while W shows

Jet stream

—N Hadley \ cell where the surface flow is predominantly westerly.

westward component, resulting in the observed easterly3 Trade winds. Ferrel therefore proposed an intermediate 'Ferrel cell' in mid-latitudes. Modern observations support this, as shown schematically in Fig. 1.5, a zonal cross-section of the tropospheric flow.

The general circulation of the lower troposphere is shown in Fig. 1.6, and the sea level pressure field for northern winter in Fig. 1.7. The ascending air of the equatorial region is shown by the low pressure. To replace this, easterly winds flow equatorwards driven by high pressure in the sub-tropics, where the air in the tropical Hadley cell descends (Fig. 1.5). This latter, relatively calm, zone has strong westerly winds on its poleward side, which, in turn, lie equatorward of another region of low pressure near 60° of latitude. This region of sub-polar low pressure forms the ascending branch of the polar Hadley cell of Fig. 1.5, with easterly winds at the surface due to polar high pressure. The most vigorous part of this system is where the tropical Hadley cell meets the mid-latitude Ferrel cell. Here the upper level convergence of air produces an extremely strong westerly jet-stream in the upper troposphere (Fig. 1.5). This often has a secondary maximum over the mid-latitudes, above the polar front (Fig. 1.5). This polar front jet-stream steers the transient pressure systems that we experience on the ground in the mid-latitudes. The latter systems are a significant mechanism in the redistribution of heat from equator to pole.

The mainly zonally symmetric structure of the general circulation is mostly due to the latitudinal distribution of the solar radiation received by the Earth; the distribution of land and sea over the Earth's surface distorts the zonality. Some aspects of this latter interaction will be discussed in Chapters 2 and 5.

1.2.1 The greenhouse effect

The vertical profile in Fig. 1.3 shows a decline in temperature of about 6.5°C per kilometre in the troposphere. It can be shown that ascending 'dry' air, i.e. air without clouds, changes temperature because of expansional cooling by 9.8°C

3 Confusingly, meteorologists and oceanographers follow different conventions when specifying the direction of fluid flow. Meteorologists use the direction from which the wind has come to describe it, while oceanographers take the direction in which the flow is going. Thus, an easterly wind to a meteorologist is a westward wind to an oceanographer! This unfortunate difference is too entrenched to be easily altered, and this book will use the convention appropriate to the fluid medium being described.

Fig. 1.6. Mean surface wind field in (a) January, and (b) July [Data from Oort, 1983.]

Fig. 1.6. Mean surface wind field in (a) January, and (b) July [Data from Oort, 1983.]

July mean surface wind 5 m/s

Fig. 1.7. Mean sea level pressure field in January. Contours are every 5 mb. The data is a mean of 17 years of National Meteorological Center model analysis fields.

for each kilometre of adiabatic vertical motion (the latter occurs if a parcel of air does not exchange any heat with its surroundings, as is a good approximation in, for example, the formation of cumulus clouds - see Appendix C). Within a cloud the decline of temperature with height in vertical motion can approach the typical value of Fig. 1.3, due to the release of latent heat upon condensation of water vapour. However, substantially less than half of the troposphere contains cloud at any one time so other processes must be lowering the environmental lapse rate below the dry adiabatic lapse rate. Diffusion and advection of heat from the stratosphere, the ground, or surrounding air masses is partially responsible but the major reason for the enhancement of tropospheric temperatures is the greenhouse effect.

A number of low concentration, or trace, gases in the atmosphere are unresponsive to illumination by short wavelength radiation from the Sun but absorb energy of infra-red wavelengths. The gas molecules do this by increasing their vibrational and rotational energies, rather than their kinetic energy. How this happens can be illustrated by the water molecule, shown in Fig. 1.8. The bond angle between the hydrogen atoms of an ordinary water molecule is 105°, but if a photon of a certain wavelength of infra-red radiation (6.27 |im)4 collides with the molecule the energy of the photon can be converted into a vibration of the hydrogen bonds, such that the angle between the hydrogen atoms undergoes rapid oscillation of a few degrees. Other forms of oscillation can be excited by wavelengths of 2.66 or 2.74 |im. The absorption spectra of H2O, shown in Fig. 1.9, is more complex than just these three wavelengths, however,

Fig. 1.7. Mean sea level pressure field in January. Contours are every 5 mb. The data is a mean of 17 years of National Meteorological Center model analysis fields.

4 1 |im (micrometre, often called a micron) = 10 6 m.

Fig. 1.8. Schematic diagram of a water molecule. H represents a hydrogen atom and the central O an oxygen atom. Solid lines show bond positions.

Fig. 1.8. Schematic diagram of a water molecule. H represents a hydrogen atom and the central O an oxygen atom. Solid lines show bond positions.

Fig. 1.9. The absorption spectrum of water vapour. Note the region 8-12 ^m, known as the 'water vapour window', where there is little absorption of infra-red radiation by the water vapour molecule.

as multiples, or harmonics, of the principal absorption wavelengths can also be absorbed. In addition wavelengths which are sums, or differences, of these three (and their harmonics) also show a degree of absorption, although generally of reduced intensity.

Equation (1.2) shows that the wavelength of electromagnetic radiation emitted by an object is inversely related to its temperature. Thus the mean wavelength of the radiation emitted by the Earth's surface, and within the atmosphere itself, will be longer than that of the incoming radiation from the Sun, as the latter has a surface temperature of about 6000K compared to a typical Earth surface temperature of 289K. Fig. 1.10 depicts a typical energy spectrum, seen from the tropopause, of the radiation from the Earth's surface, with the absorption by trace gases shown by shading. There are regions of the spectrum, such as wavelengths shorter than 8 |im and from 15 to 20 |im, where the infra-red radiation is almost totally absorbed by atmospheric gases. It is this absorption, and the associated re-emission of energy, much of which warms the troposphere, that is called the greenhouse effect. This name is a misnomer as the physical mechanism involved in keeping a greenhouse warm is totally different from this radiative physics. There is a small contribution from glass being transparent to solar radiation, but partially reflective to the outgoing infra-red radiation from the air and soil within the greenhouse. However, greenhouses are warm predominantly because the enclosed space eliminates convection, and hence mixing with cooler air.

The principal greenhouse gases, and their relative contribution to the greenhouse effect, are shown in Table 1.2. The percentages shown are not strictly additive because the absorption ranges of the different gases overlap. Table 1.2 also gives the fundamental absorption wavelengths of these molecules, but the complexity of the absorption spectra, with their harmonics and linear combinations of these fundamental wavelengths, must be remembered (see Fig. 1.9). Water vapour is two to three times as important in the total greenhouse effect as carbon dioxide. This fact is often neglected in discussions of greenhouse warming because water vapour is highly variable in concentration, both in space and

Table 1.2. The greenhouse gases and their contribution to the total greenhouse effect

Gas

Basic absorption wavelengths (im)

Contribution

Water vapour (H2O)

2.66, 2.74, 6.27

55-70%

Carbon dioxide (CO2)

4.26, 7.52, 14.99

25%

Methane (CH4)

3.43, 6.85, 7.27

5%

Nitrous oxide (N2O)

4.50, 7.78, 16.98

2%

Chlorofluorocarbons (CFCs)

typical bonds:

9.52, 13.8, 15.4

1%

Ozone (O3), sulphur dioxide (SO2),

other oxides of nitrogen, carbon

monoxide (CO), etc.

<1% each

Fig. 1.10. Earth's surface radiation spectrum, seen at the tropopause. The broken line is the black body emission for a typical surface temperature of 294K (21°C). The solid line is the observed spectrum, with the shaded region between denoting the energy absorbed by gases in the troposphere.

Fig. 1.10. Earth's surface radiation spectrum, seen at the tropopause. The broken line is the black body emission for a typical surface temperature of 294K (21°C). The solid line is the observed spectrum, with the shaded region between denoting the energy absorbed by gases in the troposphere.

time, making it difficult to isolate its global effect. It will, however, be vital to much of our later discussion.

The importance of the greenhouse effect to the heat budget of the atmosphere, and therefore the climate system, is shown in Fig. 1.11. The feedback mechanism between the radiation from the Earth's surface and the greenhouse re-radiation raises the amount of energy available to heat the surface from 70% of the incident solar radiation (in the absence of the atmosphere) to 144%. Present and future changes to the amounts, and proportions, of the trace gases that contribute to this effect may change these figures and so have implications for the global climate. A number of the greenhouse gases have increased significantly in concentration in the last 200 years. This may be linked to the rise in global average surface temperature of about 0.7°C over the twentieth century. Detailed discussion of these variations will be delayed to Chapters 6 and 7, which examine natural and anthropogenic alterations to the climate system.

Fig. 1.11. Global average pathways for energy in the atmosphere. A notional 100 units comes from the Sun. [From Bigg, 1992a with values modified from Kiehl and Trenberth, 1997.]

Fig. 1.11. Global average pathways for energy in the atmosphere. A notional 100 units comes from the Sun. [From Bigg, 1992a with values modified from Kiehl and Trenberth, 1997.]

1.2.2 Reflected radiation

Another major pathway for energy in the atmosphere is reflection from the surface, clouds or airborne particles. About 30% of the incident radiation is so lost from the climate system. The surface accounts for a quarter of this (the surface albedo) but the predominant loss is from the atmosphere. Variation in the cloud amount and type, the amount of suspended volcanic debris or solid chemical aggregates, and the characteristics of the Earth's surface can change the magnitude of this energy sink. While the ocean does not directly affect most of these processes we will see several exceptions later.

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