Brief Survey of the Atmosphere

The remainder of this chapter provides an overview of the optical properties, composition, and vertical structure of the Earth's atmosphere, the major wind systems, and the climatological-mean distribution of precipitation. It introduces some of the terminology that will be used in subsequent chapters and some of the conventions that will be used in performing calculations involving amounts of mass and rates of movement.

1.3.1 Optical Properties

The Earth's atmosphere is relatively transparent to incoming solar radiation and opaque to outgoing radiation emitted by the Earth's surface. The blocking of outgoing radiation by the atmosphere, popularly referred to as the greenhouse effect, keeps the surface of the Earth warmer than it would be in the absence of an atmosphere. Much of the absorption and reemission of outgoing radiation are due to air molecules, but cloud droplets also play a significant role. The radiation emitted to space by air molecules and cloud droplets provides a basis for remote sensing of the three-dimensional distribution of temperature and various atmospheric constituents using satellite-borne sensors.

The atmosphere also scatters the radiation that passes through it, giving rise to a wide range of optical effects. The blueness of the outer atmosphere in Fig. 1.5 is due to the preferential scattering of incoming short wavelength (solar) radiation by air molecules, and the whiteness of lower layers is due to scattering from cloud droplets and atmospheric aerosols (i.e., particles). The backscattering of solar radiation off the top of the deck of low clouds off the California coast in Fig. 1.7 greatly enhances the

Fig. 1.7 A deck of low clouds off the coast of California, as viewed in reflected visible radiation. [NASA MODIS imagery. Photograph courtesy of NASA.]

whiteness (or reflectivity) of that region as viewed from space. Due to the presence of clouds and aerosols in the Earth's atmosphere, ~22% of the incoming solar radiation is backscattered to space without being absorbed. The backscattering of radiation by clouds and aerosols has a cooling effect on climate at the Earth's surface, which opposes the greenhouse effect.

1.3.2 Mass

At any point on the Earth's surface, the atmosphere exerts a downward force on the underlying surface due to the Earth's gravitational attraction. The downward force, (i.e., the weight) of a unit volume of air with density p is given by

where g is the acceleration due to gravity. Integrating Eq. (1.4) from the Earth's surface to the "top" of the atmosphere, we obtain the atmospheric pressure on the Earth's surface ps due to the weight (per unit area) of the air in the overlying column f rn

Neglecting the small variation of g with latitude, longitude and height, setting it equal to its mean value of g0 = 9.807 m s~2, we can take it outside the integral, in which case, Eq. (1.5) can be written as

where m = j™pdz is the vertically integrated mass per unit area of the overlying air.

Exercise 1.1 The globally averaged surface pressure is 985 hPa. Estimate the mass of the atmosphere.

Solution: From Eq. (1.6), it follows that g0

where the overbars denote averages over the surface of the Earth. In applying this relationship the pressure

8 Introduction and Overview must be expressed in pascals (Pa). Substituting numerical values we obtain

The mass of the atmosphere is

4n X (6.37 X 106)2 m2 X 1.004 X 104 kg m-2 5.10 X 1014 m2 X 1.004 X 104 kg m-2 5.10 X 1018 kg6 ■

1.3.3 Chemical Composition

The atmosphere is composed of a mixture of gases in the proportions shown in Table 1.1, where fractional concentration by volume is the same as that based on numbers of molecules, or partial pressures exerted by the gases, as will be explained more fully in Section 3.1. The fractional concentration by

Table 1.1 Fractional concentrations by volume of the major gaseous constituents of the Earth's atmosphere up to an altitude of 105 km, with respect to dry air

Constituent"

Molecular weight

Fractional concentration by volume

Nitrogen (N2)

28.013

78.08%

Oxygen (02)

32.000

20.95%

Argon (Ar)

39.95

0.93%

Water vapor (H2O)

18.02

0-5%

Carbon dioxide (CO2)

44.01

380 ppm

Neon (Ne)

20.18

18 ppm

Helium (He)

4.00

5 ppm

Methane (CH4)

16.04

1.75 ppm

Krypton (Kr)

83.80

1 ppm

Hydrogen (H2)

2.02

0.5 ppm

Nitrous oxide (N2O)

56.03

0.3 ppm

Ozone (O3)

48.00

0-0.1 ppm

a So called greenhouse gases are indicated by bold-faced type. For more detailed information on minor constituents, see Table 5.1.

mass of a constituent is computed by weighting its fractional concentration by volume by its molecular weight, i.e., mi niMi

^2niMi

a So called greenhouse gases are indicated by bold-faced type. For more detailed information on minor constituents, see Table 5.1.

where mi is the mass, ni the number of molecules, and Mi the molecular weight of the ith constituent, and the summations are over all constituents.

Diatomic nitrogen (N2) and oxygen (O2) are the dominant constituents of the Earth's atmosphere, and argon (Ar) is present in much higher concentrations than the other noble gases (neon, helium, krypton, and xenon). Water vapor, which accounts for roughly 0.25% of the mass of the atmosphere, is a highly variable constituent, with concentrations ranging from around 10 parts per million by volume (ppmv) in the coldest regions of the Earth's atmosphere up to as much as 5% by volume in hot, humid air masses; a range of more than three orders of magnitude. Because of the large variability of water vapor concentrations in air, it is customary to list the percentages of the various constituents in relation to dry air. Ozone concentrations are also highly variable. Exposure to ozone concentrations >0.1 ppmv is considered hazardous to human health.

For reasons that will be explained in §4.4, gas molecules with certain structures are highly effective at trapping outgoing radiation. The most important of these so-called greenhouse gases are water vapor, carbon dioxide, and ozone. Trace constituents CH4, N2O, CO, and chlorofluorocarbons (CFCs) are also significant contributors to the greenhouse effect.

Among the atmosphere's trace gaseous constituents are molecules containing carbon, nitrogen, and sulfur atoms that were formerly incorporated into the cells of living organisms. These gases enter the atmosphere through the burning of plant matter and fossil fuels, emissions from plants, and the decay of plants and animals. The chemical transformations that remove these chemicals from the atmosphere involve oxidation, with the hydroxyl (OH) radical playing an important role. Some of the nitrogen and sulfur compounds are converted into particles that are eventually "scavenged" by raindrops, which contribute to acid deposition at the Earth's surface.

6 When the vertical and meridional variations in g and the meridional variations in the radius of the earth are accounted for, the mass per unit area and the total mass of the atmosphere are ~0.4% larger than the estimates derived here.

Although aerosols and cloud droplets account for only a minute fraction of the mass of the atmosphere, they mediate the condensation of water vapor in the atmospheric branch of the hydrologic cycle, they participate in and serve as sites for important chemical reactions, and they give rise to electrical charge separation and a variety of atmospheric optical effects.

1.3.4 Vertical structure

To within a few percent, the density of air at sea level is 1.25 kg m—3. Pressure p and density p decrease nearly exponentially with height, i.e.,

where H, the e-folding depth, is referred to as the scale height and p0 is the pressure at some reference level, which is usually taken as sea level (z = 0). In the lowest 100 km of the atmosphere, the scale height ranges roughly from 7 to 8 km. Dividing Eq. (1.8) by p0 and taking the natural logarithms yields

Because the pressure at a given height in the atmosphere is a measure of the mass that lies above that level, it is sometimes used as a vertical coordinate in lieu of height. In terms of mass, the 500-hPa level, situated at a height of around 5.5 km above sea level, is roughly halfway up to the top the atmosphere. ■

Density decreases with height in the same manner as pressure. These vertical variations in pressure and density are much larger than the corresponding horizontal and time variations. Hence it is useful to define a standard atmosphere, which represents the horizontally and temporally averaged structure of the atmosphere as a function of height only, as shown in Fig. 1.8. The nearly exponential height dependence of pressure and density can be inferred from the fact that the observed vertical profiles of pressure and density on these semilog plots closely resemble straight lines. The reader is invited to verify in Exercise 1.14 at the end of this chapter that the corresponding 10-folding depth for pressure and density is ~17 km.

This relationship is useful for estimating the height of various pressure levels in the Earth's atmosphere.

Exercise 1.3 Assuming an exponential pressure and density dependence with H = 7.5 km, estimate the heights in the atmosphere at which (a) the air density is equal to 1 kg m—3 and (b) the height at which the pressure is equal to 1 hPa.

Exercise 1.2 At approximately what height above sea level zm does half the mass of the atmosphere lie above and the other half lie below? [Hint: Assume an exponential pressure dependence with H = 8 km and neglect the small vertical variation of g with height.]

Solution: Let pm be the pressure level that half the mass of the atmosphere lies above and half lies below. The pressure at the Earth's surface is equal to the weight (per unit area) of the overlying column of air. The same is true of the pressure at any level in the atmosphere. Hence, pm = p0/2 where p0 is the global-mean sea-level pressure. From Eq. (1.9)

zm = —H ln 0.5 = H ln 2 Substituting H = 8 km, we obtain zm = 8 Km X 0.693 ~ 5.5 km

10 102 103

Fig. 1.8 Vertical profiles of pressure in units of hPa, density in units of kg m-3, and mean free path (in meters) for the U.S. Standard Atmosphere.

10 102 103

Fig. 1.8 Vertical profiles of pressure in units of hPa, density in units of kg m-3, and mean free path (in meters) for the U.S. Standard Atmosphere.

10 Introduction and Overview

Solution: Solving Eq. (1.9),we obtain z = Hln(p0/p), and similarly for density. Hence, the heights are (a)

for the 1-kg m 3 density level and (b)

1QQQ l.QQ

52 km for the 1-hPa pressure level. Because H varies with height, geographical location, and time, and the reference values p0 and p0 also vary, these estimates are accurate only to within ~10%. ■

Exercise 1.4 Assuming an exponential pressure and density dependence, calculate the fraction of the total mass of the atmosphere that resides between 0 and 1 scale height, 1 and 2 scale heights, 2 and 3 scale heights, and so on above the surface.

Solution: Proceeding as in Exercise 1.2, the fraction of the mass of the atmosphere that lies between 0 and 1, 1 and 2, 2 and 3, and so on scale heights above the Earth's surface is e_1, e~2,... e~N from which it follows that the fractions of the mass that reside in the 1st, 2nd ..., Nth scale height above the surface are e-1, e-1 (1

where N is the height of the base of the layer expressed in scale heights above the surface. The corresponding numerical values are 0.632,0.233,0.086 ... ■

Throughout most of the atmosphere the concentrations of N2, O2,Ar, CO2, and other long-lived constituents tend to be quite uniform and largely independent of height due to mixing by turbulent fluid motions.7 Above ~105 km, where the mean free path between molecular collisions exceeds 1 m (Fig. 1.8), individual molecules are sufficiently mobile that each molecular species behaves as if it alone were present. Under these conditions, concentrations of heavier constituents decrease more rapidly with height than those of lighter constituents: the density of each constituent drops off exponentially with height, with a scale height inversely proportional to molecular weight, as explained in Section 3.2.2. The upper layer of the atmosphere in which the lighter molecular species become increasingly abundant (in a relative sense) with increasing height is referred to as the heterosphere. The upper limit of the lower, well-mixed regime is referred to as the turbopause, where turbo refers to turbulent fluid motions and pause connotes limit of.

The composition of the outermost reaches of the atmosphere is dominated by the lightest molecular species (H, H2, and He). During periods when the sun is active, a very small fraction of the hydrogen atoms above 500 km acquire velocities high enough to enable them to escape from the Earth's gravitational field during the long intervals between molecular collisions. Over the lifetime of the Earth the leakage of hydrogen atoms has profoundly influenced the chemical makeup of the Earth system, as discussed in Section 2.4.2.

The vertical distribution of temperature for typical conditions in the Earth's atmosphere, shown in Fig. 1.9, provides a basis for dividing the atmosphere into four layers (troposphere, stratosphere,

100 90 80 70

160 180 200 220 240 260 280 300 Temperature (K)

Fig. 1.9 A typical midlatitude vertical temperature as represented by the U.S. Standard Atmosphere.

0.01

0.001

160 180 200 220 240 260 280 300 Temperature (K)

Fig. 1.9 A typical midlatitude vertical temperature as represented by the U.S. Standard Atmosphere.

100 300

0.001

0.01

100 300

1000

profile,

7 In contrast, water vapor tends to be concentrated within the lowest few kilometers of the atmosphere because it condenses and precipitates out when air is lifted. Ozone are other highly reactive trace species exhibit heterogeneous distributions because they do not remain in the atmosphere long enough to become well mixed.

mesosphere, and thermosphere), the upper limits of which are denoted by the suffix pause.

The tropo (turning or changing)sphere is marked by generally decreasing temperatures with height, at an average lapse rate, of ~6.5 °C km-1. That is to say, r = —~ 6.5 °C km-1 dz

where T is temperature and r is the lapse rate. Tropospheric air, which accounts for ~80% of the mass of the atmosphere, is relatively well mixed and it is continually being cleansed or scavenged of aerosols by cloud droplets and ice particles, some of which subsequently fall to the ground as rain or snow. Embedded within the troposphere are thin layers in which temperature increases with height (i.e., the lapse rate r is negative). Within these so-called temperature inversions it is observed that vertical mixing is strongly inhibited.

Within the strato-(layered)-sphere, vertical mixing is strongly inhibited by the increase of temperature with height, just as it is within the much thinner temperature inversions that sometimes form within the troposphere. The characteristic anvil shape created by the spreading of cloud tops generated by intense thunderstorms and volcanic eruptions when they reach the tropopause level, as illustrated in Fig. 1.10, is due to this strong stratification.

Cloud processes in the stratosphere play a much more limited role in removing particles injected by

Fig. 1.10 A distinctive "anvil cloud" formed by the spreading of cloud particles carried aloft in an intense updraft when they encounter the tropopause. [Photograph courtesy of Rose Toomer and Bureau of Meteorology, Australia.]

volcanic eruptions and human activities than they do in the troposphere, so residence times of particles tend to be correspondingly longer in the stratosphere. For example, the hydrogen bomb tests of the 1950s and early 1960s were followed by hazardous radioactive fallout events involving long-lived stratospheric debris that occurred as long as 2 years after the tests.

Stratospheric air is extremely dry and ozone rich. The absorption of solar radiation in the ultraviolet region of the spectrum by this stratospheric ozone layer is critical to the habitability of the Earth. Heating due to the absorption of ultraviolet radiation by ozone molecules is responsible for the temperature maximum ~50 km that defines the stratopause.

Above the ozone layer lies the mesosphere (meso connoting "in between"), in which temperature decreases with height to a minimum that defines the mesopause. The increase of temperature with height within the thermosphere is due to the absorption of solar radiation in association with the dissociation of diatomic nitrogen and oxygen molecules and the stripping of electrons from atoms. These processes, referred to as photodissociation and photoionization, are discussed in more detail in Section 4.4.3. Temperatures in the Earth's outer thermosphere vary widely in response to variations in the emission of ultraviolet and x-ray radiation from the sun's outer atmosphere.

At any given level in the atmosphere temperature varies with latitude. Within the troposphere, the clima-tological-mean (i.e., the average over a large number of seasons or years), zonally averaged temperature generally decreases with latitude, as shown in Fig. 1.11. The meridional temperature gradient is substantially stronger in the winter hemisphere where the polar cap region is in darkness. The tropopause is clearly evident in Fig. 1.11 as a discontinuity in the lapse rate. There is a break between the tropical tropopause, with a mean altitude ~17km, and the extratropical tropopause, with a mean altitude ~10 km. The tropical tropopause is remarkably cold, with temperatures as low as -80 °C. The remarkable dryness of the air within the stratosphere is strong evidence that most of it has entered by way of this "cold trap."

Exercise 1.5 Based on data shown in Fig. 1.10, estimate the mean lapse rate within the tropical troposphere.

12 Introduction and Overview

90 70 50 30 10 I 10 30 50 70 90 Equator

Summer hemisphere Winter hemisphere

201-

70° 50° 30° 10° | 10° 30° 50° 70° 90° Equator

Summer hemisphere Winter hemisphere

Fig. 1.11 Idealized meridional cross sections of zonally averaged temperature (in °C) (Top) and zonal wind (in m s~1) (Bottom) around the time of the solstices, when the meridional temperature contrasts and winds are strongest. The contour interval is 20 °C; pink shading denotes relatively warm regions, and cyan shading relatively cold regions. The contour interval is 10 m s~1; the zero contour is bold; pink shading and "W" labels denote westerlies, and yellow shading and "E" labels denote easterlies. Dashed lines indicate the positions of the tropopause, stratopause, and mesopause. This representation ignores the more subtle distinctions between northern and southern hemisphere climatologies. [Courtesy of Richard J. Reed.]

201-

70° 50° 30° 10° | 10° 30° 50° 70° 90° Equator

Summer hemisphere Winter hemisphere

Fig. 1.11 Idealized meridional cross sections of zonally averaged temperature (in °C) (Top) and zonal wind (in m s~1) (Bottom) around the time of the solstices, when the meridional temperature contrasts and winds are strongest. The contour interval is 20 °C; pink shading denotes relatively warm regions, and cyan shading relatively cold regions. The contour interval is 10 m s~1; the zero contour is bold; pink shading and "W" labels denote westerlies, and yellow shading and "E" labels denote easterlies. Dashed lines indicate the positions of the tropopause, stratopause, and mesopause. This representation ignores the more subtle distinctions between northern and southern hemisphere climatologies. [Courtesy of Richard J. Reed.]

Solution: At sea level the mean temperature of the tropics is ~27 °C, the tropopause temperature is near -80 °C, and the altitude of the tropopause altitude is ~17 km. Hence the lapse-rate is roughly

Note that a decrease in temperature with height is implicit in the term (and definition of) lapse rate, so the algebraic sign of the answer is positive. ■

1.3.5 Winds

Differential heating between low and high latitudes gives rise to atmospheric motions on a wide range of scales. Prominent features of the so-called atmospheric general circulation include planetary-scale west-to-east (westerly) midlatitude tropos-pheric jet streams, centered at the tropopause break around 30° latitude, and lower mesospheric jet streams, both of which are evident in Fig. 1.11. The winds in the tropospospheric jet stream blow from the west throughout the year; they are strongest during winter and weakest during summer. In contrast, the mesospheric jet streams undergo a seasonal reversal: during winter they blow from the west and during summer they blow from the east.

Superimposed on the tropospheric jet streams are eastward propagating, baroclinic waves that feed upon and tend to limit the north-south temperature contrast across middle latitudes. Baroclinic waves are one of a number of types of weather systems that develop spontaneously in response to instabilities in the large-scale flow pattern in which they are embedded. The low level flow in baroclinic waves is dominated by extratropical cyclones, an example of which is shown in Fig. 1.12. The term cyclone denotes a closed circulation in which the air spins in the same sense as the Earth's rotation as viewed from above (i.e., counterclockwise in the northern hemisphere). At low levels the air spirals inward toward the center.8 Much of the significant weather associated with extratropical cyclones is concentrated within narrow frontal zones, i.e., bands, a few tens of kilometers in width, characterized by strong horizontal temperature contrasts. Extratropical weather systems are discussed in Section 8.1.

Tropical cyclones (Fig. 1.13) observed at lower latitudes derive their energy not from the north-south temperature contrast, but from the release of latent heat of condensation of water vapor in deep convec-tive clouds, as dicussed in Section 8.3. Tropical cyclones tend to be tighter and more axisymmetric than extratropical cyclones, and some of them are much more intense. A distinguishing feature of a well-developed tropical cyclone is the relatively calm, cloud-free eye at the center.

8 The term cyclone derives from the Greek word for "coils of a snake."

Fig. 1.12 An intense extratropical cyclone over the North Pacific. The spiral cloud pattern, with a radius of nearly 2000 km, is shaped by a vast counterclockwise circulation around a deep low pressure center. Some of the elongated cloud bands are associated with frontal zones. The region enclosed by the red rectangle is shown in greater detail in Fig. 1.21. [NASA MODIS imagery. Photograph courtesy of NASA.]

Fig. 1.12 An intense extratropical cyclone over the North Pacific. The spiral cloud pattern, with a radius of nearly 2000 km, is shaped by a vast counterclockwise circulation around a deep low pressure center. Some of the elongated cloud bands are associated with frontal zones. The region enclosed by the red rectangle is shown in greater detail in Fig. 1.21. [NASA MODIS imagery. Photograph courtesy of NASA.]

a. Wind and pressure

The pressure field is represented on weather charts in terms of a set of isobars (i.e., lines or contours along which the pressure is equal to a constant value) on a horizontal surface, such as sea level. Isobars are usually plotted at uniform increments: for example, every 4 hPa on a sea-level pressure chart (e.g.,... 996, 1000, 1004 ... hPa). Local maxima in the pressure field are referred to as high pressure

Fig. 1.13 The cloud pattern associated with an intense tropical cyclone approaching Florida. The eye is clearly visible at the center of the storm. The radius of the associated cloud system is ~600 km. [NOAA GOES imagery. Photograph courtesy of Harold F. Pierce, NASA Goddard Space Flight Center.]

centers or simply highs, denoted by the symbol H, and minima as lows (L). At any point on a pressure chart the local horizontal pressure gradient is oriented perpendicular to the isobars and is directed from lower toward higher pressure. The strength of the horizontal pressure gradient is inversely proportional to the horizontal spacing between the isobars in the vicinity of that point.

With the notable exception of the equatorial belt (10 °S-10 °N), the winds observed in the Earth's atmosphere closely parallel the isobars. In the northern hemisphere, lower pressure lies to the left of the wind (looking downstream) and higher pressure to the right.9,10 It follows that air circulates counterclockwise around lows and clockwise around highs, as shown in the right-hand side of Fig. 1.14. In the southern hemisphere the relationships are in the opposite sense, as indicated in the left-hand side of Fig. 1.14.

This seemingly confusing set of rules can be simplified by replacing the words "clockwise" and

9 This relationship was first noted by Buys-Ballot in 1857, who stated: If, in the northern hemisphere, you stand with your back to the wind, pressure is lower on your left hand than on your right.

10 Christopher H. D. Buys-Ballot (1817-1890) Dutch meteorologist, professor of mathematics at the University of Utrecht. Director of Dutch Meteorolgical Institute (1854-1887). Labored unceasingly for the widest possible network of surface weather observations.

14 Introduction and Overview

cyclones

Fig. 1.14 Blue arrows indicate the sense of the circulation around highs (H) and lows (L) in the pressure field, looking down on the South Pole (left) and the North Pole (right). Small arrows encircing the poles indicate the sense of the Earth's rotation.

North Pole cyclones

Fig. 1.14 Blue arrows indicate the sense of the circulation around highs (H) and lows (L) in the pressure field, looking down on the South Pole (left) and the North Pole (right). Small arrows encircing the poles indicate the sense of the Earth's rotation.

"counterclockwise" with the terms cyclonic and anti-cyclonic (i.e., in the same or in the opposite sense as the Earth's rotation, looking down on the pole). A cyclonic circulation denotes a counterclockwise circulation in the northern hemisphere and a clockwise circulation in the southern hemisphere. In either hemisphere the circulation around low pressure centers is cyclonic, and the circulation around high pressure centers is anticyclonic: that is to say, in reference to the pressure and wind fields, the term low is synonymous with cyclone and high with anticyclone.

In the equatorial belt the wind tends to blow straight down the pressure gradient (i.e., directly across the isobars from higher toward lower pressure). In the surface wind field there is some tendency for cross-isobar flow toward lower pressure at higher latitudes as well, particularly over land. The basis for these relationships is discussed in Chapter 7.

b. The observed surface wind field

This subsection summarizes the major features of the geographically and seasonally varying climatological-mean surface wind field (i.e., the background wind field upon which transient weather systems are superimposed). It is instructive to start by considering the circulation on an idealized ocean-covered Earth with the sun directly overhead at the equator, as inferred from simulations with numerical models.

North Pole

^--xTropospheric ^¿J jet stream

Hadley cells

Fig. 1.15 Schematic depiction of sea-level pressure isobars and surface winds on an idealized aqua planet, with the sun directly overhead on the equator. The rows of H's denote the subtropical high-pressure belts, and the rows of L's denote the subpolar low-pressure belt. Hadley cells and tropospheric jet streams (J) are also indicated.

^--xTropospheric ^¿J jet stream

Hadley cells

Fig. 1.15 Schematic depiction of sea-level pressure isobars and surface winds on an idealized aqua planet, with the sun directly overhead on the equator. The rows of H's denote the subtropical high-pressure belts, and the rows of L's denote the subpolar low-pressure belt. Hadley cells and tropospheric jet streams (J) are also indicated.

The main features of this idealized "aqua-planet, perpetual equinox" circulation are depicted in Fig. 1.15. The extratropical circulation is dominated by westerly wind belts, centered around 45 °N and 45 °S. The westerlies are disturbed by an endless succession of eastward migrating disturbances called baroclinic waves, which cause the weather at these latitudes to vary from day to day. The average wavelength of these waves is ~4000 km and they propagate eastward at a rate of ~10 m s_1.

The tropical circulation in the aqua-planet simulations is dominated by much steadier trade winds,11 marked by an easterly zonal wind component and a component directed toward the equator. The northeasterly trade winds in the northern hemisphere and the southeasterly trade winds in the southern hemisphere are the surface manifestation of overturning circulations that extend through the depth of the troposphere. These so-called Hadley12 cells are characterized by (1) equatorward flow in the boundary layer, (2) rising motion within a few degrees of the equator, (3) poleward return flow in the tropical upper troposphere, and (4) sinking motion in the

11 The term trade winds or simply trades derives from the steady, dependable northeasterly winds that propelled sailing ships along the popular trade route across the tropical North Atlantic from Europe to the Americas.

12 George Hadley (1685-1768) English meteorologist. Originally a barrister. Formulated a theory for the trade winds in 1735 which went unnoticed until 1793 when it was discovered by John Dalton. Hadley clearly recognized the importance of what was later to be called the Coriolis force.

subtropics, as indicated in Fig. 1.15. Hadley cells and trade winds occupy the same latitude belts.

In accord with the relationships between wind and pressure described in the previous subsection, trade winds and the extratropical westerly wind belt in each hemisphere in Fig. 1.15 are separated by a subtropical high-pressure belt centered ~30° latitude in which the surface winds tend to be weak and erratic. The jet streams at the tropopause (12 km; 250 hPa) level are situated directly above the subtropical high pressure belts at the Earth's surface. A weak minimum in sea-level pressure prevails along the equator, where trade winds from the northern and southern hemispheres converge. Much deeper lows form in the extratropics and migrate toward the poleward flank of the extratropical westeries to form the subpolar low pressure belts.

In the real world, surface winds tend to be stronger over the oceans than over land because they are not slowed as much by surface friction. Over the Atlantic and Pacific Oceans, the surface winds mirror many of the features in Fig. 1.15, but a longitudinally dependent structure is apparent as well. The subtropical high-pressure belt, rather than being continuous, manifests itself as distinct high-pressure centers, referred to as subtropical anticyclones, centered over the mid-oceans, as shown in Fig. 1.16.

In accord with the relationships between wind and pressure described in the previous subsection, surface winds at lower latitudes exhibit an equatorward

North Pole

North Pole

Equator -

Fig. 1.16 Schematic of the surface winds and sea-level pressure maxima and minima over the Atlantic and Pacific Oceans showing subtropical anticyclones, subpolar lows, the midlati-tude westerly belt, and trade winds.

Equator -

Fig. 1.16 Schematic of the surface winds and sea-level pressure maxima and minima over the Atlantic and Pacific Oceans showing subtropical anticyclones, subpolar lows, the midlati-tude westerly belt, and trade winds.

component on the eastern sides of the oceans and a poleward component on the western sides. The equa-torward surface winds along the eastern sides of the oceans carry (or advect) cool, dry air from higher latitudes into the subtropics; they drive coastal ocean currents that advect cool water equatorward; and they induce coastal upwelling of cool, nutrient-rich ocean water, as explained in the next chapter. On the western sides of the Atlantic and Pacific Oceans, poleward winds advect warm, humid, tropical air into middle latitudes.

In an analogous manner, the subpolar low-pressure belt manifests itself as mid-ocean cyclones referred to, respectively, as the Icelandic low and the Aleutian low. The poleward flow on the eastern flanks of these semipermanent, subpolar cyclones moderates the winter climates of northern Europe and the Pacific coastal zone poleward of ~40 °N. The subtropical anticyclones are most pronounced during summer, whereas the subpolar lows are most pronounced during winter.

The idealized tropical circulation depicted in Fig. 1.15, with the northeasterly and southeasterly trade winds converging along the equator, is not realized in the real atmosphere. Over the Atlantic and Pacific Oceans, the trade winds converge, not along the equator, but along ~7 °N, as depicted schematically in the upper panel of Fig. 1.17. The belt in which the convergence takes place is referred to as the intertropical convergence zone (ITCZ). The asymmetry with respect to the equator is a consequence of the land-sea geometry, specifically the northwest-southeast orientation of the west coastlines of the Americas and Africa.

Surface winds over the tropical Indian Ocean are dominated by the seasonally reversing monsoon cir-culation,13 consisting of a broad arc originating as a westward flow in the winter hemisphere, crossing the equator, and curving eastward to form a belt of moisture-laden westerly winds in the summer hemisphere, as depicted [for the northern hemisphere (i.e., boreal) summer] in the lower panel of Fig. 1.17. The monsoon is driven by the presence of India and southeast Asia in the northern hemisphere subtrop-ics versus the southern hemisphere subtropics. Surface temperatures over land respond much more strongly to the seasonal variations in solar heating than those over ocean. Hence, during July the

13 From mausin, the Arabic word for season.

16 Introduction and Overview

Fig. 1.17 Schematic depicting surface winds (arrows), rainfall (cloud masses), and sea surface temperature over the tropical oceans between ~30 °N and 30 °S. Pink shading denotes warmer, blue cooler sea surface temperature, and khaki shading denotes land. (Top) Atlantic and Pacific sectors where the patterns are dominated by the intertropical convergence zone (ITCZ) and the equatorial dry zone to the south of it. (Bottom) Indian Ocean sector during the northern (boreal) summer monsoon, with the Indian subcontinent to the north and open ocean to the south. During the austral summer (not shown) the flow over the Indian Ocean is in the reverse direction and the rain belt lies just to the south of the equator.

Fig. 1.17 Schematic depicting surface winds (arrows), rainfall (cloud masses), and sea surface temperature over the tropical oceans between ~30 °N and 30 °S. Pink shading denotes warmer, blue cooler sea surface temperature, and khaki shading denotes land. (Top) Atlantic and Pacific sectors where the patterns are dominated by the intertropical convergence zone (ITCZ) and the equatorial dry zone to the south of it. (Bottom) Indian Ocean sector during the northern (boreal) summer monsoon, with the Indian subcontinent to the north and open ocean to the south. During the austral summer (not shown) the flow over the Indian Ocean is in the reverse direction and the rain belt lies just to the south of the equator.

subtropical continents of the northern hemisphere are much warmer than the sea surface temperature over the tropical Indian Ocean. It is this temperature contrast that drives the monsoon flow depicted in the lower panel of Fig. 1.17. In January, when India and southeast Asia are cooler than the sea surface temperature over the tropical Indian Ocean, the monsoon flow is in the reverse sense (not shown).

The reader is invited to compare the observed climatological-mean surface winds for January and July shown in Figs. 1.18 and 1.19 with the idealized flow patterns shown in the two previous figures. In Fig. 1.18, surface winds, based on satellite data, are shown together with the rainfall distribution, indicated by shading, and in Fig. 1.19 a different version of the surface wind field, derived from a blending of many datasets, is superimposed on the climatologi-cal-mean sea-level pressure field.

By comparing the surface wind vectors with the shading in Fig. 1.18, it is evident that the major rain belts, which are discussed in the next subsection, tend to be located in regions where the surface wind vectors flow together (i.e., converge). Convergence at low levels in the atmosphere is indicative of ascending motion aloft. Through the processes discussed in Chapter 3, lifting of air leads to condensation of water vapor and ultimately to precipitation. Figure 1.19 provides verification that the surface winds tend to blow parallel to the isobars, except in the equatorial belt. At all latitudes a systematic drift across the isobars from higher toward lower pressure is also clearly apparent.

The observed winds over the southern hemisphere (Figs. 1.18 and 1.19) exhibit well-defined extratropical westerly and tropical trade wind belts reminiscent of those in the idealized aqua-planet simulations (Fig. 1.15). Over the northern hemisphere the surface winds are strongly influenced by the presence of high latitude continents. The subpolar low-pressure belt manifests itself as oceanic pressure minima (the Icelandic and Aleutian lows) surrounded by cyclonic (counterclockwise) circulations, as discussed in connection with Fig. 1.16. These features and the belts of westerly winds to the south of them are more pronounced during January than during July. In contrast, the northern hemisphere oceanic subtropical anticyclones are more clearly discernible during July.

c. Motions on smaller scales

Over large areas of the globe, the heating of the Earth's surface by solar radiation gives rise to buoyant plumes analogous to those rising in a pan of water heated from below. As the plumes rise, the displaced air subsides slowly, creating a two-way circulation. Plumes of rising air are referred to by glider pilots as thermals, and when sufficient moisture is present they are visible as cumulus clouds (Fig. 1.20). When the overturning circulations are confined to the lowest 1 or 2 km of the atmosphere (the so-called mixed layer or atmospheric boundary layer), as is often the case, they are referred to as shallow convection. Somewhat deeper, more vigorous convection gives rise to showery weather in cold air masses flowing over a warmer surface (Fig. 1.21).

Under certain conditions, buoyant plumes originating near the Earth's surface can break through the weak temperature inversion that usually caps the mixed layer, giving rise to towering clouds that extend all the way to the tropopause, as shown in Fig. 1.22. These clouds are the signature of deep convection,

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  • kauko
    What planet's atmosphere most resembles that of earth?
    9 years ago

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