0). Despite the stable conditions, potential energy from the flow can be converted into kinetic energy. Baroclinic instabilities drive the development of mid-latitude cyclones and associated frontal systems.
Figure 2.13 Baroclinic instability illustrated by parcel trajectories in the latitude–altitude plane with respect to isentropic surfaces. The isentropes slope upward and poleward. For quasi-horizontal parcel trajectories with slopes shallower than those of the background isentropes (solid arrows), unstable conditions arise even though the background atmosphere is stable (∂θ/∂z > 0). By contrast, trajectories with slopes steeper than the isentropes (dashed arrows) are suppressed by stability.
2.9 General Circulation of the Troposphere
Solar heating of the Earth must be balanced on a global basis by emission of terrestrial radiation to space. Both of these terms are a strong function of latitude (Figure 2.14). On an annual mean basis, the tropics receive much higher solar radiation than higher latitudes. Terrestrial emission to space peaks at about 20° latitude and drops at higher latitudes. On balance, the tropics have a surplus of radiative energy and the high latitudes have a deficit. Energy balance requires that heat be transported from the tropics to the poles by atmospheric and oceanic motions. The general circulation of the atmosphere refers to the global wind systems that carry out this transport of energy.
Figure 2.14 Annual mean balance between average net shortwave solar and longwave terrestrial radiation as a function of latitude.
Reproduced with permission from Pidwirny (2006).
In the absence of planetary rotation, the latitudinal gradient in surface heating would drive hemispheric circulation cells with upward motion in the tropics, high-altitude poleward flow, subsidence over the poles, and return equatorward flow near the surface (Figure 2.15, (a)). Planetary rotation complicates this simple picture due to the Coriolis force acting on air parcels as they travel meridionally. The poleward flow is deflected to the east and the equatorward flow is deflected to the west. The high-altitude poleward flow originating from the Equator becomes fully zonal at a latitude of about 30° and at that point no further meridional transport takes place. Thus the meridional circulation cells only extend from the Equator to 30° (Figure 2.15, (b)). These are called the Hadley cells after George Hadley (1685–1768), who first recognized the effect of planetary rotation on the general circulation of the atmosphere. Convergence between the southern and northern cells near the Equator defines the intertropical convergence zone (ITCZ) as a band of persistent precipitation. The ITCZ moves seasonally with solar declination, which is the angle between the Sun’s rays and the equatorial plane of the Earth. Solar declination varies from +23° on June 21 to –23° on December 22. This defines the wet seasons of the tropics. Subsidence at 30° produces hot and dry conditions over land; the major deserts of the world are in that latitudinal band. Deflection to the west of the return equatorward flow near the surface produces the persistent tropical easterlies known as the trade winds.
Figure 2.15 General circulation of the atmosphere. (a): Expected circulation in the absence of planetary rotation. (b): Circulation of the rotating troposphere. The Hadley cells feature rising air in the tropics and sinking air in the subtropics. The higher-latitude meridional cells are mainly conceptual as most of the meridional transport in the extratropics is driven by waves traveling longitudinally.
From NASA, courtesy Barbara Summey, NASA Goddard VisAnalysis Laboratory; and from Lutgens and Tarbuck (2000).
Poleward of 30°, the movement of air masses is considerably modified by the Earth’s rotation as the Coriolis force becomes stronger. The Coriolis force imposes a strong circumpolar flow, so that air travels zonally around the Earth on a timescale of weeks. The mid-latitude troposphere is strongly baroclinic, resulting in dynamical instabilities that spawn mid-latitude cyclones (Section 2.8). In the extratropics, most of the meridional heat exchange takes place by wave systems manifested by traveling weather disturbances (cyclones, anticyclones, and associated fronts between warm and cold air masses). The polar regions are characterized by cold and dry air with small weather disturbances and rare precipitation. Meridional mixing of air within a hemisphere takes place on a timescale of about three months, while mixing of air between the two hemispheres across the ITCZ takes place on a timescale of one year. The ITCZ is a major dynamical barrier for atmospheric mixing because of the weak thermal contrast across the Equator. Many long-lived gases such as CO2 are well-mixed within each hemisphere, but feature an interhemispheric gradient maintained by the ITCZ.
The general circulation of the atmosphere is further influenced by the geographic distribution of continents and oceans. In the tropics, differences in surface heating between warm continents and cooler oceans drive zonal asymmetries in the circulation. Deep tropical convection takes place over the continents and the western equatorial Pacific (the warm pool, where sea surface temperatures are the highest in the world). Subsidence prevails over most of the tropical oceans, particularly where ocean currents maintain relatively cold surface temperatures (East Pacific, South Atlantic). Seasonal variations in land heating and cooling produce monsoon circulations, as illustrated in Figure 2.16 for South Asia. During winter the cold continental surface air flows toward the ocean, producing dry conditions over land. During summer the moist ocean air flows over the heated land, resulting in heavy convective precipitation. Yet another effect of land on the general circulation is friction and topography. Thus the extensive land masses at northern mid-latitudes promote weather disturbances and meridional flow, facilitating the transport of heat to the Arctic.
Figure 2.16 Surface circulation in Southeast Asia during northern winter (a) and summer (b), featuring the seasonal monsoons. The seasonal shift in the ITCZ is a consequence of the monsoon.
Reproduced with permission from Lutgens et al. (2013), Copyright © Pearson Education, Inc.
Figure 2.17 illustrates the mean climatological distributions of surface pressure and winds in January and July. The seasonal shift of the ITCZ is apparent. Easterly trade winds prevail on both sides of the ITCZ. The subtropics are characterized by semi-permanent anti-cyclonic conditions that reflect the downwelling branches of the Hadley cells. Mid-latitude westerlies develop on the poleward side of these subtropical anti-cyclones and are far more steady in the southern hemisphere than in the north due to lack of ocean–land contrast. Meridional pressure gradients (shown by the isobars) are generally stronger in winter than summer, due to the greater meridional heating gradients, and this results in stronger winds.
Figure 2.17 Climatological mean surface pressures (hPa) and winds in January (a) and July (b). The location of the intertropical convergence zone (ITCZ) is indicated. Major centers of high (H) and low (L) pressure are also shown.
Reproduced from Lutgens and Tarbuck (2000).
Figure 2.18 shows global climatological distributions of precipitation in January and July. The band of intense precipitation near the Equator corresponds to the ITCZ. Seasonal shift in the ITCZ drives the wet and dry seasons in the tropics; in January the northern tropics are dry while the southern tropics are wet, and this is reversed in July. Subtropics are dry while mid-latitudes generally experience moderate precipitation in all seasons. Prominent storm tracks off the east coasts of Asia and North America play an important role in transport from northern mid-latitudes to the Arctic.
Figure 2.18 Precipitation rates [mm day–1] in January (a) and in July (b), averaged between 1988 and 1996, based on data from the Global Precipitation Climatology Project (GPCP).
From Xie and Arkin (1997), copyright © American Meteorological Society, used with permission.
Different modes of interannual climatic variability are superimposed on this mean climatological description of the atmospheric circulation. The dominant mode in the tropics is the El Niño–Southern Oscillation (ENSO), a pattern of reversing ocean temperatures between the eastern and western tropical Pacific that takes place every 3–8 years (Figure 2.19). During the normal co
ld phase of ENSO (also called La Niña), sea surface temperatures are cold in the eastern Pacific and warm in the western Pacific. There results strong subsidence and dry conditions in the east, and deep convection and wet conditions in the west. During the warm phase (also called El Niño), warm waters move from the western to the central and eastern Pacific, modifying considerably the tropical circulation with droughts over Oceania, precipitation over eastern South America, and weakened trade winds. Beyond the Pacific, ENSO affects the climate of other regions of the world through complex teleconnections.
Figure 2.19 El Niño–Southern Oscillation (ENSO) mode of climatic variability, featuring La Niña conditions (cold phase, (a)) and El Niño conditions (warm phase, (b)).
Reproduced with permission from Cunningham and Cunningham (2010).
At higher latitudes, the major mode of interannual climate variability is the Arctic Oscillation (AO), characterized by changing meridional pressure gradients between northern mid-latitudes and the Arctic. The North Atlantic Oscillation (NAO) is a regional manifestation of the AO (Figure 2.20) and its phase is measured by the pressure difference between the Azores high and the Icelandic low (positive phase when the pressure difference is large, negative phase when it is small). In the positive phase of the AO/NAO, high pressure at northern mid-latitudes pushes the jet stream northward, maintains strong surface westerlies, and restricts exchange of air with the Arctic. This leads to relatively warm and wet conditions in northern Europe and Alaska, and dry conditions in the eastern USA and Mediterranean region. In the negative phase of the AO/NAO there is more meandering of the jet stream and cold Arctic air can penetrate deep into northern mid-latitudes.
Figure 2.20 Schematic representation of the main dynamical patterns over the North Atlantic during negative and positive phases of the North Atlantic Oscillation.
From the National Oceanic and Atmospheric Administration NOAA (www.climate.gov).
2.10 Planetary Boundary Layer
The planetary boundary layer (PBL) is the layer of the atmosphere that interacts with the surface on a timescale of a day or less (Figure 2.21). It typically extends up to 1–3 km above the surface. The air above the PBL is called the free troposphere. The free troposphere has a general slow sinking motion, balancing the few locations where deep convection or frontal lifting injects PBL air to high altitudes. The compressional heating from this sinking air produces a semi-permanent subsidence inversion (Section 2.6.2) that caps the PBL and sharply restricts mixing between the PBL and the free troposphere.
Figure 2.21 Diurnal evolution of the planetary boundary layer (PBL) over land (a) and implications for chemical concentrations in surface air (b).
PBL dynamics plays an important role in determining the fate of chemicals emitted at the surface and the resulting concentrations in surface air. Vertical mixing driven by solar heating of the surface can drive large diurnal cycles of concentrations within the PBL. Venting of the PBL to the free troposphere is critical for global dispersal of chemicals.
Vertical mixing within the PBL is driven by turbulent eddies. These eddies are generated at the surface by the action of the wind on rough surface elements (mechanical turbulence) and by buoyancy (buoyant turbulence). Over land, sensible heating of the surface during the day generates buoyant plumes that may rise up to the base of the subsidence inversion. Conversely, nighttime cooling of the land surface produces stable conditions that dampen the mechanical turbulence. Over the oceans, the large heat capacity of the ocean minimizes this diurnal cycle of heating and cooling and the PBL remains neutral throughout the day.
Figure 2.21 shows the diurnal evolution of the PBL structure over land. At night, mechanical turbulence usually maintains a shallow, well-mixed layer typically 10–100 m deep called the surface layer. Above that altitude, the atmosphere is stable because of surface cooling; this is the residual layer. After sunrise, surface heating erodes the stable residual layer from below, producing an unstable mixed layer that grows over the morning hours to eventually reach the full depth of the PBL. Clouds may develop in the upper part; these are the familiar fair-weather cumuli and the corresponding layer is called the convective cloud layer (CCL). The CCL tends to have moderate stability due to the latent heat release from cloud condensation, resulting in some separation from the mixed layer. The depth of the mixed layer (excluding any CCL) is called the mixing depth. Suppression of surface heating at sunset causes rapid collapse of the mixed layer and the nighttime conditions return.
The diurnal variation of PBL structure has important implications for the diurnal evolution of chemical concentrations in surface air, as shown in Figure 2.21. An inert chemical continuously emitted at the surface will accumulate in surface air over the course of the night, leading to high concentrations. During morning the concentration will decrease as growth of the mixed layer causes dilution. By contrast, a chemical originating in the free troposphere and removed by deposition to the surface will be depleted in surface air over the course of the night, and replenished during morning by entrainment from aloft as the mixed layer grows.
Over the ocean there is no diurnal cycle of surface heating and cooling, and neutral conditions prevail where vertical mixing is driven by mechanical turbulence. The mixed layer is called the marine boundary layer (MBL) and typically extends to about 1 km altitude with no diurnal variation. It is often capped by a shallow cloud layer, either cumulus clouds or stratus, capped in turn by the subsidence inversion.
Entrainment of air from the free troposphere into the PBL and ventilation of PBL air to the free troposphere are important processes for atmospheric chemistry, connecting the surface to the global atmosphere. Ventilation generally takes place by weather events, such as frontal systems or deep convective updrafts that force boundary layer air to the free troposphere. Entrainment, by contrast, generally takes place as a slow, steady process involving the large-scale sinking of the atmosphere to compensate for the convective updrafts. Typical downward entrainment velocities at the top of the PBL are of the order of 0.1–1 cm s–1, and this replaces the PBL air on a timescale of days to a week.
2.11 Middle Atmosphere Dynamics
Vertical motions in the stratosphere are strongly suppressed by the temperature inversion resulting from absorption of solar UV radiation by ozone. A first approximation of the thermal structure of the stratosphere can be made by assuming radiative equilibrium conditions, where the heating rate from UV absorption by ozone and O2 is balanced by the cooling rate from emission of IR terrestrial radiation by CO2, water vapor, and ozone. The resulting temperatures increase with latitude from the winter to the summer pole (Figure 2.11). Based on the thermal wind equation, the zonal wind is easterly in the summer hemisphere and westerly in the winter hemisphere. The polar stratosphere in winter features strong zonal winds that form a polar vortex, isolating it from lower latitudes.
Departure from radiative equilibrium conditions is induced by the dissipation of upward propagating waves generated at the Earth’s surface. Wave breaking occurs when the amplitude of the wave becomes sufficiently large to render the disturbance unstable. This dissipation process tends to mix the medium through which the wave is propagating. Further, the momentum deposited by these waves as they break produces a torque that tends to decelerate the zonal wind and generate a meridional circulation. The resulting mean meridional temperature distribution arises from a balance between the net radiative heating/cooling described previously and the adiabatic heating/cooling associated by the compression/expansion of air produced by the wave-generated vertical motions.
Different types of waves are observed in the middle atmosphere. Rossby waves are planetary-scale disturbances in the zonal atmospheric flow that owe their existence to the latitudinal variation of the Coriolis effect. A familiar example is the meandering jet stream. These waves are generated by baroclinic instability and the forcing action of zonally asymmetric heating and topography. Upward wave propagation to the middle atmosphere is possible only when the wind is
westerly (during winter) and for the longest waves with wavenumber 1–3 (“wavenumber” is the number of complete wave cycles along the longitude around the entire Earth). Shorter Rossby waves are confined to the troposphere, where they contribute to the formation of weather systems. Rossby wave breaking in the stratosphere takes place in a relatively large “surf zone” characterized by intense quasi-horizontal mixing of chemical species. The mean circulation produced by dissipation of the waves in the winter hemisphere is directed from the Equator to the pole (Figure 2.22) and is called the Brewer–Dobson circulation since it was inferred from observations of water vapor (by Alan Brewer) and of ozone (by Gordon Dobson) in the lower stratosphere. Occasional large amplification of Rossby waves in the northern hemisphere disrupts the stratospheric circulation and causes sudden warming events in the Arctic stratosphere that disrupt the polar vortex.
Figure 2.22 Schematic representation of the upward propagation of planetary waves in the winter middle atmosphere (thick black lines) and of gravity waves (thin black lines). The meridional circulation resulting from the dissipation of these waves in the stratosphere and mesosphere is shown by the large arrow. The circulation is directed from the tropics to the pole in the winter stratosphere and from the summer to the winter pole in the mesosphere. Radiative equilibrium prevails in the summer stratosphere. The position of the jets near the tropopause (dotted line) is shown.
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