prevailing winds
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General global circulation

As the Earth does continue to rotate at a constant rate, and the winds do continue, the transfer of momentum between Earth/atmosphere/Earth must be in balance; and the angular velocity of the system maintained. (The atmosphere is rotating in the same direction as the Earth but westerly winds move faster and easterly winds move slower than the Earth's surface. Remember winds are identified by the direction they are coming from not heading to!)

The broad and very deep band of fast-moving westerlies in the westerly wind belt, centred around 45°S (but interrupted at intervals by migrating cyclones moving east but not shown in the schematic above) lose momentum to the Earth through surface friction; resulting in the Southern Ocean's west wind drift surface current. The equatorial easterlies or trade winds, and to a lesser extent the polar easterlies, gain momentum from the Earth's surface. That gain in momentum is transferred, to maintain the westerlies, via large atmospheric eddies and waves – the sub-tropical high and the sub-polar low belts.

These eddies and waves are also a part of the mechanism by which excess insolation heat energy is transferred from the low to the higher latitudes.

Globally the equatorial low pressure trough is situated at about 5°S during January and about 10°N during July. Over the Pacific the trough does not shift very far from that average position, but due to differential heating it moves considerably further north and south over continental land masses.

The low level air moving towards the trough from the sub-tropical high belts at about 30°S and 30°N is deflected by Coriolis and forming the south-east and north-east trade winds. Coriolis effect deflects air moving towards the equator to the west and air moving away from the equator to the east.

Cross section of tropospheric circulation

The intertropical convergence zone and the Hadley cell

The trade winds converging at a high angle at the equatorial trough, the "doldrums", form the intertropical convergence zone [ITCZ]. The air in the trade wind belts is forced to rise in the ITCZ and large quantities of latent heat are released as the warm, moist maritime air cools to its condensation temperature. About half the sensible heat transported within the atmosphere originates in the 0 – 10°N belt; and most of this sensible heat is released by condensation in the towering cumulus rising within the ITCZ

A secondary convergence zone of trade-wind easterlies, the South Pacific convergence zone, branches off the ITCZ near Papua-New Guinea extending south-easterly and showing little seasonal change in location or occurrence.

Over land masses the trade winds bring convective cloud which develops into heavy layer cloud with embedded thunderstorms when the air mass is lifted at the ITCZ.

The ITCZ is the boiler room of the Hadley tropical cell which provides the circulation forming the weather patterns, and climate, of the Southern Hemisphere north of 40°S. The lower level air rises in the ITCZ then moves poleward at upper levels – because of the temperature gradient effect – and is deflected to the east by Coriolis, at heights of 40 000 – 50 000 feet, while losing heat to space by radiative cooling.

The cooling air subsides in the sub-tropic region, warming by compression and forming the sub-tropical high pressure belt. Part of the subsiding air returns to the ITCZ as the south-east trade winds thus completing the Hadley cellular cycle. (The system is named after George Hadley [1685-1768], a British meteorologist who formulated the trade wind theory)

At latitudes greater than about 30°S the further southerly movement of Hadley cell air is limited by instability due to conservation of momentum effects and collapses into the Rossby wave system described in section 4.7 below. The Hadley cell and the Rossby wave system, combined with the the cold, dry polar high pressure area over the elevated Antarctic continent, dominate the Southern Hemisphere atmosphere. Fifty per cent of the Earth's surface is contained between 30°N and 30°S so the two Hadley cells directly affect half the globe.

The sub-tropical anticyclones

The subsiding high level air of the Hadley cell forms the persistent sub-tropical high pressure belt, or ridge, encircling the globe and usually located between 30°S and 50°S. Within the belt there are three semi-permanent year-round high pressure centres in the South Indian, South Pacific and South Atlantic Oceans.

In winter the high pressure belt moves northward.

The Indian Ocean centre produces about 40 anticyclones annually which, as they develop, slowly pass from west to east with their centres at about 38°S in February and about 30°S in September. The anticyclones, or warm-core highs, are generally large, covering 10° of latitude or more, roughly elliptical, vertically extensive and persistent, with the pressure gradient weakening towards the centre. The anticyclones are separated by lower pressure troughs each containing a cold front.

Winds move anticlockwise around the high, with easterlies on the northern edge and westerlies on the southern edge. Air moving equatorward on the eastern side is colder than air moving poleward on the western side. The high level subsiding air spreads out chiefly to the north and south of the ridge due to the higher surface pressures in the east and west.

Rossby waves and the westerly wind belt

Upper westerlies blow over most of the troposphere between the ITCZ and the upper polar front but are concentrated in the westerly wind belt where they undulate north and south in smooth broad waves with one, two or three semi-stationary, long wave, peaks and troughs occurring during each global circumnavigation and a number of distinct mobile short waves; each about half the length of the long waves.

The amplitude of these mobile Rossby waves, as shown on upper atmosphere pressure charts, varies considerably and can be as much as 30° of latitude. Then the airflow rather than being predominantly east/west will be away from or towards the pole. The gradient wind speed in the equatorward swing will be super-geostrophic and the speed in the poleward swing will be sub-geostrophic. The poleward swing of each wave is associated with decreasing vorticity and an upper level high pressure ridge and the equatorward swing associated with increasing vorticity and an upper trough.

Downstream of the ridge upper level convergence occurs, with upper level divergence downstream of the trough. This pattern of the Rossby waves in the upper westerlies results in compensating divergence and convergence at the lower level, accompanied by vorticity and the subsequent development of migratory surface depressions – lows or cyclones (cyclogenesis) and the development of surface highs or anticyclones (anticyclogenesis).

The long waves do not usually correspond with lower level features; being stable and slow moving, stationary or even retrograding. However they tend to steer the more mobile movement of the short waves which, in turn, steer the direction of propagation of the low level systems and weather.

The swings of the Rossby waves carry heat and momentum towards the poles and cold air away from the poles. The crests of the short waves can break off leaving pools of cold or warm air, assisting in the process of heat transfer from the tropics. Wave disturbances at the polar fronts perform a similar function at lower levels.

An upper level pool of cold air, an upper low or cut-off low or upper air disturbance, will lead to instability in the underlying air. The term cut-off low is also applied to an enclosed region of low surface pressure which has drifted into the high pressure belt, i.e. cut off from the westerly stream, or is cradled by anticyclones and high pressure ridges. Similarly the term cut-off high is also applied to an enclosed region of high surface pressure cut off from the main high pressure belt (refer 'blocking pairs' section 5.2) and to an upper level pool of warm air which is further south than normal – also termed upper high.

The upper air thickness charts, used in aviation flight planning, show the vertical distance between two isobaric surfaces, usually 1000 hPa is the lower, and the upper may be 700 hPa, 500 hPa or 300 hPa. The atmosphere in regions of less thickness, upper lows, will be unstable and colder whereas regions of greater thickness, upper highs, tend to stability. On these charts winds blow almost parallel to the geopotential height lines.

Upper Air Winds and the Jet Streams

Winds at the top of the troposphere are generally poleward and westerly in direction. The figure below describes these upper air westerlies along with some other associated weather features. Three zones of westerlies can be seen in each hemisphere on this illustration. Each zone is associated with either the Hadley, Ferrel, or Polar circulation cell.

Simplified global three-cell upper air circulation patterns.

The polar jet stream is formed by the deflection of upper air winds by coriolis acceleration. It resembles a stream of water moving west to east and has an altitude of about 10 kilometres. Its air flow is intensified by the strong temperature and pressure gradient that develops when cold air from the poles meets warm air from the tropics. Wind velocity is highest in the core of the polar jet stream where speeds can be as high as 300 kilometres per hour. The jet stream core is surrounded by slower moving air that has an average velocity of 130 kilometres per hour in winter and 65 kilometres per hour in summer.

Associated with the polar jet stream is the polar front. The polar front represents the zone where warm air from the subtropics (pink) and cold air (blue) from the poles meet. At this zone, massive exchanges of energy occur in the form of storms known as the mid-latitude cyclones. The shape and position of waves in the polar jet stream determine the location and the intensity of the mid-latitude cyclones. In general, mid-latitude cyclones form beneath polar jet stream troughs. The following satellite image, taken from above the South Pole, shows a number of mid-latitude cyclones circling Antarctica. Each mid-latitude cyclone wave is defined by the cloud development associated with frontal uplift.

Satellite view of the atmospheric circulation at the South Pole. (Source: NASA)

The subtropical jet stream is located approximately 13 kilometres above the subtropical high pressure zone. The reason for its formation is similar to the polar jet stream. However, the subtropical jet stream is weaker. Its slower wind speeds are the result of a weaker latitudinal temperature and pressure gradient.

 Polar and subtropical jet streams.

 

surface winds

An air parcel initially at rest will move from high pressure to low pressure because of the pressure gradient force (PGF). However, as that air parcel begins to move, it is deflected by the Coriolis force to the right in the northern hemisphere (to the left on the southern hemisphere). As the wind gains speed, the deflection increases until the Coriolis force equals the pressure gradient force. At this point, the wind will be blowing parallel to the isobars. When this happens, the wind is referred to as geostrophic.

 

Geostrophic wind blows parallel to the isobars because the Coriolis force and pressure gradient force are in balance. However it should be realized that the actual wind is not always geostrophic -- especially near the surface.

The surface of the Earth exerts a frictional drag on the air blowing just above it. This friction can act to change the wind's direction and slow it down -- keeping it from blowing as fast as the wind aloft. Actually, the difference in terrain conditions directly affects how much friction is exerted. For example, a calm ocean surface is pretty smooth, so the wind blowing over it does not move up, down, and around any features. By contrast, hills and forests force the wind to slow down and/or change direction much more.

As we move higher, surface features affect the wind less until the wind is indeed geostrophic. This level is considered the top of the boundary (or friction) layer. The height of the boundary layer can vary depending on the type of terrain, wind, and vertical temperature profile. The time of day and season of the year also affect the height of the boundary layer. However, usually the boundary layer exists from the surface to about 1-2 km above it.

In the friction layer, the turbulent friction that the Earth exerts on the air slows the wind down. This slowing causes the wind to be not geostrophic. As we look at the diagram above, this slowing down reduces the Coriolis force, and the pressure gradient force becomes more dominant. As a result, the total wind deflects slightly towards lower pressure. The amount of deflection the surface wind has with respect to the geostrophic wind above depends on the roughness of the terrain. Meteorologists call the difference between the total and geostrophic winds ageostrophic winds.

land and sea breezes

As the day dawns, coastal skies are cloudless or nearly cloudless, and the wind induced by large-scale weather patterns is light. As the sun rises, increased solar energy heats the surface of the earth which, in turn, heats the lowest layers of the atmosphere. At sea, however, the radiant energy received is rapidly dispersed by a combination of turbulent mixing due to winds. waves, currents and the capacity of the water to absorb great quantities of heat with only slight alteration of its temperature. Thus. the air over land warms faster than that over the sea surface. Since warmer air is lighter air, the pressure over land becomes less than that over water, the average value of this difference being, during the sea breeze regime, about 1 millibar. [1013 millibars = 1 atmosphere of pressure]

A few hours after sunrise, the pressure gradient will have built up sufficiently to allow the sea breeze to begin moving inland. As the sea breeze moves inland, the cooler sea air advances like a cold front characterized by a sudden wind shift, a drop in temperature and a rise in relative humidity. A temperature drop of 2 to 10 C degrees (3.6 to 18 F degrees) within 15 to 30 minutes is not an uncommon occurrence as the sea breeze front advances.

Thus, in the tropics, the sea breezes make coastal areas more comfortable and healthy for human habitation than the inland regions.

From the time of the sea breeze front passage until late afternoon. the wind will blow inland at speeds of 13 to 19 kilometres per hour (8 to 12 miles per hour), occasionally as strong as 40 kilometres per hour (25 miles per hour). At first, the wind blows perpendicular to the shore, but as the day wears on, friction and Coriolis effects act to veer the wind until it parallels the coastline. The landward penetration of the sea breeze reaches 15 to 50 kilometres (9 to 30 miles) in the temperate zones and 50 to 65 kilometres (30 to 40 miles) in the tropics. By late afternoon, the strength of the sea breeze slowly diminishes as the influx of solar energy lessens. The decay of the circulation pattern occurs first at the shoreline and then proceeds further inland.

The Land Breeze

As the sun sets, cooling begins along the surface of the land and sea. Like daytime heating, cooling occurs at different rates over water and land. The rapidly cooling land soon has a higher air pressure over it relative to that over the sea, and the air begins to flow down the pressure gradient seaward. This is the land breeze. It too is influenced by the roughness of the coastline, strength of the large-scale winds, and coastal configuration. Unlike the sea breeze, the land breeze is usually weaker in velocity and less common. The land breeze is often dominant for only a few hours and its direction is more variable. Nevertheless, the land breeze can penetrate the marine atmosphere for 10 kilometres (6 miles) seaward.

Climatology of the Sea and Land Breeze

The sea breeze is most common along tropical coasts, being felt on about 3 out of 4 days. The warmer temperatures, increased solar radiation and generally weaker prevailing winds in the low latitudes promote the development of the sea breeze. In general, the climatic significance of the sea breeze decreases with latitude. In temperate regions, it is generally a phenomenon of late spring and summer when atmospheric conditions (higher temperatures, weaker large-scale winds) are most favourable to the formation of the thermally induced, sea-land circulation system.

The land breeze occurs less frequently. Along coasts with steep shorelines or volcanic island coasts, however, it may be the dominant partner with speeds in excess of 32 kilometres per hour (20 miles per hour). The land breeze may also occur in the temperate regions during the cold season, especially when a warm current flows along the coast.

Lake-Land Breezes

Lake may also develop a similar local wind circulation pattern. Here the inland moving wind is known as the lake breeze. Lake breezes are quite common in late spring and summer, for example, along the shorelines of the Great Lakes, providing local residents with a place of refuge during hot, humid summer days.

mountain winds

Hills and valleys substantially distort the airflow associated with the prevailing pressure system and the pressure gradient. Strong up and down drafts and eddies develop as the air flows up over hills and down into valleys.  Wind direction changes as the air flows around hills.  Sometimes lines of hills and mountain ranges will act as a barrier, holding back the wind and deflecting it so that it flows parallel to the range.  If there is a pass in the mountain range, the wind will rush through this pass as through a tunnel with considerable speed.   The airflow can be expected to remain turbulent and erratic for some distance as it flows out of the hilly area and into the flatter countryside.

Daytime heating and night-time cooling of the hilly slopes lead to day to night variations in the airflow.  At night, the sides of the hills cool by radiation.  The air in contact with them becomes cooler and therefore denser and it blows down the slope into the valley.  This is a katabatic wind (sometimes also called a mountain breeze).  If the slopes are covered with ice and snow, the katabatic wind will blow, not only at night, but also during the day, carrying the cold dense air into the warmer valleys.  The slopes of hills not covered by snow will be warmed during the day. The air in contact with them becomes warmer and less dense and, therefore, flows up the slope. This is an anabatic wind (or valley breeze).

In mountainous areas, local distortion of the airflow is even more severe.  Rocky surfaces, high ridges, sheer cliffs, steep valleys, all combine to produce unpredictable flow patterns and turbulence.

eddies - mechanical turbulence

Mechanical turbulence is determined by both the speed of the wind and the roughness of the surface over which the air flows. As wind moves through trees or over rough surfaces, the air is broken up into eddies that make the wind flow irregular. We feel these irregularities at the surface as abrupt changes in wind speed and direction -- gusts. The eddies can either combine to form larger eddies, or cancel each other out and lessen the effect.

Thermal influences interact with mechanical influences. If there is surface heating, an eddy formed by flow obstructions may be lifted up because the air is unstable. Or the eddy created could cause instability by mixing air of different temperatures. Each influence affects the other. Next we will look at some specific examples of microscale turbulence and flow.

dust devils

Localized heating and associated convection can develop into dramatic small scale vortices. These pick up available dust and debris, creating dust devils. Localized heating and associated convection can develop into dramatic small scale vortices. These pick up available dust and debris, creating dust devils. Dust devils pose the greatest hazard near the ground where they are most violent.

tornadoes

Tornadoes are one of nature's most violent storms. In an average year, about 1,000 tornadoes are reported across the United States, resulting in 80 deaths and over 1,500 injuries. A tornado is a violently rotating column of air extending from a thunderstorm to the ground. The most violent tornadoes are capable of tremendous destruction with wind speeds of 250 mph or more. Damage paths can be in excess of one mile wide and 50 miles long.

winds speeds and direction

Wind speeds for maritime purposes are expressed in knots (nautical miles per hour). In the weather reports on US public radio and television, however, wind speeds are given in miles per hour while in Canada speeds are given in kilometres per hour.

In a discussion of wind direction, the compass point from which the wind is blowing is considered to be its direction. Therefore, a north wind is one that is blowing from the north towards the south.

veering and backing

The terms veering and backing originally referred to the shift of surface wind direction with time but meteorologists now use the term when referring to the shift in wind direction with height. Winds shifting anti-clockwise around the compass are 'backing', those shifting clockwise are 'veering'. At night, surface friction decreases as surface cooling reduces the eddy motion of the air. Surface winds will back and decrease. During the day, as surface friction intensifies, the surface winds will veer and increase.

Temperature Inversions

Temperature inversion is a condition in which the temperature of the atmosphere increases with altitude in contrast to the normal decrease with altitude. When temperature inversion occurs, cold air underlies warmer air at higher altitudes. Temperature inversion may occur during the passage of a cold front or result from the invasion of sea air by a cooler onshore breeze. Overnight radiative cooling of surface air often results in a nocturnal temperature inversion that is dissipated after sunrise by the warming of air near the ground. A more long-lived temperature inversion accompanies the dynamics of the large high-pressure systems depicted on weather maps. Descending currents of air near the centre of the high-pressure system produce a warming (by adiabatic compression), causing air at middle altitudes to become warmer than the surface air. Rising currents of cool air lose their buoyancy and are thereby inhibited from rising further when they reach the warmer, less dense air in the upper layers of a temperature inversion. During a temperature inversion, air pollution released into the atmosphere's lowest layer is trapped there and can be removed only by strong horizontal winds. Because high-pressure systems often combine temperature inversion conditions and low wind speeds, their long residency over an industrial area usually results in episodes of severe smog. As the inversion dissipates in the morning, the shear plane and gusty winds move closer to the ground, causing windshifts and increases in wind speed near the surface.

Surface Obstructions. The irregular and turbulent flow of air around mountains and hills and through mountain passes causes serious wind shear problems for aircraft approaching to land at airports near mountain ridges. Wind shear is also associated with hangars and large buildings at airports. As the air flows around such large structures, wind direction changes and wind speed increases causing shear.