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
www.islandnet.com

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.

the mountain wave

The U.S. Aeronautical Information Manual states, “Your first experience of flying over mountainous terrain, particularly if most of your flight time has been over the flatlands of the Midwest, could be a never-to-be-forgotten nightmare if you are not aware of the potential hazards awaiting … Many pilots go all their lives without understanding what a mountain wave is. Quite a few have lost their lives because of this lack of understanding. One need not be a licensed meteorologist to understand the mountain wave phenomenon.” 

The most distinctive characteristic of the mountain wave is the lenticular cloud. This is a "signpost of the sky" indicating that mountain wave activity is present.

there are several terms for mountain wave:-

• Mountain wave

• Standing wave

• Lee wave

• Gravity wave

• Standing lenticular

• ACSL (altocumulus standing lenticularis)

• Or just plane "wave"

The wave that forms over the mountain is more properly called the "mountain wave." The waves downwind from the mountain are the "standing wave" or "lee wave." Pilots have come to accept all of these names for wave activity, regardless of position of the lenticular clouds.

To set up a mountain wave condition three elements are needed:

• Wind flow perpendicular to the mountain range, or nearly so, being within about 30 degrees of perpendicular.

• An increasing wind velocity with altitude with the wind velocity 20 knots or more near mountaintop level.

• Either a stable air mass layer aloft or an inversion below about 15,000 feet.

Because of these elements, the weather service is able to predict the mountain wave condition with over 90-percent accuracy.

Figure 1

In figure 1, we have likened an atmosphere with low stability to a flimsy spring that offers little resistance to vertical motion. So while the lower coils move easily up and over the mountain, the jolt received at ground level is not transmitted very far upward.

Figure 2

Figure 2 represents a stable atmosphere that is similar to a tough, heavy spring. This air, when it strikes the mountains, tends to suppress internal vertical motion. It is essentially too tough for oscillations to be set up.

Figure 3

In figure 3 we have an arrangement of a strong coil sandwiched between two weaker springs to simulate an atmosphere with a stable layer sandwiched between areas of lesser stability. With this arrangement it is conceivable that the strong spring will continue to bounce up and down for some time after the parcel of air has crossed the mountain ridge. With a stable layer (or inversion aloft) the air stream is both flexible enough to be set in vertical motion and elastic enough to maintain that motion as a series of vertical oscillations.

As the air ascends, it cools and condenses out moisture, forming the distinctive lenticular clouds. As it descends, it compresses and the heat of compression reabsorbs the moisture. It goes through this up and down action many times forming a distinctive lenticular cloud at the apex of each crest, providing there is sufficient moisture present for the cloud formation.

The up-and-down action forms a trough at the bottom of its flow and a crest at the top of the flow. The distance from trough to trough (or crest to crest) is called the wave length. The wave length is directly proportional to wind wind and inversely proportional to stability.

The wave length is used for visualization. In the area from the trough to the crest is an area of updrafts. The area from the crest to the trough is predominately downdrafts.

In the intermountain west the wave length can vary from about 2 nautical miles to over 25 nautical miles. It averages 8 miles and extends downrange about 150-300 nautical miles. Satellite photos have shown the wave capable of extending over 700-nautical miles downwind from the mountain range.

With sufficient moisture three typical wave clouds will form, although there are four types of clouds associated with the wave.

Cap cloud (foehnwall)
Lenticular
Roll (rotor, arcus)
Mother-of-Pearl

The presence of clouds merely point out wave activity and not wave intensity at any particular level. Because moist air takes less vertical distance to reach its condensation level than does dryer air, the presence of a lenticular cloud is not necessarily an indication of the strength of the updrafts or downdrafts in a mountain wave.

For example, high altitude lenticulars may indicate there is sufficient moisture at that altitude to form them, when in fact the strongest wave lift and sink occurs at a lower altitude where there isn't enough moisture to form the lenticular clouds. This is one reason visualization is so important.

The mother-of-pearl or nacreous cloud is a pancake-shaped cloud that is extremely thin and visible for only a short time after sunset or before sunrise when the sky is dark. It is normally seen in latitudes higher than 50 degree north, or over Antarctica. It is best seen in the polar regions at 80,000 to 100,000 feet when the sun is below the horizon.

The lenticular cloud appears to be stationary although the wind may be blowing through the wave at 50 knots or more. The wave lift can extend into the stratosphere, more than 10 miles above sea level, so you can't escape wave effects by flying over them.

What are the flight conditions in lenticular clouds? Generally the lenticular area will be quite smooth. The only danger is the magnitude of the sustained updrafts and downdrafts. Usually individual lenticulars are composed of ice crystals, but when they are composed of super-cooled water droplets watch out for severe icing conditions.

Normally the rotor clouds is centred beneath the lenticular cloud. Most often it extends anywhere from ground level to mountaintop level, but is frequently observed up to 35,000 feet. Destructive turbulence from the rotor rarely exists more than 2,000-3,000 feet above mountaintop level.

The rotor is described as a "dark, ominous-looking cloud with a rotating appearance." If it forms near the ground where it can pick up dust and debris, it is dark and ominous looking, but more often it looks similar to a fair-weather cumulus. Turbulence is most frequent and most severe in the standing rotors just beneath the wave crests at or below mountaintop level (visualization is helpful where there is insufficient moisture to form the rotor or the lenticular).

The rotor area forms beneath the lee wave where a large swirling eddy forms. Sometimes with an inversion (normally stable air), turbulence succeeds in overturning the air in the stable layer. Once warm air is suddenly forced beneath colder and denser air a vigorous convection is set up in an attempt to restore normal equilibrium. This makes the roll cloud a particularly turbulent hazard. If the top of the cloud is rotating faster than the bottom, avoid the area like the plague.

The most dangerous characteristic of the standing wave is the rotor. The rotor can be assumed to exist whenever a mountain wave forms, but a cloud will not always form to alert you to its presence. Avoid the area where the rotor will form with visualization.

Often the three conditions that must exist to form a mountain wave will exist (perpendicular wind flow, increasing wind velocity with altitude, and a stable air mass layer or inversion) ... but there is insufficient moisture for the wave clouds to form. This is called a dry wave. All of the updrafts, downdrafts and rotor turbulence exists, you just can't see the clouds. You must use visualization.

Just because a mountain wave exists, it is not a sure sign that your flight must be delayed or cancelled. The degree of stability can be determined from pilot reports or by a test flight.

Mountain wave safety practices

Altitude 50 percent above the terrain - Turbulence caused by extreme mountain waves can extend into all altitudes that you might use, but dangerous turbulence can usually be avoided by clearing the mountains at least half again as high as the height of the mountain. In Colorado there are 54 peaks over 14,000-foot elevation. Does this mean we have to fly at 14,000 plus one-half (7,000) or 21,000 feet? No, use the base of the terrain to begin measuring. For example, if the surrounding terrain is 10,000 feet and the mountaintop is 14,000 feet, use one-half of the 4,000-foot value and fly 2,000 feet above the mountaintops.

Approach at a 45-degree angle - The rule-of-thumb of flying half again as high as the mountain is designed to reduce the risk of entering the turbulent rotor zone, but it does not necessarily give you a sufficient margin to allow for height loss due to downdrafts. You must have an escape route.

Avoid ragged or irregular-shaped lenticulars - Ragged and irregular-shaped lenticulars can contain the same turbulence as the rotor area.

Climb in lift - Dive in sink - By diving in sink, rather than trying to maintain altitude, the airplane is exposed to the effects of the downdraft for a lesser amount of time. Even though the rate of descent will likely be double or more the rate of climbing at the best rate-of-climb airspeed, the airplane will loose less altitude overall.

Avoid the rotor - If rotor clouds are not present, visualize the area of the rotor and avoid it.

Visualize the wave length - When flying parallel to the wave, fly in the updraft area.

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.

wind shear

Air flow in the boundary layer is normally turbulent to some degree but such turbulence does not significantly alter the aircraft’s flight path. (Bear in mind that what is a minor variation in flight path at a reasonable altitude may be hazardous when operating at low altitude and slower speed – particularly in take-off, landing and 'go-around' operations.) Practically all turbulence hazardous to flight is a result of wind shear, a sudden “variation in wind along the flight path of a pattern, intensity and duration, that displaces the aircraft abruptly from its intended path and sufficiently that substantial control action is needed.” The shear is the rate of change of wind speed and direction and its effect on flight can range from inconsequential to extremely hazardous.

Vertical shear is the change in the (roughly) horizontal wind velocity with height. i.e. as the aircraft is climbing or descending.

Horizontal shear is the change in horizontal wind velocity ( i.e. speed and/or direction – gusts and lulls) with distance flown.

Updraught, downdraught or vertical gust shear is the change in vertical air motion with horizontal distance.

Wind shear can derive from orographic, frictional, air mass instability, wave disturbance, thermal and jetstream sources. The closer to the surface that the shear occurs the more hazardous for aircraft, and particularly for aircraft with low momentum. For an aircraft taking off or landing the shear may be large enough and rapid enough to exceed the airspeed safety margin and the aircraft’s capability to accelerate or climb. Thermals as such contribute relatively minor amounts of hazardous turbulence except when flying at low levels in a superadiabatic layer. Refer 9.9 below.

Sudden entry into a significant updraught, downdraught or vertical gust is the most hazardous form of wind shear. Such events are usually associated with large convective clouds, refer 9.4 below, and the air current will have lateral and horizontal components in addition to the vertical. The aircraft's inertia will initially maintain its flight path relative to the surface (or space) so the aircraft's angle of attack must alter – with a consequent change in the lift and drag coefficients, plus a change in wing loading. The following table shows the approximate change in aoa experienced by an aircraft flying at 60 knots [6000 feet per minute] and also 100 knots [10 000 feet per minute], encountering updraughts or downdraughts with vertical components of various speeds. The angles are roughly calculated using the 1-in-60 rule.

Approximate change in aoa

Thus an aircraft approaching to land at 60 knots and encountering a 1000 feet per minute downdraught would experience an initial reduction in aoa of 9°. Presuming that an aircraft on approach has an aoa of 8° to 10° and looking at the C_L curve in the Flight Theory basic forces module you can see that a 9° aoa reduction is going to reduce C_L from a value around 0.9 to about 0.1 which indicates a reduction in lift of at least 80% and, consequently, the aircraft will initially sink very rapidly – i.e. at a descent rate greater than the vertical speed of the airstream. The event becomes hazardous should the rate of downflow exceed the aircraft's rate of climb capability; and turbulence within the downdraft can add to the hazard.

Similarly if an aircraft flying straight and level at 60 knots encounters a 1500 feet per minute updraught the angle of attack will be increased by 15° exceeding the critical aoa and the aircraft will stall at an airspeed higher than the normal stall speed.

Wing loading will also change rapidly while the aoa is changing; refer to gust induced loading in the aircraft flight envelope discussion. The faster the aircraft's speed when encountering shear the greater the wing loading.

Generally the pilot's best option when sudden up/downdraught shear is recognised, which may in itself take a few seconds, is to hold the aircraft's attitude and not chase the air speed indicator, flying straight ahead until out of the up/downdraught – bearing in mind that opposite shear will be encountered on the other side. However should the aircraft encounter a severe downdraught at low level the only option is to immediately apply full power and either try to maintain height and allow the airspeed to decay or to maintain airspeed and let the aircraft lose height. Whichever way the pilot is in an extremely hazardous situation, indicating that recognition and avoidance of extreme shear conditions is really the only wise option.

Low level wind shear

   Generally, below 2000 feet agl and over flat terrain, the amount of horizontal and vertical shear, in both direction and speed, is largely dependent on temperature lapse rate conditions:-
Greater lapse rate » greater instability » greater vertical mixing » more uniformity of flow through layer and less shear. An exception is in extremely turbulent conditions below a Cb.. But if the environment lapse rate exceeds about 3 ºC per 1000 feet then convective thermal turbulence will be severe.

In stable conditions convective turbulence is minimised so that vertical shear in the boundary layer is enhanced with highest values in the lower 300 feet, which will affect aircraft taking off and landing. See the section in the meteorology guide on velocity change between surface and gradient wind.

High vertical wind shear values are often attained at the upper boundary of an inversion. An aircraft climbing through the inversion layer in the same direction as the overlaying wind would experience a momentary loss of air speed, and lift, through the effect of inertia, i.e. the aircraft will continue at the same initial velocity, relative to Earth or space, until thrust increases velocity and restores airspeed.

Thunderstorms. Wind shear, associated with thunderstorms, occurs as the result of two phenomena, the gust front and downbursts. As the thunderstorm matures, strong downdrafts develop, strike the ground and spread out horizontally along the surface well in advance of the thunderstorm itself. This is the gust front. Winds can change direction by as much as 180° and reach speeds as great as 100 knots as far as 10 miles ahead of the storm. The downburst is an extremely intense localized downdraft flowing out of a thunderstorm. The downburst (there are two types of downbursts: macrobursts and microbursts) usually is much closer to the thunderstorm than the gust front. Dust clouds, roll clouds, intense rainfall or virga (rain that evaporates before it reaches the ground) are due to the possibility of downburst activity but there is no way to accurately predict its occurrence.

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.