u n d e r s t a n d i n g m e t e o r o l o g y t h e a t m o s p h e r e , p r e s s u r e a n d f o r c e s our thanks to http://www.ucar.edu and www.raa.asn.au
(Copyright John Brandon)

The thin envelope of air that surrounds our planet is a mixture of gases, each with its own physical properties. The mixture is far from evenly divided. Two elements, nitrogen and oxygen, make up 99% of the volume of air. The other 1% is composed of "trace" gases, the most prevalent of which is the inert gaseous element argon. The rest of the trace gases, although present in only minute amounts, are very important to life on earth. Two in particular, carbon dioxide and ozone, can have a large impact on atmospheric processes.
Another gas, water vapour, also exists in small amounts. It varies in concentration from being almost non-existent over desert regions to about 4% over the oceans. Water vapour is important to weather production since it exists in gaseous, liquid, and solid phases and absorbs radiant energy from the earth.
Structure of the Atmosphere
The atmosphere is divided vertically into four layers based on temperature: the troposphere,
stratosphere, mesosphere, and thermosphere. Throughout the Cycles unit, we'll focus primarily on the layer in which we live - the troposphere.
Troposphere
The word troposphere comes from tropein, meaning to turn or change. All of the earth's weather occurs in the troposphere.

The troposphere has the following characteristics.
 It extends from the earth's surface to an average of 12 km (7 miles).
 The pressure ranges from 1000 to 200 millibars (29.92 in. to 5.92 in.).
 The temperature generally decreases with increasing height up to the tropopause (top of the troposphere); this is near 200 millibars or 36,000 ft. o The temperature averages 15°C (59°F) near the surface and -57°C (-71°F) at the tropopause. o The layer ends at the point where temperature no longer varies with height. This area, known as the tropopause, marks the transition to the stratosphere.
 Winds increase with height up to the jet stream.
 The moisture concentration decreases with height up to the tropopause. o The air is much drier above the tropopause, in the stratosphere. o The sun's heat that warms the earth's surface is transported upwards largely by convection and is mixed by updrafts and downdrafts.
 The troposphere is 70% and 21% . The lower density of molecules higher up would not give us enough to survive.
Atmospheric Processes
Interactions - Atmosphere and Ocean

In the Cycles overview, we learned that water is an essential part of the earth's system. The oceans cover nearly three-quarters of the earth's surface and play an important role in exchanging and transporting heat and moisture in the atmosphere.
 Most of the water vapour in the atmosphere comes from the oceans.
 Most of the precipitation falling over land finds its way back to oceans.
 About two-thirds returns to the atmosphere via the water cycle.
You may have figured out by now that the oceans and atmosphere interact extensively. Oceans not only act as an abundant moisture source for the atmosphere but also as a heat source and sink
(storage).
The exchange of heat and moisture has profound effects on atmospheric processes near and over the oceans. Ocean currents play a significant role in transferring this heat poleward. Major currents, such as the northward flowing Gulf Stream, transport tremendous amounts of heat poleward and contribute to the development of many types of weather phenomena. They also warm the climate of nearby locations. Conversely, cold southward flowing currents, such as the California current, cool the climate of nearby locations.
Energy Heat Transfer
Practically all of the energy that reaches the earth comes from the sun. Intercepted first by the atmosphere, a small part is directly absorbed, particularly by certain gases such as ozone and water vapor. Some energy is also reflected back to space by clouds and the earth's surface.

Energy is transferred between the earth's surface and the atmosphere via conduction, convection, and radiation.
Conduction is the process by which heat energy is transmitted through contact with neighbouring molecules.
Some solids, such as metals, are good conductors of heat while others, such as wood, are poor conductors. Air and water are relatively poor conductors.
Since air is a poor conductor, most energy transfer by conduction occurs right at the earth's surface.
At night, the ground cools and the cold ground conducts heat away from the adjacent air. During the day, solar radiation heats the ground, which heats the air next to it by conduction.
Convection transmits heat by transporting groups of molecules from place to place within a substance. Convection occurs in fluids such as water and air, which move freely.
In the atmosphere, convection includes large- and small-scale rising and sinking of air masses and smaller air parcels. These vertical motions effectively distribute heat and moisture throughout the atmospheric column and contribute to cloud and storm development (where rising motion occurs) and dissipation (where sinking motion occurs).
To understand the convection cells that distribute heat over the whole earth, let's consider a simplified, smooth earth with no land/sea interactions and a slow rotation. Under these conditions, the equator is warmed by the sun more than the poles. The warm, light air at the equator rises and spreads northward and southward, and the cool dense air at the poles sinks and spreads toward the equator. As a result, two convection cells are formed.

Meanwhile, the slow rotation of the earth toward the east causes the air to be deflected toward the right in the northern hemisphere and toward the left in the southern hemisphere. This deflection of the wind by the earth's rotation is known as the Coriolis effect.
Radiation is the transfer of heat energy without the involvement of a physical substance in the transmission. Radiation can transmit heat through a vacuum.
Energy travels from the sun to the earth by means of electromagnetic waves. The shorter the wavelength, the higher the energy associated with it. This is demonstrated in the animation below. As the drill's revolutions per minute (RPMs) increase, the number of waves generated on the string increases, as does the oscillation rate. The same principle applies to electromagnetic waves from the sun, where shorter wavelength radiation has higher energy than longer wavelength radiation.
Most of the sun's radiant energy is concentrated in the visible and near-visible portions of the spectrum. Shorter-than-visible wavelengths account for a small percentage of the total but are extremely important because they have much higher energy. These are known as ultraviolet
wavelengths.
Atmospheric oxygen
In the homosphere each gas exerts a partial pressure, the product of the total atmospheric pressure and the concentration of the gas. Thus as oxygen represents about 21% of the composite gases, the partial pressure of oxygen is about 21% of the atmospheric pressure at any altitude within the homosphere.
Interpolating from the pressure gradient graph above, oxygen partial pressure at selected altitudes is shown below. The decreasing partial pressure of oxygen as an aircraft climbs past 10 000 – 12 000 feet has critical effects on aircrew; the maximum exposure time for a fit person, without inspiring supplemental oxygen, is shown in the right hand column. Exposure beyond these times leads to unconsciousness.
Altitude O² pressure Max. exposure
Sea level 210 hPa —
7000 feet 165 hPa
10 000 feet 150 hPa
—
—
15 000 feet 120 hPa
18 000 feet 105 hPa
25 000 feet 80 hPa
30+ minutes
20–30 minutes
3–5 minutes

30 000 feet 65 hPa
35 000 feet 50 hPa
40 000 feet 30 hPa
1–3 minutes
30–60 seconds
10–20 seconds
Atmospheric density
The average density of dry air in temperate climates is about 1.225 kg/m³ at mean sea level, decreasing with altitude.
There are several gas laws and equations which relate the temperature, pressure, density and volume of a gas. However the equation most pertinent to aeronautical needs is the equation of state:
r = P/RT where:
r (the Greek letter rho) = density in kg/m³
P = the static air pressure in hectopascals
R = the gas constant = 2.87
T = the temperature in Kelvin units = °C + 273
We can calculate the ISA standard sea level air density, knowing that standard sea level pressure =
1013 hPa and temperature = 15 °C or 288 K
i.e. Air density = 1013 / (2.87 × 288) = 1.225 kg/m³
However if the air temperature happened to be 30 °C or 303 K at the same pressure then density would = 1013 / (2.87 × 303) = 1.165 kg/m³ or a 5% reduction.
By restating the equation of state: P = RrT it can be seen that if density remains constant, pressure increases if temperature increases.
The ICAO International Standard Atmosphere
The International Civil Aviation Organisation's International Standard Atmosphere [ ISA ] provides a fixed standard atmospheric model used for many purposes among which are the uniform assessment of aircraft performance and the calibration of some aircraft instruments. The model is akin to the average condition in mid-latitudes but contains the following assumptions:


 dry air is assumed throughout the atmosphere
 the mean sea level pressure = 1013.25 hPa
 the msl temperature = 15 °C [288 K]
 the tropopause is at 36 090 feet [11 km] and the pressure at the tropopause = 226.3 hPa the temperature lapse rate to 36 090 feet = 6.5 °C per km or nearly 2 °C per 1000 feet the temperature between 36 090 and 65 600 feet [20 km] remains constant at –56.5 °C.
The table below shows a few values derived from the ISA. Those pressure levels noted with a flight level designator are standard pressure levels used for aviation weather purposes, particularly thickness charts.
Pressure Flight level Temperature Air density Altitude hPa °C kg/m³ feet
1013 15 1.225
msl
1000
950
900
850 A050
14.3
11.5
8.6
5.5
1.212
1.163
1.113
1.063
364
1773
3243
4781

550
500
450
400
350
300
250
200
150
100
800
750
700
650
600
A100
FL140
FL185
FL235
FL300
FL340
FL385
FL445
-16.6
-21.2
-26.2
-31.7
-37.7
-44.5
2.3
-1.0
-4.6
-8.3
-12.3
-52.3
-56.5
-56.5
-56.5
0.747
0.692
0.635
0.577
0.518
0.457
1.012
0.960
0.908
0.855
0.802
0.395
0.322
0.241
0.161
6394
8091
9882
11 780
13 801
15 962
18 289
20 812
23 574
26 631
30 065
33 999
38 662
44 647
53 083
Not immediately apparent from the ISA table is that the pressure lapse rate is about one hPa per 30 feet up to the 850 hPa level, then slowing to 40 feet per hPa at the 650 hPa level, 50 feet at the 450 hPa level, 75 feet at the 300 hPa level and so on, however, this provides a useful rule of thumb:
Rule of Thumb #1
"An altitude change of 30 feet per hPa can be assumed for operations below 10 000 feet." station pressure, sea level pressure and altimeter setting
Station pressure is the actual atmospheric pressure at the elevation of the observing station.
QFE: The pressure corrected to the official airfield elevation. An altimeter set to the particular airfield QFE reads zero when an aircraft is on the ground (strictly the height of the altimeter above the ground). In the circuit, the height indicated is the height above official airfield datum.
QNH: The pressure 'reduced' to mean sea level, assuming ISA temperature profile from the station/airfield to MSL. An altimeter set to the airfield QNH reads the elevation of the airfield when on the ground.

pressure systems
The pressure chart shows the distribution of atmospheric pressure. Pressure systems - depressions
(LOW pressure regions) and anticyclones (HIGH pressure) are marked and Isobars are drawn on the chart to link areas with the same pressure. Isobar lines are drawn at 4mB interval (4 HPa) and weather frontal systems are marked using standard symbols.
Cold front Warm front Occluded front
Wind direction and some indication of strength can be deduced from the pressure chart. In the
Northern Hemisphere, winds blow in an anti-clockwise direction around a depression (LOW) and in a clockwise direction around an anticyclone (HIGH). The closer the isobars are together, then the greater the pressure gradient and the higher will be the wind strength.
Pressure charts are a useful help in interpreting satellite images. The satellite image shows the pattern of cloud cover and with the help of the pressure chart, frontal systems can be identified and tracked over a period of time. Typically, rain will be associated with the passage of a front - identifying and tracking the fronts can allow the forecast of rain, changes of temperature, wind direction and speed etc. low pressure areas also known as cyclones ww2010.atmos.uiuc.edu
A low pressure centre is where the pressure has been measured to be the lowest relative to its surroundings. That means, moving in any horizontal direction away from the "Low" will result in an increase in pressure. Low pressure centres also represent the centres of cyclones.
A low pressure centre is indicated on a weather map by a red "L" and winds flow counter clockwise around a low in the northern hemisphere. The opposite is true in the southern hemisphere, where winds flow clockwise around an area of low pressure. secondary low

The Secondary Low Centre within an occluded front is accompanied by rain, showers and possibly thunder.
Parameter
Precipitation
Description
Moderate to heavy precipitation and showers, some with thunder around the secondary low
Temperature
Wind (incl. gusts) Cyclonically veering winds around the low
Other relevant information
06 March 2002/06.00 UTC - Meteosat IR image; weather events (green: rain and showers, blue: drizzle, cyan: snow, black: no precipitation); position of Secondary Low indicated through of low pressure

A trough of low pressure is an extended area of low pressure. On the weather chart, it has associated with it a trough line. The pressure there is lower than that at neighbouring points on either side of the trough line.
col
A col is a neutral region between two highs and two lows. Weather conditions are apt to be unsettled. In winter, the mixing of air of dissimilar air masses frequently produces fog. In summer, showers or thunderstorms may occur. While it is quite possible for weather conditions to be fair, generally speaking, cols may be regarded as regions of undependable weather.
High Pressure Centres also known as anticyclones
A high pressure centre is where the pressure has been measured to be the highest relative to its surroundings. That means, moving in any direction away from the "High" will result in a decrease in pressure. A high pressure centre also represents the centre of an anticyclone and is indicated on a weather map by a blue "H". ridge of high pressure
An anti-cyclone ridge is a neck or ridge of high pressure with lower pressure lying on either side. The weather in a ridge is generally fine to fair.
pressure changes
Pressure readings are taken at regular intervals (usually hourly) at weather stations. Weather maps are prepared four times a day at six-hour intervals. From these readings and maps, the changes in pressure can be observed and approaching weather forecast. If a low, for example, is approaching a station, the pressure will steadily fall. Once the centre of the low has passed by, the pressure will begin to rise. This pattern of changing pressure is called pressure tendency.

forces encountered coriolis force
Since the globe is rotating, any movement on the Northern hemisphere is diverted to the right, if we look at it from our own position on the ground. (In the southern hemisphere it is bent to the left).
This apparent bending force is known as the Coriolis force. (Named after the French mathematician
Gustave Gaspard Coriolis 1792-1843). It may not be obvious to you that a particle moving on the northern hemisphere will be bending towards the right.
Consider this red cone moving southward in the direction of the tip of the cone. The earth is spinning, while we watch the spectacle from a camera fixed in outer space. The cone is moving straight towards the south.
Note that the red cone is veering in a curve towards the right as it moves. The reason why it is not following the direction in which the cone is pointing is, of course, that we as observers are rotating along with the globe.
Here we show the same image with the camera locked on to the globe.
Here we show the same image,with the camera fixed in outer space, while the earth rotates.

The Coriolis force is a visible phenomenon. Railroad tracks wear out faster on one side than the other. River beds are dug deeper on one side than the other. (Which side depends on which hemisphere we are in: In the Northern hemisphere moving particles are bent towards the right).
In the Northern hemisphere the wind tends to rotate counter clockwise (as seen from above) as it approaches a low pressure area. In the Southern hemisphere the wind rotates clockwise around low pressure areas.

w e i g h t o f a t m o s p h e r e our thanks www.raa.asn.au
(Copyright John Brandon)
Atmospheric pressure reflects the average density and thus the weight of the column of air above a given level, thus the pressure at a point on the earth's surface must be greater than the pressure at any height above it. An increase in surface pressure denotes an increase in mass, not thickness, of the column of air above the surface point. Similarly a decrease in surface pressure denotes a decrease in the mass. The gradient is the difference in pressure vertically, and horizontally.
The air throughout the column is compressed by the weight of the atmosphere above it and thus the density of a column of air is greatest at the surface and decreases exponentially with altitude as shown in the following graph which is a plot of the rate of decrease in density with increase in altitude. The plot is for dry air at mid-latitudes. ( Mid latitudes are usually accepted to be the areas between the 30° and 60° lines while low latitudes lie between the equator and 30° and high
latitudes between 60° and the pole.) The atmosphere at about 22 000 feet has only 50% of the sea level density, density decreases by about 3% per 1000 feet between sea level and 18 500 feet and thereafter the density lapse rate slows.
The dry air density gradient in mid-latitudes, refer table in 2.1
As the pressure decreases with height so, in any parcel of air, the downwards pressure over the top of the parcel must be less than the upwards pressure under the bottom: thus within the parcel there is a
vertical component of the pressure gradient force acting upward; generally this force is balanced by the gravitational force so the net sum of forces is zero and the parcel floats in equilibrium. This balance of forces is called the hydrostatic balance. When the two do not quite balance the difference is the buoyancy force which is the upward, or downward, force exerted on a parcel of air arising from the density difference between the parcel and the surrounding air.
Atmospheric pressure also varies horizontally, due to air mass changes associated with the regional thickness of the atmospheric layer. The resultant horizontal pressure gradient force, not being balanced by gravity, forces air to move from regions of higher pressure towards regions of lower pressure but the movement is modified by Coriolis effect. The horizontal force is about 1/15 000 of the vertical component.
( Advection is the term used for the transport of momentum, heat, moisture, vorticity or other atmospheric properties, by the horizontal movement of air)
The following graph plots the average mid-latitude vertical pressure gradient and shows how the overall vertical decrease in pressure – the pressure lapse rate – slows exponentially as the air becomes less dense with height. In a denser, or colder, air mass the pressure reduces at a faster rate, conversely in less dense, or warmer, air the pressure reduces at a slower rate. (The hydrostatic equation states that the vertical change in pressure, between two levels in any column of air, is
equal to the weight, per unit area, of the air in the column.) If two air columns have the same

pressure change from top to bottom the denser column will be shorter, conversely if the two columns have the same height the denser column will have a larger change in pressure from top to bottom.
In the ICAO standard atmosphere, details of which are shown in section 2.1, the rate of altitude change for each 1 hPa change in pressure is approximately: zero to 5000 feet: 30 feet/hPa or 34 hPa per 1000 feet
5000 to 10 000 feet: 34 feet/hPa or 29 hPa per 1000 feet
10 000 to 20 000 feet: 43 feet/hPa or 23 hPa per 1000 feet
20 000 to 40 000 feet: 72 feet/hPa or 14 hPa per 1000 feet
The change in altitude for one hectopascal change in pressure can be roughly calculated from the absolute temperature and the pressure at the level using the equation: = 96T/P feet.

c l o u d s
Higher level clouds
Cirrus and cirrostratus
Cirrocumulus
middle clouds
Altocumulus
Altocumulus Castellanus
Altostratus
Nimbostratus
low clouds
Stratus
Stratocumulus
Nimbostratus clouds of vertical development
Cumulus (Cu)
Towering Cumulus (TCu)
Cumulonimbus (Cb) cloud formation precipitation thunderstorms orographic_lift convection frontal lift convergence
It is essential that the pilot has a good knowledge of meteorology and that he/she understands what weather conditions are existing. Cloud formations will give a clear indication of this.
classifications of clouds
Clouds can occur at any level of the atmosphere wherever there is sufficient moisture to allow condensation to take place. The layer of the atmosphere where almost all cloud exists is the troposphere, although the tops of some severe thunderstorms occasionally pierce the tropopause.
Because of the large range in temperatures and air movement in the troposphere, clouds vary in structure and composition (a combination of ice crystal and water). Consequently, clouds are classified into three main groups: lower, middle and high level clouds.
Higher level clouds

Higher level clouds represent the cloud in the highest levels of the troposphere. They mostly appear brilliant white because of the ice crystals at that level. They tend to develop at or just above the top part of the troposphere. Higher level clouds can vary in shape, thickness and cover.
Sunlight can be observed passing through the higher level clouds most of the time. The amount of light that penetrates depends on the density and thickness of the layers. The thickness of such clouds are therefore relatively thin.
In most cases, the direction of movement of the higher level clouds do not necessarily represent the wind direction at the ground level. In fact, the wind at upper and ground levels often differ.
There are three main types of higher level clouds: cirrus, cirrostratus and cirrocumulus.
The bases of high clouds range from 16,500 feet to 45,000 feet and average about 25,000 feet in the temperate regions.
Cirrus and cirrostratus (Cl)(Cs)
Since the characteristics of cirrus and cirrostratus are similar, they can be discussed together including any differences.
Cirrus clouds are higher level clouds that develop in filaments or patches. They are virtually brilliant white attributed to their ice crystal composition. However, they lack in contrast between the top and base. They occur in flat sheets with a low height to base ratio and are usually isolated with large breaks of sky. Cirrus also vary dramatically in 'shape' or patterns they portray but these represent the fluctuating wind flow at that level both in the horizontal and vertical direction.

Cirrostratus represent clouds that are more widespread than cirrus but containing some similar features. Like cirrus, they are brilliant white and lack in contrast. Sunlight can pass through cirrostratus but this again depends on the varying thickness of the clouds.
Both cirrus and cirrostratus clouds vary in thickness. The sun can easily be observed through both types of clouds although the intensity of light that is observed depends on the thickness of their layers. In their thickest form, cirrus and cirrostratus will allow a similar intensity of light to pass through to that of thin altostratus. They do not only develop in one complete layer. It may be difficult to observe because of the lack of contrast but these clouds can consist of several thin layers.
Cirrus and cirrostratus tend to move in the direction of the wind at that level which differ to that at the surface. The most common direction of motion of these clouds are from a westerly direction. This varies with factors such as the latitude, weather conditions and time of the year. Their apparent velocities are relatively slow as compared to lower clouds.
Both cirrus and cirrostratus can occur in conjunction with any of the other cloud types. Obviously, all the lower and middle level clouds will obscure the view of the higher level clouds, appear to move faster and appear less defined. They can only be observed above other clouds when breaks in the clouds occur. Any type of higher level clouds can develop simultaneously.
Cirrus clouds tend to develop on days with fine weather and lighter winds at the surface. cirrostratus can develop on days with light winds but normally increasing in strength. Although both cirrus and cirrostratus tend to develop in fine weather conditions, they also acts as a sign of approaching changes in the weather conditions. Such changes could include any of the various types of cold front situations, thunderstorms or developing and advancing troughs of low pressure, normally with preceding cloud masses.
Except in the latter case, cirrus and cirrostratus will typically precede any other types of clouds as part of a cloud band. In fact cirrus normally precedes cirrostratus. Nevertheless, the higher level clouds will persist until the actual change in the weather occurs. The higher clouds can develop from a few hours up to a few days before an actual change in the weather conditions occurs. They may develop during one afternoon and dissipate but redevelop the next day and so on until the actual change occurs. If the amount of moisture in the lower layers of the atmosphere increases, other lower clouds may also develop changing the appearance of the cirrus or the cirrostratus clouds as well as partially or totally hiding them from view. The same situation occurs in the case where cirrus develop ahead of thunderstorms. Cirrus normally precede cirrostratus which are then followed by the anvil of the approaching thunderstorm. In fact, cirrus and cirrostratus in this case are the remnants downwind of the weakening anvil.
Both cirrus and cirrostratus can develop and persist after a change has passed through a certain location. In this situation, cloud will decrease within a few hours up to a few days following the change. If it persists for longer periods, a jet stream cloud mass may be involved.
Another situation where cirrus and cirrostratus can be observed is when lower cloud breaks or clears during days with showers or rain. This case is far less common but can indicate a few situations. The higher clouds may be the remnants of the cloud mass that produced the actual wet weather. They

can also be developing ahead of other cloud masses associated with another system, leading to the situation already discussed above. It all depends on the weather situation at that time but the observation of the movement of the higher level clouds can be critical in determining what weather may follow.
Cirrus generally does not produce precipitation except when it results from dissipating thunderstorms.
Precipitation from such cirrus usually consists of larger droplets and the cloud normally dissipates and vanishes completely. cirrostratus does not produce precipitation.
Cirrus and cirrostratus can develop and persist at any time of the day despite the perception that it tends to occur during the day. This perception arises because it is much easier to observe cirrus during the day as compared to night time. The background darkness and the fact that the stars can easily be observed through cirrus and cirrostratus as thin layers allows them to camouflage from the view of the observer.
Cirrocumulus
Cirrocumulus is a higher level cloud that is brilliant white but with a spotty appearance created by the many small turrets. The turrets indicate vertical turbulence within the cloud. Despite this spotty appearance, cirrocumulus contains many features associated with cirrostratus discussed above. It moves in directions similar to that of the other higher clouds.
This cloud can develop in conjunction with any other clouds as well as with cirrostratus clouds. In
Sydney, cirrocumulus is not as common as the other high clouds and mainly develops during the winter times with west to south westerly air streams. The development of cirrocumulus sometimes occurs in conditions similar to those associated with the development lenticular altocumulus. cirrocumulus clouds do not produce precipitation and are normally associated with fine weather. middle clouds
Middle level clouds are those clouds that develop in the middle layers of the atmosphere. These clouds are brighter and less fragmented in appearance due to their distance from the ground and the higher composition of ice crystals. Middle level clouds vary in thickness from relatively flat sheets of cloud to a more cumuliform appearance. In fact, the sun (and moon) can be observed through some thin middle level clouds.
Middle level clouds tend to have apparent speeds slower than the lower level clouds. (Recall the larger radius and associated arc length that the higher level clouds must undertake). They move in the direction of the wind at that level which does not necessarily be the same of that at the surface.

There are 3 basic types of middle level clouds: altocumulus, altostratus and nimbostratus. The bases of middle clouds range from 6500 feet to 23,000 feet.
Altocumulus (Ac)
As the name suggests, altocumulus refers to the middle level cloud that exhibit to some extent the features normally associated with cumulus. This includes cumuliform tops and bases that are usually relatively darker than the tops. This cloud type can be widespread or patchy depending on the conditions. It can vary in appearance from broken to smooth, and vary in thickness.
In its broken form, altocumulus can be confused with stratocumulus. To distinguish between them requires examining how defined the cloud appears, whether there other forms of middle level or upper level clouds present above the layer and the difference in brightness. Like stratocumulus, the breaks become more visible at a steeper angle of elevation.
If conditions are unstable in the middle level of the atmosphere, the air will tend to rise in currents allowing areas of cumuliform turrets to develop. In fact, altocumulus can develop from dissipating thunderstorms during the morning and then redevelop during the day if the air remains unstable.
Altocumulus clouds therefore in this form indicate unstable or unsettled conditions.
Altocumulus can vary in its apparent movement (speed) depending on the wind and direction at that level. However, since altocumulus (like most other cloud types) represents an ever changing system, an observer must be careful in determining cloud motion. On some days, altocumulus continuously develop as it moves in the direction of the wind. Upstream, more altocumulus may develop giving the impression that the cloud is progressing slower than its actual speed. This process can occasionally create an illusion in terms of direction. Considering an example of altocumulus observed moving to the south east, because of development on the north and north-eastern side of the cloud band, the apparent direction may be more to the east.

Altocumulus can develop in more than one layer and also in conjunction with other cloud types. The lower layer will obscure part or all of the higher altocumulus cloud layer. This situation also applies to higher level clouds. Higher level clouds will be obscured by the altocumulus. Lower level clouds, however, will obscure part or all of the altocumulus cloud layer. In fact, it may be impossible to observe altocumulus above a full stratocumulus, stratus or lower level nimbostratus cover. If a break occurs, altocumulus can only be distinguished by its different (slower) speed and direction of movement.
Altocumulus also develop within the structure of cumulonimbus (thunderstorm producing) clouds. The appearance of altocumulus within thunderstorms vary depending on the structure, severity and the amount of moisture drawn into the thunderstorm. The altocumulus usually develops after the anvil
(consisting of cirrostratus and altostratus) develops and becomes darker as the precipitation cascade approaches. However, on days where thunderstorms develop with widespread altocumulus conditions, the altocumulus obscures the thunderstorms and its development observed only through breaks in the cloud.
If altocumulus develops into thicker layers, precipitation can develop. The intensity of rainfall most often expected from altocumulus is light to moderate rainfall. If large cumulus develop amongst the rain bearing altocumulus, then heavier rainfall will develop. On days when precipitation from altocumulus becomes widespread and continuous, the cloud forms a smooth lighter-grey shaded sheet and becomes known as nimbostratus (at the middle layers).
Precipitation within altocumulus can develop rapidly at the rear even though the cloud may be moving fairly rapidly. This will obviously influence the duration of rainfall as well as the normally large cloud base. This situation often occurs before a cold front with unstable conditions.
Thunderstorms can develop amongst the altocumulus band or they may develop after the cloud band clears well ahead of the actual change.
As discussed in the case of other clouds, lower clouds may be present below the altocumulus layer but not producing the rain. The observer again must consider which cloud is producing the rain to determine in which direction it is moving.
Another form of altocumulus is the lenticular type where the altocumulus appears in the form of a lens. They appear very smooth and flat, often displaying two or more layers. This occurs due to a wave effect in the air flow. This wave effect normally develops as a result of a mountain range on windy days. The wave effect forces air to rise above the condensation level and hence allows cloud to form. Due to the rise and fall effect (peaks and troughs), the cloud may only exist in areas of peaks and therefore appear patchy. The most striking feature of this cloud is that it tends to remain relatively stationary compared with the associated wind at that level. What is actually happening is that as the air begins to rise above the level of condensation, cloud forms. When the air falls below this level, dissipation occurs and the cloud disintegrates back to clear air. So long as the peak of the air wave remains stationary as compared to the ground, the cloud will develop and dissipate almost in the same position whilst the wind conditions persist.
The direction of the wind associated with lenticular altocumulus can be determined by considering the sharpest edge as the end of the cloud where the air is flowing in and the opposite end where the air is flowing out. Sometimes this will be the longest span of cloud. The most efficient method of

determining the direction of the wind is by closely examining the direction that the patterns and ripples within the cloud base move. The cloud will also be moving in the direction of the wind within the cloud region.
Lenticular altocumulus is generally not associated with precipitation. The conditions associated with the development of this cloud involves more horizontal rather than vertical flow. The air masses are also more stable and drier.
Lenticular altocumulus mostly develop during the day when the atmosphere is most lively in terms of strong winds at that level. The wind conditions at the surface are often very similar to the direction of wind at the cloud level. In the case of Sydney, lenticular altocumulus tend to develop during the morning period and clear off the coast during the evening. Almost all lenticular altocumulus in Sydney develop under the influence of south westerly, westerly or north westerly air streams associated with cold fronts.
Altocumulus can also develop in the form of ripples. In this case, the altocumulus cloud appears broken but lined as a result of minor wind wave ripples. In fact it develops in conditions associated with the development of lenticular altocumulus. This type of cloud obviously does not produce precipitation.
Altocumulus can develop from the spreading out of the tops of cumulus. The spreading out occurs as the tops of the cumulus grows until it reaches an inversion layer (or stronger winds that cause divergence) situated in the middle levels of the atmosphere. Because the cumulus updraughts are not strong enough to pierce this layer, the tops begin to spread in the form similar to that of an anvil facing in the direction of the wind at that level. Occasionally, this situation may further develop into thunderstorm or thundery shower conditions.
Altocumulus Castellanus (Acc)
Altocumulus with a turreted appearance. Instability is a characteristic. Altocumulus castellanus may develop into cumulonimbus. (below)
Altostratus (As)
Altostratus refers to middle level cloud that appears as a flat, smooth dark grey sheet. These clouds are most often observed as large sheets rather than isolated areas. However, in the process of development, altostratus may develop in smaller filaments and rapidly develop to larger sheets.
These types of clouds in certain conditions normally indicate an approaching cloud mass associated with a cold front, a trough system or a jet stream.

Altostratus can develop into a thick or thin layer. As a thin layer, the sun can be observed through the cloud. In its thinner form, the developing altostratus can sometimes be confused with approaching cirrostratus. In its thicker form, the sun can only occasionally be observed through the thinner sections if not at all. Obviously, the thicker the altostratus, the darker it becomes. When observed closely, it becomes apparent that altostratus is not just one complete layer but a composition of several thin layers.
Altostratus can produce precipitation. It will normally develop and then thicken. The precipitation is observed as relatively thick dark sections since precipitation cascades are very difficult to observe with the same colour in the background. In this situation, rain will develop as a light shower and gradually increase to showers, light rain or moderate rain. If the precipitation becomes persistent, the cloud then becomes known as nimbostratus. The duration of the precipitation is influenced by factors similar to those discussed with other types of clouds.
In certain conditions, altostratus will develop during the afternoon period and increase to cover most or all of the sky. By late afternoon, evening or during the night, precipitation will develop. This situation is the most common observation that occurs in Sydney. However, altostratus can develop at any time as well as the associated precipitation.
As discussed above, altostratus can develop in conjunction with other clouds such as cirrostratus, altocumulus and stratocumulus. Obviously, the lower clouds will obscure the view of altostratus, appear to move faster and appear less defined. Although altocumulus is a middle level cloud, it will develop below altostratus. Sometimes, altocumulus can be observed developing from dissipating altostratus. cirrostratus can often be observed above altostratus when it does not cover the sky. On days where altostratus is observed above a stratocumulus cover, it may indicate a trough with possible rain or even thunderstorms either during the afternoon or within the next few days.
Like altocumulus, altostratus also forms part of thunderstorms normally within or below the lower part of the anvil region. Of course this depends on the height of the thunderstorm anvil. Different structures of thunderstorms display various forms of altostratus. As the anvil of the thunderstorm passes overhead, the altostratus begins to appear normally with a grey base but becoming increasingly dark.
Some altostratus develop in situations similar to the development of lenticular altocumulus.
Altostratus in this form develops in large sheets and has a patchy base appearance. The cloud seems to be moving rapidly but because of its development at the rear actually progresses very slowly in the direction of the wind at that level. This type of cloud does not produce any rain.
Nimbostratus
Nimbostratus can be described as a widespread light grey or white sheet of cloud that produces persistent rain or showers. Because of its light colours, nimbostratus lacks contrast and in fact is quite difficult to photograph. Being sufficiently thick to produce precipitation, the sun or moon can rarely

be observed through nimbostratus. The cloud may be more than 15,000 feet thick. It is generally associated with warm fronts.
Because of its lack of contrast, it is difficult to determine the apparent speed and direction of nimbostratus. This speed can sometimes be determined by observing the movement of a break in the cloud or observing the cloud's motion against the occasional glimpse of the sun or the moon that is relatively motionless. Another method involves the observation of approaching intermittent showers although patterns of precipitation can sometimes change dramatically.
Generally, precipitation associated with nimbostratus is long in duration. The intensity can vary from light to heavy depending on the associated conditions. Normally, light to moderate rain is associated with nimbostratus. However, the passage of strong lows and cold fronts can produce moderate to heavy precipitation. In Sydney, weather associated with flooding rains often contains thick nimbostratus layers.
As discussed in earlier cases, nimbostratus can develop or occur with most other types of clouds.
Stratus and stratocumulus will often develop below nimbostratus in its middle level form and obscure the view of the whole cloud base. With approaching precipitation regions, the lower clouds may appear darker or lighter than the nimbostratus creating some contrast. This depends on the intensity of the background nimbostratus. The movement of the lower clouds do not necessarily have to be the same as the nimbostratus.
Although stratocumulus clouds can develop below nimbostratus, they can also thicken to develop into a nimbostratus layer with precipitation. This refers to nimbostratus in its lower levels of the atmosphere. It can be difficult to distinguish this from nimbostratus in the middle levels of the atmosphere. It often requires observation of the initial cloud (stratocumulus or altostratus) or the cloud that follow. Another useful method is measuring the apparent speed of the cloud if it can be observed. Of course, the lower the cloud, the less likelihood that lower clouds will be observed below the nimbostratus.
Nimbostratus can develop from altostratus if it becomes sufficiently thick to produce precipitation. In fact, increasing altostratus cloud tends to lead to nimbostratus. Generally, the altostratus will become darker and gradually rain will develop. This sometimes leads to a lighter appearance of the cloud base although the cloud still remains reasonably thick.
Lower level nimbostratus can develop below altostratus and partially or completely obscure it from view. However, if the altostratus layer develops into nimbostratus itself, the lower level nimbostratus will most probably become difficult to see especially if precipitation begins to fall.
The weather conditions that produce middle level (and sometimes lower level) nimbostratus also lead to the development of higher level clouds. Nimbostratus developing or occurring below higher level clouds will obscure most or all of it from view. The higher clouds can only be observed through breaks of the nimbostratus if and when they occur. These breaks often occur when the cloud is decreasing in intensity and conditions are beginning to clear.

low clouds
Lower level clouds consist of those clouds in the lower layers of the atmosphere. Because of the relatively low temperatures at this level of the atmosphere, lower level clouds usually reflect lower amounts of light and therefore usually exhibit low contrast. The clouds at this level also appear not as well defined. When observed closely, it is easy to observe the turbulent motions and hence the everchanging structure.
Being closer to the ground, lower level clouds appear to move or progress faster than other clouds.
The clouds generally move in the direction of the wind very similar to the direction of the wind on the ground.
The most efficient method used to recognise lower clouds is when observed in conjunction with other clouds. The lower clouds will obscure part or all the view of the upper level clouds if they pass in between the observer's line of sight. In other words, all the observer can see is the lower clouds as well as parts of the higher level clouds through breaks of the lower clouds. What is observed will vary due to the different directions and relative wind speeds associated with the different layers of clouds.
There are 3 main types of lower level clouds: cumulus, stratocumulus and stratus.
The bases of low clouds range from surface height to about 6500 feet.
Stratus (St)
Stratus is defined as low cloud that appears fragmented and thin. It can also occur in the form of a layer or sheet. The sun, moon and generally the sky can usually be observed through stratus clouds, especially at a steep angle of elevation. Stratus lacks the vertical growth of cumulus and stratocumulus, and therefore lacks the contrast. This is more evident when observed as one layer as compared to patchy stratus. Being closest to the ground, stratus clouds normally move fairly rapidly in the direction of the wind depending of course on the wind speed.
Like stratocumulus, stratus develops under several conditions or weather situations. Stratus mostly develop under the influence of wind streams where moisture condenses in the lower layers of the atmosphere. Wind changes during the summer months often lead to the development of stratus as the wind evaporates moisture from the ocean and condensing as turbulence mixes the surface air with the cooler air above. In these conditions, stratus develop in patches and gradually may become widespread forming into stratocumulus.
On days with nimbostratus and rain, stratus cloud develop simply due to the amount of moisture in the air. With light winds, stratus are normally observed in sheets. In stronger wind conditions, stratus develops in patches, similar in appearance to stratocumulus. Both the direction and appearance of stratus can change rapidly with changing weather conditions. It can clear and redevelop several times

during certain conditions usually appearing when rain approaches, and clearing as the rain clears.
Being the lowest cloud layer, it obscures at least partially the view of stratocumulus or other types of clouds above.
Stratus, like stratocumulus, can develop in weather conditions associated with thunderstorms and thunderstorm development. In this case, stratus is observed moving rapidly towards the storms and thickening in the region of the updraughts, especially those of severe thunderstorms. The stratus is only the visible condensed water vapour feeding into the thunderstorm. One good example of a thunderstorm illustrating this behaviour is the violent hailstorm that occurred on the 18th of March,
1990 in Sydney (This storm is not illustrated here). Earlier in the day, stratus had developed with a south to south-easterly change and was moving rapidly with the air stream. As the thunderstorms developed and approached, the stratus thickened to form stratocumulus. As the storm (which was a supercell) with the updraught region moved almost overhead, the stratocumulus cleared rapidly. The major rain band then moved through with strong winds, heavy rain and medium to large hail in some areas.
Stratus can develop in the various types of weather conditions associated with stratocumulus discussed above. However, the characteristics of stratus do not vary as much as stratocumulus and therefore they are easily distinguishable. Therefore, there is no real need to discuss further the weather conditions associated with stratus clouds.
Stratocumulus (Sc)
Stratocumulus are low clouds that generally move faster than cumulus and are not as well defined in appearance (recall the techniques of observing clouds). They tend to spread more horizontally rather than vertically. Like cumulus, the bases of stratocumulus are normally darker than the tops. However, they can vary in terms of characteristics.
Depending on the weather conditions, stratocumulus can appear like cumulus since stratocumulus can develop from cumulus. They may also appear as large flat areas of low, grey cloud. Sometimes stratocumulus appear in the form of rolling patches of cloud aligned parallel to each other.
Stratocumulus can also appear in the form of broken clouds or globules. The sun, moon and generally the sky can be observed through the breaks in broken stratocumulus clouds. Of course, this depends on how large the breaks are, how high the clouds rise and the angle of elevation of the breaks with respect to the observer. This generally applies to all clouds but is more notable with clouds in broken form.
Stratocumulus mostly develop in wind streams moving in the direction of the wind similar to the direction of the wind at the surface. The friction created by the earth causes turbulence in the form of eddies. With sufficient moisture, condensation will occur in the lower layers of the atmosphere visible as clouds. The amount of stratocumulus covering the sky depends on the amount of moisture concentrated at that level of the atmosphere. The speed that the cloud moves varies according to the wind speed at that level.

Stratocumulus cloud also can develop in the form of lenticularis. The only method that can be used to distinguish between these clouds is that stratocumulus will not appear as well defined, will tend to move more quickly. Sometimes they develop below cumulus or cumulonimbus which means that it must be low cloud.
Nimbostratus (Ns)
A low layer of uniform, dark grey cloud. When it gives precipitation, it is in the form of continuous rain or snow. The cloud may be more than 15,000 feet thick. It is generally associated with warm fronts.
Little turbulence occurs in stratus. The low cloud bases and poor visibility make VFR operations difficult to impossible.
clouds of vertical development
The bases of this type of cloud may form as low as 1500 feet. They are composed of water droplets when the temperature is above freezing and of ice crystals and supercooled water droplets when the temperature is below freezing.
Cumulus ( C u )
Cumulus are cauliflower-shaped low level clouds with dark bases and bright tops. When observing cumulus, you are actually observing the condensation process of rising thermals or air bubbles at a certain level in the atmosphere known as the condensation level.

The air rising above this level condenses and cloud is observed. Since the height of this level is fairly constant at any particular time, then the bases of cumulus are usually flat.
The appearance of cumulus like other clouds can be illusive. If stratus formed at the same level as cumulus, the cumulus will appear different observed from different perspectives with respect to the sun's position. (If light from the sun must reflect to get to the observer, then the cloud will tend to appear brighter and display more contrast than cloud reflecting very little direct sunlight. In fact, the latter case indicates that the shadow area of the cloud is facing the observer). A similar situation may occur when observing cumulus below a much darker background such as a thunderstorm. The cumulus clouds appear as a uniform white or at least much lighter with little or no contrast. The same cumulus clouds observed away from this cloud band will appear darker, with more contrast.
With practice, an observer can easily determine the size of cumulus clouds (or any clouds in general) by considering the following factors; their apparent distances, coverage of the sky (density), their angle of elevation (how much of their base can be observed), how much overlapping occurs, and their base to height ratio. Cumulus often occurs in conjunction with other clouds and may vary in appearance. If cumulus is observed below other clouds, the shadow effect of other clouds can decrease contrast of the cumulus.
Towering Cumulus (TCu)
Cumulus clouds that build up into high towering masses. They are likely to develop into cumulonimbus. Rough air will be encountered underneath this cloud. Heavy icing may occur in this cloud type. (below)

Cumulonimbus (Cb)
Heavy masses of cumulus clouds that extend well above the freezing level. The summits often spread out to form an anvil shaped top that is characteristic of thunderstorm. (below) cloud formation
Generally upward motion of moist air is a prerequisite for cloud formation, downward motion dissipates it. Ascending air expands, cools adiabatically and, if sufficiently moist, some of the water vapour condenses to form cloud droplets. Fog is likely when moist air is cooled, not by expansion but by contact with a colder surface.
The diameter of the condensation nuclei is typically 0.02 microns but a relatively small number may have a diameter up to 10 microns. Maritime air contains about one billion nuclei per cubic metre, polluted city air contains many more. The diameter of a cloud droplet is typically 10 to 25 microns and the spacing between them is about 50 times diameter, perhaps one mm, with maybe 100 million droplets per cubic metre of cloud. The mass of liquid in an average density cloud approximates 0.5 gram per cubic metre.

Above the freezing level in the cloud some of the droplets will freeze if disturbed by contact with suitable freezing nuclei, or an aircraft. Freezing nuclei are mainly natural clay mineral particles, bacteria and volcanic dust, perhaps 0.1 microns in diameter. There are rarely more than one million freezing particles per cubic metre thus there are only sufficient to act as a freezing catalyst for a small fraction of the cloud droplets. Most freezing occurs at temperatures between –10 °C and –15 °C.
The balance of the droplets above freezing level remains in a supercooled liquid state, possibly down to temperatures colder than –20 °C, but eventually, at some temperature warmer than –40 °C, all droplets will freeze by self-nucleation into ice crystals, forming the high level cirrus clouds. In some cases fractured or splintered ice crystals will act as freezing nuclei. The ice crystals are usually shaped as columnar hexagons or flat plate hexagons, refer 3.5.2 below and 12.2.2.
Condensation of atmospheric moisture occurs when: the volume of air remains constant but temperature is reduced to dewpoint, e.g. contact cooling, mixing of different layers the volume of an air parcel is increased through adiabatic expansion evaporation increases the vapour partial pressure beyond the saturation point a change of both temperature and volume reduces the saturation vapour partial pressure.
precipitation
Rain [RA] and drizzle [DZ]
Cloud droplets tend to fall but their terminal velocity is so low, about 0.01 metres/sec, that they are kept aloft by the vertical currents associated with the cloud construction process, but will evaporate when coming into contact with the drier air outside the cloud. Some of the droplets are larger than others and consequently their falling speed is greater. Larger droplets catch up with smaller and merge or coalesce with them eventually forming raindrops. Raindrops grow with the coalescence process reaching maximum diameters, in tropical conditions, of 4 – 7 mm before air resistance disintegrates them into smaller raindrops, which start a self perpetuating process. It takes about one million cloud droplets to form one raindrop.
The terminal velocity of a 4 mm raindrop is about 9 metres/sec. Only clouds with extensive depth,
3000 feet plus, will produce rain (rather than drizzle) but very small high clouds, generating heads, may produce trails of snow crystals which evaporate at lower levels – fall streaks or virga.
Drizzle forms by coalescence in stratiform clouds with depths possibly less than 1000 feet and with only weak vertical motion, otherwise the small ( 0.2 – 0.5 mm) drops would be unable to fall. It also requires only a short distance or a high relative humidity between the cloud base and the surface, otherwise the drops will evaporate before reaching the surface. Terminal velocity approximates 1 – 2 metres/sec.
Light drizzle [–DZ]: visibility greater than 1000 metres
Moderate drizzle [DZ]: 500 to 1000 metres
Heavy drizzle [+DZ]: less than 500 metres
Light rain showers: precipitation rate under 2.0 mm/hour
Moderate rain showers: 2.0 to 10 mm/hour
Heavy rain showers: more than 10 mm/hour
Light rain [–RA]: under 0.5 mm/hour, individual drops easily seen
Moderate rain [RA]: 0.5 to 4 mm/hour, drops not easily seen
Heavy rain [+RA]: more than 4 mm/hour, rain falls in sheets
Weather radar reports precipitation into six reflectivity levels: light precipitation light to moderate rain moderate to heavy rain

heavy rain very heavy rain, hail possible very heavy rain and hail, large hail possible
Scotch mist is a mixture of thick cloud and heavy drizzle on rising ground, formed in conditions of weak uplift of almost saturated stable air.
Snow [SN]
At cloud temperatures colder than –10 °C where both ice and supercooled liquid water exist, the saturation vapour pressure over water is greater than that over ice. Air that is just saturated with respect to the supercooled water droplets will be supersaturated with respect to the ice crystals, resulting in vapour being deposited onto the crystal. The reduction in the amount of water vapour means that the air is no longer saturated with respect to the water droplets and, to achieve saturation equilibrium again, the water droplets begin to evaporate. Thus ice crystals grow by sublimation and water droplets lessen, i.e. in mixed cloud the ice crystals grow more rapidly than the water droplets. Snow is frozen precipitation resulting from ice crystal growth and falls in any form between small crystals and large flakes. This is known as the Bergeron-Findeison theory and probably accounts for most precipitation outside the tropics. Snow may fall to the surface or, more often, melt below the freezing level and fall as rain.
Snowflakes are built by snow crystals colliding and sticking together in clusters of several hundred – aggregation. Most aggregation occurs at temperatures just below freezing, the snow crystals tending to remain separate at colder temperatures.
Hail and other ice forms
The growing snow crystals acquire a fall velocity relative to the supercooled droplets and growth also continues by collision and coalescence with supercooled droplets forming ice pellets [PE], the process being termed accretion,or opaque riming if the freezing is instantaneous incorporating trapped air, glazing if the supercooled water freezes more slowly as a clear layer. The ice pellets in turn grow by coalescence with other pellets and further accretion and are termed hail [GR] when the diameter exceeds 5 mm. The size reached by hailstones before falling out of the cloud depends on the velocity and frequency of updraughts within the cloud. Hail is of course an hazard to aviation, particularly when it is unexpected, for example hail falling from a Cb anvil can appear to fall from a clear sky. Snow grains [SG] are very small, flattened, opaque ice grains, less than 1 mm and equivalent to drizzle. Snowflakes that, due to accretion, become opaque, rounded and brittle pellets,
2 – 5 mm diameter, are called snow pellets or graupel [GS]. Sleet is transparent ice pellets less than
5 mm diameter that bounce on impact with the ground. Sleet starts as snow, partially melting into rain on descent through a warm layer, then refreezing in a cold near-surface layer. The term is sometimes applied to a snow/rain mixture or just wet snow. Diamond dust [IC] is minute airborne ice crystals that only occur under very cold (Antarctic) conditions.
When raindrops form in cloud top temperatures warmer than –10 °C the rain falls as supercooled drops. Such freezing rain or drizzle striking a frozen surface, or an aircraft flying in OAT at or below zero, will rapidly freeze into glaze ice. Freezing rain is responsible for the ice storms of North
America and northern Europe, but the formative conditions differ from the preceding.
The seeder – feeder mechanism
Any large scale air flow over mountain areas produces, by orographic effect, ice crystals in cold cloud tops. By themselves the falling crystals would cause only light drizzle at the ground. However as the crystals fall through the low level mountain top clouds they act as seed particles for raindrops formed by cloud droplet coalescence with the falling crystals, producing substantial orographic rainfall in mountain areas.
Aerial cloud seeding involves introducing freezing nuclei (silver-oxide crystals with a similar structure to ice crystals) into parts of the cloud where few naturally exist, in order to initiate the Bergeron-
Findeison process.

fog
Fog is defined as an obscurity in the surface layers of the atmosphere which is caused by a suspension of water droplets, with or without smoke particles, and which is defined by international agreement as being associated with visibility less than 1000 metres. If the visibility exceeds 1000 metres then the obscurity is mist – met. code BR.
Radiation fogs are the prevalent fogs in Australia, with occurrence peaking in winter; caused by lowering of ground temperature by re-radiation into space of absorbed solar radiation from the earth’s surface. Radiation fogs mainly occur in moist air on cloudless nights within a high pressure system, particularly after rainfall. The moist air closest to the colder surface will quickly cool to dewpoint with condensation occurring. As air is a poor conductor a light wind, 2 – 6 knots, will best facilitate the mixing of the cold air throughout the surface layer, creating fog. The fog itself becomes the radiating surface in turn, encouraging further cooling and deepening of the fog. An increase in atmospheric pollution products supplies extra condensation nuclei to enhance the formation of fog or
smog.
A low level inversion forms containing the fog which may vary from scattered pools in surface depressions to a general layer 1000 feet in depth. Calm conditions will result in a very shallow fog layer or just dew or frost. Surface winds greater than 10 knots may prevent formation of the inversion, the cooled air is mixed with the warmer air above and not cooling to dewpoint. If the forecast wind at 3000 feet is 25 knots plus the low level inversion may not form and fog is unlikely. In winter radiation fog may start to form in the evening and persist to mid-day, or later if the sun is cut off by higher level cloud and/or the wind does not pick up sufficiently to break up the low level inversion.
Advection fog may occur when warm, moist air is carried over a surface which is cooler than the dewpoint of the air. Cooling and some turbulence in the lower layer lowers temperature to dewpoint and fog forms. Sea fogs drifting into coastal areas are advection fogs forming when the sea surface temperature is lower than the dewpoint but with a steady breeze to promote air mixing. Dewpoint can be reached both by temperature reduction and by increased water vapour content through evaporation. Advection fogs will form in valleys open to the sea when temperature falls in the evening combined with a sea breeze of 5 to 15 knots to force the air upslope. Thick advection fogs may be persistent in winter, particularly under a mid-level cloud layer.
Shallow evaporation fogs or steaming fogs result from the immediate condensation of water vapour that has just evaporated from the surface into near saturated air. Steaming from a sun warmed road surface after a rain shower demonstrates the process. Sea smoke or frost smoke is an evaporation fog occurring in frigid Antarctic air moving over relatively warm waters and prompting evaporation into the cold air which, in turn, quickly produces saturation.
Freezing fog is a fog composed of supercooled water droplets which freeze on contact with solid objects, e.g. parked aircraft. When near saturated air is very cold, below –24 °C at sea level to –45 °C at 50 000 feet, the addition of only a little moisture will produce saturation. Normally little evaporation takes place in very cold conditions but release of water vapour from engine exhausts, for instance, can quickly saturate calm air, even though the engine exhaust heat tends to lower the relative humidity, and will produce ice fog at the surface or contrails at altitude. If the temperature is below –40 °C ice crystals form directly on saturation. Contrails persist if relative humidity is high but evaporate quickly if low. Distrails occur when the engine exhaust heat of an aircraft flying through a thin cloud layer dissipates a trail.
Frontal fog or rain-induced fog occurs when warm rain evaporates at surface level in light wind conditions and then condenses forming fog.
orographic lift
An airstream reaching a mountain barrier is forced to rise, both at the surface and the upper levels, and cools adiabatically. If the lift and the moisture content are adequate condensation occurs at the lifting condensation level and cloud is formed on or above the barrier. Stratus is formed if the air is stable, cumulus if the air is slightly unstable. If there is instability in depth, coupled with high moisture, CB may develop. Refer 3.6 below. Solar heating of mountain ridges causes the adjacent air to be warmer than air at the same level over the valleys, thus the ridge acts as a high level heat source, increasing buoyancy and accentuating the mechanical lifting.

Orographic cloud – cap cloud – in stable conditions tends to form continuously on the windward side, clearing on the lee side. Lenticular cloud may also form a high cap above a hill when there is a layer of near saturated air aloft, orographic lifting causing condensation, descent causing evaporation. A mountain wave may form, particularly in a sandwiched stable layer resulting in the formation of a series of lenticular clouds.
convection
Warm air rises. Owing to the heating of the ground by the sun, rising currents of air occur. The upward movement of air is known as convection. (The downward movement of air is known as subsidence. ) As currents of air rise due to convection, they expand. The expansion is accompanied by cooling. The cooling produces condensation' and a cumuliform cloud forms at the top of each rising column of air. The cloud will grow in height as long as the rising air within it remains warmer than the air surrounding it. The height of the cloud, however, is also dependent on the stability of the air in the mid levels of the troposphere. Convection also occurs when air moves over a surface that is warmer than itself. The air is heated by advection and convective currents develop. Warming of air by advection does not depend on daytime heating. Convection will, therefore, continue day or night so long as the airflow remains the same.
frontal lift
When a mass of warm air is advancing on a colder mass, the warm air rises over the cold air on a long gradual slope. This slope is called a warm frontal surface. The ascent of the warm air causes it to cool, and clouds are formed, ranging from high cirrus through altostratus down to thick nimbostratus from which continuous steady rain may fall over a wide area.
When a mass of cold air is advancing on a mass of warm air. The cold air undercuts the warm air and forces the latter to rise. The slope of the advancing wedge of cold air is called a cold frontal surface.
The clouds which form are heavy cumulus or cumulonimbus. Heavy rain, thunderstorms, turbulence and icing are associated with the latter.
convergence
Synoptic scale atmospheric vertical motion is found in cyclones and anticyclones, mainly caused by air mass convergence or divergence from horizontal motion. Meteorological convergence indicates retardation in air flow with increase in air mass in a given volume due to net three dimensional inflow. Meteorological divergence, or negative convergence, indicates acceleration with decrease in air mass. Convergence is the contraction and divergence is the spreading of a field of flow.
If, for example, the front end of moving air mass layer slows down, the air in the rear will catch up – converge, and the air must move vertically to avoid local compression. If the lower boundary of the moving air mass is at surface level all the vertical movement must be upward. If the moving air mass is just below the tropopause all the vertical movement will be downward because the tropopause inhibits vertical motion. Conversely if the front end of a moving air mass layer speeds up then the flow diverges. If the air mass is at the surface then downward motion will occur above it to satisfy mass conservation principles, if the divergence is aloft then upward motion takes place.
Rising air must diverge before it reaches the tropopause and sinking air must diverge before it reaches the surface. As the surface pressure is the weight per unit area of the overlaying column of air, and even though divergences in one part of the column are largely balanced by convergences in another, the slight change in mass content (thickness) of the over-riding air changes the pressure at the surface.
The following diagrams illustrate some examples of convergence and divergence:

Note: referring to the field of flow diagrams above, the spreading apart (diffluence) and the closing together (confluence) of streamlines alone do not imply existence of divergence or convergence as there is no change in air mass if there is no cross isobar flow or vertical flow. (An isobar is a curve along which pressure is constant and is usually drawn on a constant height surface such as mean sea
level.)
Divergence or convergence may be induced by a change in surface drag, for instance when an airstream crosses a coastline. An airstream being forced up by a front will also induce convergence.
For convergence / divergence in upper level waves. Some divergence / convergence effects may cancel each other out e.g. deceleration associated with diverging streamlines.
Developing anti-cyclones – “highs” and high pressure ridges, are associated with converging air aloft and consequent wide area subsidence with diverging air below . This subsidence usually occurs between 20 000 and 5000 feet typically at the rate of 100 – 200 feet per hour. The subsiding air is compressed and warmed adiabatically at the DALR, or an SALR, and there is a net gain of mass within the developing high. Some of the converging air aloft rises and, if sufficiently moist, forms the cirrus cloud often associated with anti-cyclones.
As the pressure lapse rate is exponential and the DALR is linear the upper section of a block of subsiding air usually sinks for a greater distance and hence warms more than the lower section and if the bottom section also contains layer cloud the sinking air will only warm at a SALR until the cloud evaporates. Also when the lower section is nearing the surface it must diverge rather than descend and thus adiabatic warming stops. With these circumstances it is very common for a subsidence inversion to consolidate at an altitude between 3000 and 6000 feet. The weather associated with large scale subsidence is almost always dry, but in winter persistent low cloud and fog can readily form in the stagnant air due to low thermal activity below the inversion, producing ‘anti-cyclonic gloom’. In summer there may be a haze layer at the inversion level which reduces horizontal visibility at that level although the atmosphere above will be bright and clear. Aircraft climbing through the inversion layer will usually experience a wind velocity change.

Developing cyclones, “lows” or "depressions", and low pressure troughs are associated with diverging air aloft and uplift of air leading to convergence below. There is a net loss of mass within an intensifying low as the rate of vertical outflow is greater than the horizontal inflow, but if the winds continue to blow into a low for a number of days, exceeding the vertical outflow, the low will fill and disappear. The same does not happen with anti-cyclones which are much more persistent.
A trough may move with pressure falling ahead of it and rising behind it giving a system of pressure tendencies due to the motion but with no overall change in pressure, i.e. no development, no deepening and no increase in convergence.
thunderstorms
Like CU, surface heating, may provide the initial trigger to create isolated CB within an air mass but the initial lift is more likely to be provided by orographic ascent or convergence effects.
In the formative stages of a CB the cloud may have an updraught pulse of 1000 – 2000 feet/min, the rising parcel of air reaches altitudes where it is much warmer than the surrounding air, by as much as
10 °C, and buoyancy forces accelerate the parcel aloft possibly reaching speeds of 3000 – 7000 feet/min. Precipitation particles grow with the cloud growth, the upper levels of the cloud gaining additional energy from the latent heat released from the freezing of droplets and the growth of snow crystals and hailstones. When the growth of the particles is such that they can no longer be suspended in the updraught, precipitation, and its associated drag downdraught, occurs.

If the updraught path is tilted, by wind shear or veer, rather than vertical, then the precipitation and its downdraught will fall away from the updraught rather than back down through it (consequently weakening, or stopping, the updraught) and a co-existing updraught/downdraught may become established. An organised cell system controlling its environment and lasting several hours may evolve.
Middle level dry air from outside the cloud is entrained into the downdraught of an organised cell.
The downdraught is further cooled by the dry inflow air evaporating some of its water and ice crystals and tends to accelerate downwards in vertical gusts and, at the same time, maintaining the higher horizontal momentum it gained at upper levels from the higher forward speed of the storm at that height. When the cold plunging air nears the surface the downburst spreads out in all directions as a cold gust front or squall, strongest at the leading edge of the storm, weakest towards the trailing edge.
Anvils may take several forms:
Cumuliform: forms when a very strong updraught spreads rapidly and without restriction.
Incus: a severe storm attains maximum vertical development when the updraught reaches a stable layer which it is unable to break through, often the tropopause, and the cloud top spreads horizontally in all directions forming an overhanging anvil.
Back-sheared: the cloud top spreads upwind, against the high level flow and indicating a very strong updraught.
Mushroom: a rollover or lip around the underside of an overhanging anvil indicating rapid expansion.
Overshooting top: a dome-like protusion through the top of an anvil indicating a very strong updraught pulse. The overshooting top in large tropical storms has been known to develop into a
'chimney' towering maybe 10 000 feet into the stratosphere with an extensive plume cloud extending downwind from its top. Such clouds transfer moisture to the stratosphere.
Each organised cell system contains an updraught / downdraught core beneath which is the outflow region containing the rain shield and bounded by the downdraught gust front, a flanking line with a dark flat base underneath which is the inflow region of warm moist air and a spreading anvil. The CU and TCU generated by the inflow within the flanking line are the genesis of new cells. Within the core the condensation of moisture from the inflow region produces rain, hail and snow and the associated downdraught to the outflow region. When the cool air outflow exceeds and finally smothers ,or undercuts and chokes off, the inflow the storm dissipates.
High moisture content in the low level air with dry mid level air plus atmospheric instability are required to maintain CB development. The amount of precipitation from a large storm is typically
200 000 tonnes but severe storms have produced 2 million tonnes.

t h u n d e r s t o r m h a z a r d s our thanks www.raa.asn.au
(Copyright John Brandon) hazards encountered in and near thunderstorms turbulence
The turbulence associated with clouds types is:
St – slight
Ci, Cs, Cc, Ac, As – nil or slight except when Ac cas or when merging into Cb
Sc – moderate
Ns – moderate but may be severe near base
Cu, TCu, Cb – Generally severe but may be catastrophic and include the downbursts described below and the internal up/downdraughts.
Thunderstorms may be classified in four generalised types – single-cell, isolated multicell cluster, multicell squall line (refer 9.6 below) and supercell, although supercells are also multi-cellular. Their associated surface winds may be both high velocity and extremely turbulent, originating from the downdraughts of cold, dense air. When thunderstorms are about light aircraft should not be
airborne.
Single-cell storms are usually isolated storms moving with the mid-level wind, common in summer and occurring in conditions where the wind velocity, relative to the cell motion, does not change markedly with height. (Cb development has to be strong to overcome the detrimental effects of vertical wind shear). A single cell storm may last less than 30 minutes, its life being limited to the growth and collapse of a single updraught pulse. The diameter of the storm may be less than one nautical mile and it will not move very far in its lifetime – less than 3 nm in light winds. Such storms do not usually produce violent wind shear near the surface although microbursts may descend from even a mild looking Cb.
Isolated single cell storms, embedded in low level cloud layers, commonly form in cold winter air streams. They are generally frequent, but short-lived, with soft hail and shallow wind gusts, and caused by de-stabilisation of the cold air mass. They can be accentuated by orographic effects.
Multicell cluster storms, the most common, consist of a series of updraught pulses (cells) which may be separated by time and/or distance and be closely or widely spaced but moving as a single unit.
They may cycle through strong and weak phases, strength being indicated by closeness of pulses.
Frontal, pre-frontal, heat-trough and convergence zone systems may produce very vigorous storms several miles wide and, by continually propagating new cells, last an hour or more before the cool outflow finally undercuts and chokes off, or smothers, the warm inflow and the system collapses.
Each new cell is usually formed in the zone of maximum convergence where the gust front directly opposes the low level wind.

Weaker multicell storms advance with or to the left of the prevailing mid-level wind ( i.e. that about
base height plus one third of the cloud depth) at an average rate of 10 knots or so, but the strongest storms may turn almost at right angles to the wind. Cells will move with the prevailing mid-level wind but because of the wind change with height the new cells form on the left, if the wind backs with height, and on the right if it veers. Thus the storm turns towards the flank where the new updraughts are building – the flanking line, which is a line of Cu or TCu stepped up to the most active cell. If the new cells are forming on the upwind side, usually to the west or north-west – a back-building storm – it may appear to move slowly, possibly staying in one place for considerable time.
Strong updraught / weak downdraught storms often form in conditions where there is moist air at most levels. Such storms produce heavy rain and may produce severe hail but, because of the lack of dry air inflow, severe low level shear is unlikely.
In severe storms, with strong updraughts and downdraughts, updraught velocities increase with height, typically 1500 feet per minute at 5000 feet and 3000 feet per minute at 20 000 feet.
Updraughts of 5000 feet per minute in the upper part of a storm are not unusual. Downdraught velocities tend to be slightly less at corresponding altitudes. Vertical acceleration loads of 2g to 3g may be experienced in horizontal flight.
The areas which most concern light aircraft are the low level outflow regions where downburst gusts of 50 knots plus may be reached in the line squall. The spreading density current of the outflow may last for 10 to 30 minutes and be 1500 to 6000 feet deep, forcing the warm, moist low level air up and so continuously regenerating the updraught. Thus an area up to 25 – 30 nm from a large storm, and
15 – 25 nm for a medium storm, should be regarded as absolutely a 'no-go' area for light aircraft.
An intense narrow initial microburst may sometimes be produced, bringing short-lived wind gusts of possibly 100 knots or more.
There is an area of extreme low level shear at the leading edge of the storm between the nose of the shelf cloud and where the gust front has reached, possibly 1 – 3 nm ahead of the rain curtain. Shear at the trailing edge is not quite as severe, as velocity there equals gust speed minus the speed of storm advance.
Vertical wind shear is usually detrimental to early development of Cb cells, however if there is:

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 strong vertical wind shear, backing and strengthening with height,
 associated with a deep surface layer of warm moist air, below a mid-level layer of dry air, for each 3000 feet.
with an inversion separating the layers, and a rapid decrease in temperature with height above the inversion, then the ideal conditions are created for a severe multicell storm. Or a supercell storm if the surface wind is greater than 20 knots and the vertical wind shear exceeds about five knots
The capping inversion keeps the lid on development until the lifting force builds up sufficiently to burst through the inversion and great buoyancy develops in the colder upper layer. Upper level divergence and a jetstream will also enhance the vertical motion.

Strong wind shear both tilts the updraught and provides the means to rotate it (storm updraughts
usually do not rotate) leading to the development of a supercell storm. A supercell is a severe storm with a continuing, organised strong main updraught with usually slight rotation (helicity) and coexisting strong downdraughts, controlling and directing the inflow ( which may have a velocity of 30 –
50 knots) into the cell from the surrounding atmosphere. It will usually diverge to the left of the prevailing mid-level wind.
There may be broad anti-clockwise rotation – as viewed from below – of the cloud base beneath the main updraught. Humid, rain cooled air from the downdraught may also be pulled into the normal inflow (which is often visible as scud beneath the Cb) causing part of the cloud base to lower, forming a circular wall cloud at the updraught base, and if vorticity increases within the cloud, a tornadic funnel may form. A gustnado may form under a shelf at the leading edge of the gust front.
Broadscale rotation of a storm cell forms a mesocyclone one to ten nm in diameter with a surface pressure drop of a few hPa at the centre although a 30 hPa drop has been recorded. Supercells may last for several hours as organised systems and commonly form in warm, moist north / north-east flow into a surface trough and along a dividing range during summer.
A microburst is a strong concentrated plunge of cold dense air from a convective cloud. Peak wind gusts usually last less than ten minutes, often 3 to 5 minutes, but extremely hazardous vertical and horizontal shear results. It may be “dry” or associated with precipitation ranging from virga showers to heavy rain showers – “wet”. A curling outflow foot of dust or precipitation from the surface touchdown point may be visible near the surface.
Microbursts are generally associated with hot and relatively dry conditions at low levels (such as found in inland Australia), convectively unstable moist air aloft with high (5000 to 15 000 feet) based
Cu or TCu. If the cloud is forming when the surface temperature/dewpoint spread is 15 °C to 25 °C then the microburst potential is high. The high spread means the atmosphere can hold much more water vapour. Rain falling in, and from, the cloud is evaporating (virga), thus cooling the entrained air and resulting in downward acceleration of the denser air. Consequently flight through, or under
or near, precipitation from a large Cu involves risk. Significant hail is unlikely. The most dangerous area is the horizontal density current vortex ring close to the touchdown point. The ring moves outward from the contact point at high velocity (up to 150 knots) until it disintegrates into several horizontal roll vortices spread around the periphery and which may continue to provide extreme shear for several minutes. The maximum horizontal winds occur about 100 feet above ground level.
Microbursts occur under 5% – 10% of Cb (refer 9.5 below) but a less concentrated, longer lasting gust front macroburst is normally associated with the entire cold air outflow of larger storm cells. The

severe gust fronts from a microburst extend for less than 4 km, those from a macroburst extend for more than 4 km. The vertical gusts within the downburst, perhaps with a velocity twice the mean, may produce a microburst within the macroburst.
Squall Lines
The precipitation downdraught associated with an individual cell tends to be concentrated towards the leading edge of the storm where the cold heavy outflow spreads out at ground level forming a small high pressure cell, a meso-high, 10 – 15 nm across. The dense air lifts the warmer, moist air in its path and may initiate a self amplifying convective complex, in which neighbouring storm cells consolidate into a towering squall line of large thunderstorm cells ranged across the prevailing wind direction. At locations in the path of the squall line the resultant line squall occurs as a sharp backing in wind direction, severe gusts, temperature drop, hail or heavy rain and possibly tornadoes. If the squall line is formed in an environment of strong mid-level winds the surface gusts may exceed 50 knots.
Squall lines vary in length, some of the longest being those which develop in a pre-frontal trough 50 -
100 nm ahead of a cold front. These squall lines may be several hundred nautical miles in length and
10 – 25 nm wide moving at typically 25 knots. The pre-frontal lines form ahead of the front as upper air flow develops waves ahead of the front; downward wave flow inhibiting and upward wave flow favouring, uplift.
During daylight hours the squall line may appear as a wall of advancing cloud with spreading cirrus plume but the most severe effects will be close to each of the numerous Cb cells. The convective complex releases a tremendous amount of latent heat and moisture which may be sufficient to generate a warm core mesoscale cyclone lasting several days.
heavy rain

microburst with heavy rain
Flight through rain causes a water film to form over the wings and fuselage; if the rain falls at a rate exceeding perhaps 20 mm per hour the film over the wings is roughened by the cratering of drop impacts and the formation of waves. The effect, which increases with rainfall rate, is a lowering of the lift coefficient value at all angles of attack, with laminar flow wings being most affected and fabric wings least affected. The stall will occur at a smaller angle of attack i.e. the stalling speed increases; which is further compounded by the increased weight of the aircraft.
The water film will increase drag and the encounter with falling rain will apply a downward/backward momentum which may be significant to a light aircraft. Propeller performance is degraded and water ingestion will affect engine output.
Thus the rain effect can be hazardous when operating in conditions of low excess aircraft energy – typically when taking off, landing or conducting a go-around. Visibility through a windscreen may be zero in such conditions, so a non IFR-equipped aircraft will be in difficulties.
Lightning
The electrostatic structure within Cb, or Cu con, is such that pockets of different charge exist throughout the cloud but, in 90% or more, with a main net positive charge residing on the cloud ice crystals in the upper part of the cloud and a main net negative charge, of similar magnitude, centred near the middle or lower part of the cloud at the sub-freezing level, the charge mainly residing on supercooled droplets. A smaller positive charge centre may exist at the bottom of the cloud where temperatures are above freezing. The electrostatic forces of repulsion / attraction induce secondary charge accumulations outside the cloud, a positive region on the earth’s surface directly below the cloud. Above the cloud positive ions are transferred away from and negative ions are transferred toward the cloud.
One favoured theory for the charge separation mechanism is the 'precipitation' theory which suggests

that the disintegration of large raindrops and the interaction between the smaller cloud particles and the larger precipitation particles in the up / downdraughts causes the separation of electrical charge, with downward motion of negatively charged cloud and precipitation particles and upward motion of positively charged cloud particles.
Discharge channels
Lightning is a flow of current, or discharge, along an ionized channel that equalizes the charge difference between two regions of opposite charge, occurring when the charge potentials exceed the electrical resistance of the intervening air. These discharges can be between the charged regions of the same cloud (intra-cloud), between the cloud and the ground (cloud-to-ground), between separate clouds (cloud-to-cloud) or between the base of a cloud and a charge centre in the atmosphere underneath it (cloud-to-air). The discharge channels, or streamers, propagate themselves through the air by establishing, and maintaining, an avalanche effect of free electrons which ionize atoms in their path. Lightning rates, particularly intra-cloud strokes, increase greatly with increase in the depth of clouds. Cloud-to-cloud and cloud-to-air discharges are rare but tend to be more common in the high based Cb found in the drier areas of Australia. Discharges above the Cb anvil into the stratosphere and mesosphere also occur.
When intra-cloud lightning – the most common discharge – occurs, it is most often between the upper positive and the middle negative centres. The discharge path is established by a “stepped leader”, the initial lightning streamer which grows in stages and splits into more and more branches as it moves forward seeking an optimal path between the charge centres. The second, and subsequent, lightning strokes in a composite flash are initiated by dart leaders, streamers which generally follow the optimum ionized channel established by the stepped leader. The associated electrical current probably peaks at a few thousand amperes. A distant observer cannot see the streamers but sees portion of the cloud become luminous, for maybe less than 0.5 seconds, hence 'sheet lightning'.
Cloud-to-ground discharges
Most cloud-to-ground discharges occur between the main negatively charged region and the surface, initially by a stepped leader from the region which usually exhibits branching channels as it seeks an optimal path. When the stepped leader makes contact, directly with the earth or with a ground
streamer, which is another electrical breakdown initiated from the surface positive charge region and which rises a short distance from the surface, the cloud is short-circuited to ground and to complete each lightning stroke a return streamer, or return stroke, propagates upwards. (The return streamer starts as positive ions which capture the free electrons flowing down the channel and emit photons.
The streamer carries more positive ions upward and their interaction with the free flowing electrons
gives the impression of upwards movement.) The charge on the branches of the stepped leader that have not been grounded flow into the return streamer. Subsequent strokes in the composite flash are initiated by dart leaders with a return streamer following each contact. The return streamer, lasting
20 – 40 microseconds, propagates a current carrying core a few cm in diameter with a current density of 1000 amperes per cm² and a total current typically 20 000 amps but peaks could be much greater.
A charged sheath or corona, a few metres in diameter, exists around the core. The stroke sequence of

dart leader / return streamer occurs several times in each flash to ground, giving it a flickering appearance. Each stroke draws charge from successively higher regions of the Cb and transfers a negative charge to the surface. Return streamers occur only in cloud-to-ground discharges and are so intense because of the earth’s high conductivity. Some rare discharges between cloud and ground are initiated from high surface structures or mountain peaks, by an upward moving stepped leader and referred to as a ground-to-cloud discharge. Rather rarely an overhanging anvil-to-ground discharge can be triggered by heavy charge accumulation in the anvil and the high magnitude strike can move many kilometres from the storm – a 'bolt from the blue'.
The temperature of the ionised plasma in the return streamer is at least 30 000 °C and the pressure is greater than 10 atmospheres, causing supersonic expansion of the channel which absorbs most of the dissipated energy in the flash. The shockwave lasts for 10 – 20 microseconds and moves out several hundred metres before decaying into the sound wave – thunder – with maximum energy at about 50 hertz. The shock wave can damage objects in its path. The channel length is typically 5 km and channel length can be roughly determined by timing the thunder rumble after the initial clap, e.g. a rumble lasts for 10 seconds x 335 m/sec = 3.3 km channel length. When a lightning stroke occurs within 150 m or so the observer hears the shockwave as a single high pitched bang.
Effect on aircraft instruments
The lightning discharges emit radio waves – atmospherics or ‘sferics – at the low end of the AM broadcast band and at TV band 1, which are the basis for airborne storm mapping instruments such as
Stormscope and Strikefinder. The NDB/ADF navigation aids also operate near the low end of the AM band so that the tremendous radio frequency energy of the storm will divert the radio compass needle. Weather radars map storms from the associated precipitation.
Strike effect on aircraft
When most aeroplanes, excluding ultralights, are struck by lightning the streamer attaches initially to an extremity, such as the nose or wing tip then re-attaches itself to the fuselage at other locations as the aircraft moves through the channel. The current is conducted through the electrically bonded aluminium skin and structures of the aircraft and exits from an extremity, such as the tail. If an ultralight is struck by lightning the consequences cannot be determined but are likely to be very unpleasant. Ultralights particularly should give all Cbs a wide berth but supercells and line squalls should be cleared by 25 – 30 nm at least.
Although a basic level of protection is provided in most light aeroplanes for the airframe, fuel system and engines, damage to wing tips, propellers and navigation lights may occur and the current has the potential to induce transients into electrical cables or electronic equipment. The other main area of concern is the fuel tanks, lines, vents, filler caps and their supporting structure, where extra design precautions prevent sparking or burn through. In heavier aircraft radomes, being constructed of nonconductive material, are at risk.
Red sprites and blue jets

When large cloud to ground lightning discharges occur below an extensive Cb cluster, which has a spreading stratiform anvil, other discharges are generated above the anvil. These discharges are in the form of flashes of light lasting just a few milliseconds and probably not observable by the untrained, naked eye but readily recorded on low light video.
Red sprites are very large but weak flashes of light emitted by excited nitrogen atoms and equivalent in intensity to a moderate auroral arc. They extend from the anvil to the mesopause at an altitude up to 90 km. The brightest parts exist between 60 – 75 km, red in colour and with a faint red glow extending above. Blue filaments may appear below the brightest region. Sprites usually occur in clusters which may extend 50 km horizontally. Blue jets are ejected above the Cb core and flash upward in narrow cones which fade out at about 50 km. These optical emissions are not aligned with the local magnetic field.
St. Elmo's fire is a plasma (i.e. a hot, ionized gas) that forms around the tips of raised, pointed conductors during thunderstorms. It is known as a corona discharge or point discharge to physicists.
The few people that have had the privilege of viewing an actual St. Elmo's fire have given various descriptions. It has been seen with different physical characteristics depending on the conditions of the viewing. It could be blue to bluish-white, silent to emitting a hissing sound, and ghostly to solid.
St. Elmo's fire occurs during thunderstorms - generally after the most severe part of the storm has passed - when the air reaches a very high voltage. These conditions are necessary to accumulate a charge large enough to create the phenomenon. It is always found attached to a grounded conductor with a sharp point; the most common are masts of sailing ships, church steeples, airplane wings or propellers, or even horns of cattle. The non-attached version of St. Elmo's fire is known as Ball
Lightning.
St. Elmo's Fire forms on aircraft flying through heavily charged skies, often as a precursor to a lightning strike. The glow can be seen concentrated on wing tips, antennae, the tail, nose and propeller blades when the potential difference is large enough. St. Elmo's Fire can be heard "singing" on the craft's radio, a frying or hissing sound running up and down the musical scale, according to some pilots.
A British Airways 747 flight was flying at night in the South Pacific just after a volcanic eruption of Mt.
Galunggung, in western Java. St. Elmo's fire was observed extending from the engines and sparking across the instrument panel, while smoke was smelled inside the passenger cabin. In minutes, the number four engine shut down, followed quickly by the remaining three. The aircraft glided from
37,000 feet to 12,000 feet before the crew was able to restart the engines and steer for an emergency landing at Jakarta. Ash appeared as smoke as it was sucked into the air conditioning system. The static electricity created by the ash hitting the windscreen and wings had created the St.
Elmo's fire.
Hail
Hail can cause considerable damage to aircraft and is usually encountered between 10,000 and 30,000 feet. At times it can also be found in clear air near thunderstorms.
Icing
High humidity and low winter freezing levels provide likely conditions for icing at low levels.
Hopefully it is unlikely that a VFR GA pilot would venture into possible icing conditions but pilots may be tempted to fly through freezing rain or drizzle. Aircraft cruising in VMC above the freezing level and then descending through a cloud layer may pick up ice.
The prerequisites for airframe icing are:
The aircraft must be flying through visible supercooled liquid, i.e. cloud, rain or drizzle
The airframe temperature, at the point where the liquid strikes the surface, must be sub-zero.
The severity of icing is dependent on the supercooled water content, the temperature and the size of the cloud droplets or raindrops. The terms used are:

Light: less than 0.5 grams per cubic metre of supercooled water in the cloud – no change of course or altitude is considered necessary for an aircraft equipped to handle icing. Very few light aircraft are equipped to handle any form of airframe ice.
vModerate: between 0.5 and 1.0 g/m³ – a diversion is desirable but the ice accretion is insufficient to affect safety if anti-icing / de-icing used, unless flight continued for an extended period.
Severe: more than 1.0 g/m³ – a diversion is essential. The ice accretion is continuous and such that de-icing / anti-icing equipment will not control it and the condition is hazardous.
The diagram below shows the ice accretion in mm on a small probe for the air miles flown, in clouds with liquid water content varying from 0.2 g/m³ to 1.5 g/m³.
T he small, supercooled droplets in stratiform cloud tend to instantaneous freezing when disturbed and form rime ice – rough white ice, opaque with entrapped air. In the stable conditions usually associated with stratiform cloud, icing will form where the outside air temperature [OAT] is in the range 0 °C to –10 °C . The continuous icing layer is usually 3000 to 4000 feet thick.
The larger supercooled droplets in convective cloud tend to freeze more slowly when disturbed by the aircraft; spreading over the surface and forming glossy clear or glaze ice . In unstable air moderate to severe icing may form where the OAT is in the range –4 °C to –20 °C . Where temperature is between –20 °C and –40 °C the chances of moderate or severe icing are small except in CB CAL i.e newly developed cells. Icing is normally most severe between –4 °C and –7 °C where the concentration of free supercooled droplets is usually at maximum, i.e. the minimum number have turned to ice crystals.. Mixed rime and clear ice can build into a heavy, rough conglomerate.
Flying through snow crystals or snowflakes will not form ice but may form a line of heavy frosting on the wing leading edge at the point of stagnation, which could increase stalling speed on landing.
Flying through wet mushy snow, which is a mixture of snow crystals and supercooled raindrops, will form pack snow on the aircraft.
The degree and type of ice formation in cloud genera are:
Ci, Cs and Cc; icing is rare but will be light should it occur
Ac, As and St; usually light to moderate rime
Sc; moderate rime
Ns; moderate to severe rime, clear ice or mixed ice. As the vertical extent of Ns plus As may be
15 000 or 20 000 feet the tops of the cloud may still contain supercooled droplets at temperatures as low as –25 °C
TCu and Cb; rime, clear or mixed ice, possibly severe.
Freezing rain creates the worst icing conditions, occurring when the aircraft flies through supercooled rain or drizzle above the freezing level in Cu or Cb. The rain striking an airframe, at sub-zero temperature, freezes and glaze ice accumulates rapidly, as much as one cm per four miles.
Freezing rain or drizzle occurring in clear air below the cloud base is the most likely airframe icing condition to be encountered by the VFR pilot and, as it is unlikely to occur much above
5000 feet amsl, descent choices are possibly limited.

Effect of airframe ice
Ice accretion on the wing leading edge is a major concern for aircraft not equipped with anti-icing or de-icing. Airflow disruption will reduce the maximum lift coefficient attainable by as much as 30% –
50%, thus raising the stalling speed considerably. and, because the aircraft has to fly at a greater angle of attack to maintain lift, the induced drag also increases and the aircraft continues to lose airspeed, making it impossible to sustain altitude if the stall is to be avoided. Fuel consumption will also increase considerably.
The weight of 25 mm of ice on a small GA aircraft would be about 30 to 40 kg but the increased weight is usually a lesser problem than the change in weight distribution. Also accretion is often not symmetrical, which adds to increasing uncontrollability.
Forward visibility may be lost as ice forms on the windshield.
Icing of the propeller blades reduces thrust and may cause dangerous imbalance.
Ice may jam or restrict control and trim surface movement or may unbalance the control surface and possibly lead to the development of flutter.
Communication antennae may be rendered ineffective or even snapped off.
Extension of flap may result in rudder ineffectiveness or even increase the stalling speed.
Aircraft operating from high altitude airfields in freezing conditions may be affected by picking up runway snow or slush which subsequently forms ice possibly causing engine induction icing, frozen brakes etc.
Engine air intake icing
Impact icing may occur at the engine air intake filter. If 'alternate air' (which draws air from within the engine cowling) is not selected, or is ineffective, power loss will ensue. When air is near freezing movement of water molecules over an object such as the air filter may sometimes cause instantaneous freezing. Ice may also form on the cowling intakes and cause engine overheating.
Pitot or static vent icing
Pitot or static vent blockage will seriously effect the ASI, VSI and altimeter, as shown in the table below, but be aware that blockage of the static vent tubing from causes other than icing, water for example, will render the ASI, VSI and altimeter useless, unless the aircraft is fitted with an alternative static source.
If the static vent is totally blocked by ice –
Flight stage Altimeter reading VSI reading ASI reading
During climb constant
During descent constant
During cruise +constant
On take-off constant zero zero zero zero under over
OK under
If the pitot tube is totally blocked –
Flight stage Altimeter reading VSI reading ASI reading
During climb no effect
During descent no effect
During cruise no effect
On take-off no effect no effect no effect no effect no effect over* zero* under* constant*
If the pitot tube is partially blocked –
Flight stage Altimeter reading VSI reading ASI reading
During climb constant
During descent constant
During cruise +constant
On take-off constant zero zero zero zero under* under* under* under*

Very few light aircraft are equipped for icing conditions. Devices include inflating rubber boots on the leading edges, heating coils and surfaces that exude antifreeze. Many aircraft however are equipped with storm scopes that give position and warning of approaching storm cells.
Avoiding thunderstorms
a. Above all, remember this: never regard any thunderstorm lightly, even when radar observers report the echoes are of light intensity. Avoiding thunderstorms is the best policy. Following are some do's and don'ts of thunderstorm avoidance:
Don't land or take off in the face of an approaching thunderstorm. A sudden gust front of low level turbulence could cause loss of control.
Don't attempt to fly under a thunderstorm even if you can see through to the other side. Turbulence and windshear under the storm could be disastrous.
Don't fly without airborne radar into a cloud mass containing scattered embedded thunderstorms.
Scattered thunderstorms not embedded usually can be visually circumnavigated.
Don't trust the visual appearance to be a reliable indicator of the turbulence inside a thunderstorm.
Do avoid by at least 20 miles any thunderstorm identified as severe or giving an intense radar echo.
This is especially true under the anvil of a large cumulonimbus.
Do circumnavigate the entire area if the area has 6/10 thunderstorm coverage.
Do remember that vivid and frequent lightning indicates the probability of a severe thunderstorm.
Do regard as extremely hazardous any thunderstorm with tops 35,000 feet or higher whether the top is visually sighted or determined by radar.
b. If you cannot avoid penetrating a thunderstorm, following are some do's BEFORE entering the storm:
Tighten your safety belt, put on your shoulder harness if you have one, and secure all loose objects.
Plan and hold your course to take you through the storm in a minimum time.
To avoid the most critical icing, establish a penetration altitude below the freezing level or above the level of -15 °C.
Verify that pitot heat is on and turn on carburettor heat or jet engine anti-ice. Icing can be rapid at any altitude and cause almost instantaneous power failure and/or loss of airspeed indication.
Establish power settings for turbulence penetration airspeed recommended in your aircraft manual.
Turn up cockpit lights to highest intensity to lessen temporary blindness from lightning.
If using automatic pilot, disengage altitude hold mode and speed hold mode. The automatic altitude and speed controls will increase manoeuvres of the aircraft thus increasing structural stress.
If using airborne radar, tilt the antenna up and down occasionally. This will permit you to detect other thunderstorm activity at altitudes other than the one being flown.
c. Following are some do's and don'ts DURING the thunderstorm penetration:
Do keep your eyes on your instruments. Looking outside the cockpit can increase danger of temporary blindness from lightning.
Don't change power settings; maintain settings for the recommended turbulence penetration airspeed.
Do maintain constant attitude; let the aircraft "ride the waves." Manoeuvres in trying to maintain constant altitude increase stress on the aircraft.
Don't turn back once you are in the thunderstorm. A straight course through the storm most likely will get you out of the hazards most quickly. In addition, turning manoeuvres increase stress on the aircraft.

p r e v a i l i n g w i n d s our thanks www.raa.asn.au
(Copyright John Brandon)
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 subpolar 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 subtropical 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.
surface winds
Polar and subtropical jet streams.

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]
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Warm air over land rises
Sea Breeze moves inland
Cumuli develop aloft and move seaward
Upper level return land breeze
Cool air aloft sinks over water
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Sea Breeze (meso-cold) Front
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 largescale 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.
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Cool air over land sinks
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Land Breeze moves out over water
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Relatively warmer water heats air which then rises
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Upper level return sea breeze
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Cool air over land sinks
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
perpendicular wind flow increasing wind velocity stable layer or inversion
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

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 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.
 directly proportional to wind speed
 Inversely proportional to stability
 Intermountain West - averages
4 miles
 Appalachia Wave - averages 10 miles
 Diurnal variation: in the summer early morning or late afternoon is best for formation
 Seasonal variation: winter is the best time for formation
(jet stream, snow covered ground = no convection, stable layer aloft)
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.

Cap cloud of the Teton mountain range This cloud is mostly on the windward side of the mountain.
Foehngap
The foehngap exists because moisture is reabsorbed during the down rush of air.
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.
Lenticulars over Montana Rotor cloud in Alaska
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.
 altitude 50% above terrain
 approach at 45 degree angle
 avoid ragged & irregular lenticulars
 climb in lift
 dive in sink
 avoid the area of the
 rotor visualise the wave length
Line of rotors - Calgary
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
Vertical component of air current
500 fpm
1000 fpm
1500 fpm
60 knots
[6000 fpm]
5°
9°
15°
100 knots
[10 000 fpm]
3°
6°
11°
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.

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(Copyright John Brandon)
Insolation
The earth’s surface and the atmosphere are mainly warmed by insolation – incoming solar electromagnetic radiation. The amount of insolation energy reaching the outer atmosphere is about
1.36 kilowatts per m². About 10% of the radiation is in the near end of the ultraviolet range ( 0.1 to
0.4 microns), 40% in the visible light range ( 0.4 to 0.7 µm ), 49% in the short wave infra-red range (
0.7 to 3.0 µm ) and 1% is higher energy and X-ray radiation. Refer 1.8 below. The X-rays are blocked at the outer atmosphere and most of the atmospheric absorption of insolation takes place in the upper stratosphere and the thermosphere; with little direct insolation warming in the troposphere, which is mostly warmed by contact with the surface and subsequent convective and mechanical mixing: refer 1.7.4 below.
On a sunny day 75% of insolation may reach the earth’s surface; on an overcast day only 15%. On average 51% of insolation is absorbed by the surface as thermal energy – 29% as direct radiation and
22% as diffused radiation; i.e. scattered by atmospheric dust , water vapour and air molecules, refer
12.1. About 4% of the radiation reaching the surface is directly reflected, at the same wavelength, from the surface back into space. Typical surface reflectance values (albedo) are shown below:
Soils 5–10% Snow, dependent on age 40–90%
Desert 20–40% Water, sun high in sky 2–10%
Forest 5–20% Water, sun low in sky 10–80%
Grass 15–25%
In the insolation input diagram shown below it can be seen that about 26% of insolation is directly reflected back into space by the atmosphere but 19% is absorbed within it as thermal energy with much of the UV radiation being absorbed within the stratospheric ozone layer. Clouds reflect 20% and absorb 3%, atmospheric gases and particles reflect 6% and absorb 16%.
Altogether some 70% of insolation is absorbed at the earth’s surface and in the upper atmosphere but eventually all this absorbed radiation is re-radiated back into space as long wave ( 3 to 30 µm ) infrared. The result of radiation absorption and re-radiation is that the mean atmospheric surface temperature is maintained at 15 °C.
Terrestrial radiation
The surface/atmosphere radiation emission diagram below shows that some 6% of input is lost directly to space as long wave IR from the surface. Atmospheric O², N², and Ar cannot absorb the long wave radiation, also there is a window in the radiation spectrum between 8.5 µm and 11 µm where IR

radiation is not absorbed to any great extent by the other gases. About 15% of the received energy is emitted from the surface as long wave radiation and absorbed by water vapour and cloud droplets within the troposphere and by CO² in the mesosphere. This is actually a net 15%, the total being much greater but the remainder is counter balanced by downward long wave emission from the atmosphere.
Radiation emitted upwards into space, principally nocturnal cooling, is re-radiated from clouds (26%) plus water vapour, O³ and CO² (38%). The atmosphere then has a net long wave energy deficit, after total upwards emission (64%) and absorption (15%), equivalent to 49% of solar input and a short wave insolation excess of 19% (16% + 3% absorbed) resulting in a total atmospheric energy deficit equivalent to 30% of insolation.
Energy balance
The surface has a radiation surplus of 30% of solar input, 51% short wave absorbed less 21% long wave emitted. This surplus thermal energy is convected to the atmosphere by sensible heat flux (7%) and by latent heat flux (23%). The latent heat flux is greater because the ratio of global water to land surface is about 3:1 and over oceans possibly 90% of the heat flux from the surface is in the form of latent heat. Conversely over arid land practically all heat transfer to the atmosphere is in the form of sensible heat.
Overall the earth-atmosphere radiation/re-radiation system is in balance but between latitudes 35°N and 35°S more energy is stored than re-radiated, thus an energy surplus, while between the 35° latitudes and the poles there is a matching energy deficit. There is also a diurnal and a seasonal variation in the radiation balance. The average daily solar radiation measured at the surface in
Australia is 7.5 kW hours/m² in summer and 3.5 kW hours/m² in winter.
All substances emit electromagnetic radiation in amounts and wavelengths dependent on their temperature. The hotter the substance the shorter will be the wavelengths at which maximum emission takes place. The sun, at 6000 K gives maximum emission at about 0.5 µm in the visible light band. The earth at 288 K gives maximum emission at about 9 µm in the long wave IR band.
Tropospheric transport of surface heating and cooling
The means by which surface heating or cooling is transported to the lower troposphere are: by conduction – air molecules coming into contact with the heated (cooled) surface are themselves heated (cooled) and have the same effect on adjacent molecules, thus an air layer only a few centimetres thick becomes less (more) dense than the air above. by convective mixing – arising when the heated air layer tries to rise and the denser layer above tries to sink, thus small turbulent eddies build and the heated layer expands from a few centimetres to a layer hundreds, or thousands, of feet deep depending on the intensity of solar heating.
Convective mixing is more important than mechanical mixing for heating air and is usually dominant during daylight hours.

by mechanical mixing – where wind flow creates frictional turbulence. Mechanical mixing dominates nocturnally when surface cooling and conduction create a cooler, denser layer above the surface thus stopping convective mixing. If there is no wind mechanical mixing can’t occur.
The term (planetary) boundary layer is used to describe the lowest layer of the atmosphere, roughly
1000 to 6000 feet thick, in which the influence of surface friction on air motion is important. It is also referred to as the friction layer or the mixed layer. The boundary layer will equate with the mechanical mixing layer if the air is stable and with the convective mixing layer if the air is unstable.
The term surface boundary layer or surface layer is applied to the thin layer immediately adjacent to the surface, and part of the planetary boundary layer, within which the friction effects are more or less constant throughout, rather than decreasing with height, and the effects of daytime heating and night time cooling are at a maximum. The layer is roughly 50 feet deep, varying with conditions.
1.7.5 Heat advection
Advection is transport of heat, moisture and other air mass properties by horizontal winds.
 Warm advection brings warm air into a region.
 Cold advection brings cold air into a region.
 Moisture advection brings moister air and is usually combined with warm advection.
 Advection is positive if higher values are being advected towards lower, negative if lower values are being advected towards higher, e.g. cold air moving into a warmer region.
Advection into a region may vary with height, e.g. warm, moist advection from surface winds while upper winds are advecting cold, dry air.
Electromagnetic wave spectrum
The electromagnetic spectrum stretches over 60 octaves, the wavelengths double 60 times from the shortest to the longest. In a vacuum electromagnetic waves propagate at a speed close to 300 000 km/sec. The frequency can be calculated from the wavelength ( frequency x wavelength = 108 m/sec
) thus:
 Frequency in kHz = 300 000/wavelength in metres
 Frequency in MHz = 300/wavelength in metres or 30 000/ wavelength in centimetres
 Frequency in GHz = 30/wavelength in centimetre
The very high frequency [VHF] band used in civil aviation radio communications lies in the 30 to 300
MHz frequency range thus the 10 metre to 1 metre wavelength range. The other civil aviation voice communications band is in the high frequency [HF] range; 3 – 30 MHz or 100 –10 metre.
The amplitude of the wave is proportional to the energy of vibration. The table below shows the wave length ranges – beginning in nanometres [nm] and progressing through micrometres, millimetres, metres and kilometres – and the associated radiation bands.

Tropospheric global heat transfer
Precipitation is less than evaporation between 10° and 40° latitudes, the difference being greatest at about 20°. Polewards and equatorwards of these bands precipitation is greater than evaporation. The transfer of atmospheric water vapour, containing latent heat, is polewards at latitudes greater than
20° and equatorwards at lower latitudes. Most of the vertical heat transfer is in the form of latent heat but possibly 65% of the atmospheric horizontal transfer is in the form of sensible heat following condensation of water vapour. Horizontal latent heat transfer occurs primarily in the lower troposphere.
The general wind circulation within the troposphere ( refer 4.1 ) and the water circulation within the oceans transfer heat from the energy surplus zones ( refer 1.7 ) to the energy deficit zones thereby maintaining the global heat balance. About 70% is transferred by the atmosphere and 30% by the oceans. The large mid-latitude eddies, the cyclones and anti-cyclones in the broad westerly wind band that flows around the Southern Hemisphere, play a particularly important part in the transfer of the excess heat energy from low to high latitudes and in the mixing of cold Antarctic or arctic air into the mid-latitudes.
Temperature lapse rates in the troposphere
The temperature lapse rates in the troposphere vary by latitude, climatic zone and season, varying between less than 0 °C/km (i.e. increasing with height) at the winter poles to more than 8 °C/km over a summer sub-tropical ocean. In the mid-latitudes the temperature reduces with increasing height at varying rates but averaging 6.5 °C/km or about 2 °C per 1000 feet, although within any tropospheric layer temperature may actually increase with increasing height. This reversal of the norm is a temperature inversion condition. Should the temperature in a layer remain constant with height then an isothermal layer condition exists. At night, particularly under clear skies, the air in the mixed layer cools considerably but the long wave radiation from the higher levels is weak and the

air there cools just 1 °C or so. Consequently a nocturnal inversion forms over the the mixed layer, the depth of which depends on the temperature drop and the amount of mechanical mixing.
Tropospheric average temperature lapse rate profile
The altitude of the tropopause, and thus the thickness of the troposphere, varies considerably.
Typical altitudes are 55 000 feet in the tropics with a temperature of –70 °C and 29 000 feet in polar regions with a temperature of –50 °C. Because of the very low surface temperatures in polar regions and the associated low level inversion, the temperature lapse profile is markedly different to the mid-latitude norms. In mid-latitudes the height of the troposphere varies seasonally and daily with the passage of high and low pressure systems.
In the chart above an exaggerated environmental temperature lapse rate profile has been superimposed to illustrate the temperature layer possibilities starting with a superadiabatic lapse layer at the surface, a normal lapse rate layer above it then a temperature inversion layer and an isothermal layer.
Adiabatic processes and lapse rates
An adiabatic process is a thermodynamic process where a change occurs without loss or addition of heat, as opposed to a diabatic process in which heat enters or leaves the system. Examples of the latter are evaporation from the ocean surface, radiation absorption and turbulent mixing.
An adiabatic temperature change occurs in a vertically displaced parcel of air due to the change in pressure and volume occurring during a short time period, with little or no heat exchange with the environment. Upward displacement and consequent expansion causes cooling, downward displacement and subsequent compression causes warming. In the troposphere the change in temperature associated with the vertical displacement of a parcel of dry ( i.e. not saturated ) air is very close to 3 °C per 1000 feet, or 9.8 °C / km, of vertical motion; this is known as the dry
adiabatic lapse rate [DALR]. As ascending moist air expands and cools in the adiabatic process the excess water vapour condenses after reaching dewpoint and the latent heat of condensation is released into the parcel of air as sensible heat thus slowing the pressure induced cooling process. This condensation process continues whilst the parcel of air continues to ascend and expand. The process is reversed as an evaporation process in descent and compression. The adiabatic lapse rate for saturated air, the saturated adiabatic lapse rate [SALR], is dependent on the amount of moisture

content which in itself is dependent on temperature and pressure. The chart below shows the SALR at pressures of 500 and 1000 mb and temperatures between –40 °C and +40 °C.
The chart shows that on a warm day the SALR near sea level is about 1.2 °C / 1000 feet while at about 18 000 feet, the 500 mb level, the rate doubles to about 2.4 °C / 1000 feet.
The environment lapse rate [ELR] is ascertained by measuring the actual vertical distribution of temperature at that time and place. The ELR may be equal to or differ from the DALR or SALR of a parcel of air moving within that environment. In the atmosphere parcels of air are stirred up and down by turbulence and eddies that may extend several thousand feet vertically in most wind conditions. These parcels mix and exchange heat with the surrounding air thus distorting the adiabatic processes.
If the rate of ground heating by solar radiation is rapid the mixing of heated bubbles of air may be too slow to induce a well mixed layer with a normal DALR. The ELR, up to 2000 – 3000 feet agl, may be much greater than the DALR. Such a layer is termed a superadiabatic layer and will contain strong thermals and downdraughts.
Atmospheric stability
Atmospheric stability is the air’s resistance to any disturbing effect but might be defined as the ability to resist the narrowing of the spread between air temperature and dewpoint. Stable air cools slowly with height and vertical movement is limited. If a parcel of air, after being lifted, is cooler than the environment, the parcel being more dense than the surrounding air will tend to sink back and conditions are stable.
The temperature of unstable air drops more rapidly with increase in altitude i.e. the ELR is steep. If a lifted parcel is warmer, and thus less dense than the surrounding air, the parcel will continue to rise and conditions are unstable. Unstable air, once it has been lifted to the lifting condensation level ( refer 3.3 ) keeps rising through free convection. Instability can cause upward or downward motion.
When saturated air containing little or no condensation is made to descend then adiabatic warming causes the air to become unsaturated almost immediately and further descent warms it at the DALR.
If the ELR lies between the DALR and the SALR a state of conditional instability exists. Thus if an unsaturated parcel of rises from the surface it will cool at the DALR and so remain cooler than the environment and conditions are stable. However if the parcel passes dewpoint during the ascent it will then cool at a slower rate and, on further uplift, become warmer than the environment and so

become unstable. High dewpoints are an indication of conditional instability. The figure below demonstrates some ELR states with the consequent stability condition:
 ELR #1 is much greater than the DALR (and the SALR) providing absolute instability. This condition is normally found only near the ground in a superadiabatic layer.
 ELR # 2 between the DALR and the SALR demonstrates conditional instability. It is stable when the air parcel is unsaturated, i.e. the ELR is less than the DALR, and unstable when it is saturated, i.e. the ELR is greater than the SALR.
 ELR #3 indicates absolute stability, the ELR is less than the SALR (and the DALR).
 Neutral equilibrium would exist if the ELR equalled the SALR and the air was saturated or if the ELR equalled the DALR and the air was unsaturated.
The following diagram is an example of atmospheric instability and cloud development, comparing environment temperature and that of a rising air parcel with dewpoint of 11 °C.

The amount of energy that could be released once surface based convection is initiated in humid air is measured as convective available potential energy [CAPE]. CAPE is measured in joules per kilogram of dry air and may be assessed by plotting the vertical profile of balloon radio-sonde readings for pressure, temperature and humidity on a tephigram and also plotting the temperatures that a rising parcel of air would have in that environment.. On the completed tephigram the area between the plot for environment temperature profile and the plot for the rising parcel temperature profile is directly related to the CAPE, which in turn is directly related to the maximum vertical speed in a Cb updraught.
A tephigram is a thermodynamic graph used by meteorologists for plotting atmospheric temperature and moisture profiles. The name is a combination of T, for temperature and the Greek letter phi, for entropy, the latter roughly meaning, in this context, the potential energy of a gas. A simplified tephigram is shown below with just isobars – the horizontal lines and isotherms – the diagonal lines, and a plot of dewpoint on the left. The observed temperature profile is in the centre and the expected rising parcel temperature profile is to the right of it with the amount of CAPE related to the area between the plots.
Convergence, divergence and subsidence
Synoptic scale atmospheric vertical motion is found in cyclones and anticyclones, mainly caused by air mass convergence or divergence from horizontal motion. Meteorological convergence indicates retardation in air flow with increase in air mass in a given volume due to net three dimensional inflow. Meteorological divergence, or negative convergence, indicates acceleration with decrease in air mass. Convergence is the contraction and divergence is the spreading of a field of flow.
If, for example, the front end of moving air mass layer slows down, the air in the rear will catch up – converge, and the air must move vertically to avoid local compression. If the lower boundary of the moving air mass is at surface level all the vertical movement must be upward. If the moving air mass is just below the tropopause all the vertical movement will be downward because the tropopause inhibits vertical motion. Conversely if the front end of a moving air mass layer speeds up then the flow diverges. If the air mass is at the surface then downward motion will occur above it to satisfy mass conservation principles, if the divergence is aloft then upward motion takes place.
Rising air must diverge before it reaches the tropopause and sinking air must diverge before it reaches the surface. As the surface pressure is the weight per unit area of the overlaying column of air, and even though divergences in one part of the column are largely balanced by convergences in another, the slight change in mass content (thickness) of the over-riding air changes the pressure at the surface.

The following diagrams illustrate some examples of convergence and divergence:
Note: referring to the field of flow diagrams above, the spreading apart (diffluence) and the closing together (confluence) of streamlines alone do not imply existence of divergence or convergence as there is no change in air mass if there is no cross isobar flow or vertical flow. (An isobar is a curve along which pressure is constant and is usually drawn on a constant height surface such as mean sea
level.)
Divergence or convergence may be induced by a change in surface drag, for instance when an airstream crosses a coastline. An airstream being forced up by a front will also induce convergence.
For convergence / divergence in upper level waves. Some divergence / convergence effects may cancel each other out e.g. deceleration associated with diverging streamlines.
Developing anti-cyclones – “highs” and high pressure ridges, are associated with converging air aloft and consequent wide area subsidence with diverging air below . This subsidence usually occurs between 20 000 and 5000 feet typically at the rate of 100 – 200 feet per hour. The subsiding air is compressed and warmed adiabatically at the DALR, or an SALR, and there is a net gain of mass within the developing high. Some of the converging air aloft rises and, if sufficiently moist, forms the cirrus cloud often associated with anti-cyclones.
As the pressure lapse rate is exponential and the DALR is linear the upper section of a block of subsiding air usually sinks for a greater distance and hence warms more than the lower section and if the bottom section also contains layer cloud the sinking air will only warm at a SALR until the cloud evaporates. Also when the lower section is nearing the surface it must diverge rather than descend and thus adiabatic warming stops. With these circumstances it is very common for a subsidence
inversion to consolidate at an altitude between 3000 and 6000 feet. The weather associated with large scale subsidence is almost always dry, but in winter persistent low cloud and fog can readily form in the stagnant air due to low thermal activity below the inversion, producing ‘anti-cyclonic gloom’. In summer there may be a haze layer at the inversion level which reduces horizontal visibility at that level although the atmosphere above will be bright and clear. Aircraft climbing through the inversion layer will usually experience a wind velocity change.

Developing cyclones, “lows” or "depressions" and low pressure troughs are associated with diverging air aloft and uplift of air leading to convergence below. There is a net loss of mass within an intensifying low as the rate of vertical outflow is greater than the horizontal inflow, but if the winds continue to blow into a low for a number of days, exceeding the vertical outflow, the low will fill and disappear. The same does not happen with anti-cyclones which are much more persistent.
A trough may move with pressure falling ahead of it and rising behind it giving a system of pressure tendencies due to the motion but with no overall change in pressure, i.e. no development, no deepening and no increase in convergence.
Thermal gradients and the thermal wind concept
The rate of fall in pressure with height is less in warm air than in cold and columns of warm air have a greater vertical extent than columns of cold air. Consider two adjacent air columns having the same msl pressure; the isobaric surfaces (surfaces of constant pressure) are at higher levels in the warm air column which result in a horizontal pressure gradient from the warm to the cold air, which increases with height, i.e. the temperature gradient causes increasing wind to higher levels. The horizontal pressure gradient increases as the horizontal thermal gradient increases, the process being known as the thermal wind mechanism.

The isobaric surface contours vary with height so the geostrophic wind velocity above a given point also varies with height. The wind vector difference between the two levels above the point, the vertical wind shear, is called the thermal wind, i.e. the wind vector component caused by temperature difference rather than pressure difference. On an upper air thickness chart which indicates the heat content of the troposphere, the thermal wind is aligned with the geopotential height lines or with the isotherms on an upper air constant pressure level chart (isobaric surface
chart), and the thicker (warmer) air is to the left looking downwind.
A geopotential height line is a curve of constant height, i.e. the height/thickness contours relating to an isobaric surface, usually shown in decametres or metres above the 1000 mb surface or msl on an upper air chart. An isotherm is a curve connecting points of equal temperature and usually drawn on a constant pressure surface or a constant height surface An isopleth is the generic name for all iso-lines or contour lines.
The speed of the thermal wind is proportional to the thermal gradient, the closer the contour spacing the stronger the thermal wind. If the horizontal thermal gradient maintains much the same direction through a deep atmospheric layer, for instance there are no upper level highs or lows, and the gradient is strong with the colder air to the south, then the thermal wind will increase with height eventually becoming a constant westerly vector. The resultant high level wind will be high speed and nearly westerly.
Generally colder air is to the south so that the thermal wind vector tends westerly but if the horizontal thermal gradient reverses direction with height an easterly thermal wind will occur above that level and the upper level westerly geostrophic wind speed will decrease with height. Since the direction of the thermal gradient is reversed above the tropopause the thermal wind reverses to easterly. The horizontal thermal gradient is at maximum just below the tropopause, where the jet

stream occurs.
At latitude 45° S a temperature difference of 1 °C in 100 km will cause an increase in thermal wind of
10 m/sec, or about 20 knots, for every 10 000 feet of altitude, giving jet stream speeds at 30 000 feet, ignoring geostrophic wind. Temperature contrasts between air masses at the polar front will be greatest during winter, giving the strongest jet stream.
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air masses
Air masses are parcels of air that bring distinctive weather features to the country. An air mass is a body or 'mass'of air in which the horizontal gradients or changes in temperature and humidity are relatively slight. That is to say that the air making up the mass is very uniform in temperature and humidity.
An air mass is separated from an adjacent body of air by a transition that may be more sharply defined.
This transition zone or boundary is called a front. An air mass may cover several millions of square kilometres and extend vertically throughout the troposphere.
Weather
Phenomenon
Temperature
Atmospheric
Pressure
Winds
Precipitation
Clouds
Prior to the Passing of the Front
Warm
Decreasing steadily
South to southeast
Showers
Contact with the Front
Cooling suddenly
Heavy rain or snow, hail sometimes
After the
Passing of the Front
Cold and getting colder
Levelling off then increasing
Increasing steadily
Variable and gusty
West to northwest
Showers then clearing
Cirrus and cirrostratus changing later to cumulus and cumulonimbus
Cumulus and cumulonimbus
Cumulus
The temperature of an air mass will depend largely on its point of origin, and its subsequent journey over the land or sea. This might lead to warming or cooling by the prolonged contact with a warm or cool surface. The processes that warm or cool the air mass take place only slowly, for example it may take a week or more for an air mass to warm up by 10 °C right through the troposphere. For this to take place, an air mass must lie virtually in a stagnant state over the influencing region. Hence, those parts of the Earth's surface where air masses can stagnate and gradually attain the properties of the underlying surface are called source regions.
The main source regions are the high pressure belts in the subtropics, which produce tropical air masses, and around the poles, that are the source of polar air masses.

Polar and tropical source regions: The blue and red arrows show the polar and tropical regions respectively.
modification of air masses
As we have seen, it is in the source regions that the air mass acquires distinctive properties that are the characteristics of the underlying surface. The air mass may be cool or warm, or dry or moist. The stability of the air within the mass can also be deducted. Tropical air is unstable because it is heated from below, while polar air is stable because it is cooled from below.
As an air mass moves away from its source region towards the British Isles, the air is further modified due to variations in the type or nature of the surface over which it passes. Two processes act independently, or together, to modify an air mass.
An air mass that has a maritime track, i.e. a track predominantly over the sea, will increase its moisture content, particularly in its lower layers. This happens through evaporation of water from the sea surface. An air mass with a long land or continental track will remain dry.
Fig 2: Modification of air mass by land and ocean surfaces
A cold air mass flowing away from its source region over a warmer surface will be warmed from below making the air more unstable in the lowest layers. A warm air mass moving over a cooler surface is cooled from below and becomes stable in the lowest layers.
Fig 3: Modification of air mass due to surface temperature

If we look at the temperature profiles of the previous example, the effects of warming and cooling on the respective air masses are very different.
Fig 4: Modified vertical temperature profiles (----- line) typical of: a) tropical air cooled from below and b) polar air heated from below on its way south. Note that where the air is heated from below the effect is spread to a greater depth of the atmosphere.
weather in an air mass
Five basic types of air masses determine the weather. They can bring anything from scorching heat to bone-chilling cold depending on the type of air mass.
These air masses are:
Tropical continental
(Tc)
Polar continental
(Pc)
Tropical maritime
(Tm)
Polar
Maritime
(Pm)
Arctic
Maritime
(Am)
Returning
Polar
Maritime
(rPm)
Summer Winter
Long sea track
Short sea track
Exposed Sheltered
Temp
Humidity
Very warm or hot
Relatively dry
Average
Rather moist
Cold
Very
Moist in lowest layers cold
Very dry
Near sea temperature
Very moist
Warm
Moist
Rather cold
Moist
Cold
(colder than Pm)
Warm
(warmer than
Pm)
Fairly moist
(not as moist as Pm)
Fairly moist
(not as moist as
Pm)
Change of lapse rate
Little change
Cooled from below
Stability
Generally stable
Stable
Heated from below
Unstable
Little change
Stable
Cooled from below
Warmed in summer
Stable Stable aloft
Heated from below
Unstable
Weather
Clear, occasional thundery showers
Clear
Rain or snow showers
Clear
Low cloud, drizzle
Visibility
Moderate or poor
Moderate of poor
Good
Moderate or poor
Often poor with coastal fog
Broken cloud, dry
Variable cloud, showers
Moderate Good
Heated from below
Unstable
Heated from below
Unstable
Showers
(mainly coastal)
Very good
Showers
(mainly coastal)
Very good polar front

Several fronts and semipermanent high and low pressure systems characterize the Arctic. The "polar front" marks the boundary between cold polar air masses and warm tropical air masses. The polar front is intermittent rather than continuous around the globe. The strength of the polar front depends on the magnitude of the horizontal temperature gradient across the front. Where the temperature gradient is steep, the front is strong and is a potential site for cyclone or low pressure system development. Where temperature contrast is small, the polar front is weak.
Like the polar front, the "arctic front" is discontinuous and depends on the temperature contrast between two air masses. The arctic front is the boundary between polar and arctic air masses and lies to the north of the polar front. The arctic front can be as strong as the polar front. It is particularly prominent during summer in northern Eurasia.
Semipermanent high and low pressure systems ("highs" and "lows") are identified with particular regions and have seasonal characteristics. In winter, the Icelandic Low extends from near Iceland north into the Barents Sea, and is associated with frequent cyclone activity. The Aleutian Low is present in the Gulf of Alaska. The Beaufort-Chukchi Sea region is dominated by a ridge of high pressure linking the Siberian High and high pressure over the Yukon of Canada. In April and May arctic pressure gradients decrease. The Icelandic and Aleutian lows weaken. The Siberian High disappears, and is replaced by a wide but shallow low. The Arctic High is centred over the Canadian Arctic
Archipelago. In summer, pressure gradients are generally weak. Intermittently, however, cyclones enter the Arctic from northern Eurasia and the north Atlantic, and tend to persist over the Canadian
Basin. By October the pattern has almost returned to the winter configuration. The Icelandic and
Aleutian lows strengthen, as does the Siberian High.
Semipermanent Highs and Lows
The Arctic is characterized by "semipermanent" patterns of high and low pressure. These patterns are semipermanent because they appear in charts of long-term average surface pressure. They can be considered to largely represent the statistical signature of where transitory high and low systems that appear on synoptic charts tend to be most common.
Aleutian Low
This semipermanent low pressure centre is located near the Aleutian Islands. Most intense in winter, the Aleutian Low is characterized by many strong cyclones. Travelling cyclones formed in the subpolar latitudes in the North Pacific usually slow down and reach maximum intensity in the area of the
Aleutian Low.
Icelandic Low
This low pressure center is located near Iceland, usually between Iceland and southern Greenland.
Most intense during winter, in summer, it weakens and splits into two centres, one near Davis Strait and the other west of Iceland. Like its counterpart the Aleutian Low, it reflects the high frequency of cyclones and the tendency for these systems to be strong. In general, migratory lows slow down and intensify in the vicinity of the Icelandic Low.
Siberian High
The Siberian High is an intense, cold anticyclone that forms over eastern Siberia in winter. Prevailing from late November to early March, it is associated with frequent cold air outbreaks over east Asia.
Beaufort High
The Beaufort High is a high pressure centre or ridge over the Beaufort Sea present mainly in winter.
North American High
The North American High is a relatively weak area of high pressure that covers most of North America during winter. This pressure system tends to be centred over the Yukon, but is not as well-defined as its continental counterpart, the Siberian High.

Polar Lows
Small cyclones forming over open sea during the cold season within polar or arctic air masses are called "polar lows." Typically several hundred kilometers in diameter, and often possessing strong winds, polar lows tend to form beneath cold upper-level troughs or lows when frigid arctic air flows southward over a warm body of water.
Polar lows last on average only a day or two. They can develop rapidly, reaching maximum strength within 12 to 24 hours of the time of formation. They often dissipate just as quickly, especially upon making landfall. In some instances several may exist in a region at the same time or develop in rapid succession.
In satellite imagery polar lows show characteristic spiral or comma shaped patterns of deep clouds, sometimes with an inner "eye" similar to those seen in tropical cyclones. Convective cloud bands occupy the surroundings (see figure below). Analysis of aircraft and radiosonde data collected during field experiments reveals that polar lows may possess warm cores. This finding, coupled with their appearance in satellite imagery, has prompted some investigators to refer to polar lows as "arctic hurricanes," although they seldom, if ever, possess hurricane strength winds.
Polar lows are difficult to predict even with current high resolution and high performing operational numerical models, because they usually occur in remote oceanic regions where data are too sparse to define the model initial state on a sufficiently fine scale. However, present-day models can depict synoptic-scale patterns favourable to the development of the smaller scale systems, allowing forecasters to use the predictions in conjunction with satellite imagery and conventional observations to make subjective forecasts of their occurrence.
A NOAA-9 polar orbiter satellite image (visible band) of a polar low over the Barents Sea on 27 February 1987. The southern tip of Spitsbergen is visible at the top of the image. The polar low is centred just north of the Norwegian coast.
Image contributed by S. Businger, Department of Meteorology, University of Hawaii.

The Polar Vortex
The polar vortex is a persistent large-scale cyclonic circulation pattern in the middle and upper troposphere and the stratosphere, centred generally in the polar regions of each hemisphere. In the
Arctic, the vortex is asymmetric and typically features a trough (an elongated area of low pressure) over eastern North America. It is important to note that the polar vortex is not a surface pattern. It tends to be well expressed at upper levels of the atmosphere (that is, above about five kilometres).
fronts
A front is defined as the transition zone between two air masses of different density. Fronts extend not only in the horizontal direction, but in the vertical as well. Therefore, when referring to the frontal surface (or frontal zone), we referring to both the horizontal and vertical components of the front.
A cold front is that part (or parts) of a frontal system along which cold air is advancing and is coloured blue on the weather map.
A warm front is that part (or parts) of a frontal system along which cold air is retreating and is coloured red on the weather map.
types of front the warm front
A warm front is defined as the transition zone where a warm air mass is replacing a cold air mass.
Warm fronts generally move from southwest to northeast and the air behind a warm front is warmer and more moist than the air ahead of it. When a warm front passes through, the air becomes noticeably warmer and more humid than it was before.
Symbolically, a warm front is represented by a solid line with semicircles pointing towards the colder air and in the direction of movement. On coloured weather maps, a warm front is drawn with a solid red line.

There is typically a noticeable temperature change from one side of the warm front to the other. In the map of surface temperatures below, the station north of the front reported a temperature of 53 degrees Fahrenheit while a short distance behind the front, the temperature increased to 71 degrees.
An abrupt temperature change over a short distance is a good indication that a front is located somewhere in between.
If warmer air is replacing colder air, then the front should be analyzed as a warm front. If colder air is replacing warmer air, then the front should be analyzed as a cold front. Common characteristics associated with warm fronts have been listed in the table below.
Winds
Before Passing south-southeast
While Passing After Passing variable south-southwest steady rise warmer, then steady Temperature cool-cold, slow warming
Pressure usually falling
Clouds
Precipitation in this order: Ci, Cs, As, Ns, St, and fog; occasionally Cb in summer light-to-moderate rain, snow, sleet, or drizzle
Visibility
Dew Point poor steady rise levelling off slight rise, followed by fall stratus-type drizzle or none poor, but improving clearing with scattered Sc; occasionally Cb in summer usually none, sometimes light rain or showers fair in haze steady rise, then steady
As a mass of warm air advances on a retreating mass of cold air, the warm air, being lighter, ascends over the cold air in a long gentle slope. As a result, the cloud formation associated with the warm frontal system may extend for 500 or more nautical miles in advance of it. Warm fronts usually move at relatively slow speeds and therefore affect a vast area for a considerable length of time.
If the warm air is moist and stable, stratiform clouds develop in a distinctive sequence. The first signs of an approaching warm front are high cirrus clouds which thicken to cirrostratus and altostratus as the warm front approaches. The ceiling gradually falls and there follows a long belt of steady rain

falling from heavy nimbostratus cloud. Precipitation may lead the frontal surface by as much as 250 nautical miles.
If the warm air is moist and somewhat unstable, cumulonimbus and thunderstorms may be embedded in the stratiform layers. Heavy showers in advance of the surface front can then be expected.
Very low stratus clouds and fog throughout the frontal zone are typical characteristics of warm fronts.
The passing of the warm front is marked by a rise of temperature, due to the entry of the warm air, and the sky becomes relatively clear.
the cold front
A cold front is defined as the transition zone where a cold air mass is replacing a warmer air mass.
Cold fronts generally move from northwest to southeast. The air behind a cold front is noticeably colder and drier than the air ahead of it. When a cold front passes through, temperatures can drop more than 15 degrees within the first hour.
Symbolically, a cold front is represented by a solid line with triangles along the front pointing towards the warmer air and in the direction of movement. On coloured weather maps, a cold front is drawn with a solid blue line.
There is typically a noticeable temperature change from one side of a cold front to the other. In the map of surface temperatures below, the station east of the front reported a temperature of 55 degrees Fahrenheit while a short distance behind the front, the temperature decreased to 38 degrees. An abrupt temperature change over a short distance is a good indicator that a front is located somewhere in between.

If colder air is replacing warmer air, then the front should be analyzed as a cold front. On the other hand, if warmer air is replacing cold air, then the front should be analyzed as a warm front. Common characteristics associated with cold fronts have been listed in the table below.
Winds
Before Passing south-southwest gusty; shifting
While Passing After Passing west-northwest
Temperature warm sudden drop steadily dropping
Pressure falling steadily minimum, then sharp rise rising steadily
Clouds
Precipitation
Visibility increasing: Ci, Cs and Cb short period of showers fair to poor in haze
Cb heavy rains, sometimes with hail, thunder and lightning poor, followed by improving
Cu showers then clearing good, except in showers
Dew Point high; remains steady sharp drop lowering
When a mass of cold air overtakes a mass of warm air, the cold air being denser, stays on the surface and undercuts the warm air violently. Surface friction tends to slowdown the surface air while a sharp fall in temperature, a rise in pressure and rapid clearing usually occur with the passage of the cold front.
Sometimes, an advancing cold front will be relatively slow moving. Because it does not undercut the warm air so violently, a rather broad band of clouds develops extending a fair distance behind the frontal surface. If the warm air is stable, these clouds will be stratiform; if the warm air is unstable, they are cumuliform and possibly thunderstorms. With passage of the frontal surface, clearing is more gradual.
the stationary front
There is generally some part of a front along which the colder air is neither advancing nor retreating.
There is no motion to cause the front to move because the opposing air masses are of equal pressure.
The surface wind tends to blow parallel to the front and the weather conditions are similar to those associated with a warm front although generally less intense and not so extensive. Usually a stationary front will weaken and eventually dissipate. Sometimes, however, after several days, it will begin to move and then it becomes either a warm front or a cold front.

A noticeable temperature change and/or shift in wind direction is commonly observed when crossing from one side of a stationary front to the other.
In the map above, temperatures south of the stationary front were in the 50's and 60's with winds generally from the southeast. However, north of the stationary front, temperatures were in the 40's while the winds had shifted around to the northeast. Cyclones migrating along a stationary front can dump heavy amounts of precipitation, resulting in significant flooding along the front occluded fronts
When the progress of time as a depression advances, the cold front gradually overtakes the warm front and lifts the warm sector entirely from the ground. It is simply a case of the cold air catching up with itself as it flows around the depression. Thus only one front remains, which is called an occluded front or occlusion. An occluded depression soon commences to fill up and die away.
A developing cyclone typically has a preceding warm front (the leading edge of a warm moist air mass) and a faster moving cold front (the leading edge of a colder drier air mass wrapping around the storm). North of the warm front is a mass of cooler air that was in place before the storm even entered the region.
As the storm intensifies, the cold front rotates around the storm and catches the warm front. This forms an occluded front, which is the boundary that separates the new cold air mass (to the west)

from the older cool air mass already in place north of the warm front. Symbolically, an occluded front is represented by a solid line with alternating triangles and circles pointing the direction the front is moving. On coloured weather maps, an occluded front is drawn with a solid purple line.
Changes in temperature, dew point temperature, and wind direction can occur with the passage of an occluded front. In the map below, temperatures ahead (east of) the front were reported in the low
40's while temperatures behind (west of) the front were in the 20's and 30's. The lower dew point temperatures behind the front indicate the presence of drier air.
A noticeable wind shift also occurred across the occluded front. East of the front, winds were reported from the east-southeast while behind the front, winds were from the west-southwest.
Common characteristics associated with occluded fronts have been listed in the table below.
Winds
Temperature
Cold Type
Warm Type
Pressure
Clouds
Precipitation
Visibility
Dew Point
Before Passing While Passing southeast-south variable
After Passing west to northwest cold-cool dropping cold rising usually falling low point colder milder usually rising in order: Ci, Cs, As,
Ns
Ns, sometimes Tcu and Cb Ns, As or scattered Cu light, moderate or heavy precipitation light, moderate or heavy continuous precipitation or showers poor in precipitation poor in precipitation steady light-to-moderate precipitation followed by general clearing improving usually slight drop, especially if cold-occluded slight drop, although may rise a bit if warm-occluded
The cold air, in the distance it has travelled, may have undergone considerable change. Therefore it may not be as cold as the air it is overtaking. In this case (cool air advancing on colder air), the front is known as an occluded warm front or a warm occlusion and has the characteristics of a warm front, with low cloud and continuous rain and drizzle. It the warm air is unstable, heavy cumulus or cumulonimbus cloud may be embedded in the stratiform cloud bank.
It the cold air is colder than the air it is overtaking (cold air advancing on cool air), the front is known as an occluded cold front or a cold occlusion. A cold occlusion has much the same characteristics as a warm front, with low cloud and continuous rain. If the warm air is unstable, cumulonimbus and thunderstorms are likely to occur, with the violent turbulence, lightning and icing conditions associated with these clouds.

It will be noted that in the case of either a warm or cold occlusion, three air masses are present, a cool air mass advancing on a cold air mass, or a cold air mass advancing on a cool air mass, with, in either case, a warm air mass lying wedge shaped over the colder air. This wedge shaped mass of warm air is known as a trowel in Canada. (In some other countries, such as the US, it is called an upper front.) upper fronts
In Canada, the term upper front refers to a non-occlusion situation. Sometimes, cold air advancing across the country may encounter a shallow layer of colder air resting on the surface or trapped in a topographical depression. The advancing cold air rides up over the colder, heavier air. The cold front which is the leading edge of the advancing cold air, therefore, leaves the ground and moves along the top of the colder air. It is then known as an upper cold front.
Sometimes, the structure of the advancing cold front is such that the cold air forms a shallow layer for some distance along the ground in advance of the main body of cold air. The frontal surface If the main mass of cold air, in this situation, will usually be very steep. The line along which the frontal surface steepens is also known as an upper cold front.
On occasion, an advancing warm front rides up over a pool or layer of cold air trapped on the ground.
A station on the ground does not experience a change of air mass because the front passes overhead.
This is known as an upper warm front.
Sometimes, the surface of the cold air that is retreating ahead of an advancing warm front is almost flat for some distance ahead of the surface front and then steepens abruptly. The line along which the surface of the retreating cold air steepens sharply is also called an upper warm front.
frontal weather
The theory of the polar front, which for the sake of simplicity has been described in the form of its original conception, might leave the impression that depressions form only along some well defined line Iying somewhere midway between the poles and the equator. Air masses are in a constant state of formation over all the land and water areas of the world. Once formed, they tend to move away from the source regions over which they form. The same frontal processes and phenomena occur whenever a mass of warm air and a mass of cold air come in contact.
There is a widespread impression among pilots that fronts always bring bad weather and that all bad weather is frontal. Actually some fronts have little or no weather associated with them. A slight change of temperature and a windshift may be the only evidence that the front has gone through.
And, of course, bad weather can develop without the passage of a front. Fog, for example, generally occurs when no fronts are present and severe thunderstorms may develop in an air mass, which has no frontal characteristics.
Another common misconception is that the front is a thin wall of weather. This false idea is perhaps occasioned by the line that indicates a front on a weather map. The line on the map only shows the surface location at which the pressure change, windshift and temperature change occur. The actual weather associated with the front may extend over an area many miles in width, both well ahead and also for many miles behind the actual line on the weather map.
A front itself is actually a transition zone between two large air masses with different properties of temperature and moisture. Each individual air mass may extend over hundreds of thousands of square miles. Everywhere along the boundary of an air mass, where it overrides or undercuts the air mass upon which it is advancing and for a considerable height upward from the surface as well, there is a frontal zone. The frontal weather associated with the front, therefore, can be expected to extend for hundreds of miles along the boundary of the air mass.
Frontogenesis means a front, which is increasing in intensity.
Frontolysis means a front, which is decreasing in intensity.
If you examine the diagrams showing fronts on a weather map, you will notice that all fronts lie in regions of lower pressure. The isobars are bent sharply at a front. These two factors are

characteristic of all fronts.
weather at the cold front
Cold fronts are not all the same. The weather associated with a cold front may vary from a minor windshift to severe thunderstorms, low ceilings, restricted visibility and violent gusty winds. The severity of the weather is determined by the moisture content and stability of the warm air mass that the cold air mass is undercutting and the speed of the advancing cold front.
Fast moving cold fronts may travel across the country with a speed of 30 knots or more. If the warm air that is being undercut by the cold air mass is very moist and unstable, towering cumulus clouds and thunderstorms are likely to develop. Heavy rain or hail may be associated with the front. A slower moving cold front advancing on more stable and drier air in the warm sector will produce less severe weather conditions, stratus or altocumulus clouds with light or no precipitation.
A long line of cumulus clouds on the western horizon is usually an indication of an approaching cold front. Sometimes a deck of altocumulus cloud or decks of stratus and stratocumulus extending ahead of the front will mask the main frontal cloud from the view of the high flying or low flying pilot respectively.
weather changes
Surface Wind: The wind direction will always veer as the front passes. Gustiness may be associated with the windshift.
In flying through a cold front, the windshift may be quite abrupt and occurs at the frontal surface rather than at the front. The windshift is always such that an alteration in course to starboard is required, no matter which way you are flying through the front.
Temperature : On the ground, the temperature may drop sharply as the front passes, but usually it drops gradually. The air immediately behind the front has been warmed in passing over the warm ground. Therefore, it may be several hours before the temperature drops to the true value of the cold air mass. In flying through a cold front, there will be a noticeable temperature change when passing through the frontal surface.
Visibility: Visibility usually improves after passage of a cold front. If the front is moving fairly rapidly, the width of frontal weather generally is less than 50 miles. If the front is moving slowly, however, flight operations may be affected for many hours.
Pressure: The approach of a cold front is accompanied by a decrease in pressure. A marked rise will be noticed when the front has passed.
Turbulence: Turbulence may be associated with the cold front if it is active, although thunderstorms are not always present. Even in cases where there are no clouds, turbulence may be a problem. As a rule, flight through an active cold front can be expected to be rough.
Precipitation: The frontal rain or snow is usually narrow, especially if it is showery in character. Icing in the turbulent cumulus clouds can be severe.
l i n e s q u a l l s
A long line of squalls and thunderstorms which sometimes accompanies the passage of a cold front is called a line squall (or squall line). It is usually associated with a fast moving cold front that is undercutting an unstable warm air mass. It may form anywhere from 50 to 300 nautical miles in advance of the front itself. The line squall is a long line of low black, roller like cloud, which often stretches in a straight line for several hundred miles, and from which heavy rain or hail falls for a short time. Thunder and lightning frequently occur. The squall is also accompanied by a sudden wind change from southerly or south-westerly to north or north-westerly, together with a sudden drop in temperature and a rise in barometric pressure. The actual wind squall lasts only for a few minutes but is often extremely violent, constituting a serious menace both to shipping and to airplanes. The signs indicating the approach of a line squall are unmistakable. Airplanes on the ground should be

immediately hangared. Those in the air should at all costs avoid this violent weather phenomenon.
weather at the warm front
Warm front changes are usually less pronounced than cold front changes. The change is also generally very gradual. However, the weather at a warm front is usually more extensive and may cover thousands of square miles. A wide variety of weather characterizes warm fronts. The weather may even vary along a given front.
The degree of overrunning and the moisture content and stability of the overrunning warm air determine the seventy of the weather. If the warm air is very moist, the cloud deck forming in the overrunning air may extend for hundreds of miles up the slope of the retreating cold air. It the warm air is unstable, thunderstorms may be embedded in the cloud deck.
High cirrus cloud is the first sign of the approach of an active warm front. Cirrostratus soon follows
(the high thin cloud which causes a halo around the sun or moon). The cloud gradually thickens and the base lowers until a solid deck of altostratus/altocumulus covers the area. Low nimbostratus moves in, merging with the altostratus. With the result that a solid deck of cloud extending from near the surface to 25,000 feet or more covers the whole area. Precipitation is usually heavy.
weather changes
Windshift: With the passage of a warm front, the wind will veer, but the change will be much more gradual than in the case of a cold front.
When flying through a warm front, the windshift will occur at the frontal surface and will be more noticeable at lower levels. When flying through a warm front, the windshift is such that a course alteration to starboard is necessary.
Temperature : The warm front brings a gradual rise in temperature. A pilot flying through the frontal surface will notice a more abrupt temperature rise.
Visibility : Low ceilings and restricted visibility are associated with warm fronts and, because warm fronts usually move quite slowly, these conditions persist for considerable time.
When rain falls from the overrunning warm air, masses of irregular cloud with very low bases form in the cold air. Fog is frequently a condition 50 nautical miles ahead of an advancing warm front.
Turbulence : Cumulonimbus clouds are frequently embedded in the main cloud deck and these storms are responsible for the most severe turbulence associated with a warm front. However, these storms and the turbulence they occasion are less severe than those associated with cold fronts. The principal problem with these storms is that they cannot be located by sight since they are embedded in the main cloud cover.
Precipitation : The first precipitation begins in the region where the altostratus layer of cloud is from
8000 to 12,000 feet above the ground. As the front approaches, the precipitation becomes heavier.
Occasional very heavy precipitation is an indication of the presence of thunderstorms.
winter warm fronts
In winter, when temperatures in the cold air are below freezing and temperatures in the lower levels of the warm air are above freezing, snow and freezing rain can be expected.
Snow falls from that part of the warm air cloud that is high and therefore below freezing in temperature. From the lower cloud, where temperatures are above freezing, rain falls. However, as the rain falls through the cold air (of the cold air mass that the warm air is overrunning), it becomes supercooled and will freeze on contact with any cold object. This is known as freezing rain (ZR).
In the area ahead of the freezing rain, there is a region where the rain falling through the cold air becomes sufficiently supercooled to freeze and falls to the ground as ice pellets (IP). A pilot approaching the frontal surface at higher altitudes may not encounter the ice pellets, but the pilot

flying at quite low altitudes can expect to encounter snow, ice pellets and then freezing rain.
Icing is a problem associated with warm fronts in winter. Snow is not responsible for icing, unless it is very wet when it can stick to an airplane and form ice. Freezing rain, however, causes a rapid build up of ice. Icing will also be a problem in the cloud layers.
weather at trowals and upper fronts
The weather that occurs with a trowal is a combination of cold and warm front conditions. The cloud pattern ahead of the approaching trowel is similar to that of a warm front. Cold front cloud formations will exist behind it. Cumulus buildups and thunderstorms are likely to be interspersed with stratiform clouds, continuous precipitation and widespread low ceilings. In winter months, freezing rain and severe icing conditions are likely hazards as the rain aloft in the occluded warm air falls through the freezing temperatures of the ground based cold sectors. The maximum precipitation, convective activity and icing conditions usually occur in the northeast sector of the low and extend some 50 to 100 miles ahead of the occluded front.

c e i l i n g a n d v i s i b i l i t y
The ceiling is the height ascribed to the lowest layer of clouds or obscuring phenomena when it is reported as broken, overcast, or obscuration and not classified as "thin" or "partial." The ceiling is termed unlimited when these conditions are not satisfied.
Visibility
The greatest distance at which it is just possible to see and recognize with the unaided eye (1) in the daytime, a prominent dark object against the sky at the horizon, and (2) at night, a known, preferably unfocused, moderately intense light source.
flight visibility is the average range of visibility forward from the cockpit of an airplane in flight.
slant range visibility is the distance a pilot can see over the nose of the airplane towards the ground.
It is sometimes called approach visibility.
ground visibility is the visibility at an airport as reported by an accredited observer.
prevailing visibility is the distance at which objects of known distance are visible over at least half the horizon. It is reported in miles and fractions of miles.
Runway visual range (RVR)
The maximum distance along the runway at which the runway lights are visible to a pilot at touchdown. Runway visual range may be determined by an observer located at the end of the runway, facing in the direction of landing, or by means of a transmissometer installed near the end of the runway.
VMC and IMC
Visual meteorological conditions (VMC)indicates that visibility, distance from cloud and ceiling are equal to or better than the minimum under which flight according to the visual flight rules (VFR) may be conducted.
Instrument meteorological conditions (IMC)indicates that visibility, distance from cloud and ceiling are below minima and flight can be conducted only under instrument flight rules (IFR).

u n d e r s t a n d i n g t h e j e t s t r e a m s
'Jet streams' were first discovered during the second world war. Pilots were regularly flying between
Britain and the United States of America and they noticed that it was quicker to fly to the United
Kingdom and reported tailwinds of over 100 miles per hour. These winds blew in narrow ribbons and were named 'jet streams'.
Nowadays jet streams are closely monitored and forecast. Pilots want to know where to find them as their added push will save them time and fuel, and therefore money. But jet streams are not only important to pilots. When Breitling Orbiter 3 became the first balloon to fly non-stop around the world it used knowledge of the position of the jet streams to speed up its flight.
They are relatively narrow bands of strong wind in the upper levels of the atmosphere. The winds blows from west to east in jet streams but the flow often shifts to the north and south. Jet streams follow the boundaries between hot and cold air. Since these hot and cold air boundaries are most pronounced in winter, jet streams are the strongest for both the northern and southern hemisphere winters.
Since the earth rotates, the axis is tilted, and there is more land mass in the northern hemisphere than in the southern hemisphere, there are three global circulations...
1.
Hadley cell - Low latitude air movement toward the equator that with heating, rises vertically, with poleward movement in the upper atmosphere. This forms a convection cell that dominates tropical and sub-tropical climates.
2.
Ferrel cell - A mid-latitude mean atmospheric circulation cell for weather named by Ferrel in the 19th century. In this cell the air flows poleward and eastward near the surface and equatorward and westward at higher levels.
3.
Polar cell - Air rises, diverges, and travels toward the poles. Once over the poles, the air sinks, forming the polar highs. At the surface air diverges outward from the polar highs.
Surface winds in the polar cell are easterly (polar easterlies).

The earth's rotation is also responsible for the jet stream to move from West to East. The motion of the air is not directly north and south but is affected by the momentum the air has as it moves away from the equator. The reason has to do with momentum and how fast a location on or above the Earth moves relative to the Earth's axis.
Your speed relative to the Earth's axis depends on your location. Someone standing on the equator is moving much faster than someone standing on a 45° latitude line. In the graphic (left) the person at the position on the equator arrives at the yellow line sooner than the other two. Someone standing on a pole is not moving at all (except that he or she would be slowly spinning). The speed of the rotation is great enough to cause you to weigh one pound less at the equator than you would at the north or south pole.
The momentum the air has as it travels around the earth is conserved, which means as the air that's over the equator starts moving toward one of the poles, it keeps its eastward motion constant. The
Earth below the air, however, moves slower as that air travels toward the poles. The result is that the air moves faster and faster in an easterly direction (relative to the Earth's surface below) the farther it moves from the equator.
In addition, with the three-cell circulations mentioned previously, the regions around 30° N/S and 50°-
60° N/S are areas where temperature changes are the greatest. As the difference in temperature between the two locations increase, the strength of the wind increases. Therefore, the regions around
30° N/S and 50°-60° N/S are also regions where the wind, in the upper atmosphere, is the strongest.

The 50°-60° N/S region is where the polar jet located with the subtropical jet located around 30°N.
Jet streams vary in height of four to eight miles and can reach speeds of more than 275 mph. The actual appearence of jet streams result from the complex interaction between many variables - such as the location of high and low pressure systems, warm and cold air, and seasonal changes. They meander around the globe, dipping and rising in altitude/latitude, splitting at times and forming eddies, and even disappearing altogether to appear somewhere else.
Jet streams also "follow the sun" in that as the sun's elevation increases each day in the spring, the jet streams shifts north moving into Canada by Summer. As Autumn approaches and the sun's elevation decreases, the jet stream moves south into the United States helping to bring cooler air to the country.