As its name implies, this jet stream is associated with the marked discontinuity found at the boundary of well defined air masses - polar to the north/sub tropical to the south (in the northern hemisphere), conventionally found at the polar front. It meanders markedly in response to global/regional scale atmospheric changes but has a latitudinal 'home' roughly from 45 to 65 deg N/S. Its altitude is somewhere between 28000 ft to 34000 ft (8.5 - 10.5 km), with its own distinctive tropopause level. Speeds are of the order 80-130 knots (40-65 m/s), but may be as high as 180 knots (90 m/s), and downstream of main continental land masses in late winter/early spring, in excess of 200 knots (100 m/s). Although it is regarded as a 'single' ribbon of strength encircling the hemispheres, in reality the jet is broken and in developmental situations, can become very distorted with new jets re-forming at different levels from the 'main' baroclinic jet. (NB: during the winter half-year, jet-streams can be found at even higher latitudes in both hemispheres, roughly around 18000 ft (circa 5.5 km), which are named Arctic (or Antarctic) jets. These are again tied to a discontinuity between air masses (a frontal zone, or Arctic / Antarctic front), but this time between a very cold (and relatively shallow) 'arctic' airmass and somewhat less cold polar-maritime or polar-continental air. They are most evident (in the Northern Hemisphere), across Canada, the far north of the USA, northern Scandinavia, North Russia & other high-latitude regions. However, at times, they come further south, interacting with the PFJ, with sometimes dramatic developmental consequences.)
The average level of the core of this westerly jet lies at an altitude of about 40 000 ft/12 km, just below the tropical tropopause. It occurs in the latitude range 25-40 deg N/S, and is most marked during the winter and early spring of each hemisphere, but is not associated with any surface frontal structure. Its existence owes more to the fact that air in the high-level (poleward travelling) leg of the Hadley circulation conserves its angular momentum, being effectively 'turned' towards the east and finally concentrated in this ribbon of strong westerly wind. Because the Hadley circulation, and mid-latitude north/south linking flows are governed by seasonal heat differences north-to-south, then the STJ does vary, as noted above. Wind speeds are generally 80-150 knots (40-75 m/s), but can be much greater over eastern seaboards of large land masses, e.g. speeds of 400 kt (200 m/s) have been reported over east Asia/NW Pacific. From time-to-time, the STJ and the Polar Front Jet can interact with marked developmental consequences.
This occurs at times during the winter and early spring when the stratosphere near the poles is much colder than it is further south due to the absence of insolation at these times of the year. Its direction is westerly overall, with high variability, and has speeds in the range 100-200 kt (50-100 m/s) at altitudes around 70 000 ft/21 km and occurs on the poleward side of latitude 70 degrees.
This jet occurs in the northern summer between 10 and 20 deg N, chiefly over or just to the south of high land masses such as in Asia and Africa. Its occurrence is due to a temperature gradient with colder air to the south which produces sufficient temperature differential above 50 000 ft/15 km to give wind speeds of over 100 kt (50 m/s). Because colder temperatures at height are to the south, it is an easterly jet. (This jet is now more usually known as the Tropical Easterly Jet(TEJ) ... perhaps more correctly as it lies some distance from the Equator.)
In NW Europe, when meteorologists refer to a jetstream, it's the Polar Front Jet (PFJ), that is usually meant. As its name implies, it is associated with the classical 'Norwegian model' polar front - the surface discontinuity between cold/ex-polar latitude air, and the warm, relatively moist air originating in the sub tropical anticyclone belt.
When air masses (see "When was the concept of an air mass proposed?" and "So, how is an air mass defined?") lie adjacent to one another, the temperature difference isn't just found at mean sea level, but throughout the troposphere. Because atmospheric pressure decreases more quickly with height in cold/polar air than warm/sub-tropical air, there arises a pressure differential, which gives rise to intense pressure gradients at altitude, and hence the very strong winds observed. Because of the high wind speeds involved with jetstreams, any slight changes, in either velocity or direction, or both, leads to vertical motion in the air below the jet, and is a major player in the processes of atmospheric development.
(For more on upper air meteorology, jetstreams etc., see "Upper Air Meteorology").
First, to visualise what 'stable' and 'unstable' states mean in a physical sense, stand a round pencil on end on a level surface. From Newton's First Law of Motion, it will remain upright until a force is applied. Once displaced, the pencil falls over, failing to pass through its original (upright) position. This is the UNSTABLE state. Now lay the pencil on its side, at the bottom of an incline. Displace the pencil slightly up the incline, then remove the force of displacement. The pencil will return to its original position. This is the STABLE state.
In the atmosphere, whether air that is displaced does so in an unstable or stable environment depends upon the vertical temperature profile of the air -- its lapse rate -- and upon the moisture content of the parcel. These differences are fundamental to understanding why clouds take up the form they do.
In the atmosphere, when a 'parcel' of air moves vertically upwards (or downwards), it cools (upward motion), or warms (downward motion), in accordance with thermodynamic rules ... if the air is unsaturated (air temperature > dew point temperature), the cooling/warming will be at a rate of 3 degC per 1000 ft (or 10 degC per 1 km): This is known as the Dry Adiabatic Lapse Rate/DALR; If the air is/becomes saturated (air temperature=dew point temperature), this rate is roughly halved in the lower troposphere, due to the release of latent heat upon condensation. This rate is known as the Saturated Adiabatic Lapse Rate/SALR.
Such ascent/descent is said to be adiabatic, which means that the energy/heat changes are confined to that particular parcel.Provided the parcel is warmer (less dense) than the environmental air through which it is passing, it is buoyant, and rises. If the parcel is colder (denser) than ambient air, then it will descend, or try to descend. Because the rates of cooling (ascent), and warming (descent) of individual parcels are fixed, the important variable is the overall lapse rate (i.e. the rate of change of temperature with height) of the atmosphere. On average, this is 1.98 (call it 2 degC) per 1000 ft, or 6.5 degC per 1 km in the troposphere, but this average conceals a wide variety of cases which are important in meteorology.
Where the temperature falls off slowly with height, or indeed rises, e.g. in a slow moving anticyclone, or a tropical maritime airmass, then an air parcel subject to lifting/adiabatic cooling will readily find itself colder than its surroundings ... denser ... and try to return to its original position: The air is ABSOLUTELY STABLE. Where the temperature falls off quickly with height, e.g. in a cold/polar air mass over NW Europe in late winter/spring, then an air parcel subject to ascent, although cooling, may still find itself warmer/less dense than its surrounding air ... it will be buoyant, and tend to rise further: the air is ABSOLUTELY UNSTABLE.
Problems arise when, on ascent, the dew point of the air is reached, and the rate of cooling is therefore less - it follows the SALR figure. If, however, the parcel is still warmer/less dense, then it will continue to rise, and the condition of the air is said to be CONDITIONALLY UNSTABLE .. i.e. conditional upon whether the parcel is saturated or not. This is by far the most common situation in the 'real' atmosphere, accounting for some 65-70% of situations taking the troposphere as a whole.
Stable airmasses generally imply the absence of 'free' vertical motion, and any ascent that does occur must be forced, i.e. frontal (dynamic or mass) or orographic (mechanical) ascent, and the cloud structure is essentially layered. (NB: Forced ascent comes about in several ways: frontal ascent due to large-scale air motion within frontal systems, with of course adjacent descent; convergence into an area of low pressure - the converging air can't go down near the surface - it has to go up; and topographical forcing, that is, air is forced to rise over major upland ranges. )
Unstable airmasses imply free vertical motion (given an initial trigger action), and the cloud structure is 'heaped' or cumuliform. If the vertical motionis vigorous and deep enough, and there is sufficient moisture, then heavy showers/thunderstorms are likely. (NB: Trigger action: method of causing air to rise initially, which in the lower troposphere include not only the 'wide-area' triggers noted above under stable conditions, but also smaller/mesoscale mechanisms such as differential heat response between land and sea, coastal convergence, etc.)
For more information on these subjects, see a good textbook on meteorology, for example, Essentials of Meteorology:(Taylor and Francis/D.H.McIntosh and A.S.Thom).
'Thickness' is a measure of how warm or cold a layer of the atmosphere is, usually a layer in the lowest 5 km of the troposphere; high values mean warm air, and low values mean cold air. It would be perfectly feasible to define the average temperature of a layer in the atmosphere by calculating its mean value in degrees C (or Kelvin) between two vertical points, but an easier, practical way to measure this same mean temperature between two levels can be gained by subtracting the lower height value of the appropriate isobaric surface from the upper.
Thus one measure of thickness commonly quoted is: height (500 hPa surface) - height (1000 hPa surface)
The 500-1000 hPa value is used to define 'bulk' airmass mean temperature, and can be seen on several products available on the Web.
For more information see here.
(see also here for typical figures, extremes etc.)
The atmosphere is divided up into layers with names which describe the dynamic or thermal structure of that particular layer. The two layers which are of most interest to us are the troposphere and the stratosphere.
Troposphere: (overturning or changing sphere) - The lowest layer of the atmosphere. Positive lapse of temperature (positive lapse rate: temperature overall decrease with height). It is the most important for operational meteorology, as this layer contains almost all the water vapour, and by far the greatest part of the mass of the atmosphere. Because of its mean thermal structure, it is the region of greatest vertical motion (up and down) in the atmosphere, even without the help of vigorous thunderstorm complexes, which in themselves may occupy the entire depth of the troposphere. At some level, there is usually an abrupt change in the lapse rate from positive (decrease with height), to isothermal (no change), or a slight rise. This level is the tropopause. Typical heights of the tropopause, and therefore thickness of the troposphere, are:
In mid-latitudes, the temperate zone, which is of most interest to us in NW Europe, the tropopause is highly variable, from cold to warm season, and from cold to warm air mass. For example, it is lower in winter, and in cold/polar air masses (typically 8 to 10 km/25000 to 30000 ft), than in high summer, and in warm/sub tropical air masses (typically 12 to 14 km/35000 to 45000 ft)
Stratosphere: (the 'layered' sphere) - the next layer ascending through the atmosphere. Isothermal or negative lapse rate of temperature (i.e the temperatures rises with increasing height). Because of this temperature structure, little natural, or un-forced overturning of air takes place, either within the stratosphere, or in exchange with the troposphere. Once gases, particulates etc. penetrate to this layer, they remain there for very long periods, hence the concern regarding such substances due to both the actions of mankind (e.g. CFCs) and those of natural processes (e.g. volcanic ash). However, near the boundary with the troposphere (q.v.), marked vertical motion can occur under certain circumstances (forced by jet-stream actions), which are important in driving developments in the troposphere.
As with the troposphere, the stratosphere varies in thickness, but as an average figure the top of this layer, the stratopause, occurs around 45-48 km (148000-158000 ft).
The importance of the stratosphere (and the primary reason for its temperature structure) is that much of the atmospheric ozone is found within its lower layers - circa 18 to 30 km amsl. The selective absorption by ozone (and oxygen) of solar ultra-violet radiation leads to warming in the stratosphere - this (and other) factors give rise to its markedly stable nature. Very little water vapour is found here, nor dust (except for dust injected by major volcanic eruptions), but when the stratosphere is anomalously cold, then Polar Stratospheric Clouds (PSC) are sometimes visible.
These equivalents are based on the International Standard Atmosphere and promulgated by ICAO:
|mbar (hPa)||Nominal Altitude |
(ft to nearest 1000 ft; metres to nearest 100 m)
|100||53,000 ft / 16,200 m|
|200||39,000 ft / 11,800 m|
|250||34,000 ft / 10,400 m|
|300||30,000 ft / 9,200 m|
|400||24,000 ft / 7,200 m|
|500||18,000 ft / 5,600 m|
|600||14,000 ft / 4,200 m|
|700||10,000 ft / 3,000 m|
|850||5,000 ft / 1,500 m|
For full details, see this article. The article also has the definitions of QFE, QNH, QFF and QNE.
There are four principal types of satellite imagery used in operational and research meteorology. Each has its advantages and disadvantages. Many examples of each type can be found at meteorology related web-sites.
1. Visible Imagery (VIS)
Images obtained using reflected sunlight at visible wavelengths, in the range 0.4 to 1.1 micrometres. Visible imagery is displayed in such a way that high reflectance objects, e.g. dense cirrus from CB clusters, fresh snow, nimbostratus etc., are displayed as white, and low reflectance objects, e.g. much of the earth's surface, is dark grey or black. There are grey shades to indicate different levels of albedo (or reflectivity). Very dependent upon angle of incident sunshine, and of course, not available at night, though some military/research satellite sensors can utilise reflected moonlight to detect cloud.
2. InfraRed (IR)
These images are obtained by sensing the intensity of the 'heat' emissions of the earth, and the atmosphere/atmospheric constituents, at IR wavelengths in the range 10-12 micrometres. The earth, and its components, radiate across a wide spectrum of wavelengths, but for many of these, the atmospheric gases, of which water vapour is an important constituent, absorb a significant proportion of such radiation. Thus so-called 'windows' need to be chosen to allow the satellite sensors to detect such radiation unhindered, and the 10-12 micrometre band is one such. IR imagery is so presented that warm/high intensity emissions are dark grey or even black, and low intensity/cold emissions are white. This convention was chosen so that the output would correspond with that from the VIS channels, but there is no need to follow this scheme - indeed in operational meteorology, colour slicing is frequently used whereby different colours are assigned to various temperature ranges, thus rendering the cooling/warming of cloud tops (and thus the development/decay) easy to appreciate: warming/darkening of the imagery with time indicates descent and decay; cooling/whitening images imply ascent and development.
3. Water Vapour (WV)
This imagery is derived from emissions in the atmosphere clustered around a wavelength of 6.7 micrometre. In contrast to the IR channel, this wavelength undergoes strong absorption by WV in the atmosphere (i.e. this is not a 'window'), and so can be used to infer vertical distribution and concentration of WV - an important atmospheric constituent. WV imagery uses the radiation absorbed and re-emitted by water vapour in the troposphere. If the upper troposphere is moist, WV emissions will be dominated by radiance from these higher levels, swamping emissions from warmer/lower layers; this radiation is conventionally shown white. If the upper troposphere is dry, then the sum of the radiation is biased towards lower altitude WV bands: it is warmer/less intense radiation, and this is displayed as a shade of grey, or even black. WV imagery is very important in the study of cyclogenesis, often being displayed as a time-sequence.
4. 'Channel 3' (CH3)
Imagery from a specific wavelength of 3.7 micrometre, lies in the overlap region of the electro-magnetic spectrum between solar and earth-based/terrestrial radiation. It is sometimes referred to as 'near infrared' (NIR). CH3 images use a mixture of back-scattered solar radiation plus radiation emitted by the earth and atmosphere. It is used in fog/very low cloud studies. Interpretation is sometimes complex, especially in the presence of other tropospheric clouds.
In the troposphere (the 'weather' zone ... see here, the layers are divided up into three broad levels: (approx heights only)
|Polar latitudes||Temperate regions||Tropics|
|High||10 000 - 25 000 ft |
3 - 8 km
|16 500 - 45 000 ft |
5 - 14 km
|20 000 - 60 000 ft |
6 - 18 km
|Medium||6 500 - 13 000 ft |
2 - 4 km
|6 500 - 23 000 ft |
2 - 7 km
|6 500 - 25 000 ft |
2 - 8 km
|Low||Surface -- 6 500 ft |
up to 2 km
|Surface -- 6 500 ft |
up to 2 km
|Surface -- 6 500 ft |
up to 2 km
The heights assigned to the 'divisions' between levels should not be followed slavishly, and assignment of clouds to the various 'groups' should be made with the appearance and composition in mind.
High clouds are primarily composed of ice crystals; Medium clouds are a mixture of water droplets (usually super-cooled) and ice crystals, in varying proportion, and low clouds primarily water droplets, but in individual cases these descriptions are probably simplistic.
(NB: Super-cooled: means that although the temperature of the droplet is below 0 deg.C, it remains liquid - this is a common state in the middle part of the troposphere.)
In the 'Low' cloud classification come: Stratus (St); Stratocumulus (Sc); Cumulus (Cu) and Cumulonimbus (Cb). However, note that both Cumulus and Cumulonimbus clouds often extend well into 'medium' levels, and towering Cu, and Cb extend to 'high' levels.
In the 'Medium' cloud class come: Altostratus (As); Altocumulus (Ac) and Nimbostratus (Ns). Nimbostratus often has a base within the 'low' cloud category.
In the 'High' cloud group are: Cirrus (Ci); Cirrocumulus (Cc) and Cirrostratus (Cs).
The white trails are ribbons of ice crystals. As a by-product of the exhaust of aircraft engines, water vapour is trailed from the engine exhaust which adds to the local humidity of the air the aircraft is flying through, and which tends to super-saturation of the air. However, the exhaust gases are of course hot, and so these hot gases help to raise the temperature of the air and thus is can hold more vapour before saturation is reached. There are therefore two opposing mechanisms at work: the water vapour in the exhaust trying to saturate the air; the hot gases of the exhaust trying to decrease relative humidity. When the balance between outside air temperature (OAT) and local humidity is just right, then condensation trails will occur: usually abbreviated to CONTRAILS, and sometimes referred to, from old coding conventions, as COTRA.
Persistent condensation trails can last for many hours, gradually spreading out to form large, sometimes dense areas of cirriform cloud; they can have dimensions typically several kilometres wide and several hundred metres in depth (thus they can be seen in visible satellite imagery). They spread because of turbulence at the 'trailing' level ( enhanced by the aircraft passage), differences in wind speed along the flight-path and there is also thought to be a contribution from solar heating. Because trails can last so long and come to dominate the upper troposphere in any particular synoptic situation, the production of such form part of the debate on the overall global radiation balance.
Wake trails: As an aircraft passes through a lower troposphere having a high relative humidity, (usually during landing or take-off phases - and for military aircraft, during 'high-G' manoeuvres), very short, non-persistent 'trails' can sometimes be seen coming from the wing tips, or white 'lift-generated sheets' streaming off from the trailing edges of the main wing, control surfaces etc. Both features are due to short-term local reduction of pressure, leading to condensation, though the precise mechanism in each case is different.
(a): "lift-generated sheets": as an aircraft moves forward, air accumulates (pressure builds), at leading edges, with a compensating depletion of air (fall of pressure) across the top of the wing (generating lift) and along trailing surfaces. The reduction of temperature in the near-saturated environment, consequent upon the slight lowering of pressure, can be enough to cool the air to it's dew point, and thin sheets of water droplets are observed.
(b): "wing-tip trails": the flow of air around the wing-tips undergoes marked distortion which manifests itself as a tight-vortical (or 'twisting') motion of the airflow; the vortices are formed by, and will lead to, a local increase and decrease of pressure - in the latter case, if the atmosphere is humid enough, then white trails can be observed. In both cases, the sheets/trails (of minute water droplets) will evaporate quickly again due to mixing with the non-saturated environment in the wake of the aircraft.
Dissipation trails (DISTRAILS): In contrast to the formation of CONTRAILS (see FAQ here ), aircraft on passage at high levels can cause the dissipation of pre-existing cirriform cloud, due to the local increase in temperature consequent upon the ejection of hot exhaust gases from the aircraft engine. The passage of the aircraft will be marked by a clear lane in the cloud. However, it will be obvious from the description (above) relating to condensation trails, that the heat outflow must markedly outweigh the injection of water vapour from the spent fuel, and the phenomenon is rare. The effect may also be caused by turbulent mixing with dry air just above the cloud layer, caused by the aircraft motion, and this mechanism can lead to temporary clear lanes in other cloud forms, e.g. thin stratocumulus or altocumulus. However, beware of a similar phenomenon, whereby the shadow of a 'normal' condensation trail is cast on thin cirriform cloud below - leading to a visibly dark band in the cloud. This is not a dissipation trail.
Incidentally, whilst on the subject of 'trails', if you are looking at visible satellite imagery over the region of a slow-moving anticyclone, and notice lots of thin, white lines criss-crossing the region, which don't appear on the corresponding InfraRed image, these are ships' trails, caused by exhaust particles from the vessel acting as condensation nuclei, and 'seeding' the humid, near sea surface environment, and betraying the presence of the ship by a thin band of water droplets which are not dispersed due to the very light winds and minimal mixing in the anticyclone.
The average condition of temperature change in the Troposphere is for there to be an overall decrease of temperature with increasing height: a positive lapse rate (see here). However, in the 'real' Troposphere, frequent reversals of this 'normal' lapse are observed, particularly in the lower layers - these zones of increasing temperature with height are inversions (i.e. the inverse of the average state), and are very important for both synoptic/mesoscale meteorology (e.g. fog/stratus formation/dispersal), and pollution dispersion studies, as they cap layers of markedly stable and potentially stagnant air masses.
Examples of inversions include those due to anticyclonic subsidence; cooling land by night (nocturnal inversions); and sea-breeze inversions, where cooler sea air under-cuts warmer land air. Where the inversion is associated with an abrupt lowering of the moisture content (sharp fall of dew point), at the altitude of the temperature rise, then interesting radio-refraction conditions occur, familiar to viewers of terrestrial television in stagnant anticyclonic episodes.
In fact, if you are caught out in one, there is no difference. You can still get wet! Meteorologists however distinguish between precipitation (rain, snow, hail etc.) falling from cumuliform cloud in an unstable environment - a shower, from that falling from layer clouds in a generally stable environment which are just called rain, snow, sleet etc.
However, rain from layer cloud in a frontal situation for example can be rather hit-and-miss, especially in a weakening situation, and so forecasters will try and get around such problems by talking about 'patchy rain', 'outbreaks of rain', 'splashes of rain' etc. The opposite problem comes when a well defined trough sweeps across an area, in which the cloud structure is most likely of an unstable type: cumulus, cumulonimbus and altocumulus.
Given the definition above, the short, very sharp falls of rain might be called 'showers' (and probably coded as such by observers), but this would be misleading to members of the public caught out in such precipitation: hence the 'showery outbreaks of rain', 'showery bursts of rain', 'localised downpours' etc.
In day-to-day meteorology, the temperature of the lowest layer of the atmosphere is measured at a height of 1.25 m (about 4 feet) above local ground level. Usually, though not always, this is achieved by placing thermometers in a double-louvered screen with the bulbs of the thermometers, or the sensor heads (for distant reading thermometers), placed so that they cluster around the 1.25 m standard. The temperature so read is usually called 'the air temperature' and it is these values that appear, for example, in the World Cities reports in newspapers/teletext, or plotted on standard synoptic charts, and also it is at this level that the forecast temperatures seen on tv weather maps are based.
When the temperature as measured in this way falls below 0.0 deg C, then an AIR FROST is recorded. For other purposes though, e.g. horticulture, road gritting operations etc., we need to know what the temperature is at the surface of the ground, and most weather stations set at least two thermometers to record these values: a grass minimum thermometer, set just above/in contact with short grass, and a concrete minimum thermometer, set so that its sensor/bulb is in contact with a concrete slab of standard dimensions/composition.
When the temperature as measured by the thermometer set over grass falls below 0.0 deg C, then a GROUND FROST is recorded. (In spring and early summer, when the temperature is expected to produce a frost using the grass minimum thermometer, but not over other surfaces (due to thermal inertia of surfaces such as concrete, tarmac etc.), then the unofficial term 'grass-frost' may be heard in weather forecasts - this is to try and avoid panic by road, railway and airport operators as soon as they hear the word 'frost' but alert gardeners, growers etc., to the risk of damage).
The difference between the two levels can be considerable: On still, clear nights, with air of a low humidity content, 5 degC or more is not uncommon.
A common misconception, is that it must be coldest in the middle of the night, and warmest around midday. On some occasions, mainly due to air mass changes, this may be correct, but not usually. The lowest (minimum) temperature usually occurs a little while after sunrise, and the highest (maximum) temperature usually occurs after midday --- sometimes as late as 3 or 4 hours after midday.
To understand why, it is necessary to consider that thermal energy during the 24 hours is radiating continually from the surface of the earth (at long wavelengths), and incoming solar (relatively short wave) radiation obviously only when the sun is above the horizon. With the sun below the horizon (night), outgoing radiation allows the surface to cool, and the temperature drops. After sunrise, incoming solar radiation counteracts this loss of heat, but only after a lag - which can be up to an hour or so in winter with a low solar elevation.
The minimum temperature occurs when there is a balance between outgoing and incoming radiation. As the sun rides higher in the sky, increasing amounts of short-wave radiation are available to heat the ground, and therefore available to heat the overlying air. Although outgoing land-based radiation is also increasing, solar heating is dominant. The temperature rises, until, past noon, incoming solar radiation starts to decline again.
The highest(maximum) temperature occurs when heat gain due to incoming solar radiation, and heat loss due to outgoing terrestrial radiation balance: this occurs some time after midday.
For any particular sample of air, which is cooled at constant pressure, there will be a temperature below which water vapour condenses to form liquid water drops, assuming sufficient hygroscopic nuclei present. That temperature is known as the Dew Point and is a measure of the Absolute Humidity (see "What is the difference between Humidity and Relative Humidity?").
Absolute Humidity, often just referred to as 'the humidity', is a measure of the actual amount of water vapour in a particular sample of air: measured as a partial pressure (vapour pressure/hPa or millibars); a mixing ratio (gm water vapour/kg of dry air), dew point etc.
Relative Humidity - expressed commonly as a percentage value, is the ratio of the actual amount of water vapour present in a sample (the Absolute Humidity) to that amount that would be needed to saturate that particular sample.
The two terms are not interchangeable and can lead to confusion; e.g. on a cold, raw winter's day close to the east coast of England, the dew point might be 1 degC and an air temperature of just 2 degC. This would give a RH of 93%; a 'high' Relative Humidity, yet few would refer to such conditions as 'humid'. Conversely, on a hot summer's day, with a dew point of 18 degC, and an afternoon temperature of 30 degC, that's a RH of 49%; a 'low' Relative Humidity, but high Absolute Humidity.
Only for the special case of thunderstorms coming up from the south in summer. I have seen many thunderstorms (real crackers as well) in April with air temperatures of 8 degC and a dew point of 4 degC. What is really important is that the air must be unstable (see "stable and unstable air masses"), usually achieved by warming at the bottom or by cooling high up or both. Then you need a trigger to release the instability, usually heating and input of moist air (high dew point), but if the air is unstable enough just the heating will do. Other triggers are forced lifting of air over hills or forced lifting by convergence (e.g. sea breezes).
(thanks to Will Hand for this answer)
At mid to high latitudes in the upper part of the troposphere (above roughly 5 km ), the mean wind flow exhibits a broadly west-to-east motion - this applies in both hemispheres. On many occasions, particularly in mid-latitude/temperate zone regions, the flow is directed more or less directly from west to east, crossing few latitude zones within the same longitude range: this is a 'highly zonal' type - any short-wave disturbances embedded in the flow will be carried quickly along and the weather is ever-changing as a succession of frontal systems, interspersed with transient ridge conditions cross any one point.
However, on both average (e.g. monthly) pressure maps and on individual days, long-wave trough/ridge patterns can be found - some having large amplitude, i.e. the airflow meanders a long way north and south around the loops of the pattern, crossing many parallels of latitude in a relatively limited longitudinal range: a 'meridional' type; Usually, some west-to-east progression of the looped pattern can be seen over a 24 hr period, and the associated surface weather type changes, albeit more slowly than the zonal type described earlier.
However, if the 'loops' in the pattern become locked in one geographical area, then depending where you are in relation to the upper flow, the associated surface patterns are often little changed from one day to another, and in extreme cases, from one week to another - the pattern is said to be 'blocked'.
In, and just to the east of a slow-moving trough in the upper flow, the surface weather will tend to be of a low pressure/convective/showery type, and perhaps cool for the time of year (but not necessarily).
In, and just to the east of a static ridge in the flow, the surface pressure will tend to be high, with settled conditions lasting until the block is destroyed. This latter case is responsible for prolonged dry/hot weather in summer, but cold/sometimes grey conditions in winter, and considerable pollution build-up can occur at all seasons due to the stagnation of the lower level air and high air-mass stability encountered.
For a personal view of some aspects of upper air meteorology, and some further explanation of the terminology used, see "Upper air meteorology".
A trough on a mean sea level pressure chart, (or an upper air contour chart) can be picked out by an arrangement of isobars (contours) which are concave towards an area of low pressure (low contour height) along a particular axis, and that axis is defined so as to lie along the points of maximum curvature on the individual isobars (contours).
If this sounds complicated, it isn't really: the feature is analogous to the 'valley' on an OS map and defined in the same way - pressure, or contour heights 'fall' into the trough line. A front may have troughing along its length, but not all troughs are frontal! Indeed, not all troughs have 'weather' associated with them in the cloud/rainfall sense. Lee troughs found downwind of a major range of hills/mountains are often cloudless, and thermal troughs forming over land during the day due to mesoscale heating may only be found by careful drawing of isobars: if the air is dry and/or stable, little significant cloud will be associated with this feature.
A modern complication on charts used on the GTS is that plumes of high humidity...e.g. in the case of very humid/warm air coming northward out of France/Iberia, are also shown as 'trough' lines for want of any other identifier. Although with development pressure may become lower along this 'plume' than surrounding areas, and therefore qualify as a trough by the above definition, often the difference is small or initially non-existent.
Find the distance between adjacent isobars in the area that you are interested in - making sure that the isobaric interval is the same as that for which the scale was constructed - often 4 mbar. (Dividers can be used, but a strip of paper suitably marked is just as good.) Using the geostrophic scale for the correct latitude, put one end of your marked distance on the left-hand end of the scale, and read off at the right-hand end the geostrophic wind speed for that isobaric spacing at that latitude.
Remember though that many corrections are needed to find an approximation to the 'real' wind. See the Glossary and any good book relating to meteorology.
(see also "Thickness: what is it?": thanks to Jon O'Rourke for looking up the extreme values.)
As already noted elsewhere, the values of the (total) thickness between levels at 500hPa and 1000hPa give a useful measure of the mean temperature of that layer. In summer, values might range from 546dam (cool, showery northwesterly) to 560dam (warm, settled anticyclonic spell); in winter from 530dam( brisk, chilly, showery flow, with inland night frosts) to 550dam (mild, open-warm sector type). (The values are given for comparitive analysis only, and the weather types of course don't necessarily follow from the values); Values below 528dam in winter would herald the arrival of potentially wintry conditions, and in summer, thickness values above 564dam might be a precursor to some notably high temperatures.
I looked up some 1961-1990 average values for a couple of points across Britain (based on the RMetS 'Weather Log' charts & NOAA-CIRES/CDC Re-analysis project).
For the period 1991-2005 (15 years) based on my own records, thickness values have risen (relative to the above & for a point roughly within the CET 'domain') by about 1.2 dam (representing roughly +0.6degC through the layer; 1997 and 2003 showed an increase on the 61-90 climatology of about +3dam/+1.5C. This accords well with expected changes due to anthropogenic global warming.
The increase noted above across 'Central England' is broadly confirmed by data released by the Met Office / Hadley Centre. They show that, relative to the 1961-1990 reference period, 'lower tropospheric' temperatures have increased by about 0.4degC averaged over the years 1991-2005.
As to extremes, for the UK mainland only, the highest Jon & I could find (from rather small-scale maps) came out around 575dam in July over southeast Kent (SE England), and the lowest around 495dam on the extreme tip of NE Scotland in January. However, for the British Isles, we have a low value of 491dam over Shetland, and values below 500dam can be recorded in exceptionally cold easterly types as far south as East Anglia & SE England; for example, in January 1987, total thickness values in these latter areas were certainly below 498dam, and probably briefly around 495dam. (Record low UK day-maxima on the 12th January, 1987). For a graphical representation of maximum and minimum thickness values for 6 points around the NW of Europe, see "Thickness Extremes".
All are formed within an unstable environment (see "Stable and unstable air masses", and all require the following to be in place: (i) Instability through a reasonable depth of the troposphere; preferably (but NOT necessarily) extending above the freezing level; (ii) sufficient moisture to sustain the cloud-building process - medium level dryness will often kill shower formation unless low-level inflow of moisture is substantial; (iii) a trigger action - i.e. something to kick the whole process into life by lifting the parcel that goes on to grow into a moderate depth cumulus cloud, or a well-developed 'supercell' complex.
Once these conditions are met, then consideration of things like shear, CAPE, helicity, etc., are needed as follows:- (for definitions, see the Glossary, and in particular for helicity, see "What is helicity?")
Single-cell showers: the 'classic' growth/decay model of a Cumulus cloud , whereby a single moist convective cell develops in an airmass that is moderately unstable (CAPE values ~ 100 J/kg), provided of course that there is sufficient depth of moisture and there is an initial trigger action. When the updraught and the precipitation downdraught occupy virtually the same atmospheric column (there is little or no vertical relative wind shear to tilt the cloud), the downdraught quickly swamps the updraught - the shower soon decays (perhaps lasting only a matter of minutes - the cloud would last longer though), yielding small amounts of rain/snow. However, when there is a change of wind speed with height (but little directional change), the updraught column is tilted forward, and the resultant precipitation downdraught is held clear of the downdraught, allowing greater development and moderate intensity showers occur. The cold downdraught though soon swamps the inflow of surface air, cutting off the updraught and the shower decays after about 20 to 30 minutes. These events would be typical of Polar Maritime airmasses.
Multi-cell thunderstorms: Whenever wind shear is present in an unstable atmosphere, the developing convective clouds will be tilted to a greater or lesser extent. As seen above (single-cell showers), when only the wind speed changes, then short-lived, non-propagating showers are produced. However, given *both* change of wind speed and direction with height (relative to the storm motion), and sufficiently high CAPE (> ~ 250 J/kg), then the precipitation downdraught is skewed well to the side of the storm updraught, and does not interfere with it - allowing that storm cell to develop its full potential - other necessary factors (e.g. sufficient moisture) being in place. In addition, the downdraught will hit the surface and spread horizontally as a cold density current (gust front). At some point, this will meet the low-level inflow, and a new 'daughter' cell (see the Glossary) may be initiated which may grow into a full-scale storm cell in its own right. This usually (but not always) occurs to the right of the cloud motion, and the whole storm complex appears then to move to the right .. in fact the daughter cells take over from each successive parent to produce this effect. Large Cumulonimbus (Cb) clouds are produced with these processes; each cell lasting at least half-an-hour, and depending upon external forcing agents (e.g. coastal convergence, synoptic troughs, orographic lifting), the storm complexes may last for several hours.
Supercell thunderstorms: Although in some respects, 'supercell' storms can be regarded as a special (and intense) case of the multi-cell storm, there are important differences as well. The environment is still sheared in the vertical, indeed markedly so in the lower layers, and daughter cells are produced. However, a key distinguishing element between supercell and non-supercell events is the presence of a rotating updraught. CAPE values for supercell events will typically be ~1000 J/kg or more, and helicity will also be high - hence the tendency to rotation of the storm complex, and its individual elements. It is thought unlikely, for example, that giant hail would be possible unless the updraught were enhanced by the presence of rotation within the system.
The overall storm motion may be quite small (e.g. Wokingham storm), with the spawned cells forming close to the base of the parent cloud - often several daughter cells coinciding - these form an almost self-perpetuating system lasting several hours. These mechanisms produce the most severe late spring / summertime thunderstorms with local intense rainfall leading to flooding, plus occurrence of hail, possible tornadoes etc. (Note however that slow displacement of such storms should not be assumed - results from North America show displacements in excess of 40 knots / 74 km/hr ). Potential instability at medium levels (circa 500 hPa / 5 to 6km) is also required, as is an initial inhibiting factor (warm / dry air capping surface based instability) to allow the 'loaded gun' effect to build up.
(thanks to Will Hand for much assistance with this answer)
This is a derived parameter which quantifies the tendency for airflow in the lower levels of the troposphere to 'corkscrew' and thus encourage the formation of storms with strong mesoscale circulations, possibly leading to tornadic activity.
Helicity is related to:
(a): speed shear from surface to 3 km (about 700 hPa) - how much the wind speed changes over this altitude band.
(b): directional change of the wind over the same altitude band.
(c): the strength of the low-level wind contributing to the speed / directional shear (as above).
The numerically-greater each of these elements is, the higher is the helicity available for ingestion into a developing storm complex. (It should be noted that the storm will modify the local wind-field, often quite markedly: this means that caution should be exercised when using standard radio-sonde sounding data, or broad-scale NWP output to assess the likelihood or otherwise of severe local storms.)
Helicity has units of energy and can therefore be interpreted as a measure of wind shear energy that includes the directional shear. If there is no directional shear then the helicity is zero: if the wind backs with height then the helicity is negative; if it veers with height (more normal in storms in maritime NW Europe) then the helicity is positive.
Helicity is usually derived in a storm frame of reference, the 'storm relative' helicity, [ Hr ] between the surface and a height, [ h ] and is calculated as an integral between those limits thus: (Vh - C) x Wh x dh [units=m**2/s**2 ] Where [ Vh ] is the environmental horizontal wind velocity , [ C ] is the storm velocity and [ Wh ] is the local relative vorticity. Often [ Hr ] is calculated between expected cloud base and cloud top.
Studies in North America looked at the use of helicity (ignoring sign) for forecasting the risk of tornadoes. They found the following:
Helicity 150-299 ... weak tornadoes (possible 'supercell')
Helicity 300-499 ... strong tornadoes (favourable for 'supercell' development)
Helicity > 450 ... violent tornadoes
( These figures should be used with caution in the UK where helicity will normally lie between -200 and +200 m**2/s**2 )
(See also this FAQ entry)
(with thanks to Rodney Blackall for advice & suggestions with this and the following entry.)
And why do forecasters find it so difficult to get it right?
Whether snow penetrates to the surface as snow, or melts to rain or sleet on the way down depends upon the height of the 0 degC level (ZDL) above local terrain. It should be easy over relatively flat ground: forecast yes/no for precipitation and use a good forecast model (or dense network of boundary-layer radio-sonde ascents) to find the ZDL. If you are above this level, then expect snow, if below expect rain or sleet.
The 'air-mass' zero degree level is relatively straightforward to forecast. The problem is that snow situations in our part of the world often occur with surface temperatures 'around zero', and minor deviations from the air-mass (or synoptic-scale) ZDL are important, but difficult to predict. There are several factors that must be taken into account when assessing these potential variations in the ZDL. Among these are modification of the temperature profile in the lowest layers of the troposphere due to passage over warm or cold surfaces; cooling due to evaporation of the precipitation elements as they fall through the air and cooling due to latent heat exchanges when snow begins to melt in situations that are 'marginal'. These, and other modifying effects, are discussed this FAQ entry, but they will alter, sometimes dramatically, what type of precipitation actually reaches the surface.
In many of the countries of 'maritime' NW Europe, the major conurbations and the principal highways lie below the 200 m (circa 650 ft) contour. Variations in the low-level temperature structure, often involving changes in intensity from 'light' to 'moderate' or heavier precipitation can cause chaos, yet be difficult to predict and protect against except with very vague generalisations within forecasts. They are also difficult for road, rail & airport authorities, as it can be raining quite happily for several hours (when no precautionary measures can be taken), then all of a sudden, several cm of snow will accumulate as the precipitation intensity changes - or perhaps freezing rain is the result with obvious consequences. A ground height change of more than 30 m (around 100 ft) is quite normal within a town, so it is not uncommon for sleet to fall in one part of the town causing few problems, but snow in another spot nearby.
1. SYNOPTIC-SCALE MODIFICATION of the temperature structure of the lower troposphere. If the air passes over the sea (or similarly warm surface), then the sensible flux of heat to the air above will raise the ZDL, perhaps tipping the balance towards rain or sleet, rather than snow - windward coastal plains may miss out on the worst of the snow. ( However, these same areas may be the only places to experience moist convection in winter and provided the air is cold enough, and the sea is close and upwind, then snow showers can be frequent. ) Heat from major urban areas (provided areally extensive) can also tip the balance in highly marginal situations. If the air passes over an ice or snow-covered surface, then a flux of heat from the air to the surface occurs, modifying the ZDL structure, usually resulting in a sharp, shallow inversion. The air-mass (highest altitude) ZDL is unlikely to be affected but a secondary pair of ZDLs may form as the thermal structure of the lowest 300m is distorted and either freezing rain or ice pellets, rather than 'proper' snow is the result. This is often a difficult situation to get right after a long cold period is trying to break down.
2. EVAPORATIVE COOLING of the air through which the snowflakes are falling. Even with the most intense precipitation, there is always lots of air around the falling raindrops or snowflakes and evaporation of the precipitation elements will occur. This will lead to a microscale cooling (due to latent heat exchanges as the liquid/ice evaporates), which multiplied by the huge number of precipitation elements leads to a net cooling of the environment through which the droplets/crystals are falling. This in turn leads to a lowering of the ZDL. The effect is proportional to the precipitation intensity ( and inversely proportional to the mean wind speed through the melting layer ) and is greater when the ambient relative humidity is well under 95%.
3. PHASE-CHANGE COOLING of the air through which the snow is falling. Rain/snow situations are often marginal at low altitudes. This means that more often than not, snow is melting in the lowest 200m or so and thus the environment is cooled due to heat exchanges consequent upon the melting of the ice crystals into liquid water. Again, intensity of precipitation is a major factor - greater intensity means that there are more precipitation elements involved which means greater overall cooling. The effect compounds that at (2) above, the net effect of evaporative and phase-change cooling is quite significant - lowering the ZDL by some hundreds of metres in prolonged precipitation. This is especially pronounced in stable air and catastrophic in near-isothermal conditions in a frontal zone. Some of the worst low-level icing conditions for aircraft occur in these situations, and of course, the ground isn't very far away!
4. BULK (DOWNWARDS) ADVECTION of cold air due to drag by precipitation elements and by downdraughts in a markedly convective environment. Another effect that is related to precipitation intensity is the cold air that is dragged down by the falling elements and the associated downdraughts. Descending air warms adiabatically so this introduction of colder air from upper levels is offset somewhat and is the least effective modulator of those considered above. [ However, in such situations, the relative humidity will fall (greater separation between air and dew point temperature) so evaporative cooling will become more effective - see (2) above. ]
5. OROGRAPHIC UPLIFT COOLING. As air in a thermally stable environment is forced to rise over a range of hills or mountains then the adiabatic cooling will cause the temperature to fall with height more rapidly than in the undisturbed environment. This will lower the ZDL allowing a greater downward penetration of the snow that might otherwise be expected. (This is important in a few of our major towns and cities that rise into the 'foothills' of major hill ranges, e.g. Manchester, Sheffield & Bradford.)
6. FALLING OR SETTLING PROBLEM. Apart from the factors mentioned at the end of (1) above, snow is pretty well guaranteed to fall and settle if the surface temperature is at or below 0 degC. Rain is almost guaranteed if the surface temperature is above 4 degC. In between there is a degree of uncertainty and quite small changes in intensity can switch between sleet and snow, and between snow thawing faster than it falls or vice-versa.
In the 17th century, when the concept of the barometer was first developed and refined by Torricelli & Pascal (amongst others), atmospheric pressure values were noted in terms of the height of the mercury column supported within a tube which had one end closed and the other end immersed in a bath of the liquid exposed to the atmosphere. Barometers continued to be marked in units of length (inches or mm of mercury) long after aneroid barometers became a common instrument - hence even today it is not unusual to see barometers marked in either inches (British / Imperial and US sources) or millimetres (European / Continental sources).
1 millibar (or hectopascal/hPa), is equivalent to 0.02953 inches of mercury (Hg). It is therefore only necessary to multiply a reading in millibars by the latter figure, to achieve the required conversion. E.g. for 1023 mbar, multiply by 0.02953=30.21 inches. To go the other way, the relationship is 1 inch of mercury=33.8639 mbar; again as an example, to convert 29.45 inches, multiply by 33.8639=997 mbar. (For those of you reading this on the continent, you are more likely to be dealing with millimetres, and the appropriate conversions are: 1 mbar=0.750062 mm Hg and 1 mm Hg=1.333224 mbar. )
The dew point depression (often abbreviated to DPD), is the difference between the air temperature and dew-point of a sample. It can refer to surface (i.e. screen) temperature values, or to measurements in the upper air. The larger the value (wider separation between air temperature and dew-point), the lower the relative humidity.
Surface values of DPD (screen or psychrometer measurements) are often used in empirically derived algorithms to forecast overnight minimum temperatures, fog-point temperatures, height of convective cloud bases, likely stratus base due to turbulent-mixing etc. For example, a high afternoon DPD value would suggest low relative humidity for that air-mass (afternoon values usually being assumed to be representative due to good mixing of the boundary layer air), and assuming no air-mass change, the night minima will be relatively low, as would the fog-point temperature. Conversely, if the DPD value was small in the afternoon then, other factors being right (light wind, clear skies etc.), then mist/fog would be a high risk for the coming night - or low cloud if the wind were just a touch stronger. Daytime cumulus bases will be lower with small DPD values, than with a greater separation between air temperature and dew point.
Upper air values are often used to assess likely layered cloud amounts, again other things being equal - i.e. sufficient uplift to lead to condensation and a stable environment. On a thermodynamic diagram, if the 700 hPa DPD is less than about 3 degC, then expect thick, layer cloud to be present. Dew point depression values are also used in the study and forecasting of severe convective storms.
The speed of sound in air is considerably less than the speed of light. It is therefore possible to calculate the distance of a storm, provided you can observe the associated lightning, and the thunderstorm is close enough for the sound (of the thunder generated) to reach you (*).
A 'thunder-clap', or the noise our ears hear (ranging from a sharp crack for nearby, short path-length lightning strikes to a long, low rumble for a distant, long path-length discharge), is caused by the rapid expansion (due to heating) of air in the lightning channel that the stroke passes through. ( If you are frightened of thunder, it is useful to remember that by the time that noise has reached you, the really dangerous lightning strike has already occurred. )
Very roughly, the speed of sound in the lower atmosphere is 330 metres / second (or 1 mile / 5 seconds). If you observe the 'flash' of the lightning stroke, then begin counting (or timing) in seconds until the sound of the thunder just reaches you, using the relationships above will give a rough guide to the distance of the storm. Thus a time difference of 3 seconds will give you three-fifths of a statute mile, or about a kilometre; 6 seconds=> 2 km etc.
((*) thunder, under average atmospheric conditions, should be audible up to 8, possibly 10 km away. Strong winds & low-level temperature variations may though alter these values considerably, with values quoted up to 20 km in some cases; lightning at night can be seen up to 100 km away, depending upon your observing location, obstructions etc.)
[ Also see this FAQ entry ]
In the 'free' atmosphere on our rotating earth, the movement of air is forced by differences in atmospheric pressure between one location and another: this difference, over a specified distance, is known as the PRESSURE GRADIENT.
It might be assumed that once there is a pressure gradient, that air would travel directly from high pressure to low: this doesn't happen, because as soon as it begins to move, it undergoes an apparent deflection owing to the fact that we live on a rotating planet. In the Northern Hemisphere, the 'deflection' is towards the right of air motion; in the Southern Hemisphere it is towards the left. A balance is achieved whereby the force due to the Pressure Gradient (PRESSURE GRADIENT FORCE, or PGF) exactly equals the deflection due to planetary motion (the CORIOLIS DEFLECTION or ACCELERATION [CA]).
The wind direction is that summarised in Buys Ballots Law (q.v.)
The wind resulting from these ideal conditions is known as the GEOSTROPHIC WIND - a theoretical wind (as measured on a chart using a GEOSTROPHIC SCALE) that assumes the following:
(b): the wind is not speeding-up (accelerating) or slowing-down (decelerating) along the line of travel (isobars/contours are parallel).
(c): the pressure pattern is not changing (no atmospheric development).
(d): there are no frictional forces at work, either molecular, or due to passage over tangible obstacles (e.g. the surface of the earth).
(e): the wind blows horizontally - i.e. there is no vertical motion involved.
Looking at these factors in order:
(a): Unless the air motion is gentle, this factor must be allowed for by finding the curvature of the isobars/contours and subtracting (cyclonic curvature) or adding (anticyclonic curvature) a correction depending upon the strength of the geostrophic wind & the radius of curvature of the isobar/contour. For a tightly curved isobaric/contour pattern around a small-scale low pressure area in mid-latitudes, the value will be in the range 10 to 30 knots at geostrophic wind speeds of over 80 knots. Equally important is the anticyclonic correction: even for a gently curved isobaric pattern around a slow-moving anticyclone, an extra 5 to 10 knots can be added to the theoretical calculation - enough for example to tip the balance between a 'Force 7' and a 'Gale'! (Unlike the cyclonic correction, there is a theoretical maximum for the anticyclonic correction of twice Geostrophic value).
If the pattern is curved, then the corrections applied above will give rise to the term known as the GRADIENT WIND, which is loosely taken to be the 'free-air' value (very roughly around 900 mbar/900 m), from which the wind at the surface can be obtained - (see below).
However, even this is problematic, as it assumes that the pressure system is not moving, i.e. the air takes a path exactly governed by the pattern on the chart. In reality, especially with developing depressions, the movement of the low itself will mean that the air-path adopts a differently curved trajectory to that of the feature generating the winds. In particular, for fast-moving depressions in mid-latitudes, the geostrophic (theoretical) wind speed is probably a better approximation to the 'free-air' flow than trying to apply a correction due to curvature.
(b): When air speeds up or slows down due to changing pressure gradients along the line of travel, then motions are imparted which cause the air to deviate from the ideal path. This flows naturally from the fact that the ideal wind is trying all the time to achieve a balance between the PGF and the Coriolis Acceleration (CA). As the latter is proportional to the wind speed, any change in same will alter the deflection; the 'flow' will become unbalanced, and a 'correction' needs to be achieved, but not before the air has deviated somewhat offline (so-called 'CROSS-CONTOUR FLOW'). The effects are very important at jet-stream levels, where they lead to cross-contour movements of sufficient magnitude to cause development in the column of air below. (e.g. Jet exits, entrances), and also give rise to the potential for clear air turbulence (CAT).
(c): This is important near the surface where there are tight gradients of pressure tendency, i.e. as found in advance of, and particularly to the rear of some rapidly developing cyclonic disturbances in mid-latitudes. These ISALLOBARIC effects can be to add (rapidly rising pressure), or subtract (rapidly falling pressure) well over 10 knots to the theoretical geostrophic wind, and can be as much as 40 knots in the most 'damaging' case, where a explosively deepening low has passed by, and pressure rises sharply behind due to events in the upper air (confluent trough, q.v.).
Couple this effect with the possible downward penetration of momentum from the upper atmosphere associated with the dry intrusion (q.v.), and some highly damaging wind-gusts can result.
(d): Of all the 'controlling' additional forces, this one is the most evident from day-to-day. It is the reason why the wind that we observe blows across the isobars at an angle (from high to low pressure), rather than directly along the isobars. As friction due to the roughness of the earth's surface takes hold, the PGF gains dominance over the CA, and 'turns' the wind away from the beloved 'tramlines' of the tv weather presenters. The greater the roughness (i.e. towns, cities) the greater the effect; for this reason, wind directions over the open sea are more closely allied to the isobaric direction averaging just 8 degrees departure, as opposed to some 20 degrees or more over land.
Additionally, overland the stability of the air must be considered. Highly stable air will damp vertical exchange (or mixing) between the friction-less/'free' air (roughly above 900 mbar/900 metres) and the air in direct contact with the earth's surface (within the atmospheric boundary layer). In such conditions, the surface wind can be backed from the isobaric flow by as much as 40 degrees (perhaps more), and this effect is particularly enhanced at night in conditions of light gradient. Conversely, in unstable conditions, where vertical mixing is particularly effective, then the 'real' wind direction is sometimes only 10 or 15 degrees away from that taken from the isobars.