SPECIAL WEATHER PHENOMENA AND EFFECTS ON AIRCRAFT OPERATIONS OVER INDIAN REGION

Chapter Objectives

After reading this chapter, you should be able to:-

Understand the conditions favourable for CB formation and various

weather hazards associated with thunderstorm.

Understand the conditions favourable for duststorm and various weather

hazards associated with it.

Understand the conditions favourable for CAT and mountain waves and

various weather hazards associated with them.

Appreciate the differences between different types of icing and the

favourable met conditions for formation of the same.

Understand importance of mintra level in military flying.

Appreciate the classification of between hydrometeors and lithometeors.

Understand different types of jet streams seen over India

Understand the effect of wind shear on aircraft

Understand the various techniques in detecting microburst

Structure

1. Introduction

2. Thunderstorm

3. Dust Storm

4. Cloud Burst

5. Atmospheric Obscurity

6. Jet streams over India

7. Low level Wind Shear and Aircraft Operations

8. Microburst and its Detection Techniques

Introduction

1. Meteorology is the science that deals with the atmosphere and weather processes that occur in it. The physical state of the atmosphere at a given place and time is weather, and is described in terms of instantaneous values or short-period mean values of meteorological elements (weather elements). Weather plays an

important role in all facets of life and is a vital factor in civil as well as military aviation. Meteorological phenomena and weather processes occurring in the atmosphere are of an infinite variety. Knowledge on special weather phenomena like thunderstorm, dust storm, vertical wind shear and microburst that affect the aircraft operations helps the meteorologists to forecast present and forecast conditions at and around an operating area and enroute.

Thunderstorms

2. Cumulonimbus is a cumulus cloud, which develops upto great heights due to instability and a high degree of humidity in a deep layer of air. The release of energy due to the over-turning of air in the unstable layers gives rise to a storm. The electrical charges developed in the cloud give rise to lightning and thunder. Thunderstorms are one or more convective cells in which electrical discharges are seen as lightning or heard as thunder.

3. The thunderstorm is by far the most dangerous weather phenomenon from the point of view of aviation. A thunderstorm, resulting from vigorous convective activity, is one of the most spectacular weather phenomena in the atmosphere. As electrical charges are separated in a convective cloud the potential gradient between various regions of the cloud increases and eventually exceeds that which the air can sustain. The resulting dielectric breakdown assumes the form of a lightning flash. The temperature of conducting channels, through which the electrical discharges occur, rises to above 30,000oK in such time that the air has no time to expand. Therefore, the pressure in the channel increases almost instantaneously to 10, or perhaps 100 atmosphere. The high-pressure channel then expands rapidly into the surrounding air and creates a very powerful shockwave which travels faster than the speed of sound and, further out, a sound wave which is heard as thunder. Over dry areas with loose soil, thunderstorm wind raise considerable amount of dust/sand before the rain reaches the ground. Such phenomena are known as dust storm / sandstorm.

Conditions Favourable for Cumulonimbus Cloud Formation

4. The necessary conditions for the formation of Cumulonimbus (Cb) clouds are as follows:-

(a) Lapse rate steeper than the SALR throughout a layer at least 5 to 6 kilometres in depth, permitting development of clouds to heights at which the temperature is below 0ºC.(b) Adequate supply of moisture from below.(c) A process which produces saturation in the region of the steep lapse rate or a triggering mechanism.

5. As the instability increases cloud grows upwards, some of the surrounding unsaturated air is entrained into the cloudmass. Consequently, some of the cloud droplets evaporate. If the humidity of the surrounding air is very low, the evaporation becomes dominant and arrests further growth of the cloud. Well-developed Cbs are thus possible only when the humidity aloft is sufficiently high.

6. The triggering mechanism, which sets off a thunderstorm, may be:-

(a) Local convection (insolation).(b) Orographic lifting.(c) Convergence.(d) Frontal lifting.(e) Radiational or katabatic cooling.

7. Cellular Structure. A Cb or thundercloud is composed of several cells, each of which behaves as a unit of convective circulation and goes through its own life cycle of short duration, more or less independently of adjacent cells. The diameters of individual cells vary from 2 to 10 kilometres, while between neighbouring cells there are cloud-filled lanes upto 1 to 2 kilometres in width.

Life Cycle of a Cell

8. Extensive studies indicate that thunderstorms go through a cycle of development from birth to maturity and to decay. The first stage is known as the cumulus stage. As humid air rises, it cools and condenses into a single cumulus cloud or a cluster of clouds. At first the cumulus clouds grow upward only a short distance, and then they dissipate. This is because the cloud droplets evaporate as the drier air surrounding the cloud mixes with it. However, after the water drops evaporate, the air is moister than before. So, the rising air is now able to condense at successively higher levels, and the cumulus cloud grows taller, often appearing as a rising dome or tower. As the cloud builds, the transformation of water vapour into liquid or solid cloud particles releases large quantities of latent heat. This keeps the air inside the cloud warmer than the air surrounding it. The cloud continues to grow in the unstable air as long as it is constantly fed by rising air from below. In this manner, a cumulus cloud may show extensive vertical development in just a few minutes. During the cumulus stage, there is insufficient time for precipitation to form, and the updrafts keep water droplets and ice crystals suspended within the cloud. Also, there is no lightning or thunder during this stage.

9. As the cloud builds well above the freezing level, the cloud particles grow larger. They also become heavier. Eventually, the rising air is no longer able to keep them suspended, and they begin to fall. While this phenomenon is taking place, drier air from around the cloud may be drawn into it in a process called entrainment. The entrainment of drier air causes some of the raindrops to evaporate, which chills the air. The air, now being colder and heavier than the air around it, begins to descend as a downdraft. The downdraft may be enhanced as falling precipitation drags some of the air along with it.

10. The appearance of the downdraft marks the beginning of the mature thunderstorm. The downdraft and updraft within the mature thunderstorm constitute a cell. In most storms, there are several cells, each of which may last for an hour or so.

11. During its mature stage, the thunderstorm is most intense. The top of the cloud, having reached the stable stratosphere, begins to take on the familiar anvil shape, as strong upper-level winds spread the cloud’s ice crystals horizontally. The

cloud itself may extend upward to an altitude of over 12 kilometres (40,000 feet) and be several kilometres in diameter near its base. Updrafts and downdrafts reach their greatest strength in the middle of the cloud, creating severe turbulence. In some storms, the updrafts may intrude into the stable stratosphere, a condition known as overshooting. Lightning and thunder are also present in the mature stage. Heavy rain (and occasionally small hail) falls from the cloud. The rainfall may or may not reach the surface, depending on the relative humidity beneath the storm.

12. After the storm enters the mature stage, it begins to dissipate in about 15 to 30 minutes. The dissipating stage occurs when the updrafts weaken and downdrafts tend to dominate throughout much of the cloud. Deprived of its rich supply of warm, humid air, cloud droplets no longer form. Light precipitation now falls from the cloud, accompanied by only weak downdrafts. As the storm dies, the lower-level cloud particles evaporate rapidly, sometimes leaving only the cirrus anvil as the reminder of the once mighty presence. A single ordinary thunderstorm may go through its three stages in an hour or less. The reason it does not last very long is that the storm’s downdraft may cut off the storm’s fuel supply by destroying the humid updrafts.

13. Generally, not only do thunderstorms produce summer rainfall but they also bring with them momentary cooling after an oppressively hot day. The cooling comes during the mature stage, as the downdraft reaches the surface in the form of a blast of welcome relief. Sometimes, the air temperature may lower as much as 10oC in just a few minutes. Unfortunately, the cooling effect is short-lived, as the downdraft diminishes or the thunderstorm moves on. In fact, after the storm has ended, the air temperature usually rises; and as the moisture from the rainfall evaporates into the air, the humidity increases, sometimes to a level where it actually feels more oppressive after the storm than it did before.

14. Upon reaching the surface, the cold downdraft may force warm and moist surface air upward. This rising air then condenses and gradually builds into a new thunderstorm. Thus, it is entirely possible for a series of thunderstorms to grow in a line, one next to the other, each in a different stage of development. Thunderstorms that form in this manner are termed multicell storms. Most ordinary thunderstorms are multicell storms, as are most severe thunderstorms.

15. For a thunderstorm to develop there must be raising moist air in a conditionally unstable atmosphere. The ingredient necessary to start the air rising may be the unequal heating of the surface, a frontal boundary, a mountain range, or the leading edge of a sea breeze. Most of the thunderstorms that form in this manner are not severe, and their life cycle usually follows the pattern described for ordinary (airmass) thunderstorms.

16. Thunderstorm Structure. The cellular structure and life cycle described above envisages three stages of evolution. The life cycle of an individual cell lasts about an hour. However, observations show that in severe storms, the mature stage may last much more than an hour. The cell at the mature stage gets continually revived, developing into a super-cell with updraught and downdraught co-existing in a more or less steady state for periods lasting over an hour. There is strong vertical wind shear in such storms. Energy is provided by the absorption of latent heat from

the partial evaporation of the precipitation within the dense, cool dry high-level air which sinks as the downdraught and undercuts the warm, moist updraught at the rear of the storm. The updraught extends some kilometres beyond the tropopause and gives rise to unusually high-domed tops. Steady state model is shown in Fig. 4.

17. Electric Field. In clear air of normal density a critical electrical field of about 3,000,000 volts per metre is required for breaking down the insulation and for a spark to pass. In cloud it is reduced to about 1,000,000 volts per metre. Such intense electrical fields are built up locally within a Cb cloud. The upper portions of a Cb cloud acquire a positive charge and the lower portions a negative charge. The exact cause of the generation of these charges is not yet fully understood. However, the electrical field is sufficiently strong to permit a lightning discharge. Discharges may take place within the same cloud, from one cloud to another and less frequently from cloud to the earth. The lightning discharge is in the form of stepped strokes, 1 to 4 km long with a channel diameter of 1 to 10 m. The air through which the discharge has passed is rendered white hot and expands suddenly, giving the clap of thunder. If the lightning stroke is long, the thunder may be in the form of peals as the sound from different parts of it takes longer time for travel and reach the ground.

18. Movement of Thunderstorms. Although large thunderstorms modify the airflow in their immediate neighbourhood, most thunderstorms drift slowly with the wind prevailing in the layers in which they are embedded. As a first approximation, the direction of movement may be deduced from the prevailing winds at 3 to 5 kilometres above sea level. The speed of movement is less than the speed of the wind at these levels. When the winds aloft are weak or variable, thunderstorms show little movement. In such cases the shower from the cloud is confined to a limited area, resulting at times in exceptionally heavy falls of short duration, popularly known as cloudburst. If the thunderstorm has appreciable movement in the rain stage, the shower gets distributed over a larger belt.

19. Regeneration of Thunderstorms. The cold downdraught from one thunderstorm can activate a nearby cumulus cloud and cause the latter to build-up rapidly into a cumulonimbus. This process is known as regeneration. Regeneration can happen successively, giving the impression of movement of the same Cb or thunderstorm cloud. In reality it is a chain reaction by means of which thunderstorms occur in quick succession over an extended belt covering some hundreds of kilometres. The violent summer thunderstorms known as Norwesters or Kalbaisakhis which affect Bengal and Bihar owe their origin to regeneration from the primary thunderstorms over the Chotanagpur hills which build-up in the afternoon.

20. Vertical Extent of Cumulonimbus Clouds. From radar surveillance as well as aircraft report it has been found that Cb clouds over the Indian region reach great heights, at times upto the tropical tropopause. Average heights of tops are between 10 and 15 kilometres. Cb clouds have, however, been reported to reach heights of 18-20 kilometres. The base is usually over 1 kilometre above ground, but in heavy showers may lower to less than 300 metres above ground in the monsoon season.

Flying Hazards of Thunderstorms

21. Severe thunderstorms are capable of producing large hail, strong and gusty surface winds, flash floods and tornadoes. Just as the ordinary thunderstorm, they form as moist air is forced to rise into a conditionally unstable atmosphere. But, severe thunderstorms also form in areas with a strong vertical wind shear.

22. Strong winds aloft may cause the updrafts in a severe thunderstorm to tilt in its mature stage. The moving storm and the upper-level winds cause the system to tilt so that the updrafts move up and over the downdrafts. This situation allows the updraft to remain strong for an extended period of time. However, there are many severe thunderstorms that do not have tilted updrafts. In these storms, the updrafts may be so strong (sometimes up to 100 knots) that precipitation–size particles do not have enough time to form. Apparently, this type of thunderstorm becomes severe when a strong vertical wind shear causes the storm to rotate. It is the rotational aspect of thunderstorms that may lead to the formation of tornadoes.

23. The updrafts in a severe thunderstorm may be so strong that the cloud top is able to intrude well into the stable stratosphere. In some cases, the top of the cloud may extend to more than 18 kilometres (60,000 feet) above the surface. The violent updrafts keep hailstones suspended in the cloud long enough for them to grow to considerable size. Once they are large enough, they either fall out the bottom of the cloud with the downdaft or a strong updraft may toss them out side of the cloud, or even from the base of the anvil. Aircraft have actually encountered hail in clear air several kilometres from a storm. Also, downdrafts within the anvil may produce beautiful mammatus clouds.

24. As some of the falling precipitation evaporates, it cools the air and enhances the downdraft. The cool air that reaches the ground may act like a wedge, forcing warm, moist surface air up into the system. Thus, the downdraft may help to maintain the updraft and vice-versa, so that a severe thunderstorm with this type of updraft and downdraft configuration is able to maintain itself (for many hours, in some cases).

25. The downdraft spreads laterally after striking the ground. The boundary separating this cold downdraft from the warm surface air is known as a gust front. To an observer on the ground, the passage of the gust front resembles that of a cold front. During its passage, the wind shifts and becomes strong and gusty, with speeds occasionally exceeding 55 knots; temperatures drop sharply and, in the cold heavy air of the downdraft, the surface pressure rises. Sometimes it may jump several millibars, producing a small area of high pressure called a mesohigh (meaning mesoscale high). The cold air may linger close to the ground for several hours, well after the thunderstorm activity has ceased.

26. Along the leading edge of the gust front, the air is quite turbulent. Here, strong winds can pick up loose dust and soil and lift them into a huge tumbling cloud. As warm moist air rises along the forward edge of the gustfront, a shelf cloud (also called an arcus cloud) may form. These clouds are especially prevalent when the air is very stable near the base of the thunderstorm. Occasionally, an elongated ominous looking cloud forms just behind the gustfront. These clouds, which appear

to slowly spin about a horizontal axis, are called roll clouds. Sometimes the leading edge of the gustfront forces warm moist air upward, producing new thunderstorms.

27. Beneath a severe thunderstorm, the downdraft may become localised so that it hits the ground and spreads horizontally in a radial burst of wind, much like water pouring from a tap and striking the sink below. Such downdrafts are called downbursts. A downburst with winds extending only upto 4 kilometres or less is termed a microburst. Inspite of its small size, an intense microburst can induce damaging winds as high as 146 knots. A larger downburst with winds extending more than 4 kilometres is termed a macroburst. Since a microburst is an intense downdraft, its leading edge can evolve into a gustfront.

28. Microbursts are capable of blowing down trees and inflicting heavy damage upon poorly built structures. In fact, microbursts may be responsible for some damage once attributed to tornadoes. Moreover, microbursts and their accompanying wind shear appear to be responsible for several airline crashes. When an aircraft flies through a microburst, it first encounters a headwind that generates extra lift. However, in a matter of seconds, the headwind is replaced by a tailwind that causes a sudden loss of lift and a subsequent decrease in the performance of the aircraft.

29. Microbursts can be associated with severe thunderstorms, producing strong, damaging winds. But studies show that they can also occur with clouds that produce only isolated showers – clouds that may or may not contain thunder and lightning.

Supercell and Squall Line Thunderstorms

30. The supercell storm is an enormous rotating thunderstorm whose updrafts and downdrafts are sufficiently structured so that it is able to maintain itself as a single entity for hours till end. Storms of this type are capable of producing updrafts that can exceed 90 knots, hail of the size of grapefruit, damaging surface winds, and large, long-lasting tornadoes.

31. The squall line forms as a line of thunderstorms. Sometimes they are right along a cold front but often they form in the warm air 100 to 300 kilometres out ahead of it. These prefrontal squall line thunderstorms of the middle latitudes represent the largest and most severe type of squall line. The line of storms may extend for over 1000 kilometres (600 miles), with huge supercell storms causing severe weather over much of its length.

32. There is still debate as to exactly how prefrontal squall lines form. Models that simulate their formation suggest that initially convection begins along the cold front then reforms further away. Moreover, the surging nature of the main cold front itself, or developing cumulus clouds along the front, may cause the air aloft to develop into waves that are much like the waves that form downwind of a mountain chain. Out ahead of the cold front, the rising motion of the wave may be the trigger that initiates the development of cumulus clouds and a prefrontal squall line.

Mesoscale Convective Complexes

33. Where conditions are favourable for convection, a number of individual thunderstorms will occasionally grow in size and organise into a large convective weather system. These convectively driven systems, called Mesoscale Convective Complexes (MCCs), are quite large – they can be as much as 1,000 times larger than an ordinary thunderstorm. In fact, they are often large enough to cover an entire state, an area in excess of 1,00,000 square kilometres.

34. Within the MCC, the individual thunderstorms apparently work together to generate a long lasting weather system that moves slowly (normally less than 20 knots) and often exist for periods exceeding 12 hours. The circulation of the MCC supports the growth of new thunderstorms as well as a region of widespread precipitation.

Tropical Thunderstorms

35. So far, discussions related to the well-researched temperate latitude convective storms. In contrast limited observational data are available for the convective storms of tropical regions. Observational studies indicate that these storms tend to display more signs of organisation, often occurring in lines (squall lines). These are very similar in dimensions to the organised bands of thunderstorms that occur ahead of cold fronts in temperate latitudes. There are, however, significant differences in internal structure though strong updrafts and downdrafts are still an important characteristic. In particular the anvil streams off behind the storm rather than in front as it does in temperate-latitude storms. The anvil in tropical squall line is thicker and longer than the anvil of mid-latitude squall lines. There is sustained upward motion within the anvil providing continued source of water substance in the anvil, which requires a mesoscale circulation in the rear of a tropical squall line. In tropical squall lines, there are two distinct types of downdrafts– an intense cloud scale downdraft, 10 – 20 km wide in the narrow zone of heavy precipitation of the squall line itself, and a more gentle but broader (100–500 km wide) downdraft which forms in the precipitation region below the anvil. The low-level air entering a tropical squall line is conveniently more unstable than that in the mid- latitude squall lines.

Aviation Hazards Associated with Thunderstorms

36. The carriage of weather radar on the flight deck has enabled pilots to have early recognition of thunderstorms near the flight path. It is prudent, and always appreciated by the passengers, to avoid thunderstorms rather than to fly in or close to them if it is operationally feasible to do so. All thunderstorms are potentially dangerous and their external appearance is no guide to the severity of the hazards that may be expected. The hazards, not in any order of priority, may be summarised as follows:-

(a) Base Height. Many large cumuliform clouds have a base height of some 1,500 feet or more above ground level so there is space for light aircraft to operate beneath them if necessary. Once thundery precipitation starts, the cloud base lowers and terrain clearance beneath the clouds

becomes difficult to maintain. Inside the thunderstorm the visibility may be down to 20 meters.

(b) Hail. Although the period involved in transiting the hail area of a thunderstorm may be relatively short, damage from large hailstones can be considerable especially to radomes, transparencies, upper and leading surfaces (including de-icer equipment). The greater the aircraft’s speed, the greater the damage and hail can be met anywhere in or under the storm– even under an overhanging anvil.

(c) Icing. Thunderstorm clouds give icing problems both in the engine(s) and on the airframe. In a piston engine even at temperatures above 0oC there can be a serious loss of power while a turbine engine may suffer a flameout. On the airframe, the large supercooled water drops freeze to give a rapid build-up of clear ice with its attendant problems of increased weight, disturbed relative airflow decreasing lift and increasing drag, risk of control surfaces becoming less effective, etc.

(d) Instrument Errors. Turbulent airflow around the aircraft and the localised variations in pressure can produce rapid and serious errors in the readings of the altimeter, airspeed indicator and vertical speed indicator. If the pressure head suffers serious rain ingress or, worse still, ices up, the pressure instrument readings are useless. Remember too that a lightning strike on the aircraft will probably rearrange the aircraft’s magnetism so the compass heading will be suspect.

(e) Lightning. This can be used, especially at night, to supplement the aircraft’s weather radar for pinpointing the most active storm areas. Because of metal-to metal bonding and screening of vulnerable equipment, an aircraft struck by lightning is unlikely to suffer more than a scorch mark on the aircraft’s skin. Nevertheless the sudden brilliance of the flash, the noise, and sometimes the burning smell, can be distracting to say the least.

(f) Squalls, Windshear and Microbursts. There are three possible circumstances which present problems to aircraft taking-off or landing: (i) effect of decreasing headwind when on the approach, (ii) effect of gust front ahead of a thunderstorm of an aircraft on the approach, (iii) change in flight path due to microburst. Sometimes the normal variation of wind speed with height becomes greatly accentuated, perhaps decreasing from 40 knots or more at 1,000–2,000 feet to less than 10 knots near the surface. If the pilot does not intervene, the actual flight path will divert progressively from that intended. A varying crosswind component complicates the matter. The gust front may be 15 to 20 miles ahead of the storm which generated it. It may be marked by newly forming cumuliform cloud or by a line of duststorms in desert countries, but equally it may be quite invisible. Microbursts are particularly intense and localised– probably not more than half a mile across– and especially hazardous when they are overhead or near the runway in use. Again, if on the approach in thunderstorm conditions it is found that abnormal levels of power are necessary to maintain the airspeed, attitude and glideslope.

(g) Precipitation. Forward visibility is always seriously reduced in heavy precipitation. For departing and arriving aircraft there are the hazards of wet runways to contend with.

(h) Static. This is often more of a nuisance than a hazard, as it builds up noise on the radio and can affect communications. It can be an infuriating distraction just when it is least wanted.

(j) Turbulence. It makes the aircraft difficult to handle and indeed, incorrect handling can lead to loss of control. If the aircraft is flown at a higher speed than that specified as its ‘rough air’ or ‘thunderstorm penetration speed’ in the flight manual then there is a risk of structural failure. On the other hand, speeds lower than that recommended may lead to the aircraft stalling.

37. Because of the handling problems in rough air, pilots should make sure before entering the turbulence that the ‘Fasten seat belts’ sign is switched on, the aircraft positioned at the correct safe altitude, power set and the aircraft trimmed to fly at the recommended penetration speed and at the correct attitude. The temptation to chase the small-scale fluctuations of the instrument readings must be resisted and instead, pilots should observe the basic instrument flying principles of maintaining power and attitude.

Duststorm

38. When the instability conditions occur over desert and semi-arid areas and the humidity conditions aloft are not very favourable convective clouds do not build up to great heights; but such clouds can still give rise to storms if their tops extend above freezing level. These storms are responsible for raising loose dust or sand from the ground upto heights of 10,000 ft or so. They are hence known as “duststorms” or “sandstorms”. In northern India summer duststorms or this type are known by the local name of “andhi” (blinding storms). In the Sudan they are known as “haboobs”.

39. The mechanism for the formation of a duststorm is essentially the same as in a thunderstorm. Due to low humidities aloft the vertical growth of cloud is arrested. The down-draught is initiated quickly by the fall of supercooled water drops from a level a few thousand feet above the freezing level. The water drops evaporate completely before reaching the ground due to the high temperatures prevailing. The up-draughts in the cloud are quite vigorous and carry the dust or sand through a large part of the cloud.

40. When the humidities aloft are reasonably high the Cb clouds may grow to higher levels. In such cases the storm may initially start as a duststorm but subsequently thunder and light shower may occur. This usually happens a little before the rainy season sets in.

41. From radar observations it has been found that the tops of the Cb cloud which cause pure duststorms are of the order of 25-30,000ft. When the tops as observed on radar exceed 35,000ft. they culminate in thunderstorms although, initially, due to the dry nature of the ground dust may be raised by the leading edge of the squall.

42. Duststorms can be picked up by weather radar. However, since the storms commence even when the cloud has built up to about 25,000ft only the time interval between the initial cumulus stage and the mature storm stage is smaller than in the case of thunderstorms. For providing adequate warning the radar set has to be kept in almost continuous operation and a close watch kept on the formation of cells.

43. Movement of Duststorm. Duststorms move with the prevailing low level winds upto about 10,000ft. Their life cycle is shorter than in the case of thunderstorms. Regeneration of incipient convective cells can also occur as in the case of thunderstorms.

44. Duststorm Squalls. Surface squalls associated with duststorms can at times be severe. They are however, usually 30-40 kts in strength. Duststorms which later develop into thunderstorms give stronger squalls than pure duststorms.

45. Diurnal and Seasonal Variation. Duststorms are most common in the afternoon in the summer months over the desert and semi-arid areas of northwest India. The intense heating and the steep lapse rate favour their formation. Usually they occur only if there is concentrated fall of pressure in the region due to the movement of a low pressure area. Nearer the hills mild duststorms may occur at night and the early hours of the morning.

46. Distinction between Duststorms and other Disturbances. Duststorms are instability phenomena and belong to the family of thunderstorms. They occur over limited areas and affect any location for a short period of time. They should be distinguished from “dust-devils” and “dust-raising winds”. The former are small whirls of dust moving at random. The latter are persistent strong winds (due to a steep pressure gradient) which raise and carry loose dust over large areas and last, at times, for many days at a stretch.

47. Flying Hazards in Duststorm. Flying through duststorms involves almost the same hazards as in thunderstorms. In addition the ground visibility is reduced to very low values. If no rain-drops have fallen the ground visibility does not improve substantially until a long time after the duststorm has moved away.

Cloudburst

48. A cloudburst is sudden copious rainfall. It is a sudden aggressive rainstorm falling for a short period of time limited to a small geographical area. Meteorologists say the rain from a cloudburst is usually of the shower type with a fall rate equal to or greater than 100 mm (4.94 inches) per hour. Generally cloudbursts are associated with thunderstorms. The air currents rushing upwards in a rainstorm hold up a large amount of water. If these currents suddenly cease, the entire amount of water descends on to a small area with catastrophic force all of a sudden and causes mass destruction. This is due to a rapid condensation of the clouds. They occur most often in desert and mountainous regions, and in interior regions of continental landmasses.

49. During a cloudburst, more than 2 cm of rain may fall in a few minutes. They are called 'bursts' probably because it was believed earlier that clouds were solid masses full of water. So, these violent storms were attributed to their bursting. During a cloudburst, more than 20 mm of rain may fall in a few minutes. The results of cloudbursts can be disastrous. Cloudbursts are also responsible for flash flood creation. Rapid precipitation from cumulonimbus clouds is possible due to the Langmuir precipitation process in which large droplets can grow rapidly by coagulating with smaller droplets which fall down slowly. It is not essential that cloudbursts occur only when a cloud clashes with a solid body like a mountain. They can also occur when hot water vapour mingles into the cold resulting in sudden condensation.

Jet Streams

Clear Air Turbulence

50 The turbulence could occur at low levels as will as high levels. The essential difference between low level and high level turbulence is related to the dissipative force and the production and growth of turbulence within a limited time scale. The dissipative force is quite strong in the low level and a steady state can be achieved for considerable period. Thus the character of turbulence can be described with parameters such as mixing length and diffusion coefficient. In the upper air, however, a parcel of air moves through a dynamic pattern quite rapidly and thus the time scale for growth is quite small. The basic mechanism of CAT at high level is not unlike the turbulence generated by friction near ground level. For the generation of turbulence of this type marked changes of wind speed along vertical and fairly steep lapse rate are two essential conditions.

51 Air flow over mountainous terrain is more disturbed than over the plains. The flow pattern is complex. An understanding of the characteristics of the flow pattern and the extent, scale and degree of turbulence associated with it, is vital for the safety of aircraft flying over the terrain. When an air stream with suitable stability conditions and wind profile strikes a mountain range, the stream over and on the lee of the range forms a train of waves extending vertically to a great height and horizontally downwind to a considerable distance. The flow pattern is extremely complicated and is dependent on the shape, length and height of the range on the one hand and the lapse rate and wind gradient of the air stream, on the other.

Classification of Turbulence

52. Turbulence according to the causes can be classified as:

(a) Thermal or Convective. This type is caused by local vertical currents induced by insolation and unstable lapse rates or due to movement of cold air mass over warm surface at low levels and at high levels due to convection in a cirrus shield, due to the release of static or conditional instability in the high troposphere.

(b) Mechanical. This type is caused by strong winds over rough and uneven terrain.

(c) Frontal. This type is caused by lifting at cold fronts or squall lines.

(d) Wind Shear. This type is caused by strong vertical and horizontal wind shears. It is related to Reynold's stresses, Richardson’s number, sloping baroclinic zones, jet streams, vertical motion and its horizontal derivatives and temperature changes.

(e) Mountain wave. This type is caused by the breakup of vertical oscillations induced by strong flow across a ridge of mountains, during the interaction of stable baroclinic zones. Three factors static stability, wind flow across the ridge and slope of the mountain ridge, are important.

(f) Convective Cloud. This type is caused by thermal instability and eddies produced by the large shears between updraughts and downdraughts.

Clear Air Turbulence (CAT)

53. CAT can be defined as an imbalance in the atmosphere in which eddies and gravity waves are formed. In general the term CAT is used to denote turbulence at high level (3 km and above) outside Cb clouds. CAT can occur in cirrus clouds, haze layers or clear air.

54. Typical dimensions of areas of CAT are 30km long and 5 to 15 km across, between heights 4 km and 5 km with a thickness of the order of 300m to 500m, lasting from 30 minutes to three hours. However, U2 flights have noted CAT at 20km and X15s at still greater heights. Individual turbulent zones or patches often occur in close proximity to one another, both vertically and horizontally.

Types of CAT

55. CAT can be divided into three separate categories. The first is the most common type, namely, Wind Shear Turbulence. This is associated with strong vertical and horizontal wind shears. The second class of CAT is associated with mountain wave activity. The breakdown of the waves produces the greatest amount of turbulence of all types. Some measurements suggest that gust velocities in mountain waves exceed the design criteria of heavy multi-jet aircraft. The components of mountain wave CAT consist of three variables – two of them being atmospheric viz. wind and static stability and one orographic, viz. the slope of the ridge. The third type of CAT is convective. CAT is associated with clear or dry convective cells and also convection in a cirrus shield. This is caused by the release of static or conditional instability in the high troposphere. There is yet another type of CAT, referred to as trough turbulence. However, this can be considered as either shear or convective turbulence, depending on its position relative to the jet core since it is connected with the atmospheric variable associated with the trough.

Energetics

56. The main sources of turbulent energy in the high troposphere and the low stratosphere are:-

(a) Vertical Shear. This factor makes turbulent energy available through Reynold's stresses. This can produce turbulence at a rapid rate under low Richardson Numbers (low static stability).

(b) Horizontal Wind Gradient. Horizontal wind shear and diuence can create turbulence through Reynold's stresses. This by itself is not significant and whatever correlation has been found is believed to be due to a combination with other factors.

(c) Static Stability. Static instability by itself is probably an insignificant factor in the upper atmosphere, since in most cases light turbulence would dissipate the energy as quickly as it is generated. Static stability since it constitutes a strong turbulence sink during the growth of turbulence, however, is of considerable importance.

(d) Gravity waves. These waves can be strong enough to cause severe turbulence over mountainous regions. Again, this factor by itself may not normally constitute a source of strong turbulence though it contributes significantly in increasing the probability and perhaps the intensity of CAT (Panofsky, 1968). Gravity wave in combination with vertical shear, however, produces severe turbulence. Gravity waves themselves provide a significant source of energy and the inclination of streamlines associated with gravity waves combined with strong vertical shear permits violent turbulence to be attained due to local Reynold's stresses.

(e) Dynamic Instability. Strong anticyclonic shear and curvature are favourable for dynamic instability. Any region of decreasing negative relative vorticity will thus from an area susceptible to turbulence, particularly if associated with low Richardson Number. These can be located in the temperature or thickness field. Dynamic instability can act as a source of turbulent energy and can create mesoscale perturbations. These factors with

strong vertical shear will tend to create intense shears locally and effective release of energy. Reduced static stability will allow the energy to appear in the vertical and horizontal whereas static stability will act as a turbulent energy sink, thus constituting a very favourable region for occurrence of turbulence.

CAT Analysis

57. The main difficulty in the analysis of CAT zones is the scale of motion involved. Individual turbulent eddies or waves have a microscale structure. The size of eddies which cause aircraft bumpiness varies from 300 to 1000 metres, depending on the speed of the aircraft. The distance between radiosonde observations is extremely wide, of the order of hundreds of miles. The data based on these observations do not reveal the microstructure of the waves. Thus the structure of eddies and waves are missed on the chart.

58. Important CAT features are:-

(a) CAT is a predominantly a patchy phenomenon and its intensity is generally less than the turbulence encountered in thunderstorms.

(b) High level turbulence is frequently anisotropic with stronger horizontal gusts. The bump frequency is greater at higher levels than at lower levels.

(c) There is a strong association between jet stream and CAT.

(d) Over land, CAT occurs mostly on the cyclonic side of the jet stream, while over water, CAT is mostly encountered on the anticyclonic side.

(e) CAT is significantly more at the tropopause.

(f) Terrain has an additive influence on CAT.

Forecasting of CAT

59. The synoptic features generally associated with CAT are curved, meandering and strong jet streams and sloping stable baroclinic zone. Wind shear turbulence occurs in regions of cyclonic wind shear. According to Banon (1961) turbulence is more severe near the level of the maximum winds in the jet stream and decreases sharply at higher levels. According to Rieter (1963), CAT occurs in and near the regions of strong horizontal temperature gradients at 250hPa (which indicate the presence of a baroclinic stable zone). Cold air advection enhances the probability of CAT, as in the case of the stratospheric warm tongue and confluent regions of two jet streams near an upper trough. The other feature strongly associated with CAT is the isentropic trough north of a strong jet stream. According to Endlich (1963) the turbulence in a sharp trough associated with a stable layer with strong shear, cold advection, and convergence of the normal component of the wind. In the case of a ridge, turbulence occurs in the thermally stable layer associated with the tropopause and above and to the left of the jet core. The stable layer is associated with strong warm advection and convergence of the normal component of the wind. He suggests the following analysis scheme for CAT:-

(a) Identify the location of progressive upper-air troughs, ridges and jet stream by utilizing 12 hour changes in wind velocity, height and temperature.

(b) Utilise the lapse rate in the 300 to 200 hPa layer to aid in jet stream analysis. Jet core lies along the southern edge of the largest gradient in lapse rate.

(c) Identify the tropopause, upper stable layers, layers of strong vertical shear and advection.

60. Other synoptic features which are associated with turbulence are Jet fingers deflected from the main Jet core and easterly portion of the jet maxima of middle and high latitudes.

61. Horizontal and Vertical Shear. For shear turbulence, Harrison (1959) considers four locations in respect of the jet stream:

(a) Left of jet stream

(b) Right of jet stream

(c) Between two jets

(d) Sub-tropical jet stream

62. He suggests the use of a critical value of horizontal shear 40 kt in 150 N miles as a parameter. The probability of occurrence of turbulence in the four types as related to synoptic trough and ridges is shown in the figure above. In cases of type (b) and (c) vertical shear of 6 kt/300 m may be substituted in the case of horizontal wind shear. In the case of (c) confluence seems to be an important factor. CAT occurrences also seem to be associated with large gradient of shear.

63. George's Method. George (1960) has developed the following method for CAT prognosis: -

(a) Obtain all data where vertical shear values exceed 5 kt/300 m and plot them.

(b) Draw vertical wind shear isotachs for intervals of 3 kt/300m beginning with vertical shear value of 6. Hatch the area enclosed by the two innermost isopleths to indicate the centre of the vertical wind shear.

(c) Draw isolines at 1.5 km intervals for the altitude of the base of the vertical shear and note the regions of significant gradients.

(d) Inspect the upper wind/constant pressure charts immediately above the altitude of the base of vertical shear layer to determine if horizontal shear exceeds 40 kt/200 km. Enclose this area with a red line.

(e) Note the areas where vertical and strong horizontal wind shears overlap. Forecast severe CAT in these regions for a depth of 1000 m below to 2000m above the shear layer.

64. Tropopause and Turbulence. Tropopause is a favourable region where gravity waves, in regions of strong vertical shear, may become unstable and cause turbulence. These regions of strong vertical shears may manifest themselves as positive vertical shear as the case of polar tropopause and negative vertical shear in the case of sub-tropical tropopause. CAT also occurs at the tropopause break or where the tropopause slope is maximum.

65. Vertical Velocity and Turbulence. Variation of vertical motion perpendi-cular to the streamline of the jet axis (w/n) has been suggested as a forecasting parameter. Downstream spreading of isotherms constitutes a favourable area which causes a negative w/n.

66. Temperature Change and Turbulence. A daily temperature change chart at 200 hPa provides an easy tool for locating areas of turbulence, since CAT has been associated with temperature change of 3°C to 5°C/24 hour and to the gradient of temperature of over 5°C/3½° latitude.

67. Richardson's Number and Turbulence. CAT occurs in regions of vertical wind shear and large horizontal temperature gradient. There is also considerable observational evidence that CAT is associated with baroclinic layers in which the vertical shear is so strong as to overcome static stability. Reiter believes that CAT in these layers is due to gravity wave, which becomes unstable due to strong vertical shear. The dimensionless Richardson Number which takes into account both the vertical stability of the air and the shear is thus a convenient forecasting parameter. The Richardson Number is defined as:

Ri=

gθ

∂θ∂ z

/(∂V∂ z )2

where z is height, V the wind velocity, θ potential temperature and g gravity. Both perturbation theory and energy considerations lead us to the conclusion that there exists a critical value of the Richardson number below which perturbations can grow and produce turbulence. Taylor found this critical value to be 0.25. Panofsky et al (1968) found this value to be near 0.5. It has, however, been shown by Philips (1967) that a layer with greater than critical Richardson number can be made unstable by a wave with considerable vertical amplitude. In spite of the difficulty of assigning a critical value to the Richardson number by which turbulence is generated, there is no doubt that this can be used qualitatively at least for the separation of areas of turbulence and no turbulence. Colson and Panofsky (1965), have devised an index of clear air turbulence defined by.

I = (∇ V)2 (I - Ri / Ri critical)

Where Ri critical has been taken as 0.75. The severity of turbulence may be classified for different values of the index as below:

I >25 = Severe turbulenceI 0 to 25 = ModerateI 0 to - 25 = LightI >- 25 = Nil

The value on the right hand side can easily be evaluated from available radiosonde on which upper wind data have been plotted. For Calculation of Ri assuming linear variations, it can be written

Ri=g ΔZ

( ΔU2 )+(ΔV 2)Δθθ

=g ΔzK 2

= Δθθ

Where k2 = (ΔU 2 )+ (ΔV 2)

ΔZ ,Δθ and θ can be obtained from T-ϕ gram; to obtain k a polar diagram is used. Once these values are obtained Ri can be obtained from two other nomograms. If constant value of 9.4 is assumed for ΔZ /θ , Ri can be determined directly from Table 2. For obtaining Ri for layers between the jet maximum wind level and a standard isobaric level a correction is to be applied. This value of Ri assuming θ to be 3400 A and Z = 3250m can be written as Ri = (Ris x Δz 1)/3250, where Ris is the value given in Table 2. Table 1 and 3 gives the Ri-values for different values of ΔZ + Ris. Δz 1 is the thickness of the layer between the level of maximum wind and 300/200 hPa.

Table 1. Ri-values

Table 2.

Table 3.

68. CAT Index. Another Index relates the Laplacian of the horizontal wind shear and vertical derivative of the normal wind with the intensity of turbulence. The relationship is

I=∇ 2(Horizontal wind shear)

- 10∂∂ z (thermal wind)

Large negative values of this index are associated with severe CAT.69. Surface Feature and CAT. CAT frequently occurs above the north and northwestern section of a developing low. The turbulence in such cases is primarily due to a geostrophic field in this region of the atmosphere.

70. Turbulence and Satellite. Satellite pictures help us to identify jet stream location and configuration. An inference regarding the height of the tropopause level is also possible from such data - thus it is possible to utilise such data for delineating areas of possible turbulence. It is also possible to identify mountain waves and areas of turbulence from the satellite cloud pictures, particularly when lenticular clouds are observed Lee waves can also be observed as parallel cloud.

71. Forecasting Aids. Some useful tools for forecasting CAT are:-

(a) The 300 hPa pressure level analysis and forecast wind field. These are also used to determine the current and forecast position of jet stream core.

(b) The forecast tropopause height and smoothed vertical wind shear fields. These are used to identify the regions where the tropopause is steeply sloped and to locate the height of jet cores.

(c) The 500 hPa pressure level analysis and forecast velocity and wind fields which are used to identify areas of cold air advection, region of possible cyclogenesis and the areas where components are normal to mountain ridges.

(d) The 250 hPa and 200 hPa level analysis and forecast temperature fields, regions of strong isotherm packing and temperature trough.

(e) The current and forecast surface fronts and pressure centers which are inspected for signs of cyclogenesis and position of jet stream core.

(f) The current and forecast pressure gradient which is checked to identify areas of possible mountain waves activities.

(g) The current and forecast wind fields at the ridge top level in mountain regions which identify the areas of occurrence of mountain waves.

(h) The aircraft observation which are used to determine the relationship between CAT observation and existing synoptic scale parameter.

(j) Satellite photographs and NWP forecast which are used especially in data sparse area to identify upper air short waves, the jet stream core position, surface fronts centre and standing mountain wave’s clouds.

(k) Plotted radiosonde ascent which are used to locate area of strong vertical wind shear, stable layers and tropopause height.

72. CAT Forecasting Hints. Forecasting of CAT is still in the stage of evolution. The main difficulty in the analysis of CAT zones is the scale of motion involved. Individual turbulent eddies or waves have a micro-scale structure. The size of eddies which cause aircraft bumpiness varies from 300 to 1000 meters, depending on the speed of the aircraft. The distance between radiosonde observations is extremely wide, of the order of hundreds of miles. The data based on these observations do not reveal the microstructure of the waves. Thus the structure of eddies and waves are missed on the chart. The difficulties in this regard are enhanced due to its patchy nature and the fact that CAT zones appear to develop and dissipate with an irregular life cycle.

73. CAT can be expected in the regions where:-

(a) Vertical wind shear greater than 4 kt per 300 m.

(b) Horizontal temperature shears of 5oC per 150 km.

(c) Horizontal wind shears greater than 25 kt per 150 km (moderate CAT) and 50 kt per 150 km (severe CAT).

(d) Left or polar side of jet stream at all altitudes around the jet stream and just below the tropical tropopause.

(e) In the convergence zone of two jet streams.

(f) Inversions especially near and downwind of mountain ranges.

(g) Over sharp troughs.

(h) One can encounter CAT when flying from a slow moving air mass of about 10 to 20 kt into or near a jet stream with speed of well above 100 kt.

74. It is also found that CAT could be predicted over the areas and near the levels of large inversions (Table 4).

Table 4.

Inversion Criteria Type of CAT

1.5°C/1000' No CAT

1.5°C/1000' to 2.5°C/1000' Light CAT

2.5°C/1000' to 4.0°C/1000' Moderate CAT

4.5°C/1000' Severe CAT

75. According to published material, horizontal and vertical temperature gradient are considered to be better indications of turbulence in the 50000-70000 ft range than any other routine parameters. It has been emphasised that this includes large lapse as well as large inversion rates. A suggested relation is given in Table 5. Table 6 shows the CAT occurrence with regard to different types of waves.

Table 5.

Temp Gradient Horizontal Vertical Turbulence

Small 1°C /25 nm 1.5 ° C/1000' Very light to smooth

Medium 1.5°C/25 nm to 1°C/12 nm 1.5°C to 2.5°C /1000' Light to moderate

Large 1.5 °C/1000' 1.5 ° C/1000' Moderate to Severe

Table 6.

Type of Wave Dimension Turbulence

Large Wave Horizontal wave heights/ 1/2 wavelength 4/3 Moderate or Severe

Medium Wave Horizontal wave heights/ 1/2 wavelength 4/3 to 3/4 Light to 3/4 Mod

Small Wave Horizontal wave heights/1/2 wavelength 3/4 Light or Less CAT

76. CAT relies more on pilots' abilities to extract information from their weather charts and meteorological forecasts, the reports of other aircraft in the vicinity, and good old fashioned experience. Pilots noting the curling cloud formations would likely conclude CAT was in the vicinity.

77. Forecasting techniques used are basically extrapolation through association with synoptic features. Although one cannot see CAT visually, a close scrutiny of the weather charts or the forecast turbulence factor on the flight plan, could usually warn pilots of possible affected areas on the route. Most of the knowledge of CAT is based on actual reports from aircraft. Every forecaster should, therefore obtain post-flight debriefing whenever aviator encounters CAT.

78. Checklist for CAT. It would be apparent from above that our knowledge about CAT is still incomplete and forecasting methods depend upon high level wind and temperature features, terrain parameters and gravity waves. A Check list prepared by Clodman et al (1960) for forecasting turbulence is a convenient aid, and is given on the next page.

79. Problems in CAT Forecasting. There are several notable problems with CAT forecasting:-

(a) It cannot always be foreseen so there is no warning. (b) It is usually felt at its mildest in the flight deck and is generally more severe in the aft section. (c) It can occur when no clouds are visible. (d) Aircraft radars can't detect it. (e) It is common at high altitudes, where cruising airline suddenly enter turbulent areas.

80. To overcome the above mentioned inherent problems in CAT forecasting, it is always advisable for a field forecaster, specially posted in the valley to make maximum use of airborne reports & debriefs obtained from the aviators.

CAT Occurrence over India

81. The incidence of CAT over India has been studied from a large number of reports from aircraft on routine as well as non-routine high level flights. Some of the more important results of the study are given below:-

(a) CAT frequency is highest from October to May over central and northern India and in July-August over southern India.

(b) Maximum incidence is from December to February, coinciding with the peak activity of the sub-tropical jet stream.

(c) Cat zones are usually of patchy nature. Average dimensions of distinct zones are 150km in the north-south direction and twice this in the east-west direction.

(d) Vertical extent of a CAT zone may be about 300 to 600 metres but in many cases CAT extends through a deep layer.

(e) Most of the encounters are of feeble/moderate CAT. In a few cases they are reported as severe. Severe CAT is restricted to the period December to May.

(f) Greatest chances of CAT are in a zone about 300km to the south of the sub-tropical jet stream axis.

(g) In Western Himalayas, CAT is most frequent in October when the sub-tropical jet stream makes its appearance over the area.

(h) In Eastern Himalayas, CAT is frequency is high in the mid-winter months.

(j) In south peninsula, Cat is associated with the easterly jet stream of the south west monsoon season.

Mountain Waves

82. Though flights through mountain waves have often been reported to be remarkably smooth, violent turbulence can often occur in association with some waves. The transition can be quite and dangerous. The turbulence associated with airflow over mountains is turbulence in frictional layer, rotor zone, waves and lee waves:-

(a) Turbulence in Frictional Layer. The terrain over the mountainous region being uneven gives rise to a variable depth of frictional layer and the

variation and irregularity of the terrain gives rise to turbulent layer (sometimes indicated by stratocumulus formation). With extensive upslope, the turbulence region may be visible by the cloud fall or the Fohn wall. The factors which determine the turbulence are same as that over a level country viz., stability, and surface roughness speed of the wind. The forecast of turbulence for this region has been done subjectively. Apart from the frictional turbulence, there often develops on the leeside of mountain semi-organized mechacalturbulence, which is quite severe and dangerous. This is known as rotor streaming and arises when very strong winds with high static stability extend through a restricted vertical depth (no more than 1½ times the height of the hills). Turbulence can also occur when the wind directions varies markedly with height. A schematic diagram showing laminar, standing eddy wave and rotor streaming with schematic wind profiles is given alongwith.

(b) Rotor zone. Rotor zone or low cloud zone is often an integral part of the wave flow over moderate and large sized mountains. These zones lie below the crest of the lee waves, and the most vicious rotors form below the first lee wave crest, downstream of the mountain ridge. The zone, often indicated by rotor clouds which at times appear as large stationary rolls, and often appear as a line of cumulus or stratocumulus, give rise to the most severe turbulence and vertical accelerations over 7g have been encountered. These form in a line parallel and downwind to the mountain with the base near the level of the crest. Rotors also form when the wind changes directions.

(c) Waves. The smooth wave clouds often, quite suddenly become ragged indicating sudden breakdown of smooth flow to vigorous turbulence. This breakdown is associated with the air stream characteristics when conditions are near critical for wave formation. Turbulence in waves are more likely with a short wave length less than 2 km. The turbulence c a n not be

predicted with reliability but the height at which the turbulence is likely to be encountered, where the wave amplitude is changing most rapidly with height; the turbulent layer is the region immediately above strong inversions where the wave amplitude is changing most rapidly with height. It is interesting to note that onset of nocturnal cooling often favours the development of waves especially if the wind is not strong enough to prevent the marked increase in stability. Morning heating has opposite effect of inhibiting wave activity and sing a decrease in amplitude and

increase in wave length before the disappearance of waves. For similar reasons waves have a greater frequency in winter.

(d) Lee Waves. Two of the most important factors which characterised the formation of lee waves are static stability and wind profile. The most favourable conditions are:-

(i) The flow at and above the mountain top level is at right angles and approximately constant in direction.

(ii) The wind at mountain top level exceeds 25 kt.

(iii) There is a rapid increase in wind speed with altitude upto several km above the mountain top. Strong wave development takes place with stronger winds with a peak in the vertical wind profile somewhat above the mountain top level.

(iv) The existence of a stable layer at or a certain height above the mountain top with lesser stability aloft. Topography and stability characteristics play an important part in the evolution of these waves. The effects of these features are shown in schematic diagrams below:

83. Theoretical computation can be used to determine the formation of lee waves. The occurrence and properties of lee depend upon the parameter l

2, which

neglecting the rate of change of wind shear with height which is of a smaller order of magnitude, can be written in the form:

l2=(g/U2 )( lθ ∂θ∂ z )= S

U2 where g = acceleration due to gravity

U = horizontal wind speed θ= potential temperature

Z = height

S = static stability =

gθ ( ∂ θ

∂Z )84. Two methods are given below:-

(a) In order to compute the l - profile, Foldvik (1962) suggests a theoretical profile of the from

l = l o e-cz where c is a constant.

If the atmosphere is represented by two layers viz. 1000 to 700 hPa and 700 to 300 hPa, we can obtain an approximate profile for l centered at 850 hPa and 500 hPa level. To evaluate l 850 and l 500, take the smoothed sounding or ascertain the temperature lapse (T1000-T700)and(T700—T300)from air mass characteristics and enter them on the appropriate nomogram vide figure to obtain l 850 against appropriate prevailing wind speeds. The method to obtain a smoothed sounding is shown.

The process of smoothing. The broken lines represent smoothed wind and temperature data from which the parameter is computed. It is seen from the smoothed temperature data that the corresponding stability will vary continuously with height

To obtain maximum vertical velocities (Casewell 1966) use is made of the-following formulae:

(w1)max = Hu0 c1 (w2)max = Hu0c2 where c1 and c2 are defined by

(2.5 +0.7)c (( ρ0 / ρ1 )1 /2=c1 and

CL1

3.2C ( ρ0 / ρ1 )1 /2

=c2

H the height of the mountain barrier u0 = the wind speed at the ridge height across the mountains c1 and c2, h1, h2 and L1, L2 are found from the graphs given below:

L850 (kilometers-1)Graphs Used to Obtain L1, h1 and C1

--------- Values of C1 -- -- -- -- Values of h1 (meter)------- Values of L1 (Kms) Level Of Tropopause (hPa)

h1, h2 represent heights at which the maximum vertical velocity occurs) L1 and L2 are corresponding wave length. No lee wave is forecast if u0 is less than 20 kt or blows at an angle greater than 30 degrees from the perpendicular to the ridge line. If the point on the graph lies above the pressure line t the level of

tropopause, then no wave is to be forecast. If the height of the maximum vertical velocity is higher the level of the tropopause then the values are not reliable and marked waves are not expected. If there is marked decreases of wind speed at any level marked change in the direction say more than 30 o, then wave activity is to be considered to have been suppressed at this level and above.

(b) Scorer (1953) suggests the following:

(i) Obtain an estimate of the wind temperature profiles of the air and plot them on a T-φ gram upto at least 400 hPa.

(ii) For each 100 hPa layer, obtain the thickness Δ z and the difference in θ between the top and bottom of the layer in degrees °C.

(iii) From the figure below, read off values of (Ul )-1 multiply this by U1 the average wind speed in the layer (kt) and obtain l -1 in km.

Waves are more likely if (i) l -1 increases rapidly with height (ii) when in a layer near the ground (at least 200 hPa thick) l -1 is considerably less than the layer above (at least 300 hPa thick) (iii) If l -1 is very large near the ground with a thicker layer above in which it is small (with another layer in which it is large above that). l -1 does not change much in the lowest 500 hPa, it may be necessary to continue the calculation upto 300 or 200 hPa. The lee waves have maximum amplitude at the top of the layer in which l -1 is small i.e. at the level where l -1 begins to rise substantially with height. The wavelength will be less than 2π /l in the upper layers and more than 2π /l as measured in the lower layers. If the ridge is well defined and fairly narrow, the first lee wave is only 3/4 of wavelength from the ridge crest. If during the day, because of decrease in the lapse rate in the lower layers, l -1 increases, the wave length also increases but as it cannot exceed 2π /l as measured in the upper layers waves may become impossible, only to return in the evening with a slowly shortening wave length. In the figure, isopleths of l /UL are shown l -1 is given in miles when l is given knots.

85. Conditions favourable for Mountain Waves formation. The meteorological conditions required for the formation of MW are:-

(a) Marked stability in the lower layer with comparability low stability aloft.

(b) Wind speed at the level of the summit exceeding a minimum which varies from 8 – 13 m/sec (15kt – 30 kt) depending upon the ridge generating the waves and either increasing or at least remaining constant upto the tropopause.

(c) Wind direction within 30° to the normal of the ridge not substantially changing in direction.

86. Forecasting Hints. Gulati gives these following hints to forecast Mountain Waves:-

(a) Determine whether standing waves are possible in airstreams under consideration. Scorer parameter l 2 should be more at lower levels and less at higher levels.

(b) Check if the wind blowing is perpendicular across the ridge or not. Speed should be sufficient comparable to the height of the peak.

(c) Presence of Jet stream is favourable condition.

(d) Diurnal variation and seasonal variation to be kept in mind. Dusk is more favourable as stability sets in lower levels.

(e) It is been established that larger the mountain, stronger the waves necessary to produce maximum effect.

87. Mountain Waves over India.

(a) Mountain waves occur NE India during winter months with wave lengths 20 – 30km. Wave lengths of 5 – 10 km have also been observed.

(b) Over Western Ghats waves are observed during winter season with wave length 26 to 78.5 km. The vertical velocities were between 0.6 m/sec and 5.6 m/sec, the amplitude is found to increase with wave length.

(c) Western and central Himalayas are prone to mountain waves during monsoon and winter seasons with a systematic shift eastward during winter and a westward shift during monsoon. The observed wavelengths are 15-22 km and the longest observed was 38 km.

(d) Jammu & Kashmir experiences maximum mountain waves during winter season and in the wake of a western disturbance or westerly wave.

(e) Severe Mountain Waves are possible over Sikkim during winter after the passage of a western disturbance.

88. Forecasting of CAT is still in the stage of evolution. The difficulties in this regard are enhanced due to its patchy nature and the fact that CAT zones appear to develop and dissipate with an irregular life cycle. Most of our knowledge of CAT is based on actual reports from aircraft. CAT is generally associated with Jet stream and mostly occurs near the tropopause. Mountain barriers are responsible for the creation of Mountain Waves induced turbulence. Wind fields over mountainous terrains are disturbed than that over plain land. These disturbances are found over and to the lee side of the mountain extending upto stratosphere.

Icing

89. Formation of ice on the parts of an aircraft in flight is termed as Ice accretion or icing. Icing can affect the aerodynamics of the aircraft, leading even to loss of control. It is thus a serious hazard requiring careful study of the different types of icing, the meteorological conditions under which they occur and the techniques of flying to avoid or minimise the risk of icing. Although many types of aircraft are fitted with de-icing equipment, this provides only partial protection and its successful operation is facilitated by knowledge of the type of icing and the possible rate of accumulation.

90. Types of Icing. Ice accretion can be of two types. The processes involved in these two types of icing are different.

(a) Airframe icing(b) Engine icing

91. Airframe Icing. Airframe icing can be classified into three main types:—

(a) Hoar frost. Hoar frost occurs on an airframe in the clear air when its temperature is below the frost point of the ambient air. It is in the nature of a feathery deposit of ice crystals formed due to the sublimation of water vapour on to the airframe. Hoar frost is liable to form when an aircraft which has been flying at high altitudes rapidly descends to lower levels into warm and moist air. It is not, however, of a serious nature because it dissipates quickly when the aircraft warms up. Part of the deposit may also get shaken off due to the vibration of the airframe.

(b) Opaque rime. This consists of a white opaque deposit of ice having light porous texture. It is formed by super cooled water droplets freezing on the airframe individually when an aircraft is flying through clouds above the freezing level. It accumulates on the leading edges of wings, struts etc. A large amount of air is entrapped between the particles. As the accumulation builds out from the leading edges towards the airstream, the mechanical impact of the air tends to consolidate the ice, which thereby exhibits a white opaque appearance. Rime of this type usually does not have much weight. This form of icing may cause an alteration in the aerodynamic characteristics of the wing.

(c) Translucent rime or glaze ice. This is also known as clear ice because of its glassy appearance on the surface. It results from large supercooled cloud or rain drops freezing on the airframe on spreading after impact. Glaze ice formed due to freezing of rain drops is also known as rime ice. Since the drops unite on spreading, very little air is entrapped; hence the glassy appearance. The deposit is not confined to the leading edges, because some water remains in liquid form for some time and is spread back by the airflow. The initial deposit may have a flat surface, but with a substantial deposit the surface is generally uneven, though still smooth. Ice of this type is sticky and cannot be easily shaken off. If it breaks off at all, it comes away in lumps of dangerous size. The danger from this type of icing is not merely aerodynamic; its weight of accumulation as well as the vibration set up by unequal loading of wings, struts and propeller blades pose a serious hazard. It should be remembered that types of rime described above are extreme forms. In actual flight, aircraft may encounter conditions which lead to accretions intermediate between these three types.

92. Temperature Range for Airframe Icing. With the exception of hoar frost, airframe icing takes place below 0°C, but the probability of occurrence decreases progressively below —20°C, because at lower temperatures the proportion of supercooled water drops in a cloud decreases. Serious icing, which is usually in the form of glaze ice, is more common at temperatures a few degrees below freezing point. At lower temperatures the larger supercooled water drops spontaneously freeze; hence only smaller drops remain which may give rise to opaque rime. Cumulonimbus clouds are an exception to this. In these clouds the larger drops may be carried upto high levels in the strong upward currents. Except for cumulonimbus clouds, the optimum temperature ranges in which different degrees of icing is likely are as follows:—

(a) Severe : 0°C to - 7°C(b) Moderate : — 7°C to -12°C(c) Light : — 12°C to -20°C(d) Very light : —20°C to- 40°C

93. Kinetic Heating of Airframe. The temperatures referred to above apply to the temperature of the airframe on which freezing takes place. This temperature is different from the free air temperature at the level of flight due to two reasons:—

(a) Adiabatic compression or expansion of air close to the airframe and consequent heating or cooling on different parts of the airframeby conduction.

(b) Heating of the airframe due to friction while the aircraft is moving through air.

Both these effects depend on the speed of the aircraft. As they affect the temperature gauges provided on the aircraft, the readings of these gauges are more reliable in assessing the chances of ice accretion than free air temperatures given in Met reports or forecasts.

94. Airframe Icing in Relation to Cloud Forms. The rate of accumulation of ice on the airframe is dependent primarily on the size and the number of the drops. Even at a particular level both these factors vary within the same cloud. Nevertheless, some useful generalisations can be made as regards the probable icing conditions in different kinds of clouds.

(a) Cirrus, Cirrostratus and Cirrocumulus: These consist almost wholly of ice crystals. Icing hazard is therefore, negligible.

(b) Altostratus, Nimbostratus: They consist of supercooled water drops and ice crystals in varying proportions. Light or moderate icing is possible.

(c) Altocumulus: Light to moderate icing is likely but severe icing is possible in mountainous areas.

(d) Cumulus of large vertical development: Icing may range from light to severe type at least down to — 20° C.

(e) Cumulonimbus: Icing may range from light to severe type down to - 20 °C. Below this temperature severe icing is not significant. Liquid water content is an important factor in icing. Maximum liquid water concentration takes place around - 15°C and therefore, maximum ice formation in clouds is to be expected around - 15 °C.

95. Effect of Airframe Icing on Performance. Icing of the airframe can affect the flying characteristics in many ways: -

(a) Lift decreases. This is caused by a change in aerofoil shape when ice accumulates on the leading edges.(b) Stalling speed increases appreciably.(c) Drag is increased due to formation of ice on leading edges and on protuberances.(d) Propeller efficiency decreases. Uneven ice deposits on the blades cause vibration and blade distortion and consequent loss of effective power.(e) Hinges of ailerons, elevator and trimming tabs may not be capable of smooth and free movement and may get jammed.(f) Icing of piton tube may give erroneous readings of the ASI and icing of aerials may render communication difficult.

96. Precautionary or Remedial Measures. A few hints on precautionary or remedial measures in icing situations are listed below:-

(a) Plan your flight so as to be in the icing region for the shortest possible duration.

(b) Exercise caution when flying through rain or wet snow at flight levels near freezing temperatures.

(c) When flying into cloud above the crest of ridges or mountains, maintain clearance of about 1 to 2 Km above the ridge if the temperature is below 0°C. Icing is more probable over the crest than over the adjacent valleys.

(d) When ice has formed, avoid manoeuvres which increase the wing loading.

(e) If icing is observed, descend immediately to a level where the temperature is above 0°C Or ascend above cloud level if this is possible.

(f) Remember that fuel consumption increases due to icing because of increased drag and additional power requirement.

(g) If de-icing equipment is available, use this intelligently after assessing the type of icing and possible rate of accumulation.

97. Engine Icing. Engine icing is of two main types:-

(a) Impact icing. This occurs due to the impact of supercooled water drops on the air intake and induction system. The airflow becomes restricted, reducing engine power directly.

(b) Carburettor icing. When air passes through the carburettor choke and past the throttle butterfly, its pressure is reduced and the temperature falls due to adiabatic cooling. The cooling may be enhanced by the evaporation of fuel. If the humidity of the air entering the carburettor is sufficiently high, icing may occur due to sublimation. This type of icing can take place in clear air even when the ambient air temperature is more than 0°C. In fact the hazardous region of air temperature for this type of icing is around 13°C. Below - 10°C this type of icing is negligible unless liquid water is present. Carburettor icing is unlikely when the relative humidity is less than 60%.

Contrails

98. Clouds form due to natural processes. Condensation or sublimation can also occur in the atmosphere by the injection of water vapour into air and thereby inducing saturation. This type of artificial condensation or sublimation occurs in the wake of an aircraft and is important in high altitude flying of military aircraft. It (Condensation trails (contrails) is defined as visible streaks of condensed or sublimated water vapour formed in the wake of moving aircraft.

99. Importance to Aviation. When an aircraft leaves a. visible trail behind it, both the position and heading of the aircraft are revealed to an observer on the ground. The disadvantage of this to military aviation is self-evident. In mountainousAreas where radar surveillance is difficult, contrails may be the principal means by which enemy observers may locate an aircraft. It is, therefore, necessary that the meteorological conditions favouring their formation and the action to be taken to avoid them are well understood.

100. Types of Contrails. There are two types of contrails:-

(a) Wing tip trails. These are thin and transient contrails which are formed near wing tips and propeller edges. They are formed due to the aerodynamic reduction of pressure, expansion of air and consequent adiabatic cooling near wing tips and propeller edges. If the air temperature is already low, sufficient water may not be condensed to produce a visible trail; also if the air is very dry, the adiabatic cooling may not be sufficient for the dew point to be reached. Such trails are thus common only in mild damp weather at low altitudes. These trails are short and evanescent and are relatively unimportant.

(b) Exhaust trails. These form due to the moisture content in the exhaust gas. They may be long and persistent and are visible to a ground observer even when the aircraft is at great heights.

101. Physics of Contrail Formation. Combustion of hydrocarbon fuel (petrol or ATF) in the engine of an aircraft results in two end products, which are of importance in contrail formation:—

(a) Water vapour (b) Heat

Both these are delivered through the exhaust to the ambient air. Addition of water vapour increases the relative humidity of the ambient air. On the other hand, addition of heat raises its temperature and hence reduces the relative humidity. The two factors thus act in opposite directions. At high temperatures of the ambient air (i.e.at low levels) the second effect is predominant. For saturation to occur, the rise of relative humidity due to the addition of water vapour outweighs the reduction due to addition of heat. Contrails can, therefore, form only when the temperature of the ambient air is low, i.e. at high levels in the atmosphere.

101. Mintra Level. From calculations of the release of water vapour and heat from the exhaust, it is possible to determine the critical temperature at which contrail formation takes place. This critical temperature varies to some extent according to the relative humidity as well as the pressure of the ambient air. In any given situation, the critical temperature at which contrail formation is possible as per theoretical calculations is known as minimum trail (MINTRA) temperature and the level at which this critical temperature is found is known as the mintra level. Though the critical temperature for trail formation is reached at the mintra level, trails do not necessarily form at this height. But trail formation does not occur below this level.Since the combustion products of petrol and ATF are different in terms of composition by quantity, mintra temperatures in respect of piston engine and jet aircraft are slightly different. The mantra levels for jet aircraft are higher than those for piston-engine aircraft, other conditions being the same. When a single type of fuel is used, the variation of mintra level between one type of aircraft and another is negligible. The flight level at which the temperature is — 45°C gives a reasonably good estimate of the mintra level (at 100% humidity) for jet aircraft.

102. Drytra Level. When the temperatures are very low, contrails form even when the ambient air is absolutely dry (0% humidity). In such cases the moisture

from the exhaust is by itself sufficient to produce saturation and sublimation. Above this level contrails form irrespective of the humidity of the ambient air. This is known as the dry trail (DRYTRA) level. Thus trail formation necessarily occurs above thislevel. Drytra level is usually about 2 kilometres above the mintra level for saturated air.

103. Maxtra Level. In the lower tropical stratosphere the temperature increases slowly with height. At a certain level the temperature may become higherthan the critical temperature for contrail formation. Above this level no contrail will form. This level is known as maximum trail (MAXTRA) level. In the Indian area, except perhaps in the extreme north in winter, Maxtra levels are usually above the operational ceilings of aircraft in present-day use. Figure below explains the different layers in which contrails can form. In the case of jet aircraft, contrails in the Indian area rarely form below 7 kilometres. The usual height band of formation is 11-14 km. Mintra levels are usually lower in winter than in summer; they are lower in the extreme north than elsewhere. In the extreme north, the mintra level in winter may occasionally lower below 9 kilometres.

104. Persistence of Contrails. Long and persistent contrails occur under the following conditions:-

(a) High relative humidity of the ambient air. A state of supersaturation with respect to ice enables quick sublimation of the water vapour on the sublimation nuclei which are present in the exhaust gas.

(b) Moderate changes of wind speed along the horizontal and vertical to permit broadening of the contrail by diffusion. If the diffusion is too vigorous, then the trail dissipates quickly.

(c) Temperature more than 14°C below the mintra temperature. At these low temperatures any addition of water vapour from the exhaust is sufficient to induce saturation in the ambient air.

(d) Higher throttle setting than normal resulting in greater quantity of water vapour issuing from the exhaust. In such cases trails may form a little below the calculated mintra level because the calculations assume normal throttle setting.

105. Avoidance of Contrails. Evasive action in regard to contrail formation: -

(a) Fly below the forecast mintra level if the exercise or operation permits this.

(b) If possible climb into the stratosphere. Even if you have not reached the Maxtra level, contrails may be short and evanescent because the stratosphere is mostly dry.

(c) Fly above a layer of cirrus or cirrostratus. Apart from the cloud inversion at the top of the cloud layer, most probably the air is drier.

(d) Reduce the throttle setting as far as practicable.

106. Distrails. When an aircraft flies through a cloud below the mintra level, theheat from the exhaust may temporarily reduce the relative humidity of the ambient air to such an extent that the cloud droplets in the wake of the aircraft evaporate completely. If the cloud is thin, a furrow or lane is created in the cloud and is visible to a ground observer. This is known as a dissipation trail (DISTRAIL).

Atmospheric Obscurity: Lithometeors and Hydrometeors

107. A meteor is a phenomena, other than cloud observed in the atmosphere or on the surface of the earth, which consists of precipitation, suspension or deposit of aqueous/non-aqueous liquid or solid particles, or phenomenon of the nature of an optical or electrical manifestation. Meteors may be classified into four groups, namely hydrometeors, litho meteors, photo meteors and electro meteors. 108. Atmospheric obscurity is the name given to any distribution of solid or liquid particles in the surface layers of the atmosphere which renders surrounding objects notably indistinct or altogether invisible according to their distance. They can be classified as follows:-

(a) Hydrometeor. A hydrometeor is a meteor consisting of liquid or solid water particles, falling through or suspended in the atmosphere, blown by wind from earth’s surface, or deposited on objects on the ground or in the free air. Rain, drizzle, snow, showers, ice, mist, dew, rime, hoar-frost, glaze are the various types of hydrometeors. The hydrometeors affecting the visibility on ground are:-

(i) Fog. It is composed of suspension in the air of very small droplets reducing the horizontal visibility at the earth’s surface to less than one kilometer and relative humidity is more than 75%.(ii) Mist. It is a suspension in the air of microscopic water droplets or wet hygroscopic particles, reducing the horizontal visibility at earth’s surface.

(b) Lithometeors. A lithometeor is a meteor consisting of an ensemble of particles most of which are solid and non-aqueous. The particles are more or less suspended in the air, or lifted by the wind from the ground. The hydrometeors affecting the visibility on ground are:

(i) Haze. A suspension in the air of extremely small, dry particles invisible to the naked eye and sufficiently numerous to give the air opalescent appearance is called haze. The effect is mainly a result of scattering of light by the haze particles. In haze surface visibility is between 2000m and 5000m and RH is 75% or less.

(ii) Dust Haze. A suspension in the air of dust or small particles raised from the ground, by the dust storm or sandstorm is called dust haze. Surface visibility is 5000m or less and RH is less than 75%.

(iii) Smoke Haze. A suspension in the air of small particles produced by combustion is called smoke haze. Surface visibility reduces to 5000m or less.

(iv) Dust Raising Wind. Strong surface wind as of speed exceeding 20kt blowing and raising dust, reducing visibility to 5000m or less and RH less than 75% is called as dust raising wind.

(v) Volcanic Ash. Volcanic ash is suspended volcanic particles in the vicinity of volcanic eruptions.

(vi) Smog. Smog has been coined from a combination of the words fog and smoke. Smog is a combination of various gases with watervapour and dust.

Genetical Classification of Hydrometeors

109. From a genetical point of view the hydrometeors may be divided into three main groups.

(a) Hydrometeors of Frontal Cloud Systems. These hydrometeors are formed when huge bodies of air rise slowly because of the upglide movement due to a general convergence in the horizontal air flow along frontal surfaces; they originate from clouds of the altostratus-nimbostratus type and comprise rain, snow, or sleet as defined in Section 22.3. None of these hydrometeors is by definition indicative of frontal phenomena. What is characteristic of frontal (upglide) precipitation is the even falling of rain, snow, or sleet from a stratiform and extensive cloud system. The precipitation may be intermittent or of variable intensity, but its variations are never so pronounced as when it falls as showers of squalls (see below). As mentioned in Section 22.2, the upper portion of the altostratus system may be in an unstable state, and showers will then be superimposed on the more even frontal precipitation.

(b) Hydrometeors of Unstable Air Masses. These hydrometeors, which are formed when small bodies of air rise rapidly through the atmosphere, comprise all kinds of showery or squally precipitations. The showers or squalls may consist of rain, snow, sleet, or any kind of hail, as defined above. Showery or squally precipitation is characterized by the suddenness with which it starts and stops and by its rapid changes in intensity; also, by the aspect of the sky, which exhibits rapid variations in space and time from dark and threatening clouds (cumulonimbus) to clearings of short duration, often with an unusual blueness of the sky. Sometimes no definite clearing occurs between the showers, and the precipitation may not stop entirely between them; but, in such cases, the showery character is revealed by the sudden changes in intensity of precipitation and rapid alterations between lighter and darker clouds.

(c) Hydrometeors of Stably Stratified Air Masses. These hydrometeors, are characterized by the smallness of their dimensions and the slowness of

their fall, originate from clouds of the stratus or the fog type. r he principal types are drizzle, granular snow, and ice needles. Grains of ice are not necessarily typical of the stably stratified air masses, because they form as raindrops in warm air aloft and solidify in the cold and stable ground layer of air. Figure below shows a typical example of the variations in intensity of warm- front rain, warm-sector drizzle, cold-front rain, and showers in the rear of a depression that passed the southern part of Norway on May 9, 1938. The left-hand portion of the diagram shows the even increase in the rain intensity as the warm front approaches Lista (58°G’N, 6°34’E), the sudden but not discontinuous change to warm-sector (Irizzle at the passage of the warm front, and the similar change to cold-front rain, which in this case was of moderate intensity. The right-hand portion of the diagram shows the simultaneous rain intensities recorded at Sauda (50°39’N, 6°21’E) in the unstable northerly current of polar air in the rear of the passing depression.

Figure shows rain intensities during the passage of a Mid latitude depression.

Condensation Nuclei

110. Observation as well as theory show that condensation of water vapour into water droplets takes place on certain hygroscopic particles, and these are called condensation nuclei. Laboratory experiments, in which pure air is cooled by rapid expansion, have shown that it is possible to produce supersaturation exceeding about 500 per cent without condensation occurring. By further cooling it is found that condensation is initiated on certain ions which then serve as nuclei. In natural processes the rate of cooling (whether adiabatic or nonadiabatic) is very slow. The most rapid rate of cooling is found iii the updrafts of clouds of the cumulonimbus type, hut even iii these (buds the amount of supersaturation (in respect of water) is unlikely to exceed a fraction of one per cent relative humidity. r1his then, rules out ions and all non hygroscopic particles as nuclei of condensation. Although definite proofs are lacking, much evidence’ has accumulated in support of the view that the principal condensation nuclei iii the atmosphere consist of sea salt, with sulphurous and nitrus acids (associated with products of combustion) taking a secondary place. The number of sea- salt nuclei varies considerably; the estimates range from

10 to l00 cm3. All evidence indicates that the nuclei are sufficiently numerous and active to prevent any measurable amounts of supersaturation (in regard to water) from occurring when the air is cooled. As shown in the following paragraph, the size of the nuclei is important in considerations of the initiation of condensation. The best estimates indicate that the spherical diameters of sea-salt nuclei vary from about 0.05 to about 1µ, and there is some indication that giant salt nuclei may be present occasionally. The combustion nuclei are generally smaller than the salt nuclei, and their number varies considerably with industrial activities.

Condensation

111. The term saturation is defined as a state of equilibrium with respect to a plane surface of pure water, and the term relative humidity was defined with this saturation as reference level, if the water surface is curved, or if the water is not pure, other equilibria may be attained. In effect, this means that a droplet, such as a cloud particle, which is small and impure, may be in equilibrium with the surrounding air when the relative humidity differs somewhat from 1 00 per cent. From physical chemistry it is known that the equilibrium vapour pressure of water is reduced when salt is dissolved in liquid water, the reduction being expressed by:

e

e s=1−CM

Here, es is the saturation vapour pressure of pure water, e is the equilibrium vapour pressure of the molar aqueous salt solution containing M moles of solute per litre of solution, and C is a constant which depends upon the temperature of the solution and the particular electrolyte used. Now, if water vapour condenses on the solution, ill will decrease, and the equilibrium vapour pressure will approach the saturation vapour pressure. If an average sea-salt nucleus were dissolved in a spherical drop of water with radius larger than 1µ, the solution would be so diluted that, for all practical purposes, the effect of the salt would be unnoticeable. Hence, the solution effect is of importance only at the initiation of the condensation process, which is when the droplets are very small. For sufficiently small drops, the equilibrium vapour pressure depends not only on the concentration of the solution but also on the surface tension. This latter effect, which was discovered by Thompson (Lord Kelvin), may be expressed by the approximate formula:

e−e se

= Kr

where r is the radius of the drop and K is constant for any given temperature. Since K is positive, the equilibrium vapour pressure of a drop of pure water is greater than the saturation vapour pressure, and consequently a certain amount of supersaturation is required to cause the drop to grow. Since the concept of surface tension has no known meaning for “drops” which consist of, say, two or three molecules, the argument cannot be extended to molecular dimensions. Nevertheless, it is evident that immense supersaturation would be required to cause condensation on drops of near-molecular dimensions, and this explains the need for

the presence of nuclei to initiate the formation of cloud droplets in such conditions as obtain in the atmosphere.

112. It will be seen from the foregoing formula that the Thomson effect decreases as the radius increases. The effect becomes negligible when the radius exceeds two or three microns. The initial growth of cloud droplets was considerably clarified by Kohler who combined the effects of curvature and solute with the thermodynamic equations of the condensation processes. The combined effect may be represented schematically as shown in Figure below. Suppose that a condensation nucleus of average size (for example, 0.5µ) is introduced into moist air. The nucleus will then condense moisture on its surface and grow to an equilibrium size, say at B. If the relative humidity of the air now is increased (e.g., by adiabatic cooling) the droplet will continue to grow to a new equilibrium at, say C. If, again, the air is cooled so that the relative humidity is increased to the value indicated by D, the drop will grow along the curve CEF, without reaching equilibrium with the surrounding air. Along the path EF the solution is so diluted and the drop so large that the equilibrium vapour pressure is very nearly that of a plane surface of pure water. If a very small nucleus were introduced, the drop would soon become so diluted that the curvature effect would outweigh the effect of the solute. The drop would then develop along the path SPS’ in Fig.2 and the peak (P) would he higher than for a larger nucleus.

Figure illustrating the combined effect of solute and curvature.

Thus, when the air is cooled (e.g., in adiabatic ascent) the larger nuclei would first become active and absorb the superfluous water vapour, and many of the smaller (or less active) nuclei may not be activated at all. Computations of the rate of growth of droplets by diffusion of water show that it takes a nucleus about 1 sec to grow to 10µ, about 2 min to grow to 100µ, about 3 hr to grow to 1 mm, and about a day to grow to 3 mm. These data, as well as other evidence, rule out condensation as the mechanism that produces raindrops of the sizes observed. From the foregoing discussion it may be concluded that the condensation nuclei in the atmosphere are sufficiently active and numerous to cause condensation to occur, and clouds to form, when the air is cooled to a very slight degree of supersaturation, and that the rate of

condensation is far too slow to account for the formation of raindrops, except, perhaps, fine drops of drizzle.

Precipitation Processes

113. Recent progress in research on these processes owes much to a classical paper by Bergeron who suggested a mechanism which, at least in certain circumstances, would suffice to cause certain cloud particles to grow to the size of the observed precipitation elements within reasonable intervals of time.

114. The process visualized by Bergeron may be described briefly as follows. In the first place, the cloud droplets may be cooled far below 0°C before solidifying. Figure below shows the frequency of occurrence of water clouds and ice clouds grouped according to the temperature at the top of the cloud. It will be seen that ice clouds are very rare when the temperature at the top of the cloud is above about —8°C. When the clouds extend to greater heights (lower temperatures), their upper portions will often consist of a mixture of water droplets and ice particles, and at very low temperatures ice-crystal clouds (cirrus) are found.

Figure shows Percentage frequency of water clouds and ice cloudsin intervals of temperature at the top of the cloud

115. Second, the saturation vapour pressure over ice is lower than that over water. The difference varies with the temperature as shown in Figure below. It will be seen that the difference has a maximum at about —12°C and is appreciable

Figure shows the difference in saturation over water and ice (curve ΔE); andrelative humidity (curve R) of air that is saturated over ice.

within the interval from about —5 to about — 25°C. In a cloud that consists of a mixture of water droplets and ice crystals the vapour pressure is likely to be a compromise between the two saturation pressures; there will then be supersaturation over the ice particles and sub saturation over the water droplets, with the result that the former will grow at the expense of the latter. Once the ice particles have become large enough to fall through the cloud droplets at an appreciable rate, other processes, such as direct capture, may become important.

116. Bergeron visualized the ice as forming through the freezing of cloud droplets, although direct crystallization from water vapour on suitable nuclei was also envisaged. The criterion for release of precipitation was, according to Bergeron, the formation in a supercooled water cloud of ice crystals, and this required that the top of the cloud should reach up to such heights that sufficiently low temperatures would he encountered. The essential features of Bergeron’s model are reproduced in Figure below. The same model would apply to clouds of the cumulonimbus type, except that the cloud would be more or less vertical and that the gentle upglide motion would be replaced by strong and irregular convective currents. The physical aspects of Bergeron’s model were further developed by Findeisen who appealed to the presence of sublimation nuclei for the formation of ice crystals.

Figure shows the idealised model of precipitating cloud

117. Observations show that size of the droplets within a non-precipitating cloud varies through a fairly wide spectrum which may be characterized by its width, the median droplet size, and the number of droplets (or the liquid water content) per unit volume. In a non-raining cumulonimbus or nimbostratus the median drop diameter is about 40 to 50µ, with some drops as large as 80 to 100 µ. The liquid-water content of such clouds, though varying considerably from one case to another, is generally of the order of 1 to 2 g m-3. In thin and stable layer clouds the median drop diameter is normally about 10 to 30µ with a liquid-water content of about 0.1 to 0.3 g m-3.

118. The Bergeron-Findeisen mechanism met with considerable acceptance by meteorologists in middle latitudes, where the 0°C isothermal level normally is so low that almost all deep clouds extend to subfreezing temperatures. Computations by Houghton and others have shown that the ice—crystal effect is significant, and there can be little doubt that it is effective, provided that the temperature in the upper portion of the cloud is sufficiently low and that the cloud is sufficiently deep to ensure that the falling elements can grow further by capture of a sufficient number of other cloud elements. On the other hand, irrefutable evidence has accumulated in recent years to show that warm clouds, particularly in tropical latitudes, are capable of producing rain although their tops do not reach up to the freezing level. It is evident, therefore, that the ice-crystal effect is not the only one that can initiate the growth of cloud particles into sizable precipitation elements. The processes that determine the drop-size distribution in clouds are not well understood, but it is likely that the larger specimens result from occasional collisions of small droplets. Because of their varying size, the droplets will fall with different velocities relative to the air. One has then to consider two aspects of the relative fall. In the first place, a drop that falls through a cloud may capture smaller drops on its forward side. We shall refer to this as direct capture. Second, drops that get into the wake of a falling drop may overtake the leading drop. We shall refer to this as wake capture.

119. When a drop falls through the cloud, it will sweep through a cylinder containing a certain number of droplets. The smallest of these will be swept aside by the divergent flow in advance of the drop, but the larger elements will be captured. In this process the collecting efficiency’ is obviously less than unity. The collection efficiency varies with the size distribution. Drops in the range from 1 to 2 mm in diameter falling through droplets of 40µ have a collection efficiency of about 0.95. The efficiency decreases and tends to zero when the drops are of the order of 10µ. The direct capture is important in clouds where some elements are already much larger than others. Recent experiments by Telford Thorndike and Bowen have shown that the wake capture is highly effective in causing coalescence of drops of almost equal size. While the air motion in front of the falling drop is strongly divergent, the convergence in the rear is slow, and a wake of air trails behind the moving drop. When a neighbouring drop of almost equal size comes near the wake, it becomes accelerated toward the wake and overtakes the leading drop when it enters the wake. The acceleration of the trailing drop commences at a distance of about 40 times the diameter of the leading drop. The experiments show that the collection efficiency in this process is of the order of 10 for drops larger than about 150 in diameter.

120. The precipitation processes may now be outlined as follows:-

(a) In middle and high latitudes where the clouds often reach to temperatures where freezing takes place, the irrepressible effect of ice particles will cause certain cloud elements to grow at the expense of others. Whether or not precipitation results depends upon whether the cloud is sufficiently deep to ensure further growth by coalescence and related processes.

(b) In tropical latitudes the clouds do not normally reach up to freezing temperatures. The growth can then be initiated only by coalescence of some

kind or another. Whether or not precipitation will result depends on whether the cloud is sufficiently deep to sustain the growth to raindrop sizes.

(c) In clouds with appreciable updrafts, such as building cumuli, the initially large drops will be carried upward, although slower than the smaller drops some of which will be captured. If the updraft is sufficiently deep, the drops will grow to such sizes that they can no longer be supported; a downward trajectory begins, and the drops continue to grow until they fall out of the cloud, or reach the critical size.

(d) In middle and high latitudes, where the freezing level is relatively low, the ice-crystal effect and capture combine to make the cloud yield precipitation. Once precipitation begins to fall, other growth processes come into play. Thus, wet snowflakes will aggregate and form large flakes which later melt and form large drops. Furthermore, snowflakes (and hail) falling below the melting level will retain a temperature at 0°C, while the temperature of the ambient air may be several degrees higher. To such large temperature differences corresponds a water-vapour gradient of 1 to 2 hPa, with the result that large amounts of supersaturation are present at the surface of the melting flakes.

(e) As has been shown by Cunningham [14] appreciable amounts of water are collected when precipitation elements produced by clouds aloft fall through stratus and scud at low levels. Thus, much of the precipitation at the ground may come from low cloud layers which, by themselves, could not release precipitation.

Fragmentation

121. Lenard observed that falling water drops larger than about 5.5 mm in diameter would be unstable and break into smaller drops. Later investigations by Blanchard and others have shown that the critical diameter at which the drop breaks depends on the intensity of the turbulence. Although drops as large as 7 mm may be stable when falling through calm air, 5 mm appears to be an approximate upper limit in air with ordinary turbulence. Most frequently, an unstable drop disintegrates into a few (2 to 10) drops and a number of droplets. This fragmentation process is important since it provides a mechanism for a “chain reaction” through which drops of critical size may multiply into a large number of drops which, in turn, may grow and produce new chain reactions. It will be seen from Figure below that the terminal velocity of raindrops near the critical size is about 9 m sec-1. This velocity is comparable with the updraft velocity in cumulonimbus clouds. Drops near the critical size may, therefore, be carried upward and disintegrate in the updraft.

Figure illustrating the terminal velocity of freely falling raindrops and strength of sharp-edged gust necessary to disrupt raindrops.

122. It has been shown by Hochschwender that drops smaller than the critical size may disintegrate when exposed to the shock of a sharp-edged gust. The minimum sharp-edged gust necessary for disintegrating a raindrop is shown in Fig. 6 It will be seen that a sharp-edged gust of 9 m sec-1 suffices to disrupt a drop of 3.5 mm in diameter. Undoubtedly, such gusts may occur in cumulonimbus clouds. The effect of such disintegrations (either by gusts or by critical size) is to prevent raindrops from falling out of the updraft and to accumulate large amounts of liquid water in the cloud. As liquid water accumulates, the frictional drag retards the updraft, and a heavy downpour results. Splintering of ice crystals has been observed by Findeisen and Palmer who suggested that the process is important in forming large numbers of ice-crystal nuclei.

Some Observational Evidence

123. The results of cloud census off the coasts of Puerto Rico by Byers and Hall are shown in Figure below. In all cases the 0°C level was well above the cloud tops. It will be seen that the percentage probability for tropical cumuli to produce rain echo increases almost linearly with the height of the cloud top. Since the height of the cloud base varied only within a narrow range, the diagram may be interpreted as showing that the depth of the cloud is highly significant. In no case was rain echo observed when the height of the cloud top was less than 6000 ft (≈2 km). A small sample of cumulus clouds over Puerto Rico showed similar results, except that build-ups larger than those observed over water appeared to be required to obtain rain echoes. When the over-water cumuli reached beyond 11,000 ft, the frequency of rain echoes was 100 per cent. An examination of 42 individual clouds with heights between 6000 and 9000 ft showed that 18 of them produced precipitation during their life cycle.

Figure shows percentage of over-water clouds with radar echoesas a function of cloud- top height and temperature

124. Batton has shown that the first rain echo in cumulonimbus clouds in the United States during the warm season is sometimes observed to form below the 0°C level, and this suggests that the precipitation process, in these eases, is initiated in the warm part of the updraft. As soon as the echo forms, it spreads rapidly upward as well as downward, showing that the ice-crystal effect may come into play shortly after the appearance of the first echo. Undoubtedly, all precipitation-releasing processes contribute to the further development.

125. In spite of much progress in cloud physics research, the forecaster remains at a disadvantage, for the observations routinely available contain but little information related to the condensation and precipitation processes. In forecasting clouds and precipitation, the forecaster will be guided by the large-scale features of the synoptic systems. Although our knowledge of the precipitation processes is meagre, some weight must be placed on forecasting experience which indicates the deep clouds, and particularly clouds that reach up to subfreezing temperatures, readily yield precipitation.

Lithometeors

126. Dry lands occupy more than 40% of the world’s land surface. They are home to about one billion people. Dust storms are a symptom of poor land management and a constant reminder of the interaction between people, the land they use and the climate. When land management is inappropriate as a result of government policies or because the traditional technologies are no longer able to cope with burgeoning populations and the shrinking resource base, wind erosion occurs. It is well known that desertification is a consequence of natural factors (mainly climatic elements) and human elements. As one of the manifestations of desertification, sand-dust storms are both an important process of acceleration of desertification and a consequence of land desertification.

Physics, Mechanics and Processes of Dust / Sandstorms

127. Field observations and wind tunnel laboratory research have helped to explain the physical process of sand and dust blowing under the force of wind and moving over the land surface in arid and semi-arid zones. When the wind force reaches the threshold value, the sand and dust particles are transported from the surface and start to move. Soil erosion by wind has two broad dimensions: transport and accumulation. Studies on sand/dust storms cover both aspects, because each is damaging in its own way and each contributes to the problem of desertification in the world’s dry lands.

Dust Becoming Airborne

128. Field observations and wind tunnel laboratory research allow us to understand the physical process. Consider a surface made up of separate particles that are held in place by their own weight and some inter-particle bonding. At a low speed wind, there will be no indication of motion, but when the wind force reaches the threshold value a number of particles will begin to vibrate. Increasing the wind speed still further, a number of particles will be ejected from the surface into the airflow. When these injected particles impact back on the surface, more particles are ejected, thus

starting a chain reaction. Once ejected, these particles move in one of three modes of transport depending on particle size, shape and density of the particle. These three models are designated suspension, saltation and creep. Its size and density determine movement pattern of sand- dust particles. The suspension mode involving dust particles of less than 0.1 mm in diameter and clay particles of 0.002 mm in diameter are small in size and light in density. These fine dust particles may be transported at altitudes of up to 6 Km and move over distances of up to 6,000 Km.

129. Saltating particles (i.e. those between 0.01 - 0.5 mm in diameter) leave the surface, but are too large to be suspended. The remaining particles (i.e. above 0.5 mm) are transported in the creep mode. These particles are too large to be ejected from the surface and are therefore rolled along by the wind and impacting particles. Coarse sands of 0.5 - 1.0 mm in diameter move along in a rolling movement. Medium-sized sands of 0.25 - 0.5 mm in diameter encroach in the form of a jumping movement. As these particles impact upon the land surface, they initiate movement of other particles. About 50-80% of all soil being transported is carried in this mode. Due to the nature of this mode, the heights carried are rarely more than 30 cm and the distance travelled rarely exceeds a few meters. Sand particles, transported by saltation and by creep will accumulate to form new sand dunes when they are blown out, graded and transported for a distance. Sands of 2.0 mm in diameter will be left on land surface when fine materials are blown away. Dust may be raised from the ground by the wind and carried upwards to various heights, depending on the size of the particles and the prevailing conditions. The development depends on mainly three factors:-

(a) The availability of the dust. The dust source is a land surface covered by dry dust. It is usually a permanent desert or a region which has been subject to draught conditions. In some places soil may also be exposed as a result of the removal of the vegetation cover by artificial means.

(b) The wind speed. The wind needs to reach at least moderate speeds in order to disturb the dust.

(c) The stability of the air. In stable conditions the turbulence induced by the wind is damped out and the dust may reach a height of only a few meters.

130. The visibility tends to be the lowest during the period 1000 to 1600 IST because of the lapse rate conditions required for dust suspension. An improvement generally occurs towards evening with the onset of a nocturnal inversion and the diurnal fall in the surface wind speed. However a dust haze may persist due to the settling of dust particles from the higher levels. In situations where there are few or no clouds to impede solar radiation, and the ground is strongly heated ,leads to the establishment of a steep lapse rate near the earth’s surface. On other occasions dust may be raised by cumuliform clouds resulting in dust being carried aloft to cloud base level. Dust raising wind is different from dust storms. Dust storms are short lived, squally phenomena lasting from a quarter of an hour to 2 hours while dust raising winds last for considerably long periods of 12 to 72 hours, more or less unabated.

Smoke

131. Small particles of combustion products in suspension reducing visibility to 5 km or less is termed as smoke. The presence of both natural and artificial pollutants in the atmosphere constitutes a visibility hazard to aviation. The larger particles tend to settle, but much of the pollution in the form of finely divided particles remains suspended in the air. They are often comparable in size with the water droplets found in a fog or mist. Smoke from industrial and domestic fires can become problem if an inversion is present. In light winds and humid conditions, fog may also be present and the combined effect may seriously reduce visibility. The term smog is given to this phenomenon. The obscurity produced by smoke depends on:-

(a) The rate at which smoke is produced by the source.

(b) The distance of the source from the airfield.

(c) The rate at which the smoke is dispersed by wind and turbulence.

132. Strong winds improve the visibility by transporting the pollutants horizontally. Convection and turbulence may also assist in reducing the smoke concentration by dispersing the pollutants vertically. The action of rain or snow can also play an important role in cleansing of the atmosphere. Volcanic Ash

133. Volcanic ash is the term for very fine rock and mineral particles less than 2 mm in diameter that are ejected from a volcanic vent. Ash is created when solid rock shatters and magma separates into minute particles during explosive volcanic activity. Volcanic ash jams machinery. This poses a great danger to aircraft flying near ash clouds. There are many instances of damage to jet aircraft as a result of an ash encounter. Very fine ash particles may remain high in the atmosphere for years, spread around the world by high-altitude winds. This suspended material contributes to often spectacular sunsets, as well as an optical phenomenon known as "Bishop's Ring". This refers to a corona or halo effect around the sun.

Jet Streams over Indian Region

134. The jet streams are narrow bands of fast, meandering winds near tropopause that flow around the globe. The existence of strong winds in the upper troposphere in westerly direction over middle latitude was first discovered through early observations of the drift of Cirrus clouds, after the eruption of Krakatau volcano in 1883, but these observations were taken to indicate occurrences rather than to reflect a normal state.

135. During World War II, bomber sorties over Japan by the American planes gave the initial impression for study on this subject. American B-29's first began to bomb Japan when the pilots reported encountering high velocity head winds with speeds exceeding 250 mph at 20,000 to 30,000 ft that their ground speed was reduced nearly to zero and they found themselves remaining stationary over their target areas. Similarly when they turned downwind, then ground speed became nearly

double their air speed. These reports were hard for meteorologist to believe. Such surprisingly strong winds are not been visualised by the studies and models available at that time. It was then the study of the jet streams gained prominence and a vast amount of study has been made on the subject both observational and theoretical.

136. Subsequent high-altitude Weather studies disclosed the presence of narrow bands of high velocity winds of 200 to 300 miles an hour in other parts of the earth also. The first organised study of jet stream was carried out at the University of Chicago during 1946-47, under the leadership of Rossby and Palmen. These studies stated that “A broad stream meandering eastward around the hemisphere in wavelike pattern. The kinetic energy of this stream is concentrated in a narrow band of high wind speed embedded in a relatively quiescent surrounding atmosphere”. This narrow band was named the “JET STREAM”.

137. The jet stream is caused by the meeting of air masses just under the tropopause where winds are strongest. When two air masses of different densities meet here, the pressure created by the different densities causes winds to increase. As these winds attempt to flow from the warm area in the nearby stratosphere down into the cooler troposphere they are deflected by the Coriolis Effect and flow along the boundaries of the original two air masses. The results are the polar and subtropical jet streams that form around the world.

138. These two near-tropopause jet streams, the subtropical jet, imbedded in the sub-tropical upper troposphere front which is found in the pole-ward edge of the Hadley circulation, and the Polar jet, imbedded in the upper troposphere Polar front which is located above the Polar front zone (Holton, 1992; Bluestein, 1993). The positions of the jet streams are important because synoptic scale disturbances tend to form in the regions of maximum and minimum jet stream wind speed, and to propagate eastward along tracks that follow the jet axis (Holton, 1992). Changes in jet stream locations can therefore cause changes in the storm tracks, global weather patterns, temperatures, precipitation and the hydrological cycle.

Sub Tropical Jet Stream

139. The subtropical jet stream (STJ) acquires its name from the fact that it is found in the high level Westerlies above the subtropical high pressure belts. In winter, it migrates to about 22° N and is strong, while in summer it moves to North of latitude 35°N and weakens. The STJ lies close to the middle latitude tropopause, but this type of tropopause is seen only to the North. Towards the South of the core, the tropical tropopause lies aloft about 4 to 5 km above the jet stream core. The STJ affects India mainly in the non monsoon months.

140. The following methods are used to locate the position of jet stream:-

(a) Constant Pressure Chart. The charts for 250 and 200 hPa levels are basic for the location of jet stream. Where wind observations are available, the line of maximum wind speeds is the axis of jet stream. The Rawin observations at these levels may be supplemented by Pibal winds, reports from aircrafts & satellite derived upper winds. When sufficient wind

observations are not available, one may utilise winds derived from the contour charts to draw the isotach patterns.

(b) Vertical Time Section. Since the vertical time section show that the high speed isotachs around the jet core extend downwards (in an elliptical form) through fairly deep layer, even the wind distribution in levels below 200/250 hPa upto 500 hPa, can give a rough idea of the location of jet axis.

(c) Streamline Analysis. The jet streams usually tend to follow the streamline / contour. This is particularly so when the speed does not vary much along the axis. However when the speed varies appreciably along the axis, the jet may cross contour/streamlines. The axis tends to cross towards the higher contour values. When the speeds are decreasing downstream (i.e. where air is decelerating) towards lower contours where the speed increases along the axis (i.e. where air is accelerating). This feature is generally well marked when a jet stream is embedded in a trough ridge pattern.

(d) Space Cross-section. Along selected longitudes, these are very useful to locate the jet streams in three dimensions and study its day to day behaviour. So far as India is concerned, cross section along two mean latitudes that cover a fair density of stations suggests.

(i) Along a mean longitude of 75°E, containing stations within 3° to 5° longitude from this meridian (75°E) from Russian Turkistan (Almalta, Tashkent etc.) in the North to the Lakshadweep and Maldives in the South (Minicoy, Trivandrum, Colombo).

(ii) Along a mean longitude of (90°E) from East Tibet in the North to the Bay islands in South.

(e) Cross-section of Temperature Anomalies. Riehl suggested that a vertical cross section of temperature anomalies (from some mean value, usually w.r.t the mean sounding computed from all the soundings in the cross sections) along a longitude may be useful to locate the jet stream with a fair degree of confidence. The jet stream is located in the ‘COL’ region between two warm and two cold anomaly regions.

(f) Tephigram. The core of the jet stream is most often in between two standard pressure levels. We may utilise the property that the jet levels, the temperature gradient is a minimum (or zero) and obtain the level of jet core as the level, where two ascent curves meet. Where more than two radio-sonde observations are available, a mean level may be taken from the tephigram analysis of SRN, DLH, and NGP. The curves meet near about 200 hPa. Actual winds also show that the jet axis is located near 200 hPa between DLH and NGP. It is noticed that the temperature gradient reverses sign at about 200 hPa level.

141. Variation of STJ with season. The Subtropical jet stream can be seen over Indian subcontinent for about eight months of the year (Oct – May). With the establishment of SW Monsoon over the subcontinent, the subtropical jet shifts

northwards, weakens and disappears over Indian regions. In winter, the mean position of the subtropical jet is roughly at longitude 27°N, at the height of 12 km with a mean speed of 100 Knots. Southernmost positions as far south as 22°N are reached in February. The maximum wind speed attained range between 150 and 200 Kt. Sometimes there is a split in the jet stream due to Himalayas. The jet is seen upto Afghanistan thereafter splitting one branch to the South and another to the North of Himalayas. These branches later recombine to form a single stream over China/Japan.

Structure of Subtropical Jet Stream

142. Saucier (1958) used data from aircraft and regular aerological observations and emphasised that the structure varies greatly from one case to another.

143. Some more facts related to structure of jet stream are as follows:-

(a) The jet stream is associated with strong horizontal and vertical wind shear. The increase or decrease of wind with height and horizontally are largely caused by large meridional gradient of temperature. There is a concentration of temperature gradient both above & below the core.

(b) Koteswaram (1951), from the data of Feb 51 suggested that there is a marked increase in the horizontal temperature gradient below the jet and reversal of temperature gradient above the jet.

(c) Mohri (1953) noted frequent existence of a permanent stable layer below STJ, which he called subtropical front.

(d) Riehl et. al. (1954) emphasised the study of tropopause in relation to jet stream. They emphasised the existence of three tropopause.

(i) Polar Tropopause. It is at about 300 hPa level, north of Polar Front jet and has characteristic potential temperature of 310°A.

(ii) Middle Tropopause. This tropopause rises from the core of polar front jet to about 200 hPa level just South of core and then slowly and disappears at about 300 hPa and has an average potential temperature between 330 - 335°A.

(iii) Tropical Tropopause. This is located near 100 hPa level with a potential temperature of about 390°A.

(e) Defant and Taba (1957, 1958) confirmed the findings of Riehl and postulated the existence of the subtropical tropopause at about 200 hPa and asserted that STJ also lies in the break between the middle and the subtropical tropopause.

(f) Serebreny (1962) called this stable layer as layer of subsidence and stated that the Northern edge of STJ actually lies over Southern portion of middle where the middle tropopause is identified by subsidence layer rather

than a single break in the lapse rate. Serebreny did not find any evidence for the existence of subtropical tropopause.

(g) MS Singh (1962) studied the characteristics of subtropical jet structure over India & Pakistan with the help of longitudinal cross-section. It was observed that the subtropical jet is broad band of great latitudinal span in the break between the middle and tropical tropopause. Generally the jet core has two layers of maximum wind (LMW) attached to it. But it is replaced by separate cores with no LMWs, when the branching of jet stream takes place. Connected to each core and located beneath it, a layer of frontal type discontinuity, which may be called the subtropical front. The present knowledge about the westerly jet stream over India and Pakistan has been discussed in the light of above findings.

144. Mechanism of Formation of STJ. These systems are driven by radiative heat sources and sinks in the high atmosphere, which are quite, separate from the processes in the troposphere. Upper level divergence and convergence associated with waves in the atmosphere should in general be most pronounced in regions of strong winds. Some medium scale disturbances (cyclones and migratory anticyclones) are characterised by appreciable divergence and vertical motion, it is natural to expect these systems to show an affinity for jet streams. Each of the frontal system on a surface map has an accompanying jet stream. The STJ is related to the poleward boundary of Hadley circulation, where low level front tend to be obscure or absent.

145. Jet streams cannot be considered as uniform current around the globe, rather they are typified by concentration of stronger winds in the jet streaks, alternating with weaker winds. The individual streaks tend to progress along the current, however they move very much slower than the wind. Consequently air parcels move through them and gain speed as they pass through their upwind portions, losing speed in their downwind parts. This has been demonstrated directly by tracking the constant pressure balloons. Portion of balloon trajectory in a case where in the velocity maxima were in the wave crest and balloon underwent the remarkable deceleration in passing through a sharp trough, it is natural to acquire whether the dimensions of jet streaks can be connected with any general physical rule. One feature is that the streaks of greatest length and breadth appear to be those with the highest maximum speeds. Since the streaks are slowly moving, proportionality between longitudinal dimensions and wind speed would imply a preferred period of time required for the air particles in the jet stream core to move through them.

146. Utilising a dynamical model based on vorticity considerations Mohri (1959) achieved a considerable degree of fidelity in reproducing the characteristic winter upper level flow patterns over Southeast Asia. He concluded that the Himalayan-Tibet plateau “establishes two branches of jet stream (one to the North and one to the South side) and forms a strong jet stream on the lee side of the mountains by the confluence of the Northern branch and the Southern one” while this analysis indicate a general aversion to the plateau. On occasions STJ may pass over it, especially during the transitions between the monsoon seasons. The subtropical jet stream is identified with a break between the middle tropopause (about 200 hPa) and tropical

tropopause (about 100 hPa). Defant and Taba (1958) described a jet stream having these general characteristics.

147. The Confluence Theory of Jet-Stream Formation. From the relationship between confluence and increased Jet stream velocities, Namias and (1949) derived their “Confluence Theory”. Baroclinic Zones are formed in the atmosphere whenever air masses from different latitudes with different temperatures are brought together. The effectiveness of the advection mechanism depends upon the angle between isotherms and isobars, superimposed effect of vertical motions, and on the other hand on the wind speed itself. All three factors together the rate of propagation of the isotherms. From the turning of wind with height (Barbe, 1956), which usually is different on the warm and cold sides of a jet stream, one may discover the effectiveness of the confluence and difluence mechanism (Dahler, 1957, 1960).

148. Tropical Easterly Jet. Easterlies may be found in the upper troposphere all around the hemisphere, but they do not seem to reach Jet Stream intensity. However, over the Indian sub-continent they do reach Jet Stream intensity. Over Central America, only a weak maximum appears with average velocity of slightly more than 10 Kt. Further the Easterly Jet reaches considerable intensity in these regions only during summer. During winter only, the indication of two Easterly wind maxima one over Indonesia and another over Central Asia with 20 Kt speed each are found in the mean velocity distribution.

149. Origin. With the seasonally confined existence of the Easterly Jet Stream, it is interesting to see from where the Easterly Jet Stream draws its momentum. It is seen that during the months of summer monsoon, a quasi-stationary anticyclone establishes itself in the upper troposphere over Asiatic Plateau. On the Eastern slope of this anticyclone air masses are flowing towards the South carrying along the angular momentum.

150. The entrance area of the Indian maximum of Easterly Jet Stream coincides with the region where exactly the transport of momentum takes place i.e. South China region.

151. The second jet maximum which occurs over West Africa draws its energy from an anticyclone whose centre lies over Sahara Desert and which is separated by a trough of low over the Asiatic anticyclone. Easterly Jet Stream with strongest winds at 100 to 150 hPa, is well developed, especially South of 250 N. The thermal field associated with the belt of warm air which extends through the whole troposphere, favours a strong Easterly current in the upper levels.

152. Over the Indian region the jet axis is near 15° N. On travelling from 25° N to 10° N, due to conservation of angular momentum, an increase in Easterly velocity by about 100 Kt is experienced.

153. Koteswaram (1958) studied Tropical Easterly Jet (TEJ) over India in detail and came to the following conclusions about its origin:-

(a) The easterly current of TEJ forms off the East Coast of China and extends through India and Arabia, at least as far as India.

(b) The centre of easterly current lies roughly along 15° N latitude.

(c) The easterly current accelerates upto 80°E longitude, greatest strength is attained over South Indian Peninsula and then decelerates as it propagates Westward over to Africa.

(d) The core of easterlies is at higher altitude in the Eastern portion (in the East of 75° E) than in the Western half.

(e) Upper level divergence is regarded as favourable in the right entrance and left exit region of wind maxima.

(f) On the anticyclonic side constancy of angular momentum is observed. The pulsation of SW monsoon activity is related with the shift of TEJ.

(g) TEJ is seen to strengthen during weak or break monsoon and weaken during strong monsoon.

(h) It also helps in formation of monsoon depression.

154. In the upper troposphere the circulation over Asia and Africa is dominated by a ridge between 20° N and 40° N, which occupies a more Northern position over Asia than over Africa. A number of anticyclonic cells are noticed, and a permanent anticyclone is situated over the Tibetan Plateau. The anticyclone exists at levels above 500 hPa over the Tibetan Plateau whereas over the deserts of Africa, Arabia and further East, they descend into the lower troposphere i.e. 700 hPa.

155. Over the Atlantic and Pacific Oceans, the upper tropospheric flow breaks up into a number of cyclonic and anticyclonic vortices as pointed out by Riehl (1948). To the South of the high level ridge over Asia and Africa the Easterly flow concentrates into a Jet Stream which extends from the East Coast of China to West Coast of North Africa. Its core is located at an altitude of 150 to 100 hPa with a mean position of 15° N over South Asia. Over Africa it occupies a more Southerly location at 10° N latitude. Normally the Jet Stream over Asia is in an accelerating stage from China Sea to South India and decelerates thereafter.

156. During the South West monsoon season we have SW’ly winds at the surface over the Indian subcontinent. At higher levels (300 hPa) the winds are seen to be Easterlies. Thus there is a complete reversal of wind direction from Westerlies at lower levels to Easterlies as we go up in the atmosphere. In fact it is noticed that at 300 hPa most parts of Indian subcontinent come under the spell of an extensive anticyclone which lies over the monsoon wind system at lower levels. The height at which the wind reverses direction is between 500 and 400 hPa (6-7 km).

157. A feature of considerable interest is the existence of a belt of strong Easterlies along the Southern periphery of the upper anticyclone. The narrow belt of strong Easterlies is observed between 200 to 100 hPa. These Easterly winds, which

often record speeds exceeding 100 Kt, are known as Easterly Jet Stream of the tropic. The core of the Easterly Jet Stream is located about 150 hPa. The discovery of a Jet Stream blowing over Southern India and Gulf of Aden during the monsoon months was inferred in 1952.

158. Theory of Formation of TEJ. The Jet Stream lies directly above the strongest meridional temperature gradient. The Chicago group (1948) has shown that above the level of highest winds the temperature gradient should be reversed. This can be seen from the simplified wind equation

(UH-Uo)/H = g r/ fTWhere,

UH Zonal wind at height H.Uo Zonal wind at sea level.r Mean tropospheric temp gradient Northward at the latitude below level of H.T Mean tropospheric temp at the latitude below the level of H.f Coriolis parameter.H Height of tropopause.g Acceleration due to gravity.

159. It can be seen that a large increase in UH will lead to a sharp increase in H or r or both. There must be an abrupt increase in the height of the tropopause at the Jet and a marked increase in the horizontal gradient below the jet, which is in agreement with the observations, UH goes on increasing as long as r is positive. There should be a reversal of temperature gradient above the level of jet i.e. r is negative.

160. As it has already been said that on an average the strongest zonal wind components are found between high pressure ridge and inflection point upstream from it i.e. the region of increasing anticyclonic curvature. Strong Jet maxima should be expected along the NW slopes of subtropical high pressure cells, where warm air and Westerly angular momentum are carved out of the trade wind region towards North. On the South-eastern slopes of these cells, where the subtropical belt of high pressure is interrupted by a low pressure, which reaches far into tropics imparting an Easterly momentum.

161. Jerkinson (1955) had constructed charts of upper level winds purely based on the surface geostrophic winds with the addition of thermal vector winds in succession of layers based on the upper air temperatures. These charts revealed the existence of strong Easterlies over SE Asia South of 200 N indicating thereby that the strong Easterlies are thermally generated.

162. Koteswaram (1958) has suggested that Easterly jet results from the thermal effects of the elevated Tibetan Plateau on the general circulation in summer. The Plateau is an ellipsoid, 300 km by 1500 km, with an average elevation of over 5 km, individual mountain ranges extending to 6 to 8 km. It receives copious solar radiation in summer. The re-radiation from this elevated terrain goes directly to heat the middle troposphere and sets up a strong solenoidal field in the upper troposphere for driving a clockwise circulation. The resulting upper anticyclone is of thermal origin. The monsoon trough to the south of the Himalayas helps in augmenting the circulation by liberation of condensation heat. The equatorward branch of the upper

outflow from this system gains Easterly angular momentum. We thus observe two jet streams, a Westerly jet to the North of their anticyclone and an Easterly jet to the South (Fig.18). Constant angular momentum has been found in the equatorward branch of the outflow. A cyclonic vorticity is believed to exist South of jet, which is seen due to high level convergence and subsidence of air due to radiational cooling as well as descent of stratospheric air. For mass compensation a reverse circulation is setup on the lower troposphere and the low level poleward flow gains momentum.

163. The SW monsoon thus appears to be a return current on the lower levels of the meridional circulation system. It is convergent and picks up moisture from warm sea areas over which it passes. During its ascent in the monsoon at the rim of the Tibetan Plateau, latent heat is liberated and goes to augment the heat source over the plateau. The atmospheric circulation is thus completed. During autumn, the inflow of cold polar air into the subtropics quenches the heat over the Tibetan Plateau, the circulation ceases and SW monsoon withdraws from the region.

164. Raghavan (1972) has computed the Easterly components at levels 200,150,100 and 80 hPa for Trivandrum and meridional temperature gradient at 150 hPa over the place; he computed the temp gradient of 5 years. From the remarkable agreement between the computed and observed levels of vertical shear it seems certain that the temp distribution mainly controls the tropospheric winds.

Main Characteristics of TEJ

165. Level of Maximum Wind. It is an accepted fact that the height of the maximum wind varies with latitude; however, no correlation could be worked out. Mokashi (1970) has observed that the height of level of maximum wind from 13.5 km at Gananganagar increases to 14.2 km at Colombo, to 14.5 km over Trivandrum and to 15 km over Chennai. In essence, the level of maximum wind slopes upwards as we move Northwards from Ganganagar to Chennai. This feature seems to persist without variation all through the summer monsoon season.

166. It also appears that the stronger the Easterly Jet Stream, higher is the altitude of incidence of maximum wind. Here again Mokashi has brought out that the level of maximum wind in the speed ranges 60-69,70-79,100-129 Kts has frequency maxima near 14, 15 and 16 kms respectively. It is interesting to note here that this is just opposite in the case of Westerly Jet Stream.

167. Joseph (1972) has found that in general, the level of maximum wind increases with latitude from 13-15 km near Trivandrum to about 16-17 km at about 15-18° N. But towards further North it lowers down to as low to a height as 13-14 km, near about 22- 25° N. Thiruvengadathan (1970) has also brought out tendency for the height of maximum wind to increase gradually from Trivandrum to Visakhapatnam and then decrease further North. This converse shape of the level of maximum wind bet 8° N to roughly 22° N.

168. Strength of Maximum Wind. Thiruvengadathan (1970) in his analysis of maximum wind has brought out that the stations South of Nagpur fall into three categories in respect of the maximum winds occurring over them. Minicoy, Trivandrum and Chennai fall in the first group which has the strongest speeds in all

months. Visakhapatnam and Mumbai fall in the next group, which has intermediate values and Nagpur in the third that has the smallest speed values. Generally it can be seen that speeds are more over Southern regions than over Northern regions.

169. July is the month of maximum speeds and minimum in September over all places. From July to August the maximum speed values fall rather gradually whereas between August and September the decrease is much more rapid. According to Mokashi in June the highest speeds are over Trivandrum with a mean of 67.5 Kts at 12 GMT. Over Colombo and Chennai the mean are 66 and 58 Kts respectively. During July between Colombo and Chennai there is hardly any difference; the mean speeds being 64 and 65 Kts over these places. However, the trend for the axis of TEJ is to be closer to Trivandrum during July and between Trivandrum and Chennai region during August; it’s maximum position with 69 Kts as compared to Chennai and Colombo whose speeds are 64 Kts. In September Colombo is the seat of the axis of the Jet, being the only station with mean wind exceeding 60 Kts. It is interesting to note that maximum wind was noticed at Minicoy for more than 60 % of the occasions, North of Chennai this reduces to 23%.

170. Mokashi’s study (1970) has given a mean picture of the vector winds at the levels of maximum wind between June and September and found that the jet core lies between Trivandrum and Colombo.

171. Vertical Wind Shear. One of the characteristics of the jet is that the vertical wind shear should be greater than or equal to 5 meter/sec per km. Thiruvengadathan (1970) has worked out the vertical wind shear for 3 years from 1967 to 1969 for the monsoon months above and below the layer of maximum wind. The percentage frequency of occurrence of vertical wind shear below the layer of maximum wind is 5 meter/sec per km is around 90 to 100% over Minicoy, Trivandrum and Chennai. This decreases as we move Northwards over Visakhapatnam, Mumbai and Nagpur, the least being at Nagpur. The greater shear values are frequent over Minicoy, Trivandrum and Chennai in that order. It is noticed that shears above the layer of maximum winds are generally larger than those below the layer of maximum wind. The median value of the shear above the Jet axis is one and a half times that below the jet. Other characteristics regarding vertical shears are similar to both above and below the jet level.

172. Mokashi (1971) showed that vertical wind shears are of the order of 10 Kt per km, the shears above the level of the maximum wind are slightly higher than below the level of the maximum wind.

173. Horizontal Wind Shear. Thiruvengadathan (1970) has compared the horizontal shears of zonal winds between adjacent stations at the levels 300,200,150 and 100 hPa. It is to be noted here that a cyclonic shear between two stations implies that the stations to the North has stronger Easterlies compared to those to the South and anticyclonic shear implies that the stations to the South has stronger Easterlies as compared to those to the North.

174. At 150 hPa, cyclonic and anticyclonic shears occur with nearly the same frequency on Colombo — Trivandrum sector, cyclonic shear being slightly larger in July. This means that the strongest winds at 150 hPa occur in Colombo—Trivandrum

sector and it is nearer to Trivandrum in July. In other sectors, anticyclonic shears are generally prominent. However, in July in Trivandrum - Chennai sector there is only a very slight predominance of anticyclonic shear.

175. At 100 hPa cyclonic shears predominate in Colombo - Trivandrum and Trivandrum - Chennai sector while anticyclonic shears prevail generally in other sectors. This only means that at 100 hPa, the strongest winds are close to Chennai. Since the level of maximum wind slopes upwards as we go northwards, the zone of cyclonic shear should also move upwards as one goes Northwards from Trivandrum to Chennai.

176. Mokashi (1971) found out that the horizontal wind shears are very weak, being of the order of 10 Kts per 500 km. Anticyclonic shear is to the North of the wind maximum.

177. Steadiness. Easterly Jet stream exhibits a remarkable steadiness at and near the jet levels. The steadiness factor is defined by the ratio of mean vector wind to mean scalar wind. This factor ranges from 95 to 100 % between the levels 13.5 to 16 km. Even at 17 km the steadiness factor is about 90%. However, at 18.3 km it is as low as 6%, thus exhibiting a sharp decline with the height.

Various Studies on Tropical Easterly Jet

178. Mohan (1963) examined the thermal and dynamical characteristics of TEJ during July and Aug from 1956 to 1962. He studied the average circulation in the entrance and exit regions and found evidence of direct solenoidal circulation in the entrance region (East of 90°E) and indirect solenoidal cell in the exit region.

179. Ramamurthi and Swaminathan (1968) argued that if North-South temperature gradient results in TEJ, then TEJ should be stronger in active monsoon condition than in weak monsoon. But it is not so. Wind anomaly charts brought out the reason for the difference in strength of TEJ. It can be seen from the charts that: -

(a) A cyclonic perturbation characteristic of weak monsoon is affected by flow patterns at all levels. The cyclonic perturbations of strong monsoon contributed to Westerly component over South India. This resulted in a weakening of a basic current over that area.

(b) The centre of cyclonic wind anomaly associated with the perturbations shifted Southwards across Visakhapatnam between 200 and 150 hPa levels. Therefore there was a marked strengthening of the Easterly wind over that station at 150 hPa. It is also seen that south of 1 5°N there was a reversal of temperature gradient between 150 and 100 hPa. Hence the maximum wind occurred over Chennai and Trivandrum below 100 hPa.

(c) During weak monsoon, the centre of the anticyclonic wind anomaly lay over West Central Bay and shifted North-westwards above 250 hPa. Therefore, the effect of anticyclonic perturbation was the addition of an Easterly component to the basic Easterly current over South India at 150 hPa

and 100 hPa. Hence, the Easterly jet becomes strong when monsoon was generally weak over India.

180. Thiruvengadanathan (1972) found that the median value of layer of maximum wind is about 1 km over Nagpur and 1.5 km over other stations. During the other months it lies between 1.5 and 2.0 kms except over Nagpur in September where it is about 1km latitude.

181. A study carried out by Bhaskar Rao, Gopala and Saxena (1974) based on the Canberra aircraft data of Exercise ‘Storm Exchange’, showed the following: -

(a) The Easterly flow in the tropics is more variable than the Westerly flow of the middle latitude. The Easterly jet sometimes appears as a narrow core of fast winds and at other times as a broad stream covering most of Peninsular India. It is highly variable in speed.

(b) The Easterly winds are seen to have a speed of 10 to 15 Kt at 300 hPa. Easterlies with strength 80 to 100 Kt prevailed over most of Peninsular India.

(c) It was found that temperatures at all levels near Bangalore were higher than the corresponding levels at Trivandrum. The temperature gradient between Bangalore and Trivandrum was maximum at Flight Level (FL) 350 and FL 470.

(d) Small-scale variations are present in the temperature field at all levels in the tropospheric easterly over Peninsular India during Southwest monsoon season. They, however, do not seem to contribute to the vertical stratification of the atmosphere as per hydrostatic assumption. Therefore the thermal wind and the vertical wind shear between standard oceanic levels calculated over short distances sometimes do not agree and at times are out by one order of magnitude. When the same comparison is made between the thermal wind and the actual wind shear, they show a high degree of agreement, thereby confirming that geostrophic wind can be assumed to hold good in the mean even at lower latitudes, and the small scale variations in the temperature data cancel out when averaged.

182. Mokashi’s study of TEJ over India and Ceylon brings out that the number of ascents with wind speed 100 Kt in the case of Trivandrum is nearly twice that of Chennai. Also the total ascents with winds > 60 Kts in case of Trivandrum is 26% more than that of Chennai. Some of his other findings are as follows:-

(a) Winds over Trivandrum are stronger in June and August, while in July, Trivandrum and Chennai have practically the same maximum wind.

(b) In September the only station with mean wind exceeding 60 Kt is Colombo.

(c) Considering the general distribution of the horizontal shears on the cyclonic and anticyclonic sides of the jet stream, it may be concluded that the axis of the jet stream is likely to be closer to Trivandrum than Chennai.

(d) The meridional component of the Easterly jet decreases with increasing latitude. It has a maximum of 12.5 Kt to 17.7 Kt at Trivandrum and minimum of I to 2 Kt at Chennai.

(e) The highest value of vertical wind shear (19 Kt/km) is observed over Trivandrum at 16 km in August. Next highest of 17 Kt/km is observed over Trivandrum at 15.4 km in June and over Colombo at 15.2 km in September.

(f) The stronger the TEJ, the higher the altitude of maximum wind. The area of maximum wind examined over latitude and altitude suggests a diffuse character of TEJ.

(g) At times in addition to a core of maximum wind over Trivandrum, Minicoy region at 16.2 km another core has been noticed at a higher altitude in the latitude around Goa-Hyderabad-Mumbai. Thus one may suggest the simultaneous existence of two cores of TEJ, but it is yet to be conclusively proved.

183. It is difficult to determine the thermal source that produces and maintains the temperature gradient necessary to cause the tropical jet. The temperature distribution in the upper troposphere and lower stratosphere upto about 20 km is significantly influenced by vertical motion associated with the summer monsoon.

184. Ranjit Singh (1985) found that Easterly wind maxima at 200 hPa level over Bay of Bengal and the Arabian Sea are mostly observed southwest of well organised cloud systems associated with well marked lows or depressions, during Southwest monsoon season. Their formation was explained with the help of conservation of angular momentum during its Southwest motion. Maximum temperature was found to occur over the region of maximum cloud build up at 300, 250 and 200 hPa looking downstream. He said that split in the Easterly jet is observed due to the subsidence warming on the equatorward leg of the meridional monsoon cell in the vertical.

Easterly Jet Stream Studied in Other Parts of Globe

185. Hay (1953) found that high level strong Easterlies with speeds exceeding 60 Kts, at a height of above 40,000 feet prevailed over Singapore and Hong Kong. He found that above 25,000 ft, Easterlies prevail throughout the year. His other findings were as follows:-

(a) Monthly vector means show a gradual increase with height upto 45,000- 50,000 ft.

(b) Monthly vector means reach 35-60 Kt during the months of SW monsoon (Jul-Aug) and the NE monsoon (Dec-Feb). In other months the mean vector winds are smaller.

(c) Monthly vector means show a sharp decrease around 50,000 feet to 55,000 ft. Available evidence indicated that the tropopause height is approximately 55,000 ft.

186. Similar studies carried out at Darwin by AJ Group (1961) showed that: -

(a) During parts of Jan and Feb strong Easterly winds were found to occur at high levels over Darwin. The maximum wind recorded during Jan was 0800/81 Kts at 52,000 ft and during Feb 100/79 Kts at 53,000 ft.

(b) It was seen that the layer of wind in excess of 50 Kts was a comparatively shallow one, usually less than 5,000 ft in depth.

(c) At Darwin the existence of Easterly jets coincided with period of strong low level Westerlies occurring at the peak of the wet season, when the Inter Tropical Convergence Zone (ITCZ) is close.

(d) The jet streams also coincided with the presence of a low pressure trough at 500 hPa, lying WNW to ESE through Darwin.

Low Level Jet

187. There are several places where strong low level currents are observed after winds are averaged on monthly or seasonal basis. Such currents generally do not satisfy WMO recommended definition of speed ≥ 30 m/s, nor are their vertical and horizontal extents comparable to those of planetary scale upper tropospheric jet streams like polar front jet stream or subtropical Westerly jet stream. Still, these low-level strong wind currents are considered as jet streams. These low-level jet streams are generally located in the lowest (1-2 km) of the troposphere. These are strongly influenced by orographic, friction, and diurnal cycle of heating and corresponding variation of pressure gradient and static stability. Due to very limited horizontal and vertical extent of these low level jets and large temporal variability in their structure, comparatively little is known about their climatological structure and dynamical properties. Of all these low level jet streams, relatively more is known about the East African low level jet.

188. Geographical Location Favourable for the Occurrence of LLJ. The following geographical locations are favourable for the occurrence of these low level jet streams or strong air currents:-

(a) Slopes of the mountains parallel to the anticyclonic flow around the subtropical anticyclones; for example, low level jet over West Central USA, along the Peru Coast in South America and along the Namibian Coast in South African continent.

(b) North-South oriented continental coasts near cross-equatorial flow, e.g. East African low level jet.

(c) Narrow Mountain gaps, like Marsabit (North Kenya) Jet stream.

Low Level Wind Shear and Its Effects on Aviation

189. Atmospheric processes vary in terms of their origin, size, intensity and spatial and temporal scales. The phenomena affecting aviation can be classified in to the following:-

(a) Physical motion of air.

(b) Hydrometorological phenomena.

(c) Ice formation.

(d) Poor Visibility.

(e) Atmospheric Electricity

190. The best known, and now highly feared, type of air motion associated with many atmospheric processes is wind shear (Kessler, 1985; Mahapatra et al., 1982). Aircraft accident statistics consistently show numerous entries against wind shear and other wind-related factors. UK accident analysis for 1977 (Civil Aviation Authority, 1978) shows as many as eight accidents (two of which involved commercial aircraft) as being due to ‘un-favourable wind conditions’. US data list many accidents as being due to wind shear and wind-related phenomena such as ‘downdraft’, ‘high wind’, ‘mountain waves’, ‘crosswind’, ‘tail wind’, ‘unfavourable wind’, ‘gusts’, etc. In 1990, these factors together appeared as causes or factors in 16 accidents involving air carriers of various kinds (National Transportation Safety Board, 1993a) and 345 involving general aviation aircraft (National Transportation Safety Board, 1993b).

191. A general definition of wind shear refers to spatial as well as temporal rates of variation of wind speed and/or direction. Thus, a certain amount of wind shear exists at all points in the atmosphere at nearly all times. However, wind shear becomes important from the aviation point of view only when it is of certain types and strengths. The destructive role of wind shear in aviation has been recognised as early as the 1950s (Stewart, 1958; Viemeister, 1961), though the term ‘wind shear’ was not in use in this context at that time. Through subsequent years, as the phenomenon was better understood and implicated in a growing number of air mishaps, its potential for aviation disaster and disruption became clearly evident. The concern of the scientific community at large regarding the wind shear problem in the aviation context is reflected in a specific study conducted by the US National Academy of Sciences on this topic (National Academy of Sciences, 1983).

192. Wind shear is the change in wind speed or direction with height in the atmosphere. It is the spatial and temporal rates of variation of wind speed and or direction. Wind shear arises due to variation in atmospheric air motions. It ranges from small-scale eddies and gustiness to the large scale flow of one air mass layer past another one. Wind shear is always present in the atmosphere and it may be present at all levels of the atmosphere. Its occurrence in the lowest 500 m is of particular importance for aircraft landing and take-off. The types of wind shear are shown in the figure below:-

193. Spatial Classification. All turbulences are associated with wind shear, though on a very small scale. The micro scale wind shear is felt as turbulence by aircraft. Such variations are temporary and in the nature of eddies. Wind shears of large scale (synoptic scale), do not very much affect aircraft. It is the Meso-scale wind shear that poses great problems. The most dangerous Meso-scale shear is the horizontal wind shear. The responses of aircraft to wind shear are extremely complex and depend on many factors including the following:-

(a) Type of Aircraft

(b) Phase of Flight

(c) Terminal area operations

(d) En-route operations

194. Wind Profile in Lower Layers of the Atmosphere. The friction layer can be sub-divided into two:-

(a) The “Surface Boundary Layer” which extends from earth’s surface up to 100 m (330') in which air motion is controlled predominantly by friction with the earth’s surface.

(b) The “Ekman Layer” from about 100 m up to at least 600 m (sometimes 1 km) in which the effect of friction diminishes with increasing height and other controlling factors such as the Coriolis force and horizontal pressure gradient force becomes increasingly important.

195. Generally in the friction layer the wind speed tends to increase with height, with the maximum change occurring immediately above the earth’s surface. The wind direction tends to remain constant with height in boundary layer but it veers with height in the northern hemisphere in the Ekman layer

195. Causes of Wind Shear. The causes of wind shear are:-

(a) Weather fronts.

(b) Sea breeze frontal surface associated shear

(c) Low level jet

(d) Wind shear due to obstacles

(e) Valley wind shear

(f) Wind shear due to inversion

(g) Wind shear due to downbursts

196. Weather Fronts. Significant shear is observed, when the temperature difference across the front is 5 °C or more, and the front moves at 15 kt or faster. Because fronts are three-dimensional phenomena, frontal shear can be observed at any altitude between surface and tropopause, and therefore be seen both horizontally and vertically. As the warm front approaches, wind shear area decreases with decreasing height to ground level. Following the passage of cold front, maximum wind shear area rises from the ground level. Due to slow movement of warm front, wind shear condition existing ahead of warm front can affect the aerodrome longer than those behind the cold front. Shear across the front is proportional to the temperature gradient.

197. Sea breeze Frontal Surface associated Shears. The sea breeze is essentially a shallow cold front because cold air replaces warmer air. Wind shear in the sea breeze occurs predominantly at the surface along the leading edge as the front penetrates inland. Wind shear of lower magnitudes occurs at higher levels.

198. Low Level Jet. When a nocturnal low-level jet forms above the boundary layer ahead of a cold front, significant low level vertical wind shear (LLWS) can develop near the lower portion of the low level jet. This is also known as non-convective wind shear.

199. Wind Shear Due to Obstacles. A combination of strong winds & obstacles to the prevailing wind flow situated upwind of the approach or departure path can create localised area of LLWS. The most commonly encountered wind shear, particularly at smaller aerodromes, is that caused by large buildings in the vicinity of the runway. While the buildings, such as hangar & fuel storage tanks etc., are comparatively low, they present wide & solid barrier to the prevailing surface wind flow. The wind shear is normally very localised, shallow and turbulent. It is of particular concern to light aircraft operating in smaller aerodromes.

200. Narrow Valley Effect. In cases where there are hills along both sides of the runway, the funnelled wind flow may exhibit a venture like effect which results in acceleration in the surface winds.

201. Valley Wind Shear. "Valley wind shear" is a name for another natural cause of wind shear which results from a temperature inversion. This phenomenon begins by the cooling of the air in a valley. This cooling results in a stable air mass

on the valley floor. The wind blowing across the top of the mountain pushes air down the mountain slope. The air experiences adiabatic heating during this descent. However, when this air encounters the stable air mass on the valley floor, it cannot penetrate it and flows over top of it. This results in a layer of warmer air being pushed out over a layer of colder air, a temperature inversion. Thus a wind shear is developed due to the air flowing over the inversion, and stable air below it. This is shown in Figure given below:-

202. As the valley floor air is warmed the next morning, it begins to rise weakly. Air begins moving down the slope of the mountain downwind to replace this rising air. This is shown in Figure A. As this continues and increases in strength, it results in a rotary motion as shown in Figure B. This rotary motion is another form of wind shear.

Figure A Figure B

203. Wind Shear by Turbulence. Another way that mountains can create wind shear is by turbulence. As the wind blows up one side of a mountain and reaches the top, it can begin to mix turbulently. This is shown below in Figure. This turbulence on the lee (downwind) side of the mountain is a form of wind shear. The same effect takes place around the sides of buildings.

204. Wind Shear due to Inversion. When on a clear and calm night, a radiation inversion is formed near the ground; the friction does not affect wind above the inversion top. Change in wind can be 90 degrees in direction and 40 kt in speed. Even a nocturnal low level jet can sometimes be observed. Density difference causes additional problems to aviation.

205. Wind Shear due to Downbursts. The downburst is a downdraft which induces a divergent outburst of damaging winds on or near the ground. The spatial extent of such winds is of great importance for aviation. When an outflow boundary moves away from a thunderstorm due to a shallow layer of rain-cooled air spreading out at ground level, both speed and directional wind shear can result at the leading edge of the three dimensional boundary. The stronger the outflow boundary, the stronger the resultant vertical wind shear.

Microbursts

206. A downburst with its outburst wind zone exceeding 4 km in horizontal dimension is called a macroburst. Smaller downbursts, with damaging winds extending 4 km or less, are called microbursts (Fujita, 1985). More precisely, a microburst may be defined (Wilson et al., 1984) as a divergent outflow for which the differential radial velocity between maxima is 10 m/s or more, and the distance between the maxima is ≤ 4km (yielding a radial divergence of 2.5×10 -3s-1 or more). When the outflow has a radial divergence ≥2.5×10-3s-1, but the distance between the radial velocity maxima exceeds 4 km, the event is classified as a macroburst. A macroburst can produce winds as high as 60 m s-1 (120 knots or 134 mph), with the damaging winds lasting 5-30 min. A microburst may cause winds up to 75 m s -1 (150 knots or 168 mph), but the intense winds often last only 2-5 min.

207. The microburst as a hazardous atmospheric phenomenon is of relatively recent research focus, yet it is recognised as a prime hazard to aviation because of its high wind shear content and its transitory nature. It has been implicated in many aircraft accidents and incidents, most notably the well-studied accidents at New York’s John F. Kennedy International Airport in 1975 (Fujita and Byers, 1977; Fujita and Caracena, 1977), New Orleans in 1982 (Fujita, 1983a; Dietenberger et al.,

1985), and Dallas-Fort Worth in 1985(Fujita, 1986), all in the USA. Posterior analysis, discussed quite graphically by Fujita (1985), has also brought out a clear possibility of microbursts having been responsible for airliner accidents elsewhere in the world, e.g. at Kano, Nigeria, as early as 1956, Pago Pago, American Samoa, in 1974, and Doha, Qatar, in 1976.

208. Microbursts are phenomena of frequent occurrence in many parts of the world, and their spatial and temporal density can be quite high. In a systematic observational programme in the vicinity of Darwin in northern Australia, Potts (1991) reports a daily average of 5 and maximum microburst events within a circle of 40 km radius during a period of 15 days. Given that the influence zones of individual microbursts can be tens of square kilometres, such a high density and frequency of occurrence of microbursts makes it distinctly probable for aircraft to encounter them while operating in the area.

209. Microburst Types. A microburst is difficult to detect because of its small spatial dimensions and short lifetime. A number of studies have been conducted to generate a knowledge base regarding microbursts to facilitate their detection. These include dedicated major co-ordinated projects such as the Northern Illinois Meteorological Research on Downbursts (NIMROD) (Fujita, 1979), the Joint Airport Weather Studies (JAWS) (Fujita and Wakimoto, 1983), as well as numerous other studies (e.g. Eilts and Doviak, 1987). Microbursts do not always coincide with significant rainfall at the ground level. Those which produce appreciable rain on the ground are called wet microbursts; others are termed dry rnicrobursts. A comparison between the two types of microburst is made in Table.7.

Table.7.Wet and Dry Microburst

SlNo.

Characteristics of Dry Microbursts

Characteristics of Wet Microbursts

(a)

Little or no rain reaches the surface. Often associated with virga shafts containing lightly rimed snowflakes that completely evaporate before reaching the surface (i.e. generally small particles)

Associated with heavy rainfall, and precipitation core is mainly in the form of ice, e.g. melting hail (i.e. generally larger particles)

(b)

Strong surface winds caused by negative buoyancy due to evaporation, melting and sublimation of precipitation below cloud base

Strong surface winds caused by precipitation loading in addition to negative buoyancy. Downward momentum transfer and/or dynamically induced pressure gradients may also contribute, especially in strong events

(c) Strong synoptic-scale forcing not necessary

Strong synoptic-scale forcing not necessary

(d) Downdraft entrainment considered minimal

Downdraft entrainment of environmental air at level of minimum equivalent to potential temperature considered important

(e) Dry or nearly dry adiabatic sub cloud lapse rate

Pseudo-adiabatic sub cloud lapse rate early in the day, becoming dry adiabatic by the time of maximum solar heating

(f) Dry in the sub cloud layer Relatively moist in the sub cloud layer

(g) Moist at mid-levels Dry at mid-levels

(h) Relatively weak convection/updrafts Relatively strong convection/updrafts

(ix) Relatively high cloud bases Relatively low cloud bases

(x)Are functions of solar heating. Thus the time of maximum occurrence is mid-afternoon, local time

Are functions of solar heating. Thus the time of maximum occurrence is mid-afternoon, local time

(xi)Exhibit relatively small lowering of surface temperature during the event

Exhibit relatively large lowering of surface temperature during the event

210. Wet microbursts can produce intense rain which may exceed rates of 200 mmh-1 over a few minutes. The microburst blamed for the 1982 airliner crash at New Orleans (Fujita, 1983a) was accompanied by heavy rain, Fujita(1983b) studied a microburst which hit the Andrews Air Force Base airport along with extremely intense rain a few minutes after the US Air Force One, the plane used by the US president, landed on 1 August 1983. Although wet microbursts can be as hazardous as, or possibly more hazardous than, the dry ones, the latter are of greater concern as they do not offer significant visual clues to deter pilots from entering their area. This is corroborated by accident investigations. For example, in 1990, dry microbursts were held as the cause of four general aviation accidents in the USA, while only one accident was attributed to a wet microburst (National Transportation Safety Board, 1993b). As will be discussed later, the absence of precipitation also makes dry microbursts much more difficult to detect using radars.

211. Because of the projects and studies mentioned above, a fairly de knowledge base has been generated regarding microbursts within the US landmass. A good review of the characteristics of microbursts in the US has been made by Wolfson (1988). It has been found that the proportion of different types of microbursts has a geographical variability. For example, the JAWS programme showed that 83% of the microbursts around the city of Denver in the state of Colorado were dry, while the results of the NIMROD programme indicated only 36% of the microbursts in the northern parts of the state of Illinois to be dry (Fujita and McCarthy, 1990).

212. Characteristics. As wind shear is the main danger arising from microbursts, the horizontal and vertical structure of air flow within microbursts is important in the aviation context. The wind structure is also of great importance for designing and optimising sensors for microburst detection. Also of crucial importance from the detection point of view is the evolutionary lifecycle of the microburst, especially the timescales associated with its rise, persistence and decay. Wilson et al. (1984), using Doppler radar data from the JAWS project, determined that the shaft of downdraft air associated with microbursts has a typical diameter of ~1 km, and it begins to spread horizontally at a height that is normally <1 km from the ground. Temporally, from the initiation of divergence on the ground it takes a median

time of 5 min for the microburst to develop its maximum horizontal wind shear, as measured by the differential of wind velocity across the divergent flow field of the microburst. The maximum velocity differential in the study had a median value of 22 m s-1, and occurred over an average distance of 3.1 km. Further, the maximum wind shear was found to occur at a height of ~75m above the ground. However, this height was the lowest height of data collection for horizontal wind speeds; hence the actual height of occurrence of the maximum wind shear could possibly be somewhat less than 75 m.

213. The average and median spatial parameters are shown schematically in a hypothetical microburst in Figure below for visualisation purposes. All these microburst parameters are of great significance for aviation weather, surveillance. However, particular attention is drawn to the fact that the highest horizontal wind shear in a microburst occurs within the lowest 100 m of height This means that an aircraft will encounter maximum microburst wind shear just prior to landing, at heights of only tens of metres above the ground, where it is most susceptible to disturbances. Further, the occurrence of high wind speeds so close to the ground creates a thin layer of high vertical wind shear (i.e. vertical rate of change of horizontal wind speed) as the horizontal wind speed decreases to zero at the ground level due to friction with the ground. The drop in horizontal wind speed above the height of maximum differential also causes vertical shear, but its magnitude is lower than that in the layer below the maximum. Finally, the detection of winds at such low levels requires very special considerations regarding radar location, scanning strategy, sensitivity and clutter filtering for their detection within the time frame of their rapid evolution. These aspects will be considered later in the appropriate context.

214. In another detailed study, also based on JAWS data, Hjelmfelt (1988) arrived at similar microburst parameters. He noted that microburst outflows have depths varying from 300 to 1200 metres. In addition, he observed that the average time from the microburst reaching its maximum horizontal wind shear to its decay is ~8 min. Together with the 5 min build-up time of the microburst mentioned above, this gives an average total lifetime of 13 min for the JAWS microbursts. In this study, downdraft diameters were in the range of 1.5 to 3km, and the maximum downdraft speeds varied in the range of 6 to 22 ms-1, another important observation made from this study was that the outflow morphology of the microbursts was independent of their

associated precipitation rates. Thus the rainfall intensity, as observed by rain-gauges or many current and older generations of weather radars, cannot offer any significant clue to the occurrence of microbursts. Indeed, the study noted that some of the strongest microbursts (maximum differential velocity >25 ms-1) occurred with very low radar reflectivities (<0 dBZ1). Similar lack of correlation has also been reported by Wilson et al. (1984).

215. The low altitude of occurrence of dangerous wind shear due to microbursts is also borne out from observations and investigations carried out in connection with aircraft accidents and incidents caused by microbursts. In the case of the Dallas-Fort Worth accident (National Transportation Safety Board, 1986) the Lockheed L-1011 aircraft involved entered the heavy shear zone of a microburst at a height of ~750 ft (~250 m) from the ground, from which it could not recover. In an incident at the Denver Stapleton Airport, a Boeing 737 aircraft encountered a microburst during the final approach to land, and was driven down to a height <100 ft (~30 m) above the ground by the wind shear before managing to recover (Schlickenmaier, 1989). Another incident, at Atlanta Hartsfield airport, involved a Boeing 767 aircraft that penetrated a microburst just prior to landing and descended uncontrollably down to as low as ~70 ft (~20 m) above ground, short of the runway threshold, before it could recover and execute a missed approach (Lewis et al., 1994).

216. Some studies have also been conducted on microbursts outside the USA. Two important parameters of microbursts observed in northern Australia (Potts, 1991) are presented in Figure below. The peak velocity differential majority of a vast majority of the 76 microbursts detected over 15 days lay between 10 and 20 m/s, with the maximum reaching 27 m/s and the median value being 17 m/s. This median is less than the corresponding value of 22 m/s from JAWS as mentioned above, but the author ascribes the difference largely to the difference in the range resolution of the observing radars. The lifetime of the phenomenon (the interval over which the radial divergence ≥ 2.5 × 10-3 s-1) was found to vary from 5 to 55 min (mean value 15 min), with the shorter lifetimes being more probable. The results show that even in the tropical regions the microbursts have major characteristics similar to those studied in the USA under the JAWS programme. Other important observations during the study were that all the microburst events were associated with thunderstorms, coincided with reflectivity maxima (which correspond to the zones of maximum rain, which were of moderate to high intensity, and occurred during the afternoon or evening, peaking between 1500 and 1700 h.

Figure illustrating parametric distribution of 76 microburst events observed over a 15-day period in February 1989 within 40 km of a 5.33 cm Doppler weather radar located near

Darwin in northern Australia: (a) peak velocity differential and (b) Lifetime

217. Asymmetry. A characteristic of considerable importance in the description and detection of a microburst is its asymmetry. The asymmetry of divergent fields was briefly discussed in the preceding section. Microbursts very often display an asymmetry, with radial outflows being stronger and more spread out in certain directions from the microburst centre (the point at which the axis of the downdraft shaft encounters the ground) than in others. Accordingly, the strength asymmetry of a microburst is the ratio of its maximum to its minimum strength (‘strength’ is the highest differential velocity) over all aspect angles or viewing directions, and the shape asymmetry is the ratio of the longest to the shortest spatial extent of the outflow field over all such directions (Hallowell, 1990). Hjelmfelt (1988) obtained an average strength asymmetry of ~2, and shape asymmetry values of the same order. The study by Wilson et al. (1984) deduced strength asymmetry ratios up to ~6, with an average of~3, over distances of 3km, which is a spatial scale of crucial importance when considering wind shear effects on aircraft. Hallowell (1993) analysed a large number (859) of microburst observations and obtained generally lower values of asymmetry (range 1.0 to 3.0, median value 1.34) after removing the effects of radar configuration used for data gathering, temporal difference between radar observations, and residual errors of apparent asymmetry. He also found that the difference in microburst asymmetry between widely separated geographical areas (within the US) such as Orlando, FL, and Denver, CO, is minimal. The extent of asymmetry has a strong bearing on the automatic detection of microbursts and their hazard estimation based on data from single radar installations, which is the normal mode of data generation and processing.

Gust Front

218. The cold air in the downdraft of a thunderstorm spreads on the ground in all directions. The leading edge of the spreading cold air is accompanied by a large increase of wind speed (squall). The leading edge of the cold air is called a Gust Front. Charba (1974) studied the low level wind and thermal structure of one intense gust front at Oklahama, USA using data up to 444 metres above ground from an instrumented tower and also data from a surface meso-network. Goff (1976) using the same tower and meso-network studied 20 different gust front cases. The sequences of meteorological events as a gust front passes a station are:-

(a) a rise in atmospheric pressure at the surface.

(b) a change in wind direction.

(c) a sudden increase in wind speed (squall).

(d) a drop in the air temperature at the surface.

219. The following figure below gives the schematic cross section through the gust front of a thunderstorm taken from Droegemeir & Wilhelmson (1987)

Figure shows schematic cross section through the gust front of a thunderstorm. Prior to the arrival of the cold air, the wind begins to shift (WS) and the pressure increases or

jumps (PJ-NH) due to a dynamic deceleration between the cold and warm air masses. The passage of the cold air, often called the temperature break or drop (TD), is accompanied

by a hydrostatic increase in pressure (PJ-H) behind the outflow head. H is some characteristics depth of the gravity current usually taken as the height of the

following flow far behind the head. The main body of the outflow is characterised by pressure and temperature that

are a quasi-hydrostatic balance.

220. Wakimoto (1982) made investigations on the dynamics of the gust front and the different stages in its life cycle. He found that the life cycle is divided into four stages:

(a) The formative stage, when the evaporative cooled down-draft air begins to diverge near the surface.

(b) The early mature stage, when an advancing gust-front forms. It exhibits a roll structure at its leading edge.

(c) The late mature stage.

(d) The dissipation stage, when the gust front is no longer fed by the evaporative cooled downdraft air and the advancing gust-front’s height decreases.

221. The gust front thus continues even after the disappearance of the parent thunderstorm. These four stages are shown in the following figure.

Figure shows four stages of a Gust Front

222. Effect of Wind Shear on Aircraft. Aircraft fly in atmosphere and dependant on it for generation of aerodynamic forces Wind shear may be expressed quantitatively in terms of the temporal and spatial derivatives of the wind velocity vector. In studying the effect of wind shear on aircraft, complications are introduced by the spatial motion of the aircraft, which has the effect of coupling the spatial and temporal variations of the wind velocity vector as experienced by the aircraft. As an example, consider an atmosphere with different but constant horizontal wind velocities at each point, as shown in below Figure A. To all static observers, the wind field would appear steady (as opposed to being gusty or time-varying).

Figure A shows a steady but layered wind field

However, an aircraft moving through these points would experience a time-varying or unsteady wind Figure B and such wind variations would affect its flight path and induce oscillations.

Figure B shows the time-varying wind experienced by an aircraft descending through the layers

223. Since the dynamical equations of aircraft motion are well known (e.g. Etkin, 1982), using appropriate mathematical models of wind shear (e.g. Frost and Camp, 1977; Swolinsky, 1986) it is possible to determine quantitatively the effects of wind fields on aircraft motion (e.g. Brockhaus, 1986; White, 1992). It is useful to express wind shear in terms of its components. One may define the partial derivative of each component of wind speed with respect to each spatial co-ordinate as a type of wind shear. In three-dimensional space, there would be nine such spatial derivatives. It is convenient to set up a Cartesian co-ordinate system with its x- and y-axes horizontal, oriented along chosen directions (often along and perpendicular to the aircraft heading), and the z-axis vertical. The wind components along the x, y and z directions are designated Wx, Wy, and Wz, respectively. Then the wind shear types or components are given by the derivatives:-

224. Each of these shear types or components may be designated by a name. For example, if the x-axis is along the aircraft heading, then Sxx is the forward (along-track) shear of forward wind, Syx is the forward shear of lateral wind; Sxz is the vertical shear of forward wind, etc. It is often necessary to designate the absolute wind shear at a particular location or over a distance without reference to any aircraft flight. In such a case, terms such as ‘horizontal shear of horizontal wind’ (or simply ‘horizontal shear’) and ‘vertical shear of horizontal wind’ (‘vertical shear’) are used. All the types of wind shear listed in eqns. Mentioned

above do not have the same effect on aircraft. Some of them are far more important than others). Since fixed-wing aircraft, especially those used in civil aviation, fly essentially horizontally (their climb and descent angles do not usually exceed about 6° even during takeoff and landing), the derivatives along the forward direction, given by the first three of eqns. are of greater significance for flight. However, since the vertical component of speed is significant during takeoff and landing, the wind shear in the vertical direction is also of importance during these flight phases. The effects of some of the important types of wind shear are discussed in the paragraphs below. The most serious effect on flight can come from horizontal wind shear in the direction of flight (e.g. Shen et al, 1996) since such shear causes rapid variation of the aircraft’s air speed (i.e. speed relative to surrounding air) and the main forces and moments governing or affecting flight vary nearly as the square of the air speed. The effect of horizontal wind shear on a nominally straight and level flight is depicted graphically in Figure A & B below.

Figure A shows the nominal flight path, shown here to be straight and level, is deflected upward due to increasing headwind.

Figure B shows the nominal flight path, shown here to be straight and level, is deflected downward due to decreasing headwind along the flight path.

225. In Figure A, the wind is opposite to the direction of flight (i.e. there is headwind), and its speed increases along the flight path. Here, if the inertial speed and the angle of flight path. Here if the inertial speed and angle of attack of the aircraft remain constant, the lift force would increase with time, resulting in an upward deflection of the trajectory. It is, of course, possible to adaptively correct for

such deviations by varying the aircraft’s angle of attack and/or speed through pilot or autopilot action, but the time lag between the wind disturbance and the corrective measure could still leave a significant uncorrected trajectory deviation. The trajectory aberration is of an opposite nature if the headwind diminishes along the flight path. In such a case the lift decreases as the flight proceeds, resulting in the drooping trajectory shown in Figure B.

226. Similar behaviour is observed if the aircraft experiences a tailwind that becomes stronger along the flight path. Finally, a decreasing tailwind, as in the case of the increasing headwind shown in Figure A below, will cause an upwardly deviated trajectory. The effect of vertical wind shear, i.e. the vertical variation of horizontal wind, is shown in Figure A. A rapid change of wind direction from headwind to tailwind, as the aircraft passes through different layers of air during takeoff (Fig. A ) , results in an equally rapid loss of lift and a corresponding droop in the trajectory relative to the nominal.

Figure shows an aircraft taking off from ambient headwind conditions into a layer of tailwind through a shear layer would suffer a trajectory drop and possibly stall.

227. Similar trajectory perturbations can occur during landing through shear layers. Figure B, shows a situation in which an aircraft flying in an environment of strongheadwind descends through a shear layer into a zone of tailwind existing close to the ground. Here again, the rapid reduction of air speed while flying through the shear layer would result in a loss of lift (or even stall) and consequent steepening of the descent path. The loss of height may be irreversible and catastrophic under certain conditions, e.g. (i) when the wind change is strong and persistent, (ii) when the loss of air speed is strong enough to drive the aircraft to the stall condition, and (iii) if a high level of oscillatory motion is induced by the sudden change in the aircraft’s air speed. Tailwind in general was implicated in 68 aircraft accidents in the USA in 199O (National Transportation Safety Board, 1993a; 1993b), three of these involving commercial air carriers. However, the exact role of the tailwind in these accidents is not mentioned in the statistical data, which probably includes both horizontal and vertical wind shear effects.

Figure shows an aircraft descending from a layer of strong headwind toone of tailwind through a shear layer would experience a steeper than

expected descent leading to a hard landing or crash

228. Steeper-than-expected encounter with the runway in general causes hard landings. In relatively benign cases, only passenger discomfort results. However, in more severe cases of hard landing, damage to the landing gear and other parts of the aircraft can occur. If the ‘undershoot’ (shortfall with respect to the landing threshold) due to the trajectory droop is large, the aircraft may touch the ground before the start of the runway. In such cases, damage to the undercarriage is almost certain, with a high probability of more severe damage to the aircraft and passengers. Kayton (1969) mentions that manually controlled and automatic landing systems were typically designed to cope with a vertical wind shear <2 m/s per 100 ft of altitude change, while shear values as high as about 10 m/s per 100 ft altitude were reported by Kramer (1965). However, modern specifications require automatic landing systems to perform satisfactory touchdown with headwinds to 25 knots (—12 m/s), tailwinds to 10 knots (—5 m/s), crosswinds to 15 knots (—7 m/s), moderate turbulence (see Section 3.3.2), and wind shears of 8 knots (—4 m/s) per 100 ft of height from 200 ft to touchdown (Kayton and F1ied, 1997, Chapter 13). Wind shear involving changes in lateral wind is important during touchdown. During this operation, the aircraft is required to maintain its line of flight along the runway centreline, and unexpected strong changes in lateral wind may force the aircraft to veer off this line. Sudden gusts of lateral wind may also excite the oscillatory roll modes of the aircraft, making its control difficult during touchdown US data show that crosswind was involved as a cause or factor in 115 general aviation accidents (National Transportation Safety Board, 1993b) and three commercial air carrier accidents (National Transportation Safety Board, 1993a) in the year 1990. The wind shear cases mentioned in this section constitute the simple or basic types of idealised wind shear. Actual wind variations along the path of aircraft in real life are usually much more complex, involving more than one significant component. A type of powerful wind shear field that occurs commonly in nature and is of great hazard value for aviation results from outflows of storms. 3’he nature of such fields, which involve strong horizontal divergence as well a vertical air currents and their gradients, will be discussed in some detail in the next chapter. In particular, dangerous levels of wind shear at low altitudes can arise from microbursts, which cause damaging winds over a small patch of ground. Low-altitude wind shear caused by microbursts and other forms of thunderstorm-induced divergence is said to be of convective origin, and that due to other causes is

non-convective low-level wind shear. Although microbursts are the source of some of the strongest wind shear at low altitudes, a low-level wind shear alert by the control tower personnel does not automatically mean the presence of microburst(s) in the area. However, if a microburst is identified, a microburst alert may be issued irrespective of the actual value of the associated wind shear.

229. Operationally, low-level wind shear is included in an Aviation Terminal Forecast if there are pilot reports of wind shear that causes air speed gain or loss of 20 knots (~10 m/s) or more within 2000 ft of the surface, vertical shear of 10 knots (~ 5 m/s) or more per 100 ft within 2000 ft of the surface, or if meteorological conditions for wind shear are satisfied (Jackson, 1991).

230. The effect of wind shear fields on aircraft performance is quantified through a wind shear hazard index called the F-factor. In a simple case of planar wind shear where the wind variation is in a vertical plane and the aircraft also flies in the same plane, the F-factor is given as

Where = time-derivative of the horizontal wind component along the aeroplane ground track

= vertical component of wind, positive downward Va= air speed of the aircraft g = acceleration due to gravityIn a more general wind shear the F – factor takes the form

Where W is the absolute (inertial) wind vector, Va is the air speed vector, and (Va/V) is the unit vector along the air speed direction. It is possible to measure the along-

track shear component directly with forward-looking sensors mounted on the aircraft itself or the F-factor may be derived from the absolute wind field data obtained from ground-based observations. It is clear from the above two eqns that a positive value of F is produced by a performance decreasing wind shear such as a decreasing headwind or a downdraft or a combination of both. Conversely, a performance-enhancing shear such as an increasing headwind and/or updraft would yield a negative F- factor.) Physically, the F-factor refers to the rate at which the wind field changesthe energy of the aircraft. The stronger this change the more pronounced would be the effect of the shear field on the aircraft. While moving through a wind shear with a positive F-factor, an aircraft would have to generate an excess thrust (over drag) equal to the F-factor times its own weight to maintain level flight at constant speed. If the shear is so strong that positive F-factor exceeds the specific (ie. per unit weight) excess thrust capability of the aircraft, then the aircraft will suffer some loss of air speed and/or altitude in spite of the best possible pilot input. To form an idea of the numbers involved, the maximum specific thrust capability of commercial transport aircraft typically lies in the range

from 0.11 to 0.17 depending on the number of engines (Lewis ci at., 1994). If the wind shear field produces a negative F-factor, then the aircraft, to maintain its speed and altitude, must cut down its thrust to be lower than the drag by an amount equalling the F-factor times its own weight. In another, or perhaps equivalent, manner of looking, the F-factor balances with the climb capability of an aircraft. If an aircraft climbing at an angle of ө radians in still air at constant speed a wind shear with a positive F-factor equal to ө then it will cease to climb and fly level if it holds its thrust and speed unchanged. Similarly, if it encounters a wind shear with F-factor equal to - ө while descending at an angle ө in still air, its path will be restored to level flight if its speed and thrust are held constant. The F-factor as defined is a point function and thus can be defined and possibly measured at each point along the flight path of an aircraft. However, since strong wind shear in nature is usually accompanied by significant turbulence, the instantaneous values of F-factor would contain a corresponding amount of rapidly varying random components. These random variations must be filtered out over an appropriate path length in order for the F-factor to be useful for characterising the wind shear. Analysis of flight data from the Lockheed L-1011 aircraft, which suffered a fatal accident (National Transportation Safety Board, 1986) due to severe wind shear encounter while attempting to land at the Dallas—Fort Worth airport in Texas, USA, has shown that the F-factor, averaged over a 1 km length, was as high as ~ 0.30 (Lewis et al, 1994). Comparison of this magnitude of the F-factor with the maximum specific thrust capability of commercial aircraft indicated above would make it appear that the accident was inevitable if the aircraft got into the wind shear without prior warning and evasive action. On the other hand, with sufficient warning, it is possible to escape strong shear fields, as proven by the aircraft that followed the ill-fated L-1O11 aircraft in the landing sequence at the Dallas—Fort Worth airport (National Transportation Safety Board, 1986; Fujita, 1986). Visser (1997) discusses optimal guidance strategies for escaping from microburst wind shear encounters by aircraft during final approach. The technique relies on in situ (reactive) or short-range forward-look (predictive) wind shear detection.

231. Low Level Wind Shear Alert Systems (LLWAS) . It has been repeatedly pointed out that wind shear in the layers of the atmosphere constitutes perhaps the most severe and frequent source of hazard for aviation operations. Thus, even in the absence comprehensive surveillance mechanism to monitor the entire weather picture affecting the aviation system, a simpler observation system dedicated to the detection of low-altitude wind shear should greatly enhance aviation safety by minimising aircraft encounters with such shear during landing and off. This was the motivation for the development of the low-level wind shear alert system (LLWAS) before the modern generation of weather radars became available.

232. The LLWAS was conceived as a specific technological response to the weather-induced air disasters of the 1970s such as the highly visible and well investigated Eastern Airlines accident (National Transportation Safety 1976) at John F. Kennedy Airport, New York, in 1975. Many of these accidents were attributed to low-altitude wind shear, and had attracted the particular attention of the US National Academy of Sciences (National Academic Sciences, 1983).

233. Concept and Basic Configuration. The concept and design of the LLWAS system are based on the premise that phenomena which generate wind

shear in the lowest layers of the atmosphere must manifest themselves through detectable ground-level signatures. V proper sensing, processing and interpretation of such signatures should lead to conclusions regarding the existence or otherwise of hazardous levels wind shear at low altitudes.

234. The LLWAS is an instrument system based on anemometers which are insitu wind sensors. Anemometers are usually either of the vane type or cup type. The former, a wind vane with its horizontal axis aligned along the wind’ direction is driven by the local wind and rotates at a rate which is a monotonic function of the wind speed. A reading of the rotational speed of the vane therefore serves to indicate the local wind speed. The latter type anemometer has a set of cups, usually three in number, mounted cyclically with their individual axes horizontal, on a rotor with a vertical axis. Because of such an arrangement, the rotor is insensitive to the horizontal wind direction and turns at a rate dependent on the wind speed. With either type the capability to sense wind direction can be incorporated by adding weathercock mechanism. Modern anemometers have built-in transducers to convert their readings into electrical impulses which can be transmitted over wires or wireless links for remote monitoring.

235. Each weathercock—anemometer combination measures the horizontal wind vector only at one point. To sense the existence of wind shear over linear path and estimate its intensity, a minimum of two anemometers may be used one at each end of the path segment, and their difference taken. For more reliable and detailed information regarding the wind distribution along the path, additional sensors may be placed along the path. For sensing wind shear over an area, the area may be appropriately circumscribed with a number of vector anemometers, a comparison of their readings would serve to indicate the strength of wind shear across the area. A denser coverage of the area with anemometers, including some located in its interior, would enhance the reliability of the wind shear information and provide a more detailed picture of the wind field.

236. LLWAS is based on this principle, and is optimised for aviation support. It utilises an anemometer array to cover the most sensitive parts of airport areas from the point of view of low-altitude wind shear. These include the runway complex and the adjoining areas over which flights occur at very low altitudes. In the original version of the LLWAS (Goff, 1980), the array consists of one anemometer positioned close to the centre of the runway complex and five more at outlying spots located, where possible, close to the approach paths of individual runways. Each anemometer is mounted on a mast at a height ranging between 4 and 20 m from the ground depending on the local air flow quality (velocity perturbations, turbulence, etc.) and on obstruction considerations. A central processor monitors the individual anemometer readings received through a radio link at 10 s intervals. The data from the centre-field anemometer are averaged over a 2 mm period on a moving-window basis. The instantaneous readings of each outfield anemometer are compared with this averaged centre-field wind, and a wind shear advisory is issued if any of the outlying anemometers shows a difference >15 knots relative to the centre-field wind.

237. The LLWAS was originally intended to be an interim solution to the low altitude wind shear warning problem, to be eventually replaced by more elaborate

systems based on Doppler weather radars. Specifically, the initial design of the LLWAS was tailored to satisfy the requirements of detecting the wind shear associated with gust fronts and other frontal phenomena not originating from thunderstorms (US Department of Transportation, 1989b). the installation of LLWAS units started in 1977—1978, and over a hundred airports in the US have been equipped with the system. Owing to the success of the system, the LLWAS is no longer regarded as a temporary solution, but has been incorporated as an element of the 1987 Integrated Windshear program Plan of the US Federal Aviation Administration (US Department of transportation, 1987). Because of the simplicity and cost-effectiveness of the system, other countries have experimented with, designed and/or considered and or considered installation of the system or its variants.

238. Enhanced System. Although the original LLWAS does provide useful indication and warning of low altitude wind shear, this basic configuration has its own limitations (Kessler, 1990; National Transportation Safety Board, 1983). These limitations arise essentially from the small number of sensors in the basic system as well as their location. Indeed the effects of the two factors are interrelated. Because of the small number of sensors in the original LLWAS configuration, only a limited area can be covered. Thus, the outfield anemometers are located so as to cover only a short terminal segment of the aircraft glideslope. Wind shear zones outside this coverage are not sensed by the LLWAS array. This situation can be remedied to an extent by enlarging the array coverage by locating the outfield anemometers farther out along the aircraft approach paths. But it must be remembered that intense low-altitude wind shear such as that due to microbursts and gust fronts can be quite localised. Thus spreading out a limited number of sensors over an expanded boundary could lead to spatial undersampling of the wind field, resulting in missing or underestimating localised strong shear occurring within the nominal area coverage.

239. To overcome such limitations, the original LLWAS has been enhanced in many ways. These include (US Department of Transportation, 1989b; Goff and Gramzow, 1989; Jaffe, 1989):-

(a) Network expansion, i.e. increasing the number of wind sensors beyond the original six.

(b) Network and software design enhancements to identify microbursts and gust fronts and to provide runway-oriented wind shear information.

(c) Elimination of site effects on wind sensors.

240. The functional data flow in an LLWAS-NE (i.e. LLWAS with network expansion) is shown in Figure below. Two enhanced ‘test bed’ systems installed at New Orleans and Denver airports in the USA in 1984 and 1985 respectively. An enhanced array of 12 wind sensors tested at Denver, a with a wind shear and microburst detection algorithm, was found to performing appreciably better than the conventional method of merely comparing the outer sensor readings with the averaged centre-field sensor reading (Smythe, 1989). With these improvements, the scope of performance of the LLWAS has been greatly enhanced. Against its original expectation of providing only general indication of the existence of low-altitude wind

shear in the runway area and its close environs, essentially due to gust fronts, the enhanced versions of LLWAS can perform detection and recognition of major shear phenomena such as microbursts and gust fronts, and provide more quantitative estimation of the wind shear along the runway direction, which is the main parameter of interest for decisions regarding landing and takeoffs. With such improved performance, the LLWAS has emerged as a viable stand -alone system of great value in aviation safety and efficiency augmentation at airports not provided with more expensive, elaborate and versatile aviation weather surveillance systems, and a very useful supplement to such systems where they exist. The system has been credited with helping avert major accidents in real life (e.g. Hughes, 1990).

Figure shows LLWAS functional data flow diagram(from Nilsen and Starr, courtesy K.M. Starr, TRW Inc.)

241. When an aircraft encounters a wind shear field and senses gain or loss of altitude and/or speed, it must immediately initiate a recovery procedure to escape the damaging effect of the shear. Such recovery procedures are a standard part of many crew training procedures now. However, the proper in flight recovery techniques for a head-on encounter with a microburst and with non convective low-level wind shear are opposite (Jackson, 1991). The wrong choice of recovery procedure can make the difference between a successful escape and a mishap. Thus recognition of the type of wind shear phenomenon ahead of an aircraft is of paramount importance for aviation safety against low-altitude wind shear. The ability of the enhanced LLWAS to recognise the nature of wind shear fields is therefore one of its greatest advantages in promoting aviation safety.

242. Irrespective of the level of its sophistication, the LLWAS does, however, suffer from certain drawbacks which are inherent to ground-based in situ sensors. The chief one among them is that surface winds detected by the sensors may be significantly different from those along the aircraft flight paths during landing and takeoff. For example, the results of one study (Elits,1987) showed that winds at a few hundred feet altitude, even as low as hundred feet (30 m) above the ground,

were on average ~60% higher than at heights corresponding to the location of the LLWAS sensors. The difference between the true and the surface-sensed winds in terms of effects on flight is even higher since wind effects on aircraft vary as the square of the wind speed.

243. Such limitations notwithstanding, the LLWAS is a useful system for enhancing aviation safety. In particular, it can perform a very useful role in conjunction with terminal area Doppler weather radars such as the TDWR. In such a combination, the LLWAS would provide wind data very close to the level, for which heights radar data are often either unavailable or unreliable due to radar horizon limitations and ground clutter contamination.

244. Airborne Wind Shear Detection. All the systems for wind shear detection discussed hitherto in this work are ground based systems. Irrespective of whether they are based on in situ sensors or remote sensors, a common feature of all these systems is that information on wind shear is obtained at a location on the ground. If the information is to be used by the aircraft pilot for being aware of the weather scene and/or taking decisions regarding the flight path and operations, it must be transmitted to the aircraft in raw or processed form. Alternatively reports, advisories or warnings based on the ground-sensed weather data may be transmitted to the pilot.

245. A measure of autonomy can be achieved by the pilot if wind shear could sensed on board the aircraft itself, minimising the need for extensive ground support and communication links. Such an arrangement would have important advantages. First, the time involved in collecting, processing and possibly interpreting and confirming the observations, and in multiple transmissions (e.g. from the sensing system to the air traffic control centre’ and then on to the pilot) would be eliminated. This may add precious moments to the time available for the pilot to be aware of any impending danger due to wind shear and take evasive action or other crucial flight decisions. A second major advantage is that aircraft so equipped could have a high degree of protection from wind shear hazards even in areas and spaces not provided with adequate ground-based shear-warning systems. The quality of weather coverage will be non uniform over large areas (e,g continents) for a considerable period in the future, and may never uniform globally. Aircraft with on-board capability to sense wind shear would be less dependent on ground support, and can minimise wind shear hazard even while conducting flight operations in the less developed parts of the world.

246. Insitu Sensing. Wind shear may be detected on board aircraft either in situ or by remote sensing. In the former, air motion is sensed by measuring its effect on the aircraft trajectory and dynamics (McLean, 1988). Perturbations in aircraft flight paths and body attitudes are readily detected with the help of inertia sensors such as accelerometers and gyroscopes, which are now part of the standard navigational instrumentation on all but the simplest and smallest aircraft. 247. The F-factor was discussed as a measure of wind shear. This factor can be derived in flight by monitoring certain flight parameters such a speed, climb/descent rate, thrust level and angle of attack. The aircraft may be declared to have entered a strong wind shear zone when the sense F-factor exceeds a certain threshold. It may

be recalled from the earlier discussion that most commercial passenger and transport aircraft, which constitute perhaps the most significant segment of aviation activity, would inevitably lose height and/or speed if subjected to wind shear with F-factor values somewhere in the range from 0.11 to 0.17. Hence a threshold of 0.1 to 0.15 may be applied, depending on the type of aircraft, to determine entry into a strong wind shear field.

248. Measuring the local F-factor on board and applying thresholds to judge its severity however, is not without difficulties. Strong wind shear in nature does not appear as a neat streamlined flow field with spatial gradients, but is usually accompanied by strong turbulence or gustiness. The effects of these gusts generally cancel out among themselves and do not affect aircraft trajectory significantly. However, they can induce sharp local spikes in the variation of the F-factor, giving it a noise-like characteristic. Straightforward thresholding of the raw instantaneous F-factor data may then e rise to false alarms and multiple threshold crossings. The high-frequency caused by gustiness must be filtered out before the in situ measured Factor can be meaningfully used for wind shear detection and evasion. Such smoothing filter, called gust filter or turbulence filter would introduce its own time delay into the data chain, which would correspondingly delay the shear detection process. Thus, airborne in situ wind shear sensing not only does not provide any advance warning of approaching shear, but actually delays the data compared to real time. This delay is undesirable, and would have an effect opposite to advance warning, i.e. the aircraft would be deeper into the shear field before the pilot even realises the fact.

249. Experiments in insitu wind shear detection have been conducted by the US national Aeronautics and Space Administration using an instrumented Boeing737 research aircraft to penetrate microbursts. The details of the experiment and the results are discussed by Lewis et al. (1992) and Oseguera (1992). A stand-alone in situ wind shear warning system comprising a warning computer and associated cockpit displays is described by Aeronautical Radio Inc (ARINC, 1988). The system uses data pertaining to aircraft movement with respect to air to detect and annunciate a wind shear condition. Optionally it may also provide instrument guidance to the crew indicating the Optimum pitch to endure the wind shear encounter.

250. Forward-looking Remote Sensing. Insitu detection of wind shear by aircraft in flight is quite straightforward and can provide useful warning to the pilot before maximum levels of shear are encountered. It can also facilitate the incorporation of automatic compensation mechanisms to mitigate the effects of shear. However, it has the drawback that the aircraft would notice the shear only after it actually enters the shear zone. It would be preferable to sense wind disturbances along the flight path in advance of their encounter with the aircraft to provide some warning time to the pilot to take remedial measures. The advantages of such advance detection of wind shear over reactive mitigation of wind shear effects have been discussed by Hinton (1990).

251. Even very short warning times can have dramatic effects on the ability of aircraft to survive adverse wind fields. In a study related to the Integrated Wind-Shear Program undertaken jointly by the US National Aeronautics and Space

Administration, Federal Aviation Administration and industry found that the factor which most strongly affects the ability of recover from microburst-induced wind shear is the time at which the recovery is initiated. Improving the alert time by just 5 s generally provided increase in recovery performance than could be achieved by changing the recovery strategy. Forward-look alerts given 10 s prior to entry1 microburst permitted recoveries to be made with negligible altitude (Hinton, 1992). At an approach speed of 150 knots (-75 m/s), a 5 s time corresponds to a distance of -375 m, and 10 s to -750 m. In study, Hinton and Oseguera (1993) have noted that all classes of transport aircraft (i.e. those with two, three or four engines) can escape ‘worst-case’ microburst located straight ahead if given 20 s of warning ~1.5 km of warning distance. This has been proven in practice by the free escape of the aircraft that followed in the landing queue the Lockheed L-1011, which crashed after encountering a strong microburst at the Dallas Fort Worth airport on 02 Aug 1985Four other aircraft successfully passed through an equal strong microburst in Denver in 1988 and one is known to have performed the same feat in the subsequent year (Hughes, 1990).

252. Warning times of such orders can be obtained readily and autonomously through remote sensing of the wind field ahead by employing forward looking instruments. Airborne remote sensing of wind shear is an area of research. It is a complex topic that involves not only hazard definition sensor selection and optimisation, but also system integration and management (US Department of Transportation, 1989a). Disciplines studies contributing to a better understanding and solution of the problem include aircraft simulation, Meso-scale atmospheric modelling and analysis and instrumented flight tests. Besides providing the much-needed advance warning time, wind detection ahead of aircraft using forward-looking airborne sensors, another advantage. This pertains to gust filtering.

255. Lewis et al (1994) have made a comparison of the performance limit curves for transport aircraft, with F values estimated from a variety of real wind shear events, to conclude that an averaging interval of the order of 1 km is optimum for discerning hazardous shear. Lower averaging intervals will be dominated by Turbulence, which is not a performance threat (in the sense of loss of flight Altitude/speed, though turbulence by itself is a serious hazard in a different way), and significantly higher averaging intervals may result smoothing out hazardous events of small spatial extent, making them look having low F values. In each of the cases they studied, the hazardous wind shear event crossed the appropriate aircraft performance limit curve at scale lengths of ~1 km. This order of scale length provides a basis of filtering or raging the wind field detected by forward-looking wind shear sensors. The threshold for shear warning and activation of evasive manoeuvres would depend on the aircraft performance characteristics in the takeoff and landing figurations and the initial energy state of the aircraft at entry into the shear field. The US Federal Aviation Administration has specified an alert threshold boundary for reactive wind shear detection systems (US Department of Transportation, 1990).

256. Airborne Doppler radar, Lidar and Infra-Red sensors are candidate devices for remote sensing of wind shear ahead of an aircraft from the aircraft itself. Each of these sensors has its own special advantages, but the Doppler radar cores higher on the important point of all-weather wind shear detection. A small low-power Doppler radar may have a maximum detection range of the order of 2 km ahead of the

aircraft, but a more powerful (may ‘see’ up to 10 km or more, providing a longer warning time. Doppler radars also have a minimum range or blind range, which may be as much as a kilometre, within which they would not be able to observe the wind shear.

257. The basic principle of wind determination by airborne Doppler radars is the same as that for ground-based Doppler radars, but there are differences in terms of details. An obvious difference is that the airborne radar unlike its ground-based counterparts, itself moves along with aircraft relative to the surrounding air, and this motion causes a Doppler shift which is superimposed on the Doppler signal due to the absolute motion (i.e. relative to the earth) of air. This component must be removed to retrieve the true atmospheric Doppler signals which provide a picture of the fields ahead. The radar-motion-induced Doppler shift is a function of the aircraft velocity, which is available on board from the aircraft’s navigation system, and the pointing angle of the antenna relative to the aircraft velocity vector, which can be read from the antenna pointing mechanism. Cancellation of the Doppler component due to radar motion is therefore a straightforward procedure that can be carried out in real time, but completeness depends on the accuracy of the velocity and pointing data. Incomplete cancellation would leave a residual component that contaminates the atmospheric velocity estimates. The cancellation procedure may also be aided by comparing the atmospheric Doppler signals with the returns from stationary ground clutter, if the latter is available in sufficient strength.

258. Sensing wind shear ahead of aircraft using airborne Doppler radar has natural advantage compared to performing the same task with ground-based radars. As the airborne radar looks ahead along the flight direction, t, component of wind variation along the path, which is of most concern for the flight, appears to the radar as a radial component. The Doppler radar, which can sense only the radial component of wind, can therefore realistically estimate the severity of the wind shear which the aircraft is likely to encounter ahead of the current position. This is not necessarily true in the case of ground-based radars which would, in general, observe the path ahead of the aircraft from an oblique aspect, and therefore sense a different component of the wind shear than the one along the flight path. Such a situation is depicted in Figure below. The ground-based radar data may, of course, be used in an indirect way by recognising the nature of the shear field (e.g. that due to microburst or gust front), modelling the field based on its observation, and then computing the wind variation along the expected flight path.

Figure shows the difference in the geometry of wind (shear) detection byairborne (a) and ground-based (b) radars. Each measures the componentof wind in its direction, and the components are not the same in general

259. An important limitation of airborne weather radars relative to those based on the ground is with regard to the choice of operating frequency. From size and weight considerations, the antennas of airborne radars are of limited aperture, and therefore utilise higher operating frequencies to achieve a narrow beam width. Frequencies in the X-band (wavelengths of the order of 3cm are commonly used, with those in the C-band (wavelengths of the order 5cm) a possibility. These higher frequencies have their attendant problems enhanced rain attenuation and lower sensitivity for clear-air observation.

260. While operating at higher altitudes, such as during en-route flight, airborne radars have the advantage of negligible ground clutter interference. However, at the higher altitudes, where the refractive index fluctuations due to turbulent mixing are feeble because of lower water vapour concentration the detection of wind shear turbulence in clear air becomes difficult. On the other hand, these phenomena can be very well detected in the presence of precipitation.

261. At altitudes close to the ground level the reflectivity of clear-air phenomena is much better, but the problem of ground-clutter contamination becomes significant. For such operation, Doppler Lidars have an advantage in clear air (as Lidars do not have the ground-clutter problem), but are ineffective in seeing through precipitation and fog. Airborne Doppler radars would perform quite well in light rain, but heavy rain would impair their performance due to absorption of their shorter wavelengths. Sensor choice and operation as well as system configuration for airborne remote sensing of wind, shear under a wide variety of operating conditions is thus a complex task, and no easy solutions exist. The airborne Doppler weather radar has the potential to provide useful wind data in a variety of Lions of interest, and may find wide usage in this role in the future when technology is perfected.

262. Microburst is one of the major aviation weather hazards. Until 1985, on an average one Microburst related aircraft accident used to occur in 18 months. Continued efforts by aviation industry and meteorological community laid to widespread awareness about microburst among aircrew, air traffic controllers and weather forecasters. The knowledge microburst and its detection by high resolution Doppler radar, contributed to significant decline in microburst related aircraft accidents. After 1985, next accident occurred nine years later in 1994 at Charlotte, North Carolina in United States. In the absence of environment conditions favourable for the occurrence of microburst the detection by visual clues can be of great help in avoiding hazardous situations caused by microburst.

Conclusion

263. Knowledge of weather conditions likely to be encountered during a specific flight enables an aviator to plan his flight with a view towards economy of fuel, operational efficiency and flight safety. Therefore, there is a need for accurate prediction of weather phenomena such as thunderstorms, heavy rain, clear air turbulence, icing, low level wind shear, mountain waves, dust storms, squalls, gale, hailstorm, fog etc. which have the potential to adversely affect aviation activities. Most of these weather phenomena are localised and are short lived but they pose a serious hazard to aviation. Much progress has taken place in the field of weather observation and forecasting by radar, satellite, numerical computation etc. Thus, a

need exists for theoreticians to understand operational requirements and for operational forecasters to be brought up to date with the latest advancements. Scientists and meteorologists from the civil and military need to synergise their efforts to improve the reliability and accuracy of forecasting of weather phenomena. This involves intensive training, research and development.

Check Assimilation

Fill in the blanks

1. Severe thunderstorms form in areas with a ________________.

2. MCC covers an area in excess of _______________ square kilometres.

3. The intensity of CAT is generally ____________than the turbulence encountered in thunderstorms.

4. Opaque rime icing may cause ___________________________ of the wing.

5. CAT can be expected in the regions where the vertical wind shear is greater than _______________ .

Key

1. Severe thunderstorms form in areas with a strong vertical wind shear.

2. MCC covers an area in excess of 1,00,000 square kilometres.3. The intensity of CAT is generally less than the turbulence encountered in thunderstorms.

4. Opaque rime icing may cause an alteration in the aerodynamic characteristics of the wing. 5. CAT can be expected in the regions where the vertical wind shear is greater than 4 kt per 300 m

11 BPKC: METEOROLOGYASSIGNMENT: 8/24 (AVIATION METEOROLOGY)

Total Marks: 20

PART I

1. Fill in the blanks (4x1 Marks=4)

(a) Cloudburst is generally due to ___________________ .

(b) Cloudburst is generally due to a rapid condensation of the clouds. CAT has been associated with temperature change of ____________per hour and to the gradient of temperature of over _________° latitude.

(c) Two of the most important factors which characterise the formation of lee waves are static stability and wind profile.

2. State true or False and if false give the correct answer (4x1 Marks=4)

(a) A microburst may be defined as a divergent outflow for which the differential radial velocity between maxima is 20 m/s or more, and the distance between the maxima is < 4km.

(b) A negative value of F is produced by a performance decreasing wind shear such as a decreasing headwind or a downdraft or a combination of both. (c) The STJ is related to the equatorward boundary of Hadley circulation, where low level front tend to be obscure or absent.

(d) Wind shear in the lowest 500 m is of particular importance for aircraft landing and take-off.

PART II

Answer any three of the following questions (3x4 Marks=12)

3. What are the aviation hazards associated with thunderstorms and dust storms?

4. Explain the forecasting methods of CAT.

5. Explain the formation of mountain waves.

6. Explain the mechanism of formation of STJ and Tropical Easterly Jet.

7. Explain the effects of wind shear on aircraft with diagrams.