Land and sea breezes

Another thermally induced wind regime is the land and sea breeze (see Figure 6.11). The vertical expansion of the air column that occurs during daytime heating over the more rapidly heated land (see Chapter 3B.5) tilts the isobaric surfaces downward at the coast, causing onshore winds at the surface and a compensating offshore movement aloft. Typical land-sea pressure differences are of the order of 2 mb. At night, the air over the sea is warmer and the situation is reversed, although this reversal is also the effect of down-slope winds blowing off the land. Figure 6.12 shows that sea breezes can have a decisive effect on temperature and humidity on the coast of California. A basic offshore gradient flow is perturbed during the day by a westerly sea breeze. Initially, the temperature difference between the sea and the coastal mountains of central California sets up a shallow sea breeze, which by midday is 300 m deep. In the early afternoon, a deeper regional-scale circulation between the ocean and the hot interior valleys generates a 1-km deep onshore flow that persists until two to four hours after sunset. Both the shallow and the deeper breeze have maximum speeds of 6 m s-1. A shallow evening land breeze develops by 1900 PST but is indistinguishable from the gradient offshore flow.

The advancing cool sea air may form a distinct line (or front, see Chapter 9D) marked by cumulus cloud development, behind which there is a distinct wind velocity maximum. This often develops in summer, for example, along the Gulf Coast of Texas. On a smaller scale, such features are observed in Britain, particularly along the south and east coasts. The sea breeze has a depth of about 1km, although it thins towards the advancing edge. It may penetrate 50 km or more inland by 21:00 hours. Typical wind speeds in such sea breezes are 4 to 7 ms-1, although these may be greatly increased where a well-marked low-level temperature inversion produces a 'Venturi effect' by constricting and accelerating the flow. The much shallower land breezes are usually weaker, about 2 m s-1. Counter-currents aloft are generally weak and may be obscured by the regional airflow, but studies on the Oregon coast suggest that under certain conditions this upper return flow may be related very closely to the lower sea breeze conditions, even to the extent of mirroring the surges in the latter. In mid-latitudes the Coriolis deflection causes turning of a well-developed onshore sea breeze (clockwise in the

Sea Land Breezes Typical Scales

Figure 6.11 Diurnal land and sea breezes. (A) and (B) Sea breeze circulation and pressure distribution in the early afternoon during anticyclonic weather. (C) and (D) Land breeze circulation and pressure distribution at night during anticyclonic weather.

Figure 6.11 Diurnal land and sea breezes. (A) and (B) Sea breeze circulation and pressure distribution in the early afternoon during anticyclonic weather. (C) and (D) Land breeze circulation and pressure distribution at night during anticyclonic weather.

Sea And Land Breezes

Figure 6.12 The effects of a westerly sea breeze on the California coast on 22 September 1987 on temperature and humidity. (A) Wind direction (DIR) and speed (SPD). (B) Air temperature (T) and humidity mixing ratio (Q) on a 27-m mast near Castroville, Monterey Bay, California. The gradient flow observed in the morning and evening was easterly.

Source: Banta (1995, p. 3621, Fig. 8), by permission of the American Meteorological Society.

Figure 6.12 The effects of a westerly sea breeze on the California coast on 22 September 1987 on temperature and humidity. (A) Wind direction (DIR) and speed (SPD). (B) Air temperature (T) and humidity mixing ratio (Q) on a 27-m mast near Castroville, Monterey Bay, California. The gradient flow observed in the morning and evening was easterly.

Source: Banta (1995, p. 3621, Fig. 8), by permission of the American Meteorological Society.

northern hemisphere) so that eventually it may blow more or less parallel to the shore. Analogous 'lake breeze' systems develop adjacent to large inland water bodies such as the Great Lakes and even the Great Salt Lake in Utah.

Small-scale circulations may be generated by local differences in albedo and thermal conductivity. Salt flats (playas) in the western deserts of the United States and in Australia, for example, cause an off-playa breeze by day and an on-playa flow at night due to differential heating. The salt flat has a high albedo, and the moist substrate results in a high thermal conductivity relative to the surrounding sandy terrain. The flows are about 100 m deep at night and up to 250 m by day.

3 Winds due to topographic barriers

Mountain ranges strongly influence airflow crossing them. The displacement of air upward over the obstacle may trigger instability if the air is conditionally unstable and buoyant (see Chapter 5B), whereas stable air returns to its original level in the lee of a barrier as the gravitational effect counteracts the initial displacement. This descent often forms the first of a series of lee waves (or standing waves) downwind, as shown in Figure 6.13. The wave form remains more or less stationary relative to the barrier, with the air moving quite rapidly through it. Below the crest of the waves, there may be circular air motion in a vertical plane, which is termed a rotor. The formation of such features is of vital interest to pilots. The presence of lee waves is often marked by the development of lenticular clouds (see Plate 7), and on occasion a rotor causes reversal of the surface wind direction in the lee of high mountains (Plate 13).

Winds on mountain summits are usually strong, at least in middle and higher latitudes. Average speeds on summits in the Colorado Rocky Mountains in winter months are around 12 to 15 m s-1, for example, and on Mount Washington, New Hampshire, an extreme value of 103 m s-1 has been recorded. Peak speeds in

Lee Wave Diagram

Figure 6.13 Lee waves and rotors are produced by airflow across a long mountain range. The first wave crest usually forms less than one wavelength downwind of the ridge. There is a strong surface wind down the lee slope. Wave characteristics are determined by the wind speed and temperature relationships, shown schematically on the left of the diagram. The existence of an upper stable layer is particularly important.

Figure 6.13 Lee waves and rotors are produced by airflow across a long mountain range. The first wave crest usually forms less than one wavelength downwind of the ridge. There is a strong surface wind down the lee slope. Wave characteristics are determined by the wind speed and temperature relationships, shown schematically on the left of the diagram. The existence of an upper stable layer is particularly important.

Source: After Ernst (1976), by permission of the American Meteorological Society.

excess of 40 to 50 m s-1 are typical in both these areas in winter. Airflow over a mountain range causes the air below the tropopause to be compressed and thus accelerated particularly at and near the crest line (the Venturi effect), but friction with the ground also retards the flow, compared with free air at the same level. The net result is predominantly one of retardation, but the outcome depends on the topography, wind direction and stability.

Over low hills, the boundary layer is displaced upward and acceleration occurs immediately above the summit. Figure 6.14 shows instantaneous airflow conditions across Askervein Hill (relief c. 120 m) on the island of South Uist in the Scottish Hebrides, where the wind speed at a height of 10 m above the ridge crest approaches 80 per cent more than the undisturbed upstream velocity. In contrast, there was a 20 per cent decrease on the initial run-up to the hill and a 40 per cent decrease on the lee side, probably due to horizontal divergence. Knowledge of such local factors is critical for siting wind-energy systems.

A wind of local importance near mountain areas is the föhn, or chinook. It is a strong, gusty, dry and warm wind that develops on the lee side of a mountain range when stable air is forced to flow across the barrier by the regional pressure gradient; the air descending on the lee slope warms adiabatically. Sometimes, there is a loss of moisture by precipitation on the windward side of the mountains (Figure 6.15). The air, having cooled at the saturated adiabatic lapse rate above the condensation level, subsequently warms at the greater dry adiabatic lapse rate as it descends on the lee side. This also reduces both the relative and the absolute humidity. Other investigations show that in many instances there is no loss of moisture over the mountains. In such cases, the föhn effect is the result of the blocking of air to windward of the mountains by a summit-level temperature inversion. This forces air from higher levels to descend and warm adiabatically. Southerly föhn winds are common along the northern flanks of the Alps and the mountains of the Caucasus and Central Asia in winter

Askervein Wind Profile

Figure 6.14 Airflow over Askervein Hill, South Uist, off the west coast of Scotland. (A) Vertical airflow profiles (not true to scale) measured simultaneously 800 m upwind of the crest line and at the crest line. L is the characteristic length of the obstruction (i.e. one-half the hill width at mid-elevation, here 500 m) and is also the height above ground level to which the flow is increased by the topographic obstruction (shaded). The maximum speed-up of the airflow due to vertical convergence over the crest is to about 16.5 m s-lat a height of 4 m. (B) The relative speed-up (per cent) of airflow upwind and downwind of the crest line measured 14 m above ground level.

Source: After Taylor, Teunissen and Salmon et al. From Troen and Petersen (1989).

Rocky Mountains Chinook Foehn
Figure 6.15 The föhn effect when an air parcel is forced to cross a mountain range. Ta refers to the temperature at the windward foot of the range and Tb to that at the leeward foot.

and spring, when the accompanying rapid temperature rise may help to trigger avalanches on the snow-covered slopes. At Tashkent in Central Asia, where the mean winter temperature is about freezing point, temperatures may rise to more than 21°C during a fohn. In the same way, the chinook is a significant feature at the eastern foot of the New Zealand Alps, the Andes in Argentina, and the Rocky Mountains. At Pincher Creek, Alberta, a temperature rise of 21°C occurred in four minutes with the onset of a chinook on 6 January 1966. Less spectacular effects are also noticeable in the lee of the Welsh mountains, the Pennines and the Grampians in Great Britain, where the importance of fohn winds lies mainly in the dispersal of cloud by the subsiding dry air. This is an important component of so-called 'rain shadow' effects.

In some parts of the world, winds descending on the lee slope of a mountain range are cold. The type example of such 'fall-winds' is the bora of the northern Adriatic, although similar winds occur on the northern Black Sea coast, in northern Scandinavia, in Novaya Zemlya and in Japan. These winds occur when cold continental airmasses are forced across a mountain range by the pressure gradient and, despite adiabatic warming, displace warmer air. They are therefore primarily a winter phenomenon.

On the eastern slope of the Rocky Mountains in Colorado (and in similar continental locations), winds of either bora or chinook type can occur depending on the initial airflow characteristics. Locally, at the foot of the mountains, such winds may reach hurricane force, with gusts exceeding 45 m s-1 (100 mph). Down-slope storms of this type have caused millions of dollars of property damage in Boulder, Colorado, and the immediate vicinity. These windstorms develop when a stable layer close to the mountain-crest level prevents air to windward from crossing over the mountains. Extreme amplification of a lee wave (see Figure 6.13) drags air from above the summit level (4000 m) down to the plains (1700 m) over a short distance, leading to high velocities. However, the flow is not simply 'down-slope'; winds may affect the mountain slopes but not the foot of the slope, or vice versa, depending on the location of the lee wave trough. High winds are caused by the horizontal acceleration of air towards this local pressure minimum.

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Responses

  • CRAIG HERNANDEZ
    What will be the effect of global warming on California "offshore flow"?
    3 years ago
  • juhana
    Are there land breezes in california?
    3 years ago
  • P Hirvi
    How deep is sea breeze circulation?
    3 years ago
  • Kieron
    Which are land and sea breezes examples of ?
    3 years ago
  • tesmi
    How land and sea breezes effect the weather?
    3 months ago

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