Mountains have several effects on winds. Firstly, there is generally an increase of velocity at higher levels, towards the speeds characteristic of the upper troposphere (Figure 12.10). Secondly, a strong wind perpendicular to a high and long range will undulate on the lee side if there happens to be a slightly higher inversion layer to bounce against. These undulations (called lee waves) yield lines of cloud fixed parallel and downwind of the mountains wherever the upper part of an undulation is above the lifting condensation level (Section 8.1). Between them are separated bands of strong surface winds, at the lower parts of the undulations. Also, there may be downslope windstorms over the downwind foothills of the mountain, due to the strong winds prevailing at mountain-crest level being carried down in a lee wave. Lee waves are felt in the upper troposphere too, up to a height perhaps five times that of the mountain range itself. These high waves can help induce Clear Air Turbulence (Note 12.M).
Thirdly, there is the foehn effect, the warming of winds blowing down mountains when there has been rainfall on the windward side. The customary explanation (Section 7.2) provides an instructive exercise in the use of an aerological diagram but applies only to long mountain ranges. Surface winds onto an isolated mountain are normally deflected sideways around it, and do not flow upslope unless the atmosphere is almost neutral or unstable. A slight pressure reduction on the lee side of the mountain draws dry air down from the upper levels of the slope, and the air warms at the dryadiabatic lapse rate as it descends. This particular explanation of the foehn effect does not necessitate precipitation on the windward side of the mountain.
Foehn winds are most evident in winter, when their warmth contrasts most sharply with prevailing conditions. Sometimes their onset is sudden, leading to a rapid rise in temperature. There is a detectable effect in the lee of the Blue Mountains near Sydney when westerlies blow in winter.
Fourthly, there are the strong winds caused by the channelling of airflows within narrow valleys. Such gully winds are important at Adelaide in South Australia, for instance, formed within the Lofty Range behind the city.
All four effects occur when a mountain acts as a topographic barrier to winds. But now we will consider winds caused rather than obstructed by the mountain. In particular, there are thermally direct circulations due to temperature differences on mountains, as follows.
The name of katabatic winds comes from the Greek word for 'down', as these are downhill flows of heavy cold air, driven by gravity and sometimes called 'cold-air drainage'. Initially, these flows slide down the hillsides on either side of a valley, but then they turn down-valley to become valley winds. A fall wind is a cold large-scale katabatic wind from an elevated plateau. All katabatic winds behave like density currents (Note 14.D).
The winds are unaffected by the air's humidity or the surface roughness and they are shallow—typically about 5 per cent of the descent. So the wind is only 10 m deep at the bottom of a slope 200 m high, for example. Nocturnal coldair drainage over Sydney is commonly a flow about 100 m deep moving at about 4 m/s. But there are katabatic winds about 300 m deep at the coast in Antarctica, of remarkable strength and persistence (Chapter 16). Katabatic winds blow at Mawson (67°S) on about a hundred days in a year, with an average speed of 11 m/s, equivalent to a strong breeze (Table 14.1). As a result, winds flow out from the central ridges of the Antarctic plateau in such volume that they induce the polar subsidence which is part of the global circulation (Section 12.3).
The daytime counterparts of katabatic winds are called 'anabatic' winds. They flow up sunny slopes, providing lift for glider pilots and for hang-gliding. Anabatic winds are generally deeper and more gusty than katabatic winds. Speeds of 6 m/s have been measured in Papua New Guinea, in flows more than 150 m in depth.
Anabatic winds are not like density currents but arise when the PBL over mountains is warmer than the adjacent layer of free atmosphere. This causes a surface-pressure reduction over the mountains, and hence winds down the pressure gradient up the side of the mountain. They flow up both sides of broad valleys in South America, for instance, causing a central downdraught which dissipates clouds, so that the middle of a valley is more arid than on the slopes. Also the upslope anabatic winds may trigger thunderstorms, which helps explain why thunderstorms are more common around mountains in summer, especially in the tropics.
All these various kinds of wind can be found in hilly terrain. Depending on the time of day, there may be a toposcale katabatic or anabatic flow near the ground, surmounted by a mesoscale wind such as a country breeze or sea breeeze, with another quite different layer above that containing the synoptic-scale quasi-gradient wind. Thus, for example, afternoon measurements in a South African valley showed a surface anabatic wind of 1 m/s of 50 m depth, with a different wind of 3 m/ s between 50-200 m above ground, and a quasigradient wind dominant above 600 m.
Another example of the various kinds of wind near mountains is illustrated in Figure 14.13,
which shows the sequence of winds in the Parramatta Valley, which opens out into Sydney. There is cold-air drainage downhill during the early morning, with an inversion layer between the cold air and the stationary air above. Once the Sun's heat has warmed the surface sufficiently, there is convection which links the surface air to the quasi-gradient wind, which may be in any direction. However, this period ends once the sea breeze arrives from the coast, with an inversion above the cool maritime air. That dies away sometime after sunset, and the sequence repeats itself, unless there is a change of cloudiness or a strong gradient wind.
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