Content

• 6.1 Geostrophic and cyclostrophic winds
• 6.2 Effect of surface friction
• 6.3 Calculating low-level geostrophic wind speed
• 6.4 Slope and valley winds
• 6.5 Squalls and gusts
• 6.6 Tropical cyclones
• 6.7 Determining wind velocity
• 6.8 The compass rose and the wind rose

6.1 Geostrophic and cyclostrophic winds

Winds exist because of horizontal and vertical pressure gradients, so atmospheric motion can be deduced from isobaric surface charts. In the absence of surface friction, if the horizontal pressure gradient force is exactly balanced in magnitude by Coriolis effect then accelerations of the air will be relatively small and a geostrophic wind (from the Greek 'geo' = earth, strophe = turning ) will flow horizontally at a constant speed that is proportional to the isobaric spacing gradient. The flow will be perpendicular to the two opposing forces and parallel to straight isobars. Air will be accelerated to the extent that these forces are unbalanced. Transitory disturbances and vertical movement will create imbalance. When vertical motion is present the horizontal wind cannot be exactly geostrophic.

Geostrophic flow is predominant above the friction layer in very large-scale weather systems, where the pressure gradient force and the Coriolis force are nearly equal and opposite; e.g. the Southern Ocean west wind belt.

Between 15°S and 15°N latitudes there is little geostrophic flow due to weak Coriolis effect (it being zero at the equator), and winds tend to flow across the isobars. (In which case it is more useful to show wind flow on upper air charts as streamlines. A streamline arrow shows the direction of flow, whereas an isotach is a line along which the speed of flow is constant.)

At the other end of the scale in short-lived mesoscale systems, Coriolis has insufficient time to take effect or is relatively weak compared to other forces, so geostrophic balance is not present and air accelerations can be quite large.

If atmospheric circulation was always in perfect balance between geostrophic forces and pressure gradient forces, geostrophic winds would flow and there would be no change in pressure systems. In reality the pressure distribution takes the form of curved isobars resulting in a third force — the centripetal acceleration — which pushes the flow inward of the curve.

The gradient wind is the equilibrium wind for the three forces — centripetal acceleration, pressure gradient force and Coriolis (or geostrophic). It is roughly aligned with the isobars on the meteorological surface chart. The vector difference between the geostrophic and the gradient winds is the ageostrophic wind. Thus, ageostrophic movement is large for small-scale systems and small for large-scale systems.

When the centripetal acceleration becomes the major control of the gradient wind, there is an extremely strong curvature of the airflow and the winds are called cyclostrophic (Greek = circle – turning); for example, tropical cyclones and tornadoes. When a body is moving in a curved path, centripetal force is the radial inward force that constrains the body to move in that curved path and, even at constant speed, there is an inward acceleration resulting from the body's continually changing velocity. (The same applies to an aircraft in a constant-speed level turn.) The equal and opposite centrifugal force that appears to act outward on a body moving in a curved path is a fictitious force, but convenient to show the equilibrium forces for air moving in a cyclonically curved path; e.g. around a surface low pressure system, thus:

For the gradient wind to follow cyclonically curved isobars, the pressure gradient force must be slightly stronger than Coriolis to provide the centripetal force. As the magnitude of the Coriolis is directly dependent on wind speed, it follows that the wind speed around a low is less than would be expected from the pressure gradient force and the gradient wind is sub-geostrophic.

For air moving in an anticyclonically curved path (e.g. around a high), the opposite occurs, and the Coriolis provides the centripetal force.

For the three forces to be in equilibrium, the Coriolis must exceed the pressure gradient force. Consequently, the gradient wind speed must be greater than would be expected from the pressure gradient force — and thus is super-geostrophic.

Air moving within a pressure pattern possesses momentum. If the air moves into a different pressure pattern and gradient it will tend to maintain its speed and Coriolis for some time, even though the pressure gradient force has changed. The resultant imbalance will temporarily deflect the airflow across the isobars in the direction of the stronger force — Coriolis or the pressure gradient force.

6.2 Effect of surface friction

The Earth's surface has a frictional interaction with atmospheric motion that reduces the wind speed and thus the Coriolis effect. The pressure gradient force remains the same, so the wind is deflected towards the region of lower pressure. The friction effect is greatest at the Earth's surface and reduces with height until, at the top of the friction layer or boundary layer, the wind velocity is the gradient wind. This will usually occur somewhere between 1500 feet and 5000 feet above ground level — much lower over the sea.

So, the gradient wind flow tops the boundary layer and the wind profile decreases in speed and veers* (in the southern hemisphere) in direction as height decreases — and the friction effects increase — until the wind velocity at the surface has a cross-isobar component and a much lower speed. Thus, in the presence of surface friction — a force that always acts opposite to wind direction — the veering boundary layer air spirals in towards a low (clockwise rotation) and out from a high (anticlockwise rotation).

*The terms veering and backing originally referred to the shift of surface wind direction with time, but meteorologists now also use the terms when referring to the shift in wind direction with height. Winds shifting anti-clockwise around the compass are 'backing', while those shifting clockwise are 'veering'.

Velocity change between surface wind and gradient wind

Over land, the surface wind speed may be only 30–50% of the gradient wind speed. In the boundary layer, wind slants across the isobars in the direction of the gradient force; i.e. towards the lower pressure. The stability of the boundary layer affects the strength of the friction force; a very stable layer suppresses turbulence and friction is weak, except near the surface. In a superadiabatic layer, convective turbulence is strong and the friction force will also be strong (refer to sections 3.3.2 and 9.1). The following table is for a typical neutrally stable layer, and shows the daytime average angular change in wind direction for an average wind profile over various terrains and beneath a moderately strong gradient wind of 30 knots or so.

The wind is backing as height increases; the change in direction in the first 300 feet is negligible in strong winds but greatest in light winds (below 10 knots) and may be as much as 15–20° if the surface wind is less than 5 knots. The greatest change in wind speed occurs at night and early morning.

Also read the 'Wind shear and turbulence' module of the 'Decreasing your exposure to risk' guide.

6.3 Calculating low-level geostrophic wind speed

The geostrophic wind can be estimated from the isobar spacing on a surface (mean sea level) synoptic chart. The estimation is usually a reasonable approximation of the wind speed around 3000 feet agl over much of Australia. The equation applied is:

Geostrophic wind speed (knots) = 3832 G × T / P sine L
where G = horizontal pressure gradient in hPa/km
T = air temperature in Kelvin units
P = msl pressure in hPa
L = the latitude in degrees

Because the proportion T/P normally doesn't vary greatly at msl, the equation can be simplified to:

Geostrophic wind speed (knots) = 2175/D sine L
where D = the distance in kilometres between the 2 hPa isobars on the chart.

The sine of an angle less than 60° can be estimated easily without reference to tables by using the 1-in-60 rule of thumb; i.e. the sine of an angle is roughly degrees × 0.0167 [or 1/60]; e.g. sine 36°S = 36 × 0.0167= 0.601; or 36/60 = 0.6

The following table is derived from the preceding simplification and shows the estimated wind speed in knots for spacings between the 2 hPa isobars, from 40 to 600 km. If the surface chart shows 4 hPa spacing, then just halve the estimated distance between the isobars and still use the table below.

6.4 Slope and valley winds

Valleys tend to develop their own air circulation, somewhat independent of the ambient wind overflow. They have a tendency to flow up or down the valley regardless of the prevailing wind direction. This circulation is modified by solar heating of the valley slopes.

Anabatic winds form during the day when hillside slopes are heated more than the valley floor. The differential heating of contact air causes air to flow upslope. Wind speeds of 10 knots plus may be achieved.

Katabatic winds normally form in the evening. They are the result of re-radiative cooling of upper slopes, which lowers the temperature of air in contact with the slope and causes colder, denser air to sink rapidly downslope. In some circumstances, katabatic winds can grow to strong breeze force during the night but cease with morning warming. Anabatic and katabatic winds are usually confined to a layer less than 500 feet deep. However, the turbulence — and the sink — associated with a katabatic wind will adversely affect aircraft. Aircraft flying in a southern Australian valley late in a warm evening should expect the onset of katabatic winds.

Katabatic winds are density or gravity currents. They can also occur in the tropics; for example, the Atherton tablelands in northern Queensland form a plateau adjacent to the tropical coast. Winter nocturnal temperatures on the plateau can reduce to near freezing and the cold, dense surface layer air flows downslope onto the warm coastal strip. In some cases katabatic winds can persist for days; an extreme example is the large-scale diurnal katabatic winds flowing from the dome of intensely cold, dense air over the Antarctic ice plateau — average elevation 6500 feet. These winds can achieve sustained speeds exceeding 80 knots, though speeds of 160 knots have been recorded at Commonwealth Bay — the windiest place on earth.

6.5 Squalls and gusts

Squalls or 'squally winds' are a sudden onset of strong wind lasting at least a minute then dying quickly. Wind speeds exceed 22 knots, and possibly reach 70–90 knots. They may be associated with a thunderstorm (rain squall, snow squall), with a squall line, with a dry outflow from a thunderstorm in the interior (dust squall) or with an intense cyclone where the squall reinforces the strong environment wind. Gusts or 'gusty winds' are onsets of increased wind speed that exceeds the mean wind speed by at least 30% but are shorter-lived than squalls, and often complemented by matching lulls.

6.6 Tropical cyclones

Tropical cyclones form only in very moist air in ocean regions where surface water temperatures exceed 26 °C. They generally occur between November and April, and in latitudes 5° – 20° South and are prominent features on the synoptic charts. Coriolis effect within 5° of the equator is too weak to develop the initial vorticity and sea surface temperatures are too low at latitudes higher than 20°. To be named as a tropical cyclone (typhoon in South-east Asia) the storm must have sustained wind speeds exceeding 33 knots; if wind speed is less, it is a tropical depression. In the eastern Pacific and the Atlantic the tropical cyclone would be named as a tropical storm for wind speed between 34 and 62 knots, then upgraded to hurricane status when the sustained wind speed exceeds 62 knots; the hurricane is then downgraded back to tropical storm when it weakens.

Small tropical depressions (warm-core lows) form on a trough line. Warm-core lows usually become less intense with increasing height but — powered by the latent heat of condensation and if the vertical wind shear is low (below 20 knots) — some become more intense with height. They develop into a tropical storm or a monsoon low, with a very rapid updraught. This may create a cyclostrophic vortex and possibly grow, over two or three days or even less, into a full-scale tropical cyclone, with wind speeds often much greater than 62 knots. A gust of 139 knots was recorded at Mardia, near the Pilbara coast of Western Australia, in February 1975.

The pressure drop within the tropical cyclone may be 50 to 100 hPa. (TC Orson produced a msl pressure of 905 hPa at the North Rankin gas platform in April 1989.) The diameter of the vortex may be 400 km, with a central eye 20–40 km in diameter surrounded by spiral feeder bands of CB cloud reaching the tropopause. The dry air in the eye usually descends slowly and warms adiabatically; near the surface it may be 5–8 °C warmer than the surrounding cooled air. The enormous energy within a large tropical cyclone can result in a local lifting of the tropopause; the Atlantic hurricane Bonnie of August 1998 produced chimney clouds reaching 59 000 feet.

The tropical cyclones affecting Australia mainly form in the Coral Sea, Arafura Sea, Timor Sea and the Gulf of Carpentaria. They are usually more compact, but no less severe, than their counterparts elsewhere. While developing, the cyclone usually drifts to the west or south-west at about 10 knots. Sometimes it recurves and accelerates to the south-east and, unless it crosses a coastline, loses its impact by 30° S. They last about six to 10 days (although TC Justin persisted for three weeks off the Queensland coast in 1997. When a cyclone crosses a coast it loses the source of latent heat from the warm, moist ocean air, and weakens into a rain depression, which has high potential for major flooding. About nine tropical cyclones occur around Australia each year.

Wind speeds felt at the surface in the south-west quadrant, before recurving, will be much greater than those in the north-east quadrant, due to the addition or subtraction of the forward movement to the rotational movement. Wind speeds of 148 knots, with a core pressure of 877 hPa, have been recorded in Pacific Ocean tropical cyclones.

Monsoon lows are a feature of the active period of the northern Australian wet season. They develop over land from tropical depressions but don't grow into a tropical cyclone unless they move offshore. Monsoon lows bring turbulence, low cloud and heavy rain with reduced visibility over an extensive area for a considerable time; as does a tropical cyclone when it weakens into a rain depression.

Further information concerning tropical cyclones can be found at the Australian Bureau of Meteorology's tropical cyclone page.

Tropical cyclone categories

The Australian Bureau of Meteorology defines cyclone intensity in its area of responsibility, 90°E to 160°E, from category 1 to category 5, according to the expected strongest gust, as below:

Cyclone Tracy, which wrecked Darwin (24/12/1974) was category 4, but the highest recorded gust in the city was 117 knots. Cyclone Vance (22/3/99) was category 5.

The Saffir-Simpson scale, however, is used in the Atlantic and Eastern Pacific for categorising hurricane intensity:

6.7 Determining wind velocity

During pre-flight planning, pilots determine the forecast wind velocities, at various cruising levels and at aerodromes along their route, by reference to forecast information provided by an authority such as the Australian Bureau of Meteorology or Airservices Australia.

Meteorological forecast information for an area [an ARFOR] can be obtained from Airservices Australia's internet meteorological briefing system. Just click 'Area Briefing' on the opening page, uncheck the other boxes on the next page and select the ARFOR area on the map. An area briefing provides area and aerodromes meteorological information and relevant Flight Information Region meteorological information. The aerodrome meteorological information is in the form of existing meteorological reports [METARs] and aerodrome forecasts [TAFs]. For further information see 'Obtaining weather forecasts'.

The real-time weather observations, at about 190 airfields, can be obtained by telephoning the Australian Bureau of Meteorology automatic weather station at the location and listening to the audio data. See AWIS in the VHF radiocommunications guide.

As the flight progresses, the navigation techniques employed enable calculation of the actual wind velocity at cruising level. While airborne, a radio-equipped aircraft can usually obtain a report of actual weather conditions at the larger aerodromes; see 'Acquiring weather and other information in-flight' in the VHF radiocommunications guide. If a mobile phone is carried, the AWIS (if available) can be used to obtain surface wind and some other weather data. However, surface wind velocity at smaller airfields can be estimated from the probable wind profile knowing the upper level velocity (refer to section 6.2) or determined by observation.

Determining surface wind direction visually while airborne

Apart from an airfield windsock, the most obvious indicators of surface wind direction are dust from agricultural operations or moving vehicles and smoke from chimneys or smaller fires.

Wind ripples in grassland, crops or tree tops provide a reasonable indication in light to moderate winds, as does wave movement on small to larger lakes.

In lighter winds the wind shadow of still water, at the upwind edge of a small lake or dam, is usually apparent. And, of course, when the aircraft is flying at a lower level the aircraft's drift is a strong indicator of the near-surface winds.

The Beaufort wind speed scale (land)

The Beaufort wind speed scale (sea — and perhaps large lakes)

     0 – Sea is mirror-like
     1 – Ripples present but without foam crests
     2 – Small wavelets, glassy appearance and do not break
     3 – Large wavelets, crests begin to break, with scattered white horses
     4 – Small waves becoming longer, fairly frequent white horses
     5 – Moderate waves, many white horses with chance of spray
     6 – Large waves are forming with extensive white foam crests, spray probable
     7 – The sea heaps up, white foam from breaking waves is blown in streaks
     8 – The edges of crests break into spindrift with well marked, foam streak lines
     9 – High waves with tumbling crests and spray, dense foam streaks
   10 – Very high waves with overhanging crests, surface appearance white, visibility affected
   11 – Chaotic sea, large parts of waves blown into spume with foam everywhere
   12 – Air filled with foam and spray, visibility severely impaired

State of seas classification

The following table is the state of seas classification, with likely maximum wave height in metres, used in general meteorology reports and warnings for Australian coastal waters:

State of swell classification

The following table is the state of swell classifications used for reporting the wave-train height and length:

The length and speed of the wave-train can be calculated readily if the period (in seconds) is measured; i.e. the length in metres is 1.56 × the period squared and the speed in knots is 3.1 × the period.

e.g. if period = 10 seconds, then train lengths = 156 metres and propagation speed = 31 knots

6.8 The compass rose and the wind rose

In nautical terms there are 32 compass 'points' each division being 11.25° of azimuth. Winds shifting anticlockwise around the compass rose are 'backing', those shifting clockwise are 'veering'. The names of the compass points and the associated compass direction in degrees are shown in the following table. The term 'by' indicates plus or minus one point (11.25°) in the stated direction; e.g. 'nor'east by north' indicates north-east (45°) minus 11.25° = 33.75°.

The wind rose

The term 'wind rose' nowadays applies to the diagram meteorologists use to represent the wind velocity statistical data collected for a particular location. To view the wind rose for a specific location in Australia, go to the Bureau of Meteorology's wind rose page.