2.6                                   Mean values of the main meteorological elements

2.6.1                                Pressure at Mean Sea Level 

The lack of observations over the Southern Ocean and across much of the interior of the Antarctic makes the production of accurate, mean climatological fields rather difficult and there are quite large differences between the data sets that have been produced over the last few decades. Nevertheless, the modern NWP systems are capable of assimilating the many forms of satellite and in situ data that are available and can be regarded as producing reasonable analyses. However, as mentioned earlier (see Section 2.4.6) there is still a problem with the derivation of MSLP fields under the high orography of the continent, which is over 4 km in height in places and where the surface pressure is less than 600 hPa. Therefore, although MSLP is plotted across the whole of the Antarctic on many meteorological charts, it should not be regarded as meaningful where the orography is higher than about 1 km.

Figure 2.6.1.1 shows the mean fields of MSLP (1969–98) for the four seasons as produced from the USA NCEP re–analysis project (see, for example, Kalnay, 1996). Key features apparent on these charts are:

·                         The circumpolar trough that rings the Antarctic between the latitudes of 60‑70^o S. This feature is present because of the large number of depressions in this zone that have either developed at these latitudes or moved south from the extra–polar regions.

·                         Within the circumpolar trough there are three climatological low–pressure centres throughout the year located close to 30^o E, 90^o E and 150^o E;

·                         Generally high pressure over the continent, although as noted above, because of the reduction of pressures to sea level, the fields over the interior should be regarded as no more than a qualitative indicator of conditions.

The circumpolar trough exhibits a semi–annual oscillation in both its position and strength (Simmonds and Jones, 1998), a signal that can be observed in the observations from the coastal stations. As can be seen in Figure 2.6.1.1 the trough is deepest and at its most southerly position in the spring and autumn and further to the north and weaker in the summer and winter. To the north of the circumpolar trough, in the mid–latitude regions the phase of the semi–annual oscillation is reversed, with pressure maxima in the spring and autumn.

2.6.2                                The upper–air height field

Mean fields of geopotential height (1969–98) at the 500–hPa and 300–hPa levels for the four seasons as produced from the USA NCEP re–analysis project are shown in Figures 2.6.2.1 and 2.6.2.2, respectively. The 500–hPa surface is everywhere above the orography of the Antarctic and is therefore useful for forecasting for sites on the high plateau. At this level the mean flow (Figure 2.6.2.1) consists of a weak cyclonic vortex centred over the Ross Ice Shelf/South Pole region. The flow over the Southern Ocean is generally zonal, but with a weak wave number three pattern with troughs close to 20^o E, 90º E and 90^o W. At the 300‑hPa level (Figure 2.6.2.2) the vortex is stronger and centred closer to the pole and there is a weak trough/ridge structure that varies throughout the year. Of course on a day–to–day basis the height fields at both these levels show a marked trough/ridge structure and the zonal symmetry only becomes apparent when mean fields are produced.

For an appreciation of the mean vertical temperature structure and wind flow the reader is referred to Appendix 3 in which mean January and July vertical temperature profiles and wind rose data are presented.

2.6.3                                Surface air temperature

The air near the surface over continental Antarctica is intensely chilled as heat is lost to the ice surface by conduction, then, in turn, through long wave radiation from the ice surface to space. This chilling of the surface airflow continues as it drains towards the coast.

On the dome–shaped ice cap of Antarctica, temperatures drop steadily away from the coast. This is due in large part to the higher elevation, but also to the higher latitude and hence less intense insolation of the continental interior. In most coastal regions the mean annual temperature is around –10 to –15ºC, at 1,000 m it is around –20ºC, and in the highest parts near 4,000 m it falls to near –60ºC. Figure 2.6.3.1 shows estimated surface (10 m) air temperature across the Antarctic: these have been computed from the 10 m mean temperatures in ice cores, which approximates to the annual surface air temperature.

On top of this general pattern there are many variations. In some coastal regions, especially the Antarctic Peninsula during mid–summer, temperatures occasionally rise to around +10ºC, while in winter they fall into the minus forties and fifties, depending on latitude. High on the polar plateau of East Antarctica, the temperature in summer rises to around –30ºC before the cold of winter returns with extremes in the minus eighties. Unlike annual surface air temperature estimates, seasonal equivalents are not as readily estimated. At present the only reasonably accurate (to within about 5ºC) way of obtaining such estimates is through a numerical model. Figure 2.6.3.2 shows the European Centre for Medium–range Weather Forecasts (ECMWF) seasonal temperatures from the ECMWF 15–year re–analysis (Gibson et al., 1996).

On average, many Antarctic stations, in particular those on the continent, experience little variation in temperature during the winter months which has lead to the notation that the Antarctic has a "coreless" winter. Figure 2.6.3.3, for example, shows the annual temperature variation at four stations located in the Antarctic interior. The shape of each graph shows a well–defined temperature peak in the summer but a relatively flat section for the winter ("coreless winter"). (See also Figure 7.12.2.4.1 (in Appendix 2) as an example for a more coastal station, Terra Nova Bay in this case.) Part of the reason for the rapid summer temperature rise is the increase in solar radiation, but also the surface of the ice is a little less reflective after the winter. The winter onset is rapid; a small accumulation of fresh snow restores the surface albedo. The "coreless" nature of the winter is due to an approximate equilibrium state being reached in the near surface heat budget after an initial rapid loss of heat through radiation induced losses from the near surface at the onset of darkness (Schwerdtfeger, 1970, p. 276–277).

Figure 2.6.1.1     Mean (1969–98) fields of MSLP (hPa) for the four seasons as produced from the USA NCEP re–analysis project. (Courtesy of Gareth Marshall, British Antarctic Survey.)

Figure 2.6.2.1     Mean (1969–98) fields of geopotential height (m) at the 500–hPa level for the four seasons as produced from the USA NCEP re–analysis. (Courtesy of Gareth Marshall, British Antarctic Survey.)

The world's lowest surface temperature yet recorded was –89.2ºC at Vostok Station (on 21 July 1983), 1,300 km from the coast at an elevation of 3,488 m, in East Antarctica. The South Pole area is generally much colder than its northern counterpart primarily because of its elevation (2,800 m above sea level) but also because the continental ice cap reflects 80 to 90% of incoming solar radiation back to space. Moreover, relatively few maritime air masses reach this location. There is a marked surface temperature inversion of up to 40ºC in the lowest 600 m of the atmosphere over the continental interior in mid–winter, that is the temperature above this atmospheric boundary layer is much warmer than at the ice surface.

Figure 2.6.2.2     Mean (1969–98) fields of geopotential height (m) at the 300–hPa level for the four seasons as produced from the USA NCEP re–analysis. (Courtesy of Gareth Marshall,     British Antarctic Survey.)

2.6.4                                The continental surface temperature inversion

Temperatures vary markedly in the lowest few hundred metres of the atmosphere over the continent. There is a close correspondence between surface temperature and inversion strength, both for individual stations on a day–to–day or month–to–month basis; and for all stations for the winter season.

The surface temperature inversion over the Antarctic continent results from the radiative heat loss from the ice surface, particularly during the polar night of winter. Phillpot and Zillman (1970) found (see Fig 2.6.4.1) that the average strength of the surface temperature inversion over the Antarctic continent in winter was about 25^oC on the high plateau area of East Antarctica decreasing to about 5^oC near the coast. Winter mean inversion depths of 500‑700 m are found at high plateau stations while the depth at McMurdo it is in the range 400‑500 m. At coastal stations that are not on ice shelves, such as Mawson and Davis, have a depth of 300–400 m in winter. There is also a definite seasonal variation of inversion strength and depth. In some cases the surface temperature occasionally falls to near –80^oC at Vostok in winter and the strength of the surface temperature inversion exceeds 30^oC. Then the depth of the boundary layer (as defined by the height above the ice surface of the highest temperature in the troposphere) may be 1,500 m.

Schwerdtfeger (1984) noted that acoustic soundings at Amundsen–Scott Station indicated a well–defined ground–based shear zone of some 40–300 m deep, well below the height of the highest tropospheric temperature. Radiosonde measurements leave no doubt that above this pronounced surface inversion layer in the interior, normally there is a rather thick layer between say 500 and 1,500 m above the ice surface in which the temperature changes little with height.

 

  Figure 2.6.3.1     Annual mean surface temperatures over Antarctica, deduced from 10 m

  snow temperature measurements. (From King and Turner (1997, p. 82).)

2.6.5                                Cloud, white–out and surface and horizon definition

2.6.5.1                          Distribution of cloud amount

Schwerdtfeger (1970, p. 298) outlines some of the difficulties with surface cloud observations in the Antarctic; for example: blowing snow and a lack of light during winter. More recently the role of satellites has contributed significantly to the monitoring of, among other parameters, cloud characteristics. In 1982 the International Satellite Cloud Climatology Project (ISCCP) was established as part of the World Climate Research Programme to collect and analyse satellite radiance measurements (collected from the suite of weather satellites operated by several nations) to infer the global distribution of clouds, their properties, their diurnal, seasonal, and inter–annual variations. However, satellite observations also have limitations: for example, King and Turner (1997, p. 104) note that the ISCCP data underestimate the cloudiness at the South Pole when compared to conventional surface observations taken at the South Pole Station and attribute the discrepancy to poor cloud detection algorithms particularly for thin ice clouds at high southern latitudes.

Referring to Figures 2.6.5.1.1 and 2.6.5.1.2, which show mean summer and winter cloud cover isopleths from ISCCP data for the area south of about 50º S, one might infer that there is less cloud over the Antarctic continent in summer than in winter. This seems at odds with the observational/climatological study of Warren et al.(1986) that indicated increased cloudiness over the continent in summer when compared to winter and yet the ISCCP data for winter agree very well with the Warren data for winter. (See for example, King and Turner (1997, p. 102–104, their Figures 3.25 and 3.26). Perhaps the area of doubt lies with the ISCCP summer data. Intuitively one would expect increased cloudiness in summer/autumn, at least over the Antarctic coast and near inland, due to the greater availability of moisture from open sea water after the sea ice has reached a minimum due to melt.

This certainly seems to be the case for stations in East Antarctica. Figure 2.6.5.1.3 shows mean total cloudiness for the sub–Antarctic Macquarie Island and for the East Antarctic coastal stations of Casey, Davis, and Mawson (I. Barnes–Keoghan and D. Shepherd, personal communication). It may be seen that while there is little variation in cloud amount in the marine environment of Macquarie Island there are clear summer/early autumn maxima in cloud amounts at Casey, Davis and Mawson that correspond with minimum sea ice coverage in the neighbouring seas. However, even with these three stations the maximum in cloudiness appears to occur early to mid summer at Mawson and late summer to early autumn at Casey and Davis. In other words, as with most other parameters, it is important to obtain data specific to a site or area when trying to specify mean cloud amounts.

2.6.5.2                          Distribution of cloud type

King and Turner (1997, p. 105) report on the work of Warren et al. (1986) in summarising the zonal distribution of cloud types. As noted by these workers stratiform cloud (stratus/nimbostratus/altostratus/cirrus) is the most common cloud type although cumuliform clouds do occur as reported, for example, by Schwerdtfeger (1970, p. 299). Moreover, in some cases localised instability can be quite severe and lead to mammatus and small cumulonimbus development (see, for example, the comments on cloud near Davis Station in Section 7.8.4.4) particularly where local convergence occurs in coastal environments or where cold air flows over relatively warm water or rock.

2.6.5.3                          White–out and surface and horizon definition

In the context of practical weather forecasting for the Antarctic, cloud amount is arguably more important than cloud type because one particularly insidious Antarctic hazard that deserves special comment is white–out. This is an optical phenomenon that occurs in uniformly overcast conditions over a snow–covered surface. It is associated with diffuse (uniform), shadowless illumination that causes a lack of surface definition/contrast (the ease with which features on a snow surface can be discerned) and reduced horizon definition (the ease with which the horizon can be defined) (see Table 2.6.5.3.1).

Table 2.6.5.3.1     White–out, surface and horizon definition defined in terms of the obscurity of the sun.

Qualitative term	Surface definition or contrast	Horizon definition	White–out
Good	Snow surface features such as sastrugi, drifts and gullies are easily identified by shadow. The sun is usually un–obscured. Surface features are clearly defined for as far as the eye can see.	The horizon is sharply defined by shadow or contrast. The horizon is distinct with an obvious difference between land (snow) and sky.	Nil
Fair or moderate	Snow features can be identified by contrast. No definite shadows exist. The sun is usually totally obscured. Surface features become indistinct at distances of more than a few kilometres.	The horizon may be identified, although the contrast between sky and snow is not sharply defined.	Perhaps slight effect.
Poor	Snow surface features (e.g. skidoo tracks) cannot readily be identified except from close up (within 50 metres). The sun is usually totally obscured.	The horizon is barely discernable: in other words, the sky can be discriminated from land but no distinct horizon is visible.	Partial – may be more noticeable in some directions, or at a distance.
Nil	Snow surfaces cannot be identified. No shadows or contrast exist. Dark coloured objects appear to float in the air. The sun is totally obscured, although the overcast sky may exhibit considerable glare. The glare appears to be equally bright from surface reflection and from all directions.	Total loss of horizon: the snow surface merges with the whiteness of the sky.	100%

A person's ability to perceive snow–covered orographic features depends on the shadows that they cast, such forms become indistinguishable under white–out conditions. Without any visual stimulation it is common to incorrectly evaluate an incline: one may walk up and down hills without realising it. Furthermore, it is known that an individual attempting to follow a straight path unaided will veer. Judgments of the distance and orientation of objects in the field of view is severely handicapped. Such spatial disorientation is enhanced inside a moving vehicle. White–out conditions can occur while visibility (i.e. transparency of the air) remains good.

While total white–out results from nil surface and horizon definition there are degrees of this effect, for example, partially reduced horizon and surface definition can occur under a broken cloud layer. Thus figures such as Figures 2.6.5.1.1 and 2.6.5.1.2 (summer and winter mean cloud amounts) are a defacto measure of the occurrence of white–out.

Figure 2.6.5.1.2     Mean winter cloud amount (%) from ISCCP data. (Courtesy of the British Antarctic Survey.)

2.6.6                                Precipitation/Accumulation

Schwerdtfeger (1970, p. 294) notes "With reference to Antarctica, the term ‘cryometeors’ might be more appropriate than "hydrometeors, but it is not used".

Over the continent of the Antarctic air is extremely dry, (although more moist over the marginal sea ice zone and further north). Air in the upper levels of the atmosphere circulates towards Antarctica from more northern latitudes. By the time the air descends over the polar central plateau to reach the boundary layer of the atmosphere, which is immediately above the dome shaped ice cap, most of the moisture has been removed.

Although precipitation falls mainly as snow, some coastal areas, especially the Antarctic Peninsula can receive rain at any time of the year. Synoptic scale lows, mesoscale lows, and their associated frontal systems are responsible for much of the near coastal precipitation. The high plateau of Antarctica receives little precipitation and is the world's largest and driest desert: much of the precipitation that does occur results from clear sky deposition of ice crystals.

Figure 2.6.5.1.3     Mean cloud amounts (oktas) for Casey, Davis, Mawson, and Macquarie Island. (Data from 1969–97 synoptic observations – courtesy of Doug Shepherd,                           Australian Bureau of Meteorology.)

The measurement of precipitation over the Antarctic is, however, problematical with difficulties arising in the separation of falling snow from blowing snow. On the other hand, the measurement of snow accumulation is more achievable; snow accumulation being the snow cover that results from the balance between precipitation and "evaporation", where "evaporation" includes the removal of snow by the wind (ablation).

For the continent as a whole, annual snow accumulation is equivalent to about 150 mm of water. Figure 2.6.6.1 shows estimates of the mean annual accumulation of snow and it may be seen that the 50 mm water equivalent isopleth over central Antarctica accords well with the estimate of mean annual precipitation there of around 50 mm of water equivalent, given that wind speeds, and thus ablation, are relatively low over central Antarctica. On the other hand the accumulation rate on the eastern slopes of Law Dome (120 km from Casey), for example, is in excess of 700 mm water equivalent per year; making it second only to the northwest Antarctic Peninsula in accumulation rates measured at specific sites. In general there is a rapid drop in accumulation away from the coast.

2.6.7                                The wind field

The surface wind regime is one of the most characteristic features of Antarctic climatology. Nowhere on any other continent has one single element such an overwhelming influence on the climate of the continent as a whole. The model of the Antarctic wind regime described below requires validation, and certainly the relevance of this scale of motion to the general circulation of the atmosphere is still open to debate. For example, researchers are trying to determine the interrelationship, if any, between the katabatic flow and the circumpolar vortex found during winter and spring in the upper levels of the troposphere and in the lower stratosphere. More information on the nature of the surface wind in the Antarctic is given in Section 6.6.1.

2.6.7.1                          Katabatic winds

The Antarctic surface winds are significantly influenced by the strong radiative cooling of the ice sheet; the very cold (dense), high velocity airflow is confined to a layer about 600 m thick, with the highest speed winds at a height of about 200 m above the ice surface. The gravity driven flow moves very slowly at first away from high elevation areas of the ice cap, accelerating as it moves towards the coast. The configuration of the ice orography provides an extensive elevated cold air source and lower lying glacial basins which cause strong confluence of airflows (see Figure 2.6.7.1.1).

These katabatic winds blow with remarkable constancy in direction, forced to the left of the line of maximum ice slope as a result of the Earth's rotation (the Coriolis force). The wind slowly accelerates over more than a thousand kilometres in some areas, reaching a mean speed of about 11 m s^–1 (~22 kt) perhaps 200 km from the coast. In some glacial valleys where confluence is particularly strong, katabatic winds can reach around 40 m s^–1 (~75 kt) for hundreds of kilometres as the airflow makes its way towards the coast. There is often, however, a slight decrease in speed, due to ice surface roughness, within 100 km of the coastal escarpment. The flow of air down the ice slopes brings about a compensatory downward movement of dry air from the atmosphere above the katabatic wind level.

Once katabatic winds reach the Antarctic coast their downslope driving force has been lost. Rapid deceleration and dissipation occurs within a relatively short distance offshore because glacial valley convergence in the airflow is replaced by divergence of the shallow air–stream that was once katabatic. Figure 2.6.7.1.2 is a schematic of the variation of wind speed radially northwards from the continental interior.

2.6.7.2                          Pressure gradient winds

The simple conceptual model of katabatic winds given above is, in reality, complicated by the existence of migratory low–pressure systems over the seasonal sea ice zone. Low–pressure systems, which bring bad weather to the coastal regions of Antarctica, are generally born over the Southern Ocean and move on a south–eastward course toward the Antarctic low–pressure trough, which is located off the Antarctic coast. On the other hand, there are numerous instances of cyclogenesis over the seasonal sea ice zone, the smaller scale cases being described as “polar lows”, often with life times of less than 24 hours. Whatever their origin, low–pressure systems can add considerable influence to the coastal katabatic winds, producing some of the strongest winds on the face of the earth. Wind speeds often exceed hurricane force (33 m s^–1 (~64 kt)) for several days at a time, with maximum gusts of more than 70 m s^‑1 (~140 kt). This is equivalent in force to Tropical Cyclone Tracy that destroyed Darwin, Australia, in 1975. The winds from that storm, however, affected Darwin for only a few hours. Over most of the interior of the Antarctic continent, where the slope is slight and low–pressure systems normally are less frequent and less intense, the wind speed is often less than 4 m s^–1 (~8 kt).

Figure 2.6.7.1.2     Variation of mean wind speed from Dôme C to the coast. (From Parish and Wendler (1991), International Journal of Climatology. © (1991) Royal Meteorological  Society. Reprinted by permission of John Wiley & Sons Ltd.)   (See Figure 6.6.13.1 for the locations of most of these AWSs.)

2.6.7.3                          Barrier winds  

Ironically, one of the calmest places on Earth is believed to be in the Antarctic, in the area of Windless Bight that delineates the Ross Ice Shelf just south east of Mt Erebus. O'Connor and Bromwich (1988) describe how the interaction between synoptic scale flow and the Antarctic environment cause a stagnation region over Windless Bight as "barrier winds" are deflected around Ross Island.

The barrier winds themselves are formed by synoptically driven boundary layer air with strong static stability impinging on high orography. And not being able to ascend the orography the winds are forced parallel to the barrier (O'Connor and Bromwich, 1988, p. 921). In the case described by O'Connor and Bromwich (see also O'Connor et al., 1994) the high orographic barrier is caused by the Transantarctic Mountains that are west of the Ross Ice Shelf. The Ice Shelf plays a role in allowing strong low–level temperature inversions to develop thus giving rise to the high static stability. After a period of time if no other meteorological event has occurred in the area the cold air has dammed against the high barrier. In the case of the Ross Ice Shelf–Transantarctic Mountains' situation with a synoptic easterly pressure gradient, the depth of the cold air boundary layer increases from east to west, that is towards the mountains. A geostrophic balance is then set up between the Coriolis force and the pressure gradient force generated by the relatively high–pressure area in the cold air dammed against the mountains and the relatively low pressure further east over the ice shelf. Thus south to southwest winds, often quite strong, are generated below the height of the orographic barrier.

Schwerdtfeger (1975) was the first to describe this phenomenon in the Antarctic context when he discussed a possible mechanism for the strong south and southwesterly surface winds along the east coast of the Antarctic Peninsula in the absence of significant low‑pressure systems over the Weddell Sea. In a very similar manner to the Ross Ice Shelf‑Transantarctic Mountain situation described above, cold air masses may move westwards over the frozen Weddell Sea, particularly in winter, and dam against the mountains of the Antarctic Peninsula.

2.6.8                                Visibility including blizzards/blowing snow

2.6.8.1                          Blowing snow

As winds increase in speed to above about 8 m s^–1 (~15 kt) they can cause any loose snow on the surface to begin to drift. If winds strengthen to exceed about 11 m s^–1 (~21 kt) and loose snow is present, then drifting snow may be raised above eye level thus disrupting outdoor activity and being defined as blowing snow for the purpose of international reporting of weather observations. Winds stronger than about 17 m s^–1 (~33 kt) can reduce visibility to only a few hundred metres, if there is sufficient loose snow in the vicinity. In coastal Antarctica, blowing snow is slightly less of a problem during the warmest part of the summer because some melting and consolidation of the surface layer of snow occurs.

2.6.8.2                          Blizzards

Occasionally, blizzards bring all activity to a halt, often for several days at a time. They are said to occur with the combination of freezing temperatures, gale–force winds (or stronger), with blowing snow causing the visibility to be reduced to 100 m or less. Blizzards are usually, but not always, associated with a deep low–pressure systems hundreds of kilometres off the coast of Antarctica. Blizzards may or may not be accompanied by precipitation, although the difference between falling and blowing snow is difficult to discern.