Antarctic Climates

Introduction

Climate and the polar regions

Ice plays a critical role in polar climate (Bintanja, 2000). In contrast with lower latitudes, the high albedo and low thermal conductivity of snow cover and ice (sea ice, floating ice shelves, continental ice sheets) have associated negative values of the surface radiation budget (e.g. Nakamura and Oort, 1988). Accordingly, low temperatures occur at the surface and in the free atmosphere in polar regions, and the troposphere has low thickness and tends to be barotropic. The low average temperatures mean a low water-holding capacity of the air, realized as low precipitation amounts annually (i.e. “polar desert”). Because of the atmosphere’s extreme static stability over icecovered surfaces, clouds and precipitation tend not to be generated in situ; rather, they are advected in from lower latitudes by frontal cyclones. This poleward movement of moisture is a consequence of the inflow of energy into polar regions necessitated by the net-radiation deficit. Although these climatic characteristics are common to both polar regions on Earth, there are physical geographic major differences between the Arctic and Antarctic, manifest in their respective climates. Most notably, surface and free-air temperatures, along with cloud amounts, average lower in the Antarctic than the Arctic (Weller, 1982).

The Antarctic versus the Arctic

Whereas the Arctic is a landlocked ocean covered most of the year by sea ice, Antarctica is a polar continent upon which rests a thick ice sheet. The ice sheet rises to over 3000 m in East Antarctica (Figure A26), with a small area (“pole of inaccessibility”) exceeding 4000 m. Moreover, Antarctica is surrounded by the Southern Ocean: the only interruption to the zonal con- figuration of the continent occurs in the Antarctic Peninsula (AP). The extreme continentality results from the high elevation of the ice sheet and the continent’s isolation from other land masses (Carleton, 1992). In contrast, the Arctic is linked directly with middle latitudes via North America and Eurasia. The greater longitudinal variation of land and sea in boreal regions is expressed as a more meridional average pattern of the tropospheric waves and, accordingly, with enhanced transport of heat and momentum by the waves occurring between middle and high latitudes (van Loon, 1979). Over Antarctica the air is often poorly mixed, especially in austral winter and spring. This promotes a colder circumpolar vortex compared with that in the Arctic, and largely is responsible for the more extensive and intense stratospheric “ozone hole” of southern high latitudes.

Figure A26

Map of Antarctica showing major physical geographic features (coastline, Antarctic seas, mountain ranges, ice-shelf margins, ice-sheet contours) and locations of selected permanent manned bases (named). (From Parish and Bromwich, 1991.)

Antarctica virtually doubles its own area between the early fall and spring because of the growth of sea ice (Figure A27). This large annual increase in sea ice occurs because, unlike the Arctic Ocean, the equatorward margin of Antarctica is unconstrained by land, and because atmospheric and oceanic mean circulations are relatively zonal (Carleton, 2003a). Interannual variations of Antarctic sea-ice conditions (its latitude extent, ice-water concentration) are most pronounced regionally (Parkinson, 1992), impacting the surface temperatures and precipitation. In addition to its negative feedback with surface temperature, the sea ice extent influences snowfall in coastal areas: the increasing distance between the moisture source (Southern Ocean) and the coast that occurs as sea-ice extent increases during fall and winter, reduces snowfall amounts. Sea ice-climate interactions manifest fluctuations in atmospheric circulation that include shifts in the preferred “storm tracks” of synoptic cyclones, as modulated by lower-frequency climatic teleconnections. The latter involve particularly the “Antarctic Oscillation” (AAO) and El Niño Southern Oscillation (ENSO).

Figure A27

Schematic summarizing important features of Antarctic-region climate (side legend), as derived from multiple sources (map reoriented with respect to that appearing in Figure A26). (From Carleton, 1992.)

Developments in Antarctic climatology

Major advances have occurred in our understanding of Antarctic meteorology and climatology since the first Encyclopedia of Climatology, “Antarctic Climates” item was written 20 years ago. In that time the major developments have been as follows:

• 1. Improvements in the routine observational network. For the surface, automatic weather stations (AWS) provide data on pressure, temperature and winds, as frequently as every 15 minutes, and their spatial density has increased since 1980, especially in the Ross Sea area. AWS data have been used to improve the representation of 500 hPa height fields over the continent via the atmospheric thickness relationship (Phillpot, 1991); to characterize temperature-sea-ice extent associations locally (Carleton et al., 1998); and to better determine the surface energy balance (Reijmer and Oerlemans, 2002). Satellite retrievals of climate variables additional to those acquired by visible and infrared sensors, include scatterometer-derived near-surface winds over ice-free ocean, ice-sheet altitude and topography from radar altimetry, surface “skin” temperatures and sea ice extent and concentration, including open-water leads and polynyas, from microwave radiometers; stratospheric ozone from atmospheric sounders; and upper-ocean biological activity from shortwave narrow-band sensors. Crucial in these developments has been the establishment of satellite direct-readout facilities on the continent, and the ability to include Antarctic data in global real-time synoptic analyses. Thus, our understanding of Antarctic contemporary climates comes from an array of data sources; some conventional, some remotely sensed, yet with different lengths of record.

• 2. Advances in numerical modeling of Antarctic climates; not just using atmospheric General Circulation Models (GCM) but also coupled ocean-ice-atmosphere models (e.g. Hall and Visbeck, 2002). In particular, mesoscale meteorological models developed for middle-latitudes are now applied to polar environments (e.g. Polar MM5: Guo et al., 2003). More reliability can be placed in modeling the following features of Antarctic climate: the impacts of changes in sea-ice conditions on atmospheric circulations in the southern extratropics (Raphael, 2003a); the important role of ice-sheet topography and local sources of heat and moisture associated with coastal polynyas, for the surface katabatic winds and cold-air mesoscale cyclone systems (mesocyclones) generated in coastal regions of Antarctica (van den Broeke and van Lipzig, 2003); and understanding how the ENSO signal is expressed in Antarctica. These studies confirm that Antarctica interacts actively with lower latitudes, and even may influence climate in parts of the northern hemisphere at certain times of year (Hines and Bromwich, 2002).

• 3. Development of longer-term archives of daily synoptic analyses for higher southern latitudes, as compiled from operational analyses of the Australian Bureau of Meteorology (ABM) and the European Centre for Medium- Range Weather Forecasts (ECMWF), for studying the circulation climates of higher southern latitudes. More recently, the application of automated techniques for tracking cyclones and anticyclones in these data has greatly extended our understanding of Antarctic-region synoptic climatology (e.g. Simmonds et al., 2003). Reanalysis data sets — meteorological fields generated by an analysis/forecast model that is fixed rather than changing through time, and in which all available data are included — particularly are illuminating Antarctic-region circulation climate (Pezza and Ambrizzi, 2003). Composite studies of synoptic phenomena using reanalysis data have identified the typically occurring circulation environments of mesocyclones, explosively deepening cyclones, and extremely strong wind events (Carleton and Song, 1997; Murphy, 2003). However, reanalysis datasets are not equally reliable for climate change studies: the NCEP-NCAR reanalyses of SLP generally are considered inferior to those of the ERA (ECMWF reanalyses) over southern high latitudes.

• 4. Dedicated intensive observing periods and special monitoring programs to study Antarctic-region weather and climate. In the mid-1990s a concerted effort to better depict and understand Antarctic synoptic meteorology and climatology was realized with the FROST (First Regional Observing Study of the Troposphere) project (Turner et al., 1999). An outgrowth of FROST was the drive to develop an updated long-term climatology of surface and free-atmosphere conditions for the Antarctic manned stations as part of the SCAR (Scientific Committee for Antarctic Research) READER (Reference Antarctic Data for Environmental Research) project. Stringent quality control of the observation data acquired by the nations which are signatories to the Antarctic Treaty, is permitting the reliable determination of annual and seasonal trends and changes in climate. For longer time scales, recent scientific traverses, particularly the shallow ice cores drilled in West Antarctica as part of ITASE (International Trans-Antarctic Scientific Expedition), are permitting an assessment of climatic variability and changes over about the last two centuries at mesoscale spatial resolutions.

• 5. Improved determination of the variability (spatial, temporal) of Antarctic surface energy budgets, over both sea ice and continental ice. These have involved shipboard measurements in the sea ice zone, as well as satellite remote sensing (Wendler and Worby, 2001).

• 6. Refinements in depicting the spatial patterns and temporal variability of variables significant to the ice sheet mass balance (Vaughan et al., 1999); notably, precipitation and the near-surface katabatic winds (Guo et al., 2004).

• 7. Application of newer analytical and statistical techniques to determining the links between coarse-resolution climate data, such as mapped fields of geopotential height generated from reanalyses or GCM, and local climate conditions (i.e. “downscaling”) (Reusch and Alley, 2002). These methods include artificial neural networks and “self-organizing maps”.

• 8. Development of satellite-image based “climatologies” of mesocyclones for several genesis key regions around Antarctica. Moreover, there is a realization that these storms are important for the snowfall and wind climatologies of coastal areas and embayments. Accordingly, attempts are being made to more accurately forecast these developments.

• 9. The addition of another 20 years’ data has permitted identification of temporal trends and longer-term changes in Antarctic climate (atmospheric variables, sea ice conditions, ocean temperatures)). In particular, opposing trends of surface temperature between the AP and the rest of Antarctica have become evident (see “Antarctica and global change”, below).

• 10. A realization that Antarctica is not isolated from the weather and climate processes of extratropical latitudes, or even the northern hemisphere. Rather, Antarctica is intimately connected with other places via atmospheric teleconnections, primarily the ENSO.

• 11. Identification and characterization of the dominant modes of variability (subseasonal, interannual, subdecadal) in the atmospheric circulation of higher southern latitudes, and links with the upper ocean circulation. These modes include particularly the ENSO, but also the AAO, the “Pacific-South America” (PSA) pattern, and an “Antarctic Circumpolar Wave” (ACW) that was discovered in the mid-1990s.

Accordingly, this revised item emphasizes particularly the above-noted developments in Antarctic climatology.

Antarctica and global climate change

Documenting Antarctic-region climate and its temporal variations has taken on increased urgency in recent decades, in the context of “global change”. The following climate trends and changes, and their possible global associations, have spurred the increased monitoring of Antarctic weather, climate, biology, glaciology, and oceanography, particularly using satellites (e.g. Schneider and Steig, 2002).

• 1. “Global warming”. This is likely to be evident earliest in polar regions, owing to “ice-albedo feedback”. The latter involves a positive link between reduced snow/ice extent associated with warming, greater absorption of solar radiation by the surface, and continued warming. With respect to Antarctica, the following have been observed: a strong warming trend in the AP over the past several decades (Harangozo et al., 1997), especially in winter (van den Broeke, 2000a), that has been accompanied by reduced sea-ice extent west of the AP; the recent retreat and collapse of ice shelves such as Larsen-B on the eastern AP (Vaughan and Doake, 1996); an observed freshening of the upper ocean in the Ross Sea that suggests increased precipitation and/or melting from the West Antarctic ice sheet (Jacobs et al., 2002); and a subsurface warming concentrated within the Antarctic Circumpolar Current (ACC) (Gille, 2002). GCM studies (e.g. Wu et al., 1999) support a close link between recent climate warming and observed sea ice changes in Antarctica that mostly involve reduced ice-water concentration.

• 2. Stratospheric ozone depletion over the Antarctic in austral spring is a major contributor to global-scale reductions of ozone. Moreover, interannual variations in high-latitude circulation affect the configuration, spatial extent, intensity, and seasonal persistence of the ozone hole (e.g. Thompson and Solomon, 2002). The increased receipt of ultraviolet radiation in higher southern latitudes during spring and summer has deleterious consequences for organisms, including humans (e.g. Jones et al., 1998). In regions of Antarctica where cooling has occurred, a physical link with ozone depletion via the AAO has been proposed (e.g. Sexton, 2001).

• 3. The observed role of larger-scale atmospheric circulation patterns and their teleconnections for interannual variations and recent changes in climate over extrapolar as well as polar latitudes. These include surface temperature trends and summer melt periods linked to the AAO (Torinesi et al., 2003), and the sensitivity of the Antarctic ice sheet’s mass balance to fluctuations in moisture transport, especially in the Pacific sector, where there is a strong ENSO signal (Bromwich et al., 2000).

• 4. Since the mid- to late-1970s the change in the dominant Semi-Annual Oscillation (SAO) of climate variables, especially SLP, temperature and precipitation in sub-Antarctic latitudes and coastal Antarctica, has been accompanied by warming in the tropical Pacific and cooling over large areas of the ice sheet (Hurrell and van Loon, 1994). For the same period the ozone hole worsened.

• 5. The strong interactions glimpsed between biological productivity in the Southern Ocean, and circulation-climate variations expressed as changes in sea surface temperature (SST), sea ice extent, and zonal wind speed.

• 6. The role of Antarctica in long-term global-scale climate changes (glaciations, deglaciations), via its influence on the deep-water ocean circulation. Moreover, contemporary climate processes in the Antarctic appear crucial influences on deep water formation; notably via the polynyas and leads of coastal and sea ice areas generated by strong offshore (katabatic) winds, and by the spin up of polar mesocyclones (see Figure A27). Mesocyclones extract large amounts of energy from the upper ocean, increasing the salinity and density of the water, which then sinks to great depths (Carleton, 1996).

Antarctic climate(s)

No single “Antarctic climate” can be identified (e.g. Bintanja, 2000). Regional ice-atmosphere interactions change markedly with latitude, elevation, season, and exposure to heat and moisture sources (either from open ocean or within the sea ice zone), and with the passage of synoptic systems (Schwerdtfeger, 1984). The seasonal sea ice zone (SSIZ) near maximum extent is considerably further equatorward than the Antarctic Circle, but it remains relatively close to the continent in longitudes south of Australia (Figure A27). The AP frequently exhibits a pattern of climate variability different to that of the rest of Antarctica (Rogers, 1983). The influence of the southern polar region in the global oceanic and climatic environment demands a broad definition, so here “Antarctica” is taken to include the continent and its ice shelves, the SSIZ, and the adjacent ocean area including the ACC (see Figure A26). A discussion of the radiation and energy climates of these higher southern latitudes, and their influence on the atmospheric circulation, is given below in the context of the general circulation, as well as synoptically and subsynoptically. Accordingly, regional variations in these climatic factors are also discussed.

Radiation climates and the energy balance

The radiation balance is the principal climate-causing factor (World Meteorological Organization, 1967) in the Antarctic, where there are strong seasonal variations in solar radiation as well as variations in elevation and surface characteristics. The high reflectivity of the surface (i.e. surface albedo) means that very little insolation is absorbed; thus keeping surface temperatures low all year. Satellite monitoring (visible, infrared, microwave) provides synoptic information on variables crucial to an accurate determination of the radiation balance, such as the extent, thickness and concentration of sea ice; accurate albedo measurements, and cloudiness. For modeling the surface energy balance, these satellite data augment ground-level observations.

Surface energy balance equation

The energy balance equation for unit horizontal area at Earth’s surface is:

Where K↓=global (direct plus diffuse) solar radiation; α=albedo, or fraction of shortwave reflected; L*=net longwave radiation (L↓−L→) (where L↓=atmospheric downward and L→ the terrestrial upward); Q _H =turbulent transfer of sensible heat; Q _E =turbulent transfer of latent heat (E=evaporation); Q _G =subsurface heat flux through ice, soil, or water. The conductive flux (Q _G ) approximates to zero on an annual basis. Over snow- or ice-covered surfaces the radiative fluxes dominate the turbulent convective terms of Q _H and Q _E . However, over icefree ocean or areas of low sea-ice concentration or small ice thickness within the pack, the convective terms become important, at least locally.

Solar radiation (K↓)

The global radiation (K↓) varies from virtually zero in midwinter south of the Antarctic Circle to values comparable to the highest on Earth over the ice sheet in December and January. The high summer insolation is due to lack of clouds, elevation of the Antarctic plateau, a highly transmissive (low aerosol and water vapor content) atmosphere, 24 hours of daylight, and the occurrence of perihelion. These high values are not reduced much even by periods of cloud cover; a decrease in the direct (clear-sky) component is more or less balanced by an increase in the diffuse radiation.

Annual total received solar radiation is relatively high over the continent, particularly for East Antarctica, but decreases over the pack ice zone to a minimum over the sub-Antarctic oceans. This gradient is due to the increasing cloudiness associated with traveling cyclonic systems and proximity to the Antarctic Circumpolar Trough, ACT (see “Climatic variables; Cloudiness”, below).

Albedo

The Antarctic experiences an annual negative net radiation balance largely due to high surface albedo. Albedo values are highest over the permanent snow cover of the continental interior, where around 83% of the incident solar radiation is reflected, and lowest over the ocean (around 11%). Over the coastal sea ice and pack ice zones the albedo is strongly dependent on the presence of meltwater (low α), the presence of snowcover (increases α), and the amount of open water that includes substantial polynyas (Wendler and Worby, 2001). Wendler et al. (1997a) found that midsummer values of albedo for 10/10 ice cover averaged 59%, but with hourly values increasing to 76% where snow covered. Average values of albedo for 5/10 ice concentration were around 30%.

Although the surface albedo increases generally with latitude, the greater cloudiness over sub-Antarctic oceans modulates the effect of short-term variations in sea-ice concentrations on the planetary albedo. Over the continent, determination of cloud amount from space is difficult and leads to large discrepancies in computed solar radiation budget if not constrained using available surface observations (Hatzianastassiou and Vardavas, 2001). In summer the maximum variation in surface albedo occurs at about latitudes 60–70°S (78%), where it is dominated by the change from open water to ice cover. Seasonally, this constitutes about a 60% change in albedo. Between latitudes 80°S and 90°S the annual range does not exceed 10%.

Net longwave radiation (L*) and the Antarctic inversion

The net longwave radiation (L*) at the surface is a function of surface and atmospheric temperatures, and of the downward radiation from clouds, water vapor and aerosols. L* is uniformly negative at the surface over southern high latitudes in winter. In summer there are large differences between the continent and peripheral oceans (50–60°S), due mainly to cloud cover effects. Increasing cloud cover tends to increase L* over the ice sheet because L↓ becomes more important over high albedo surfaces. Over sea ice, L* varies due both to cloud amount and the ice-water concentration (L→; is reduced with increasing ice cover, Table A16). Satellite observations of the outgoing longwave radiation at the top of the Antarctic atmosphere show strong negative values in both summer and winter.

The intense radiational cooling and highly transmissive atmosphere produce a persistent low-level temperature inversion that reaches its greatest depth over the higher elevations of the ice sheet, where it is present almost the entire year (Schwerdtfeger, 1984). The inversion is no stronger at the end than at the beginning of the polar night, because an equilibrium is reached between L→ (decreases as surface temperature falls), and L↓ from the atmosphere above the inversion layer, which changes relatively little with time. Combined with the slope of the ice surface, the semipermanent inversion is the source of the persistent katabatic winds (Connolley, 1996). A clearly defined time for the occurrence of the inversion maximum temperature is absent in the Antarctic, comprising the “coreless winter”. This feature is evident for surface temperatures (T _s) at the Antarctic stations, including the AWS. This is because the rapid radiational cooling begun in the fall is arrested and, in many years, reversed in winter. Changes in the tropospheric longwave

Table A16 Mean outgoing longwave radiation (L↑) and surface temperature (T _s) for different Antarctic sea-ice concentrations, from shipboard observations in the period 24 December, 1994 to 6 January, 1995

Table A17 Mean fluxes of energy budget (EB) components for interior Adélie Land, Antarctica, in the period 20 November–22 December 1985. Values are expressed in Wm⁻² and percent, and are positive (negative) toward (away from) the surface. Q _B replaces Q _G and is the snow heat flux. I = imbalance (3.4%), resulting from measurement errors

pattern between early winter and midwinter, and the consequent poleward movement of warmer air, help explain this temperature feature, which comprises part of the Semi-Annual Oscillation (SAO) of SLP and tropospheric height (see “Climatic variables; Temperature”).

Net (all-wave) radiation (Q*) and energy balance

For the South Pole, Carroll (1982) estimated that 85–90% of the winter deficits of Q* (i.e. the net “all-wave” radiation, or net shortwave minus net longwave) are made up of sensible heat (Q _H ) losses. Table A17 summarizes the energy budget over Adelie Land in East Antarctica, for the period 20 November-22 December 1985, emphasizing the importance of the radiation terms. In the coastal ablation zone, Q* is positive from late September through late February. Q* becomes slightly positive over much of the Antarctic interior around the time of maximum solar elevation, due to increased cloud cover and a downward flux of Q _H .

The annual deficit of Q* necessitates a net influx of energy to Antarctica. This occurs primarily as eddy sensible heat advected by frontal cyclone systems migrating into high latitudes. Because the amplitudes of southern hemisphere planetary waves are reduced, on average, compared with their northern hemisphere counterparts, the waves themselves transport relatively little sensible heat into the Antarctic. Instead, the remainder of the heat transport is carried out by the ocean, whereby divergence (convergence) of air in the near-surface wind field occurs on the western (eastern) side of a standing tropospheric trough, promoting upwelling of cold (downwelling of warm) water.

Turbulent exchange and the role of sea ice in the energy balance

Seasonal and interannual variations in sea ice conditions (areal extent, ice-water concentration, presence/absence of snow cover, ice-edge latitude) profoundly influence the surface energy budget and climate. The sea ice cover modulates the oceanic fluxes of heat to the atmosphere, and thereby its own thickness (Wu et al., 2001). The average thickness of Antarctic sea ice is around 1 m or so in late winter, except where deformation occurs and it reaches thicknesses of up to 4 m. During the 1990s, which was broadly representative of the longer period 1978–2000, a trend to increased Antarctic sea-ice extent accompanied reduced ice-water concentration.

Close to the continent the sea ice tends to have higher concentration and a longer seasonal duration than further out. Moreover, when sea ice is advected equatoward (poleward) by winds and the ocean circulation, it undergoes divergence (convergence), which leads to increasing ice extent and decreasing ice concentration (decreasing ice extent and increasing ice concentration). These associations are evident particularly in the major embayments, notably the Weddell Sea, and on interannual time scales (Carleton, 1988; Turner et al., 2002b). Anomalies of this type appear related to the AAO and the ENSO). Additionally, the presence of low ice-concentration areas and open-water leads and polynyas can significantly influence the local to regional-scale climate (e.g. as increased cloud cover and precipitation). This occurs owing to the enhanced oceanic heat losses to the atmosphere as Q _H and Q _E , especially in winter. Then these fluxes can be at least one order of magnitude greater than those of adjacent areas of higher sea ice concentration.

Table A18 summarizes the radiation and energy fluxes for a ship transect through variable sea ice conditions in Antarctic longitudes of the Tasman Sea during summer (Wendler et al., 1997a). In addition to showing the very strong contribution of the radiative fluxes to the surface energy balance, and of the ocean as a heat sink, these results reveal that the fluxes of Q _H and Q _E are of approximately similar magnitude but opposite sign: Q _H is a source of energy to the surface, Q _E is a sink. In late winter, as sea ice concentration decreases below about 50%, the convective fluxes change relatively little compared with those from open ocean (Worby and Allison, 1991).

Climatic variables

Temperature

Setting the surface energy balance to zero determines the equilibrium T _s. Mean winter T _s of −60°C to −70°C occurs on the highest areas of the ice sheet. Maps of mean T _s (Figure A28) indicate that the cold pole is displaced into East Antarctica, both in January and July (Borisenkov and Dolganov, 1982). There is a general decrease of temperature with latitude, especially evident in the AP. Considerable interannual variability of

Table A18

Summary of radiation and energy luxes (Wm⁻²) through the Antarctic sea ice for the period 24 December, 1 994 to 6 January, 1 995 (variable ice concentrations). Note the dominance of the radiative fluxes. As in Table A1 7, the fluxes are positive (negative) toward (away from) the surface. Q _B = ocean or ice heat flux or phase changes

T _s occurs, notably along the western AP and for stations south of the ACC (60–70°S) because of their position within the SSIZ (the temperature gradient in the latter zone increases to above 1°C/1° latitude in winter, compared with about 0.6°C/1° latitude between 30°S and 60°S). Some of the interannual variability of temperature is related to ENSO (Schneider and Steig, 2002). SST gradients are steepest in the vicinity of the ACC, comprising an Ocean Polar Front (OPF), but strong thermal gradients also border the continent, particularly in winter. On average the meridional strong gradients of temperature near the surface disappear in the mid-troposphere (van Loon et al., 1972).

Figure A28

Map showing mean annual near-surface temperature field for Antarctica, synthesized from surface observations; isotherms in degrees below 0°C. (From Guo et al., 2004.)

Trends in Antarctic temperatures are evident over the past several decades. Although the sign and magnitude of recent temperature trends differ somewhat by study and the method of spatially extrapolating the sparse data (cf. Doran et al., 2002; Turner et al., 2002b), a strong warming of the AP region (winter on western side, summer on eastern side) is evident in many studies. The winter warming has been accompanied by decreases in sea ice extent and reduced concentration to the westward, suggesting the importance of ice-ocean-air positive feedbacks. Synoptically, warm (cold) winters in the AP are associated with a greater (reduced) frequency of northerly wind components as the Amundsen-Bellingshausen Sea mean low pressure intensifies (weakens). These SLP patterns also resemble those connected with the tropical ENSO and its teleconnection to higher southern latitudes occurring via the Pacific-South America (PSA) pattern (see “Circulation variations and climatic teleconnections; The El Niño Southern Oscillation (ENSO) in Antarctica”).

In contrast to the recent warming of the AP, a cooling over much of the continent has coincided with lowering pressures in the ACT, strengthening westerlies over sub-Antarctic latitudes, and greater sea ice extent on average. These are associated with a trend toward more positive values of the AAO.

The dominant SAO in the annual march of temperature — and related changes in wind, pressure, geopotential height and precipitation — over sub-Antarctic latitudes and coastal Antarctica, produces gradient maxima in the equinoctial (March, September) months. The SAO results primarily from the tropospheric radiative imbalances between middle and high latitudes, which are at a maximum around the equinoxes. Also, a SAO has been detected in the latitude of the speed maximum of the ACC (Large and van Loon, 1989).

The atmospheric SAO is involved in the annual patterns of sea ice freeze-up and melt (Enomoto and Ohmura, 1990), as well as ice-water concentration (Watkins and Simmonds, 1999), via its interactions with the latitude location of the ACT and associated patterns of easterly (westerly) low-level winds occurring to the south (north) of this feature. In summer and early fall, when the ACT lies equatorward of the ice edge, easterly winds produce convergence of the pack, resulting in little increase in ice area despite the lowering surface temperatures. However, as the ACT shifts poleward of the ice edge in March, westerly winds to the north encourage ice divergence and a relatively rapid expansion of the pack through the winter and early spring. In December and January the ACT again lies north of the ice edge, which has melted back as temperatures have risen. The associated easterly winds encourage compaction of the ice pack and its retreat to higher latitudes.

There has been a climate change in the SAO since the mid- 1970s to late 1970s (Hurrell and van Loon, 1994; Simmonds and Jones, 1998), involving primarily a shift in the springtime phase from September and early October into November. This change has coincided with strengthened latitude gradients of temperature originating mostly from a warming in the tropical Pacific, possibly associated with more frequent El Niño events (Mo, 2000); reduced sea-ice extent in the Amundsen- Bellingshausen seas (van den Broeke, 2000b), but significant cooling in coastal East Antarctica in early winter (van den Broeke, 2000a). At least part of the increased severity of the ozone hole over the past two decades likely is related to the longer isolation of the southern circumpolar vortex in austral spring, accompanying this change in the SAO.

Cloudiness

Satellite-observed cloud cover is at a maximum (80–100%) over ocean higher latitudes adjacent to the pack ice zone, varies little with the season, and is due to traveling cyclonic storms (see “Synoptic processes; Synoptic cyclone activity”, below). Manned observing station data indicate a cloudiness minimum over the interior plateau (Figure A29 Borisenkov and Dolganov, 1982), especially in midwinter (less than about 25%), and generally low seasonal variability. Maximum cloud cover variability occurs at Antarctic coastal stations located between these two zones (van Loon et al., 1972), arising from higher rates of alternation of high- and low-pressure systems. The contribution of lower cloud to the total cloud amount reaches a maximum over coastal regions and the pack ice zone (cyclonic), and a minimum over interior East Antarctica where subsidence of air dominates. In recent decades, cloudiness in the spring and summer seasons has increased by about 15–20% at the South Pole. These increases appear to be associated with changes in the high-latitude circulation.

Figure A29

Map showing mean annual total cloud cover from station observations (dots show observation sites used for the compilation); cloud cover isopleths in tenths. (From Guo et al., 2004.)

Precipitation

Snowfall dominates precipitation in the Antarctic. However, the contribution of clear-sky precipitation (e.g. “diamond dust”) and blowing snow to the ice sheet mass balance, increases in importance over the higher, colder and dryer regions of the interior (Giovinetto et al., 1990). The mass balance is the result of precipitation, evaporation, direct deposition (rime, frost), drifting, and runoff from melting at low elevations (Figure A30). Snow drift represents a net loss from the continent to the coastal sea ice zone. Although our confidence in measurements of Antarctic continental T _s and pressure is reasonably high, it remains low for precipitation. This results from a lack of data and the difficulty of separating real precipitation from drifting snow, especially in katabatic wind-prone areas. Accordingly, indirect estimates of Antarctic precipitation have been more forthcoming. These compute the net surface mass balance as P−E, where P=precipitation and E=sublimation, from moisture fluxes directed toward the continent. Recently, model calculations of precipitation have been undertaken. The decrease of precipitable water, relative humidity, and precipitation with increasing latitude, together with increasing distance from the Southern Ocean moisture source, means that annual precipitation on the high plateau is less than about 2 cm yr−1 water equivalent. This increases by up to at least 2 orders of magnitude at the coast (Figure A30). The lower elevations of West Antarctica facilitate moisture transport into the interior, mostly by synoptic cyclones.

Figure A30

Map showing long-term accumulation distribution; units of mm yr⁻¹. (From Guo et al., 2004.)

Interannual variations of coastal precipitation are high, in response to high variability of cyclonic activity, especially along the western AP where they manifest fluctuations in the intensity of the Amundsen-Bellingshausen mean low pressure. The latter has an ENSO component (Guo et al., 2004). There has been a statistically significant upward trend in precipitation averaged over the continent between 1979 and 1999, although weakly downward trends are evident regionally and in the interior.

Surface winds

The surface wind is a persistent katabatic (downslope) flow that is gravitational and thermal in origin, relatively shallow (100–200 m deep), and essentially decoupled from the synoptic gradient flow above. It dominates the annual wind speed and direction at many stations in West Antarctica and coastal East Antarctica (Borisenkov and Dolganov, 1982; Figure A31). At Cape Denison the mean annual windspeed exceeds 20 m s⁻¹, which is the strongest on Earth close to sea level (Wendler et al., 1997b). Extremely high wind speeds can persist for weeks or even a month or two, especially in winter. In summer, solar radiation absorption reduces the katabatic effect (van den Broeke and van Lipzig, 2003). Katabatic winds are important because they produce snowdrift and enhance coastal polynyas and sea ice production (Adolphs and Wendler, 1995), thereby affecting the surface mass and energy budgets. Also, katabatic winds promote mesoscale cyclogenesis in confluence areas near the coast (e.g. Heinemann, 1990; Bromwich, 1991; see Figure A27). Streten (1968) found that katabatic winds are best developed in coastal regions with little synoptic cyclone activity, but with generally lower pressure located to the eastward. However, individual synoptic events may enhance the katabatic winds at certain coastal locations. Modeling suggests that synoptic forcing may not be important in determining these winds over interior East Antarctica. There, they can be explained adequately from consideration of the topography (i.e. channeling effect) and the intensity of the wintertime inversion (Connolley, 1996). The katabatic winds are a critical component of the atmosphere’s meridional circulation for southern higher latitudes and are linked to the intensity of the circumpolar vortex in the mid- to upper-troposphere. Katabatic winds may even participate in high latitude-tropical interactions on subseasonal time scales (Yasunari and Kodama, 1993).

Figure A31

Antarctic near-surface winds for winter; units of m s⁻¹. (From Guo et al., 2004.)

Winds on the eastern side of the AP and the western Weddell Sea comprise a strong and persistent southwest to southerly flow, or barrier wind (see Figure A27). This wind is important for advecting ice equatorward into the westerly wind belt, thereby maintaining lower air temperatures in the southern South Atlantic than at comparable latitudes in the southeast Pacific (Schwerdtfeger, 1975, Schwerdtfeger, 1979). The barrier wind results from very cold and stable air that moves westward over the Weddell Sea, is dammed up against the eastern side of the AP, and sets up a thermal gradient similar to that forcing the katabatic wind. This explanation does not require the presence of a semipermanent low pressure in the Weddell, although such a feature often exists and may enhance the barrier wind (Figure A32). Interannual fluctuations in the intensity of the Weddell Sea low, some of which are ENSO-related (Yuan and Martinson, 2001), are strongly evident in the sea ice conditions of the embayment (Turner et al., 2002a).

Figure A32

Schematic summarizing dominant features of the synoptic climatology of southern higher latitudes (side legend), compiled from Multiple sources (map has similar orientation to that in Figure A27). (From Carleton, 1992.)

Large-scale atmospheric circulation

Mean surface pressure and 500 hPa height fields

The essential features of the atmospheric circulation around the Antarctic have been known since at least the International Geophysical Year (IGY) of 1957–1958 (e.g. Lamb, 1959; van Loon et al., 1972). Mean monthly height, temperature, and wind fields at different tropospheric levels are available in atlas form (see van Loon et al., 1972), and have been summarized by Schwerdtfeger (1984). Synoptic and dynamic climatologies of the higher latitude circulation derived from the ABM, ECMWF, and other datasets, largely confirm the IGY studies, although greater detail is now known for previously data-void ocean areas (the South Pacific). In addition, much more is known of the circulation variability on interannual and decadal time scales.

The dominant features of the SLP field of southern high latitudes are a thermal anticyclone over the continent (Figures A32 and A33), although this feature is somewhat artificial owing to the method of reducing pressures to sea level. An almost continuous belt of low pressure (the ACT) around the continent and pack ice zone has well-defined centers located near the major Antarctic embayments, and also off Wilkes Land in East Antarctica (Figure A33a,b). The Amundsen- Bellingshausen mean low (ASL) comprises a “pole of maximum variability” in the SLP field (Connolley, 1997). Modeling experiments suggest that the strong interannual variability of the ASL is explained by the displacement of Antarctica’s highest elevation areas toward the Indian Ocean (Lachlan-Cope et al., 2001).

Figure A33

Mean sea level pressure (SLP, contour interval of 5 hPa) derived from NCEP-NCAR reanalyses for (a) winter, June through August, of 1979–1999; and (b) summer, December through February, of 1980–2000. (From Simmonds et al., 2003.)

The mean lows of the ACT are the summation of individual synoptic cyclones moving in from middle latitudes (Figure A32), and represent areas of stagnation and cyclolysis Carleton, 1979; Simmonds et al., 2003). However, recent satellite- based studies show that these “cyclone graveyards” are often active sites for mesoscale cyclones forming in cold-air outbreaks just to the westward (e.g. Carleton and Song, 1997). The SLP has lowered over Antarctica since the mid-1970s to late 1970s as the surface and troposphere has cooled. This trend comprises part of the AAO teleconnection pattern.

Although the ACT undergoes little variation in intensity between summer and winter (Figure A33a,b), its latitudinal position changes markedly in connection with the SAO: more equatorward in the solsticial (January, July) months, and closer to Antarctica in the equinoctial (March, September) months. Also, the broad zonality of the Antarctic mean circulation undergoes quite marked meridionality on individual (daily) analyses, especially in the South Pacific. High-pressure ridges frequently interrupt the ACT (see Taljaard, in van Loon et al., 1972). Favored longitudes for ridges are in East Antarctica between 0–15°E, 50–60°E, and 140–150°E. The last location particularly is important in blocking events, which typically are more prevalent during ENSO warm phases (i.e. El Niño) (Marques and Rao, 2000).

In the free atmosphere the Antarctic is dominated by a circumpolar westerly vortex that expands equatorward in winter. Lowest heights (and layer thicknesses, or mean temperatures) occur over the ice sheet, and there is evidence that the intensity of this feature is coupled to the katabatic winds. On average the whole vortex is displaced slightly toward the Indian Ocean sector, or wave number 1 pattern. However, a three-wave pattern of tropospheric troughs and ridges also often occurs (Figure A32), especially during blocking events, with a trough over each of the three major ocean basins (e.g. Kiladis and Mo, 1998). In contrast to their counterparts in the northern hemisphere, the tropospheric waves have low amplitude and their positions do not change much from season to season. Interannual variations of wave number 1 occur preferentially toward either the Australian region or the Falkland Islands, with accompanying large variations in sea-ice conditions in the Scotia Sea (Carleton, 1989). This mode of variability is described by a Trans-Polar Index (TPI), which is based on the SLP anomalies at Hobart and Stanley (Pittock, 1984). However, the negative relationship of SLP between these two places is not stable in the long term (Carleton, 1989). Temporal changes in the tropospheric waves are confirmed from NCEP-NCAR reanalyses (Raphael, 2003b), showing that the largest changes have occurred since about 1975 in the southern late fall through early spring, and from the Indian Ocean eastward to the Weddell Sea.

Zonal circulation

In accordance with the limited summer-winter variation in the wave pattern (see “Mean surface pressure and 500 hPa height fields”, above), there is little seasonal change in the intensity of the zonal westerlies at 500 hPa in middle and high latitudes (Trenberth, 1979). However, there is marked interannual variability, especially in the Pacific sector, where it is dominated by the ENSO — westerly zonal winds in adjacent broad latitude zones vary out-of-phase, or so-called “split jet” flow (Bals- Elsholz et al., 2001). During ENSO warm events the STJ (PFJ) tends to be stronger (weaker) (Figure A34); but in cold, or La Niña events, the STJ (PFJ) is weaker (stronger). These variations modify the momentum and heat transports, as well as storm tracks, over southern middle and higher latitudes. However, it is important to note that there is considerable variability of the above-noted patterns within a given phase of ENSO, especially the El Niño events.

Figure A34

Schematic showing the dominant changes in mid- to upper-tropospheric circulation over the Pacific and southwest Atlantic sectors that accompany a warm (El Niño) event. Features over lower latitudes (MJO: Madden-Julian Oscillation, SPCB: South Pacific Cloud Band) mostly are representative of the summer season. Those over middle and higher latitudes (STJ, PFJ, pressure/height anomalies) apply to the winter season 4–8 months earlier. The solid (dashed) lines indicate that the represented feature is stronger (weaker) during El Niño than in La Niña, for the respective season. (From Carleton, 2003a.)

Synoptic indexes long have been used to characterize the zonal atmospheric circulation in the southern extratropics (e.g. the TPI). These are constructed using SLP observations for station pairs across widely separated latitudes (zonal index) or longitudes (meridional index). Indexes of this type capture the SAO in near-surface SLP, and confirm the recent change in the AAO to a more positive (i.e. increased zonality) mode. The increased strength of the southern westerlies that occurs between fall and spring occurs concurrently with the expansion of the sea ice (Ackley, 1981); also, greater zonal wind variability tends to accompany increased variability of the ice. Similarly, between-year regional variations of the ice near maximum extent accompany changes in the zonal index which, in turn, are related to patterns of higher latitude cyclonic activity and the teleconnection patterns of AAO and ENSO (Carleton, 1989; Renwick, 2002).

Synoptic processes

Air masses and air streams

The Antarctic ice sheet is the source region for extremely cold, dry and stable Antarctic continental (cA) air year round (Wendland and McDonald, 1986). It is recognized as inversion air only over the continent and ice shelves because it transforms

Table A19 Mean temperature T (°C), mixing ratio q (gm kg⁻¹), and pseudo wet-bulb potential temperature θ _sw (°C) of higher latitude air masses in summer (s, December–March) and winter (w, June–September) at selected stations

strongly during equatorward excursions (Andreas and Makshtas, 1985). Taljaard (in van Loon et al., 1972) identified a transitional airmass (Antarctic maritime, mA) over the pack ice which is present only during winter. It is slightly warmer and less stable than cA (Table A19), but at times may be indistinguishable from that airmass.

Polar maritime (mP) air — sometimes also known as Southern maritime (mS) — dominates the middle and higher-latitude oceans that are ice free. Developed partly over the marked SST gradients of the OPF, mP air exhibits south-north variations in temperature and moisture (Table A19), being strongly destabilized on moving equatorward. Thus, the OPF is the site for many higher-latitude polar mesocyclone (“polar low”) systems observed in satellite imagery and a high frequency of lower-level (“boundary-layer”) fronts (Carleton, 1995). Modification of mP air also occurs when it returns poleward, typically in advance of frontal cyclones. In this case the air stabilizes and may penetrate far into the interior of Antarctica via the Pacific sector of West Antarctica.

Climatic fronts

The climatic frontal zones (and, hence, major cyclone tracks) of the southern hemisphere (see Figure A32) show a close association with the OPF at middle and high latitudes (Baker, 1979). These polar fronts are manifest as tropospheric thickness gradient maxima (TGM) in association with middle- and high-latitude jet streams (Taljaard, in van Loon et al., 1972). In summer there is some correspondence between the 1000–500 hPa TGM and the summer pack ice margin, implying frontal activity. In winter the higher-latitude frontal frequency maxima occur equatorward of the sea ice margin in the South Pacific, but there are no corresponding bands of TGM at this latitude owing to higher rates of alternation of high- and low-pressure systems (Taljaard, in van Loon et al., 1972). Instead, a well-defined TGM in the 1000–500 hPa layer lies poleward of the ice border along 65–70°S. This has been considered evidence of a secondary Antarctic (as opposed to Polar) Front (see Figure A27). In East Antarctica, TGM in the 500-300 hPa layer may correspond to an Antarctic Jet related to the topography of the ice sheet, as identified during the IGY and subsequently modeled (Mechoso, 1980). Interestingly, there has been little research on the climatology of atmospheric fronts over southern high latitudes in recent years. However, useful insights can be gleaned from satellite passive microwave retrievals of cloud water, columnintegrated water vapor, and precipitation over ice-free ocean, as well as the summary maps of extratropical cyclone tracks derived using automatic tracking of systems in digital analyses.

Synoptically, fronts penetrating Antarctica often lose their identity on ascending the ice sheet, although they may temporarily disrupt the strong surface inversion. Warm fronts advancing southward ahead of mP or mA air provide overrunning precipitation, and frequently are involved in warming episodes associated with the coreless winter.

Synoptic cyclone activity

Automated cyclone-tracking routines applied to digital SLP data (Simmonds and Keay, 2000a,Simmonds and Keay, 2000b) largely confirm the mean patterns of extratropical cyclones determined in the earlier satellite-based (cloud vortex) and IGY climatologies for the Antarctic and sub-Antarctic (Carleton, 1979; Figure A32). However, they add information on the variability (frequency, spatial locations) of these systems and also their characteristic intensities. Over the Southern Ocean, including the SSIZ, the circulation is dominantly cyclonic all year (see “Large-scale atmospheric circulation; Mean surface pressure and 500 hPa height fields”, above), the mean pattern involving cyclogenesis over middle latitudes of the western South Atlantic, Indian Ocean, and mid-Pacific (Carleton, 1979). Synoptic cyclones bring clouds, precipitation and strong winds to coastal Antarctica. The maturity and dissipation (cyclolysis) of these systems occur at successively higher latitudes along well-defined climatic tracks that merge in the ACT. These “storm tracks” are evident as cloud bands in satellite VIS/IR imagery. They are also retrievable from digital daily analyses as zones of high variance of the v-wind (south-north) component and geopotential height, but low variance in the u-wind (west-east) component (Trenberth, 1987). However, there are some differences in the locations of storm tracks derived in this manner and those obtained from tracking cyclone systems (Jones and Simmonds, 1993). Relatively few lows invade interior East Antarctica, whereas many enter West Antarctica from the South Pacific. The latter region particularly shows an interannual variation in cyclone frequencies related to ENSO (Carleton and Song, 2000). Moreover, a trend to decreased (increased) cyclone frequencies over the sub-Antarctic (Antarctic) seas for the period 1960–1999 is consistent with numerical modeling studies employing anthropogenic forcings (Fyfe, 2003). These imply a poleward migration of the zone of maximum baroclinity under global warming. However, this change may depend, at least partly, on the intensity class of cyclone studied (Pezza and Ambrizzi, 2003).

Similar to earlier studies for the northern hemisphere, explosively deepening cyclones (“bombs”) have been identified and studied for the southern extratropics. Those occurring closer to Antarctica have much in common with some “polar lows”, in terms of their large-scale synoptic environments (cf. Carleton and Song, 1997, Carleton and Song, 2000; Simmonds et al., 2003).

The time-averaged centers of cyclolysis within the ACT are located just eastward of the longitudes of maximum variation of sea-ice extent; Carleton and Fitch, 1993; Yuan et al., 1999; cf. Figures A27 and A34). Therefore, the longitudinal frequency of depression centers influences ice distribution and advance/ retreat patterns via temperature and dynamical (wind-induced) advection (Cavalieri and Parkinson, 1981). On average there is no clear connection between Antarctic sea ice extent and cyclone tracks, although closer associations can be found for certain time periods and regions (Yuan et al., 1999).

Mesoscale cyclone and “polar low” activity

High-resolution meteorological satellite imagery has shown that mesoscale cyclones, or polar-air mesocyclones (approximate length scales 100–800 km), occur frequently in Antarctic and sub-Antarctic latitudes (Carleton, 1992, Carleton, 2003b). The improved detection of these systems likely is involved in a statistically significant recent increase in extratropical cyclone frequencies over the southern hemisphere that is not matched by significant changes in SLP (Sinclair et al., 1997). Mesocyclones comprise two main cloud-vortex types: inverted comma-shaped “polar lows”, dominated by positive vorticity advection associated with short waves and jet stream maxima; and spiral-shaped (multi-banded) systems generated mostly by strong sea-air interactions of heat and moisture (Carleton, 1995, Carleton, 1996). One zone of mesocyclone high incidence is located near the winter/spring sea-ice margin, especially in the South Pacific sector (see Figure A27). There, streams of cA or mA air are strongly destabilized in the vicinity of the ice edge and OPF. These storms often track east or northeast through Drake Passage to decay in the Weddell Sea area (Carleton and Song, 1997). The other major zone of meso-cyclogenesis is close to the Antarctic coastline and over the ice shelves; near the confluence areas for the katabatic winds (see Figure A27). There, “vortex stretching” as air descends the steep slope of the ice sheet, enhances cyclonic vorticity. Also, some mesocyclones develop in association with coastal polynyas because of the oceanic large heat losses to the atmosphere. These near-coastal mesocyclones may lack a cloud signature, being detectable only in the AWS data, or be composed mostly of low-level clouds.

The zonally averaged meso-cyclogenesis seems to be greatest in March and September, associated with the SAO (Carleton and Song, 1997). A number of regional-scale “synoptic climatologies” of mesocyclones in the Antarctic have been developed for time periods of varying length; for the Ross Sea, Marie Byrd Land and the Siple Coast, the Bellingshausen and Amundsen seas, and the Weddell Sea (refer to section 2.2.3 in Carleton, 2003b; also Carleton and Carpenter, 1990; Carrasco et al., 2003). In the Bellingshausen-Amundsen seas, strong differences in the ENSO-related SLP and SST anomalies (SSTA) for the southern winters 1988 and 1989 were reflected in major changes in mesocyclone activity (Carleton and Song, 2000). Mesocyclones were more (less) frequent in this sector in 1989 (1988), associated with decreasing (static) SSTA during the respective May to September periods.

Circulation variations and climatic teleconnections

Climate fluctuations in the Antarctic and sub-Antarctic occur interannually (i.e. climate variations), decadally (climate trends), and multidecadally (climate changes). Interannual variations primarily involve teleconnections to the tropical Pacific ENSO, which is, globally, the dominant coupled ocean-atmosphere interaction mode. On decadal time scales, Antarctic-region circulation-climate variations are most closely connected with an Antarctic Circumpolar Wave (ACW) of extratropical coupled anomalies of SST, upper-ocean salinity, SLP, meridional winds, and sea ice. At times the ACW may be linked to, and modulated by, the ENSO. Recent climate changes in the Antarctic, particularly as revealed in the retreat and collapse of ice shelves, and the ongoing warming of the western AP, may be evidence of “global warming”. However, a more immediate explanation for these climate changes derives from long-term shifts in the atmospheric circulation; primarily, the zonally varying mode or Antarctic Oscillation (AAO).

The El Niño Southern Oscillation (ENSO) in Antarctica

The climate teleconnections to the ENSO phenomenon also are evident in Antarctica, especially in the Pacific sector (Kwok and Comiso, 2002). Interestingly, the pressure changes that occur over southern higher latitudes in the lead-up to an El Niño may be at least as large as those in the tropics, and often precede the changes there. (In an El Niño, tropical convection shifts eastward from the “maritime continent” to the central tropical Pacific, in association with the area of highest SST.) The associated weakening of the South Pacific subtropical high, a key “center of action” in ENSO, permits the semi-permanent South Pacific Convergence Zone (SPCZ) to move eastward (Figure A34). The SPCZ is a major trough ahead of which energy and moisture are advected rapidly southeastwards into extratropical latitudes. Consequently, its forward side is evident in satellite imagery as a quasi-meridional band of deep clouds that often energizes cyclonic circulations transiting the Pacific sector of the Southern Ocean. During El Niño the strengthened latitudinal pressure and temperature gradient resulting from the higher SST in the tropical central Pacific, helps intensify the STJ in that sector. At the same time the PFJ tends to weaken. These synoptic interactions are evident on monthly time scales as a standing wave train pattern of alternating anomalies linking the South Pacific and South Atlantic sectors, or so-called Pacific-South America (PSA) pattern (e.g. Mo, 2000). The PSA is reminiscent of the North Pacific PNA pattern, although the pressure/height anomalies generally are of lower magnitude. The PSA pattern during composite El Niño events is shown in Figure A34. Note particularly the out-of-phase relationship of the pressure anomalies between the Bellingshausen-Amundsen sector (Weddell Sea), where SLP anomalies are positive (negative). Similarly, this relationship is evident as out-of-phase anomalies of temperature and sea-ice extent. Thus, El Niño events tend to produce colder conditions, more frequent southerly winds and greater sea-ice extent east of the AP.

The out-of-phase temperature and related pressure anomaly pattern between the seas west and east of the AP comprise an Antarctic Dipole Index (ADP), which correlates well with tropical Pacific SST and surface air temperature, as well as the Antarctic sea ice edge (Yuan and Martinson, 2001). Thus, the ADP appears to link interannual climate variability in the Antarctic seas with the ENSO pattern. The signal is particularly strong in La Niña events, for which it may have some predictive value.

Over most of Antarctica, El Niño events tend to be accompanied by colder conditions than normal and negative anomalies of surface pressure (Smith and Stearns, 1993). The cooling of the troposphere helps intensify the katabatic winds, leading to stronger wind speeds near the coast; at least, as occurred during the major El Niño of 1982–1983. Accordingly, sea ice may be more extensive in those regions while it is reduced in others, such as the Ross Sea.

During La Niña, or ENSO “cold events”, the regional pattern of coupled anomalies in pressure, meridional winds, temperature and sea ice over the Southern Ocean are more or less opposite those for El Niño (Sinclair et al., 1997). In the Pacific sector, La Niñas typically are expressed as a strengthened Amundsen Sea low and associated stronger PFJ (a weaker STJ), and a weaker Weddell Sea low (Yuan et al., 2001). Accordingly, the temperature anomalies along the western AP are mostly negative in La Niña, with greater sea-ice extent, in contrast to the eastern side of the peninsula.

As noted for climate variations in other parts of the world, the ENSO teleconnection is not stable through time. In the Pacific sector of West Antarctica a change in the relationship between ENSO and the poleward flux of moisture evidently occurred around 1991; from strongly positive before that time to significantly negative afterward. This change impacted regional precipitation amounts.

The Antarctic Circumpolar Wave (ACW)

Unlike the ENSO climate signal in Antarctica, the subdecadal ACW appears primarily to be of southern extratropical origin, although it likely interacts with the ENSO as a “slow teleconnection” (Peterson and White, 1998). In the atmosphere the ACW is a propagating wave number 2 pattern that makes a complete circuit about Antarctica in about 8 years (4 years per wave). It differs from the ADP, which is a quasistationary wave, although the two may interact in the Amundesen-Weddell Sea sectors (Yuan and Martinson, 2001). Because the western (eastern) sides of a trough are characterized by opposite anomalies of meridional wind and temperature, the patterns of upwelling and downwelling in the ocean — important for sea temperature and salinity differences — and of sea-ice extent, also are opposite across the wave (Motoi et al., 1998). East of a trough and west of a surface high pressure, the northerly winds advect mild air southward, leading to downwelling and decreased near-surface salinity, and a general retreat of the sea-ice extent. Conversely, west of a trough and east of a surface high pressure, the southerly winds advect cold air equatorward, producing upwelling and increased salinity, and advancing sea ice. Thus, the ACW in the upper ocean is evident as a slowly migrating (to the eastward) suite of anomalies that maintain their identity across seasons and through the different regions of the Southern Ocean. Interestingly, a number of modeling studies have successfully simulated an ACW-like phenomenon, yet its period and even wave number can be different from those evident in the data (e.g. Motoi et al., 1998). Similarly, the analysis of datasets for the periods before and after that analysed by Peterson and White (1998) suggests that the ACW is not always a prominent teleconnection (Kidson, 1999). This lack of temporal stability is reminiscent of the Antarctic teleconnection to ENSO.

The Antarctic Oscillation (AAO) and recent climate changes

The AAO, also known as the zonally varying mode, southern annular mode, and high-latitude mode, comprises the first EOF (empirical orthogonal function in a principal components analysis) of SLP and tropospheric height over southern middle and higher latitudes (Rogers and van Loon, 1982; Kidson, 1999). An intensification of the mean pressure centers respectively over middle latitudes and Antarctica results in stronger geostrophic westerly winds in the sub-Antarctic and PFJ, or the positive phase of AAO. Conversely, a weakening of the pressure meridional difference reduces the westerlies and PFJ, or negative phase of AAO. After the 1970s, the trend of the AAO has been one of increasing “positive polarity”, partly linked to rising SST in the tropical Pacific and more frequent El Niño events. The stronger westerlies bring in relatively mild and moist air to the western side of the AP, providing at least a partial explanation for the observed winter warming and reduced sea ice there. During the same period the intensification of the polar vortex over the continent accompanied downward trends of temperature there, and reduced surface melt in the summer (Torinesi et al., 2003). Stronger westerlies in the AAO positive mode tend to be associated with a zonally averaged greater extent of the sea ice (Hall and Visbeck, 2002).

Although it is most marked in the troposphere, the AAO has links with the stratosphere, and may modulate the shape and intensity of the annually occurring Antarctic ozone hole (Thompson et al., 2000). In general, greater cooling over Antarctica accompanying positive AAO, better isolates stratospheric air from that in middle latitudes, permitting the formation of more polar stratospheric clouds upon which chlorine compounds accumulate, thereby accelerating the breakdown of ozone during austral springtime (Sexton, 2001). Thus, recent Antarcticregion climate changes involving human activities may be linked to changes in frequency of “natural” teleconnection patterns.