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Meteorological regime and precipitation formation in central Antarctica
Meteorological regime of central Antarctica forms as a result of interaction of a number of factors among which the most significant are radiation balance of the underlying surface (briefly discussed in the Introduction) and atmospheric circulation that brings heat and moisture to the central part of the continent.
Atmospheric circulation in the southern hemisphere
The latitudinal temperature gradient existing all year round in the whole troposphere above Antarctica is responsible for the formation of the circumpolar vortex in the free atmosphere which is characterized by lower pressure in its center and clockwise rotation. As a result, air is descending over the most of the continent [Voskresensky, Lysakov, 1976], which is one of the main factors in forming the anticyclonic type of weather. The predominance of clear sky typical for such weather is favorable for the radiative cooling of the surface, while low temperature causes an extreme dryness of the air which leads to further cooling. Cold air is flowing down along the glacier slope, while in the free atmosphere this flux is counterbalanced by inflow of moist and warm air from ocean. The change of the direction of the meridional component of the air flux takes place at the altitude of 3.8–5 km above sea level. This circulation develops most intensively in winter when the gradient between the pole and the low latitudes is the strongest [Averianov, 1990; Schwerdtfeger, 1987]). The above picture is often disturbed by meridional invasions into the high-latitude region of the cyclones formed at polar or, less often, Antarctic fronts [Dydina et al., 1976; Savitsky, 1976]. The latter are usually smaller and less developed in height. They are formed at the latitudes of 60–65 °S and the zonal component is dominant in their movement: they move around Antarctica parallel to the main stream, i.e., from west to east. Polar cyclones are generally deeper and larger than Antarctic ones. Possessing considerable meridional component in their movement, they sometimes penetrate far into the Antarctic plateau and thus play an important role in the inter-latitudinal exchange of heat and moisture of the southern hemisphere. Approaching the boundary between comparatively warm waters of the Southern ocean and cold Antarctic coast the polar cyclones can become stronger and, provided the presence of the blocking ridges of high pressure, stationary. On the maps of long-term average cyclone system density the areas of most frequent cyclone occurrence can be clearly seen (Fig. 3): Weddell, Ross, Bellingshausen and Commonwealth seas. These areas play important role in formation of climatic regime of Antarctica [Averianov, 1990; Schverdtfeger, 1987].
Fig. 3. Climatological cyclone system density distribution derived from the NCEP (National Center for Environmental Prediction) reanalysis (1958–1997) for winter (from [Simmonds, 2003]). The contour values are 1, 2, 4, 6 and 8 • 10-3 (degrees of latitude)2.
As a whole, the atmospheric circulation in the middle and high latitudes of the southern hemisphere is governed by the following basic regimes (see the review in [Simmonds, 2003]).
First of all, this is the so-called Southern Annular Mode [Thompson, Wallace, 2000] that is characterized by opposite air pressure variations in the middle and high latitudes of the southern hemisphere. The index of its intensity is Antarctic Oscillation Index (AOI) representing the difference of mean latitudinal near-surface air pressure at 40 and 65° S [Gong, Wang, 1999].
Higher index means stronger gradient of pressure and temperature between high and middle latitudes, stronger westerly and weaker inter-latitudinal exchange, which causes cooling in Antarctic. This annular mode is related to the tropical circulation (ENSO), which is confirmed by the fact that El-Nino years are often correspond to the lower values of AOI [Maslennikov, 2002a,b]. Another important regime is the Antarctic Circumpolar Wave, which characterizes the drift of anomalies of meteorological and oceanographical parameters around Antarctica from west to east with a period of about 8–10 years [White, Peterson, 1996]. Anomalies of temperature and pressure, being born in the subtropical zone of the Pacific in relation to El-Nino, are then transferred by Antarctic circumpolar current to the east. This phenomenon is specific to the southern hemisphere, because in the northern one there is no continuous circumpolar current [Peterson, White, 1998]. The period of oscillations related to this wave is 4–5 years.
Disturbance to the two previous regimes is brought by Antarctic Dipole Mode (ADM) that is opposite oscillations of temperature, pressure and sea ice cover in east part of Pacific sector and in Atlantic sector of Antarctic [Yuan, Martinson, 2001]. The Antarctic Dipole is related to tropical circulation, too, the El-Nino years being characterized by positive anomalies of temperature in the Pacific sector and negative ones in Atlantic sector. ADM is actually one of the strongest mechanisms responsible for the transmission of the climatic signal from low to high latitudes [Liu et al., 2002]. Despite relatively weak degree of investigation of the above circulation regime, their role in forming climatic variability of interior parts of Antarctica is in general beyond doubt. In years of anomalous development of meridional processes more cyclones invade into the continent, which causes warming and increasing of precipitation. On the contrary, when zonal processes are stronger, air temperature and pressure are lower in high latitudes [Dydina et al., 1976; Zhukova, 1986; Savitsky, 1976]. In particular, in years with higher AOI index increased temperature is observed over Antarctic Peninsula and decreased over the rest of the continent, especially in East Antarctica. The influence of the tropical circulation on the Antarctic climate mainly reveals itself in reduced sea ice cover in Amundsen and Bellingshausen seas in the El-Nino years (which corresponds to negative Southern Oscillation Index) and to lesser degree in cooling of the interior part of the Antarctic [Kwok, Comiso, 2002]. Thus, cooling observed during the last 10– 20 years over the most of the continent with simultaneous warming over the Peninsula [Doran et al., 2002] is consistent with stronger Southern Annular Mode and El-Nino during the same period. Rapid warming in the area of Antarctic Peninsula is related to stronger westerly and thus to more intensive advection of warm oceanic air, as well as with destruction of sea ice in the surrounding seas [Kwok, Comiso, 2002]. Intensification of the annular mode (increasing of AOI index) is accompanied by an increased air pressure to the north of 40° S and its decrease in the high latitudes. At the same time smaller amount of cyclones are observed in the southern hemisphere. This apparent contradiction is explained by the fact that though the number of cyclones is less, they became deeper and more intense [Simmonds, Keay, 2000].
Surface temperature inversion at Vostok and wind regime.
Since the mean values of main meteorological parameters at the near-surface level were discussed in the Introduction, below we will consider the meteorological regime of troposphere using the published data of balloon-sounding observations.
Fig. 4. Vertical distribution of air temperature in the boundary layer at Vostok (1) and Mirny (2) stations in July and January (from [Voskresensky, Tsigel’nitsky, 1985]).
The most typical feature of tropospheric structure in central Antarctica is a stable thick layer of surface inversion of mixed radiation and dynamic origin [Connolley, 1996; Phillpot, Zillman, 1970; Tsigel’nitsky, 1982; Voskresensky, Tsigel’nitsky, 1985]. The mean thickness of inversion in winter is about 800 m with a temperature difference between upper and lower boundary of about 25 °C and occurrence of nearly 100 % (Fig. 4). These values are twice of those in central Greenland. Monthly means of the main inversion characteristics are listed in Table 1.
During its maximum development, in winter, the inversion layer is not homogeneous. Three sub-layers can be distinguished. The first one is about 100 m thick and characterized by the most intensive temperature changes with mean vertical gradient of -8 °C 100 m-1. For the second, that has thickness of 250 m, a weaker gradient (-2 °C 100 m-1) is typical. Finally, the third one (about 500 m) is isothermal. Just above the surface inversion a quasi-stationary layer is situated with weak positive gradients of temperature (0.3 °C 100 m-1). Thus, during the periods of maximum inversion development the normal temperature distribution typical for free atmosphere (0.6 °C 100 m-1) is established only from the altitude of 2000 m above ice surface [Tsigel’nitsky, 1982; Voskresensky, Tsigel’nitsky, 1985].
With such high values of thickness, intensity and probability surface inversion plays a role of screen preventing thermal and dynamic interaction of free atmosphere with the underlying surface. It is confirmed by the fact that the maximum amplitude of temperature in winter in the isothermal layer above inversion (11.0 °C) is twice as less as near the surface (21.2 °C) and less than in free atmosphere (13.1 °C) [Tsigel’nitsky, 1982].
In summer months because of radiation heating the thermal stability of the lower atmospheric layers above Antarctica sharply decreases. In the afternoon the surface inversion can be completely destroyed. Moreover, at this time of the day the conditions are favorable for the formation of a thin (about 100 m) layer with unstable stratification overlaid by elevated inversion or by isothermal layer [Tsigel’nitsky, 1967].
The inter-annual variability of inversion parameters (thickness and intensity) is a complex index of climate variability of central Antarctica because inversion is formed under the influence of several main climatic factors: underlying surface, radiation conditions and atmospheric circulation. During the period 1958–1982 the inversion parameters revealed significant trends which were opposite for thickness and intensity of inversion. It is explained by the fact that, with an increasing (decreasing) inversion thickness, the gradient of temperature in the inversion layer will become weaker (stronger) provided that the intensity of radiation cooling remains the same. In general, during 1958–1982 the mean thickness of winter (June–August) inversion lowered by 330 m and ∆T increased by 1.02 °C*. At the same time the surface air temperature warmed by about 1.25 °C. Since dynamic settling of air masses is important for inversion formation, reduced inversion thickness can be related with an increased intensity of vertical air movement. The growth of ∆T means a more rapid warming in the free atmosphere comparing to the near-surface air, which is explained by the screening effect of the inversion layer [Tsigel’nitsky, 1990].
The seasonal variation of temperature in troposphere, like near the ground surface, has coreless character. At all the altitudes the warmest month is January, the coldest are August and September. The temperature distribution in whole boundary layer has positive asymmetry, which is especially pronounced in winter. The reason for this is warm advection during strengthening of meridional circulation and cyclone invasions into the continent [Tsigel’nitsky, 1982].
The vertical distribution of wind is characterized by a rapid growth of wind speed in the lower inversion layer due to diminution of friction. Maximum speed is reached at the lower boundary of the isothermal layer and wind here is actually stronger than geostrophic wind. This phenomenon is called « meso-jet stream » [Vorontsov, 1967]. The origin of this wind is due to katabatic forces. During cyclonic weather situations, the speed and direction of the wind is highly variable and depends on the station’s position relative to the center of cyclone. Seasonal variations of wind speed in the boundary layer are characterized by higher values in winter due to both stronger inversion and more frequent cyclone invasions [Averianov, 1990; Tsigel’nitsky, 1982; Voskresensky, Tsigel’nitsky, 1985].
The height of boundary layer in central Antarctica is determined as 1) the mean altitude at which wind rotation stops (for dynamic boundary layer) and 2) the upper boundary of surface inversion in winter or elevated inversion in summer (thermal boundary layer). At Vostok the height of the dynamic boundary layer is 660–670 m in winter and 1250–1400 m in summer and that of thermal one is 610–650 m in winter and 1500–1800 m in summer (see Table 5 from [Tsigel’nitsky, 1982]).
Precipitation and water vapor in central Antarctica.
The influence of the main factors governing formation and precipitation of atmospheric moisture (humidity and temperature of air masses, atmospheric circulation) is controlled in * Absolute values of these changes are likely too high because of inhomogeneity in the series of balloon-sounding data (see Chapter « Experimental data », section II.1).
Antarctica by the elevation of the cold ice sheet surface, the distance from the ocean, and the location in relation to the major cyclonic paths. Different combinations of precipitation-forming factors with morphometric features of the ice sheet preventing moist oceanic air penetration into the interior of the continent are favorable for formation of three types of precipitation. They are: snow from clouds, ice crystals from clear sky and rime on the surface (see the review in [Averianov, 1990; Petrov, 1975; Schverdtfeger, 1987]).
According to the ideas formed in the first half of the last century [Shumsky, 1955] the conditions of formation, growth and precipitation of ice crystals in the atmosphere, as well as their forms and sizes, are related to temperature and humidity of air. The necessary condition of crystal formation and growth is supersaturation of air by water vapor which can be reached mainly by cooling. The nuclei of condensation are usually marine aerosols brought by marine air masses [Hogan, 1997; Golubev, 2000].
The two main types of atmospheric crystals are lamellar (growing to the direction of basic plane) and columnar (growing to the direction of main crystal axis). Both of them have a great variety of sub-types. During the snowfalls different sub-types (and types) are usually aggregated. Columnar crystals are often smaller than lamellar ones. Besides, there are two other types of crystals: needle-shaped and grains. It is assumed that needles are growing in the direction of a secondary axis and thus can be considered as asymmetric lamellar crystals. Snow grains, sleets, are formed as a result of supercooled water freezing on the surface of ice crystals [Averianov, 1990].
High supersaturation of cold air leads to the formation of crystals with complex shapes, like stars or dendrites. Low supersaturation leads to more regular shapes. The colder is the air, the higher is the ratio of the columnar crystals, and the smaller are the crystals of all the types [Averianov, 1990; Bromwich, 1988; Golubev, 2000].
In central Antarctica, up to 98 % of the total precipitation is formed by columnar crystals with a typical length of 0.025–0.6 mm and thickness of 0.01–0.08 mm. Most of the snow falls from As and Ac from the height of 1000–3000 m above ice sheet surface [Averianov, 1990].
The mean total water content of the atmosphere in the vicinity of Vostok (for clear sky conditions) for the period 1977–1981 is 0.34 mm. Seasonal changes of this parameter is comparatively simple and closely related to those of air temperature: from 0.17–0.19 in winter to 0.73–0.74 in summer (Fig. 5) [Burova et al., 1990]. The seasonal variability of atmospheric water content is confirmed by satellite observations on total water vapor content [Miao et al., 2001]. At the same time, the inter-annual variability of this parameter is practically absent. In particular, during the period 1960–1985 the mean total summer water content at Vostok did not changed while at most coastal stations it increased, which is related to warming observed during this period [Burova, Voskresensky, 1990]. At the same period of time, the precipitation rate has increased by about 5 %, which is explained by intensification of cyclonic activity [Bromwich, Robasky, 1993].
Fig. 5. Seasonal variation of total water content of atmosphere above Vostok (in kg m-2) during the days with clear sky (0–3 balls): 1–5 – correspondingly, 1977–1981, 6 – mean value for 5 years (from [Burova et al., 1990]).
Snow from clouds falls during passage of cyclones over the ice cover, so it is frequent at the coast and over the lower part of the ice sheet slope. Cyclonic precipitation at the coast forms in alto-stratus, strato-nimbus and less frequently in stratus clouds. The most frequent and intense snowfalls occur over the areas where local altitude does not exceed the level of condensation, i.e., 800–2000 m above sea level [Averianov, 1990; Aleksandrov et al., 1991; Kotlyakov, 1961]. About 20 % of Antarctica lies below this height but this part of the continent receives roughly one half of all the snow deposited on the whole ice sheet [Averianov, 1990]. To the central parts, the frontal clouds carrying precipitation penetrate quite rarely. And even if it happens, the clouds are already depleted in moisture, so the snowfalls are less intensive than over the coast. The number of days with snow from clouds per year in Antarctica is as follows: about 150 on the coast, 200 over the lower part of the slope and 25–50 in the central part [Bryazgin et al., 1990; Spravochnik po klimatu, 1977].
The question about the source of moisture feeding the area of Vostok is still under discussion. Results of General Circulation Models simulations suggest that most of the precipitation (40 %) at Vostok comes from low and middle latitudes of the Indian ocean, about 10–20 % does from the Atlantic and the Pacific, as well as from Antarctic seas, and the rest (less than 10 %) comes from the Antarctic ice shelves. The contribution of local moisture is negligibly small [Delaygue et al., 2000]. On the other hand, experimental data lead to another conclusion. The results of chemical analyses of snow cover imply that Vostok source area is in the Pacific sphere of influence [Averianov, 1969]. The same conclusion is achieved when analyzing the distribution of isotope composition of surface snow in Antarctica: on the diagram of snow isotope composition versus mean site annual temperature Vostok Station is situated on the continuation of Patriot Hills–South Pole line (Pacific sector) and aside from lines Mirny– Komsomolskaya and Dumont-d’Hurville–Dome C (Indian sector) [Ekaykin et al., 2001] (see also Fig. 8 in this work).
The precipitation of tiny ice crystals under clear sky conditions is typical for the anticyclonic weather which is dominant in the interior parts of the continent. This phenomenon is sometimes called « ice crystals » or « diamond dust ». Deposition of these ice crystals is constantly observed at the high-latitudinal Antarctic stations: 247 days per year at Vostok [Averianov, 1972] and 316 days per year at Plateau [Schwerdtfeger, 1987].
Table of contents :
INTRODUCTION
General characteristic of the area of study
The aims of the study
I. METEOROLOGICAL REGIME AND ISOTOPE COMPOSITION OF PRECIPITATION: REVIEW OF LITERATURE
I.1. Meteorological regime and precipitation formation in central Antarctica
I.1.1. Atmospheric circulation in the southern hemisphere
I.1.2. Surface temperature inversion at Vostok and wind regime
I.1.3. Precipitation and water vapor in central Antarctica
I.2. Isotope composition of precipitation and its relation to the conditions of formation: Theoretical considerations and empirical data
I.2.1. Theoretical basis of the relationship between isotope composition of precipitation and air temperature: Simple isotope models and GCMs
I.2.2. Empirical estimations of relationship between isotope composition of precipitation and temperature
Direct comparison of isotope composition and air temperature
Borehole thermometry
Use of melt layers
Correlation with snow accumulation rate
Data on gas inclusion
Isotope composition of trapped air
I.3. Factors influencing the relationship between snow isotope composition and surface air temperature
I.3.1. Moisture source conditions
I.3.2. Seasonality of precipitation
I.3.3. Microphysical conditions of precipitation formation
I.3.4. Difference between condensation and surface air temperature
I.3.5. Glaciological factors
I.3.6. Post-depositional processes
I.4. Conclusion of Chapter I
II. METHODS AND EXPERIMENTAL DATA
II.1. Experimental data
II.1.1. Meteorological data
II.1.2. Balloon-sounding data
II.1.3. Snow accumulation rate
II.1.4. Isotope composition of snow
II.2. Field works
II.2.1. Stratigraphic studies in pits
II.2.2. Snow sampling in pits
II.2.3. Sampling of precipitating and blowing snow
II.2.4. Construction of new snow accumulation-stake network
II.2.5. Snow surface leveling
II.3. Laboratory measurements
II.3.1. Isotope measurements
II.3.2. Measurements of beta-radioactivity
II.3.3. Measurements of liquid conductivity
II.4. Conclusion of Chapter II
III. METEOROLOGICAL CONDITIONS OF SNOW FORMATION
III.1. Contribution of different precipitation types in total precipitation amount
III.2. Temperature of condensation
III.3. Conclusion of Chapter III
IV. SPATIAL VARIABILITY OF SNOW ISOTOPE COMPOSITION: PLAYGROUND OF WIND
IV.1. Mega-dunes and micro-relief
IV.2. « Meso-dunes » signature in spatial and temporal series of snow build-up
IV.3. Relief-related oscillations in temporal isotope series Post-depositional changes of snow δD content in the past
IV.4. Conclusion of Chapter IV
V. MODERN TEMPORAL VARIABILITY OF SNOW ISOTOPE COMPOSITION
V.1. Seasonal variability of isotope composition of precipitation
V.2. Temporal variability of isotope composition and snow accumulation rate in the vicinity of Vostok Station over the last 50 years
V.3. The deuterium content – temperature slopes
V.4. Short-term variations of isotope composition in deep ice cores from Vostok
V.5. Conclusion of Chapter V
VI. CHANGES IN TEMPERATURE AND SNOW ACCUMULATION RATE AT VOSTOK STATION OVER THE PAST 200 YEARS
VI.1. Series of isotope composition and snow accumulation rate from deep pits
VI.2. 50-year cycle in changes of accumulation and isotope composition: A teleconnection between central Antarctica and tropical Pacific?
VI.3. Secular trends of accumulation and isotopes at Vostok: Climate or mega-dunes?
VI.4. 200-year accumulation and isotope tendencies at other East Antarctic sites
VI.5. Conclusion of Chapter V
CONCLUSION AND PERSPECTIVES
REFERENCES
ANNEXES
