Mesoscale Meteorological Conditions in Dronning Maud Land, Antarctica, during Summer: A Qualitative Analysis of Forcing Mechanisms

Transcription

1 2348 JOURNAL OF APPLIED METEOROLOGY Mesoscale Meteorological Conditions in Dronning Maud Land, Antarctica, during Summer: A Qualitative Analysis of Forcing Mechanisms RICHARD BINTANJA Institute for Marine and Atmospheric Research Utrecht, Utrecht University, Utrecht, Netherlands (Manuscript received 23 June 1999, in final form 2 February 2000) ABSTRACT Detailed concurrent summer observations of surface meteorological variables, high-resolution boundary layer profiles, and upper-air profiles were carried in the intermediate region between the high plateau and the coast in Dronning Maud Land, Antarctica, a region that includes several blue ice areas. The main goal is to find out to what extent the flow is influenced by 1) slope-inversion pressure gradients (katabatic force), 2) synoptic pressure gradients (geostrophic wind), and 3) thermal gradients (thermal wind), and how this compares with the situation near the coast. Both synoptic and katabatic forcings generally induce easterly winds in the boundary layer, with a clear diurnal cycle in their relative importance. Westerly surface winds occur less often than at nearby coastal stations Halley and Neumayer, indicating that the influence of synoptic and thermal effects decreases toward the interior of Antarctica. Hence, katabatic forcing dominates, in particular during clear-sky conditions. Under weak synoptic forcing, the boundary layer over the main blue ice area is susceptible to localscale circulations, which produces noneasterly winds at the surface. This result is caused mainly by reduced cooling over blue ice, which renders the katabatic force weak or absent, and by the typical topographic setting. 1. Introduction An important feature of the Antarctic climate is the predominant downslope katabatic wind. Continuous cooling of the air over the sloping ice sheet surface induces a downslope pressure gradient that causes inversion winds of record strength and extreme constancy (e.g., Wendler et al. 1997). Over steep slopes, the negative buoyancy of the air induces strong winds in an approximately downslope direction. Over less steep slopes, the Coriolis force becomes important and turns the wind in a more cross-slope direction. Despite former distinctions between the two cases, I will henceforth refer to all slope-inversion pressure gradient winds as katabatic winds. In the coastal regions, the winds in the boundary layer are additionally influenced by 1) the synoptic pressure gradient induced by the weather systems migrating around Antarctica the low pressure belt is associated with baroclinicity of the coastal atmosphere and 2) the thermal wind associated with the spatial (ice sheet sea) temperature gradient. It is well known that the effect of both mechanisms decreases toward the interior of the continent, because 1) the transient cyclones, associated with the circumpolar trough, Corresponding author address: R. Bintanja, Institute for Marine and Atmospheric Research Utrecht, Utrecht University, P.O. Box 80005, 3508 TA Utrecht, Netherlands. r.bintanja@phys.uu.nl remain confined to the coastal regions where they strongly affect the geostrophic winds (Heinemann and Rose 1990) and 2) the large-scale temperature gradient above the boundary layer is largest at the ice sea boundary (Kottmeier and Stuckenberg 1986). In order to gain fundamental insight into the forcing of the flow over Antarctica and for validation of regional- or continentalscale meteorological models it is important to find out at what rate these effects decrease toward the interior and thus how strongly each of these mechanisms controls the (surface) wind field over Antarctica s interior slopes, by using direct observations. The boundary layer flow over the sloping snow plains is undoubtedly influenced by the presence of steep topography and blue ice areas in the Heimefrontfjella. Hence, local effects are superimposed on the characteristics of the otherwise undisturbed mesoscale flow. In particular, barrier winds may develop in the vicinity of the mountain range. To assess these local effects, detailed meteorological measurements were taken at seven sites spread across the central Heimefrontfjella, ranging from relatively undisturbed snow-covered locations to sheltered blue ice area sites where local circulations may develop under conditions of weak largescale forcing. A comparison of meteorological variables between these sites may provide clues about the nature of the local forcing mechanisms. This is important, because the surface flows strongly affect the surface mass balance and the climate sensitivity of blue ice areas American Meteorological Society

2 DECEMBER 2000 BINTANJA 2349 The flow and thermal regime of the atmospheric boundary layer profoundly influence the surface mass balance of the Antarctic ice. In particular in regions with steep topography and blue ice areas the divergence of snowdrift and the sublimation of ice/snow are important mass balance components (Takahashi et al. 1988; van den Broeke and Bintanja 1995a; Bintanja and van de Broeke 1995a; Bintanja 1999). The surface climate plays an important role in the formation and climate sensitivity of blue ice areas (Bintanja and van den Broeke 1995b; van den Broeke and Bintanja 1995b). Bintanja (1998) stresses the importance of the sublimation of drifting snow for the surface mass balance of the Antarctic ice sheet. Therefore, a detailed assessment of the surface mass balance of the entire Antarctic ice sheet depends crucially on a proper understanding of the surface climate. One way to study the surface climate of the Antarctic is to use atmospheric models. The (surface) climate of the Antarctic has been simulated in many numerical models. These range from local diagnostic analyses of the boundary layer wind system (e.g., Parish 1982) and continental-scale 3D primitive equation models (e.g., Parish and Bromwich 1991) to global-scale general circulation model analyses (e.g., Genthon 1994; van den Broeke et al. 1997). These studies have provided valuable insight into the characteristics of the large-scale flow over the Antarctic continent and have supplemented the insight obtained from scattered observations. However, it is essential that numerical models are validated with observations (in particular with radiosonde data), but unfortunately such measurements are scarce, especially in the interior of Antarctica. The measurements performed at inland stations such as Svea may prove valuable for model validation; data from Svea gathered in were used by van Lipzig et al. (1998) to validate the limited area model that they developed to study the Antarctic climate. The other approach for studying Antarctica s surface and boundary layer characteristics is to use direct observations. Automatic weather stations (AWS), for instance, provide year-round observations of the near-surface meteorological conditions. The number of automatic weather stations for long-term monitoring of meteorological variables has increased significantly over the last few decades. However, AWSs provide only surface observations. Most of the regular weather observations, including upper-air soundings, are made at coastal stations, which are relatively accessible. Supplementary data from inland locations are therefore needed to complete the observational picture of the meteorological conditions over the entire continent. There have been only a few observational studies of upper-air flow in the intermediate region between the inland plateau and the coastal areas (e.g., Ohata et al. 1985; Sorbjan et al. 1986; Kodama et al. 1989; van den Broeke and Bintanja 1995a). Evidently, much can yet be learned from measuring and studying upper-air data in the interior of Antarctica (e.g., Kottmeier and Fay 1998). Here I will present, analyze, and discuss meteorological data measured near the inland station Svea, western Dronning Maud Land, which is located in the intermediate region between the coast and the high plateau. I will discuss various flow characteristics in relation to the various forcing mechanisms. The advantage of basing a study on this region is that the prevailing meteorological characteristics, including the meteorology of blue ice areas, have been relatively well studied (Jonsson 1992, 1995; Bintanja and van den Broeke 1995a; van den Broeke and Bintanja 1995a; Bintanja et al. 1997). These studies form the basis for further exploration of the meteorological conditions in this region, as reported in this paper. Svea is located only about 300 km from the projected deep-drilling site in Dronning Maud Land, which will be exploited within the framework of European Project on Ice Coring in Antarctica (EPICA) during the coming years. Interpretation of ice core content relies to a large extent on knowledge of the meteorological conditions around the region of the drilling location. The data gathered at Svea will contribute to our understanding of this important aspect of ice core interpretation. 2. Location and experimental setup Data presented in this paper were obtained during the Swedish Antarctic Research Programme (SWEDARP) field campaign in the austral summer (Bintanja et al. 1998). The research area was the Swedish research station Svea (74 34 S, W), located in the Scharffenbergbotnen valley in the Heimefrontfjella, and its surrounding snow plains and blue ice areas. Svea is located on the coastward side of the mountain range, at an altitude of 1250 m above mean sea level (MSL), approximately 350 km from the edge of the Riiser-Larsen ice shelf (Fig. 1). The Heimefrontfjella marks the western end of the chain of mountain ranges in Dronning Maud Land, which runs approximately parallel to the coastline and constitutes a massive obstacle for the coastward ice flow as well as for the downslope atmospheric flow. The Heimefrontfjella harbors many small blue ice areas, of which the one in the Scharffenbergbotnen Valley is the largest (3 6km 2 ). The experimental area in the central Heimefrontfjella is depicted in Fig. 2. The field campaign was centered around the blue ice area in the Scharffenbergbotnen valley near Svea. Radiosonde and cable-balloon soundings were performed near Svea, and meteorological profile stations were set up at seven predetermined locations. Four stations were placed on snow slopes (3, 4, 6, and 7), at which the flow was expected to be relatively undisturbed. The other three stations were placed on blue ice adjacent to steep topography (1, 2, and 5), where the flow was expected to be strongly influenced by local disturbances. Site 7 and to a lesser extent site 6 represent

3 2350 JOURNAL OF APPLIED METEOROLOGY FIG. 1. Map of (left) the Antarctica continent and (right) western Dronning Maud Land, showing the location of Svea in the Heimefrontfjella. Bold lines in the right panel represent mountain ranges. the relatively undisturbed conditions of the high-altitude plateau Amundsenisen. Sites 3 and 4 represent conditions of the lower plateau Ritscherflya. Given the fact that the prevailing winds are from easterly directions, sites 6 and 7 are located upwind from the Scharffenbergbotnen valley, and site 4 is situated on its downwind side. The meteorological experiment can be considered as the follow-up of a similar experiment in ; during this experiment meteorological stations were placed at sites 1 and 3 (and at 5 other sites) and cableballoon soundings were carried out near Svea (Bintanja and van den Broeke 1995a; van den Broeke and Bintanja 1995a). The seven meteorological stations carried sensors for temperature, humidity, and wind speed at five (sites 1 and 3) or three heights above the surface. In addition, we measured all radiation components (incoming and reflected shortwave and incoming and emitted longwave), subsurface temperatures at five depths using temperature probes, and wind direction. All sensors were sampled every 2 min, except at sites 6 and 7 where sensors were sampled every 5 min. After quality checks and calibration (before, during, and after the expedition) a complete dataset of 37 days was obtained (28 December 2 February), during which period all meteorological stations worked continuously. Except for the first two days, radiosonde soundings were performed during the entire period using the Vaisala, Inc., MW15 radiosonde system with global positioning system wind finding. This system is used regularly at standard meteorological stations around the world. A balloon was launched twice a day until 19 January (1200 and 0000 h UTC), after which the 1200 sounding was discontinued. The radiosonde balloons usually reached the stratosphere, yielding tropospheric profiles of temperature, humidity, wind speed, and wind direction. Unfortunately, no surface wind data near the launch site of the radiosonde were available to feed into the radiosonde software system at the start of a sounding. During the sounding, the internal software applies smoothing techniques to the raw measurements in order to obtain smooth profiles. Because of the erroneous default surface values of wind that were perforce entered, profiles of wind in the lowest few hundred meters may be inaccurate; comparison with cable-balloon data shows that easterly and southerly winds in the lowest few hundred meters are underestimated by 1 2ms 1. However, this has little effect on the qualitative analysis presented here. Additional high-resolution cable-balloon soundings could be performed only when the surface winds did not exceed 10 m s 1. In good weather, a sounding was performed every 3 h. In total 56 soundings were performed near Svea and 9 near site 1 on blue ice. The soundings near Svea were taken in three periods: 4 5, 6 8, and January. Normally, soundings went to a height of 800 m above the surface. More details of the measurements and calibration procedures and other information about the experimental set-up can be found in Bintanja et al. (1998). Table 1 shows some characteristics of the various observations. Radiosonde sounding data obtained at Halley and Neumayer stations (Fig. 1) were used to interpret the data measured at Svea in a broader perspective and to evaluate horizontal gradients. Soundings at these stations are done daily on a routine basis, and form part

4 DECEMBER 2000 BINTANJA 2351 FIG. 2. Detailed map of the central Heimefrontfjella, denoting location of Svea and of the seven meteorological stations. Altitude contours are drawn every 100 m. Light gray and dark gray areas represent exposed rock and moraine, respectively. Adapted from the map Heimefrontfjella Nord (Maudheimvidda), Sheet D8, Norsk Polarinstitutt (Oslo 1988), 1: Notice that in some areas away from nunataks the elevation contours are very inaccurate, in particular near site 7 that is actually at an altitude of 2100 m MSL and on a relatively steep slope. Average wind direction at each site is indicated by arrows, whose length indicates the average wind speed.

5 2352 JOURNAL OF APPLIED METEOROLOGY TABLE 1. Overview of the various observations that were performed during the experiment. Observation Location Sampling Period Surface layer profiles of wind, temperature and humidity, wind direction, radiation budget, subsurface temperatures Sites 1, 2, 3, 4, and 5 Sites 6 and 7 every 2 min every 5 min 28 Dec 2 Feb 28 Dec 2 Feb Cable-balloon soundings Svea every 3 h 4 5 Jan 6 8 Jan Jan Radiosonde soundings Svea twice daily 29 Dec 19 Jan once daily 20 Jan 5 Feb Synoptic observations Svea every 3 h 28 Dec 2 Feb of the standard meteorological observations carried out over a longer period. 3. General meteorological description The meteorological conditions in the Heimefrontfjella region consist of a mixture of 1) quite distinct coastal weather influences, such as the high variability on a timescale of a few days due to the influence of eastward migrating coastal low pressure systems, and 2) the more constant inland climatic features such as the quasi-continuous katabatic winds. Parish (1982) shows that high variability on a day timescale is associated with coastal influence due to transient cyclones, and that this typical variability is largest in the coastal areas. Jonsson (1995) analyzed annual meteorological variables near Svea and found a strong variability on a 3 4 day timescale. He could link this to synoptic low pressure systems passing along the coast and an associated periodic build-up and drainage of cold air from the interior. Figure 3 shows the temporal variation of the basic meteorological parameters at three selected sites during the measuring period. In particular the variability in relative humidity reflects the different origins of the air masses advected toward the Heimefrontfjella: moist air from the coast and dry air from the inland. The exceptionally warm period at the beginning of January with occasional above-zero temperatures at all sites and extensive surface melt over the blue ice areas was followed by a severe storm on 8 10 January. This storm, with hourly mean winds up to 20 m s 1 at site 2, advected cold air toward the region. Temperatures remained quite low for the rest of the period, with a brief warm interruption near the end of January. Winds over the snow sites were generally from the east, whereas winds over blue ice at site 1 blow regularly from the west. Interestingly, these westerly winds at site 1 occur simultaneously with more southeasterly winds over snow for reasons discussed later. These observations in the Heimefrontfjella are put into a broader perspective in Fig. 4, which shows European Centre for Medium-Range Weather Forecasts (ECMWF) analyzed fields of surface wind, sea level pressure, and surface air temperature averaged over the measuring period. A synoptic low pressure area is centered at 20 E off the coast of Dronning Maud Land. Interestingly, its presence produces two distinct belts of maximum easterly winds, one just off the coast that is entirely synoptically forced and one over the steepest slopes of the ice sheet where katabatic and synoptic forcing interact to produce strong easterly winds. The Heimefrontfjella is located at the western tail of the latter region with moderate easterly surface winds. Furthermore, Fig. 4c shows that the study area is located in the coastal region with strong cross-shore gradients in surface air temperature. Table 2 shows average values of the various meteorological variables for the entire measuring period. Differences in temperatures between the sites are caused by differences in elevation and by the surface type: blue ice is heated more effectively in daytime because of its lower albedo. For instance, temperatures over blue ice at site 1 are 2.5 C higher than at site 3 at the same elevation. Figure 5 depicts the average potential temperature at the various sites, along with the average potential temperature profile as obtained from the radiosonde soundings. It is obvious that the blue ice sites are warmer than their snow-covered surroundings because of diabatic heating. The snow sites are significantly colder than the free atmosphere at the same elevation, indicating that the katabatic force B in the downslope direction induces a downslope katabatic flow at these sites. Here, B is defined as g B sin, (1) 0 where is the slope angle, 0 is a reference potential temperature and is the deviation of the surface potential temperature from the background atmospheric temperature (also referred to as the temperature deficit). The average temperature deficit is largest at the highelevation sites 7 and 6 and somewhat smaller at the lowlevel sites 3 and 4. The magnitude of the surface slope is largest at sites 6 and 7, which further indicates that katabatic winds will be most pronounced at sites 7 and 6. The maximum (nocturnal) temperature deficit is 9 C at site 7, where we may expect the strongest katabatic winds. In fact, sites 6 and 7 experience katabatic forcing for most of the day on average as the temperature deficit just reaches zero in the afternoon (as opposed to sites 3 and 4). This is independently confirmed by calculating

6 DECEMBER 2000 BINTANJA 2353 FIG. 3. Temporal variation in daily mean (a) temperature, (b) relative humidity, (c) wind speed, and (d) wind direction at sites 1, 3, and 7 during the measuring period. Values were measured at 2 m above the surface.

7 2354 JOURNAL OF APPLIED METEOROLOGY FIG. 4. ECMWF analyzed fields of (a) sea level pressure, (c) surface air temperature at 2 m, and (b) zonal wind speed at 10 m averaged over the measuring period (28 Dec to 2 Feb). In (b) negative values indicate easterly winds. The marks the location of Svea.

8 DECEMBER 2000 BINTANJA 2355 TABLE 2. Meteorological data of the seven sites averaged over the 37-day measuring period from 28 Dec 1997 to 2 Feb Values of the directional constancy are based on half-hourly wind values. Site 1 Site 2 Site 3 Site 4 Site 5 Site 6 Site 7 Surface type Blue ice Blue ice Snow Snow Blue ice Snow Snow Elevation (m MSL) Air temperature 2 m ( C) Relative humidity 2 m (%) Specific humidity 2 m (g kg 1 ) Wind speed 2 m (m s 1 ) Wind direction ( ) Directional constancy Albedo sensible heat fluxes from the measured vertical temperature profiles at each site, which indicate that stable stratification and consequent cooling of the surface air persist throughout the day (on average). This is an unusual situation, because solar heating apparently is not able to warm the surface to create a well-mixed boundary layer in the afternoon as was found in Adélie Land by Kodama et al. (1989). The katabatic forcing is weak or absent over blue ice, because surface potential temperatures are comparable with the background temperature (i.e., the free atmosphere temperature at the same altitude). This means that other forcing mechanisms may dominate the boundary layer flow over blue ice areas, as will be discussed in later sections. FIG. 5. Average potential temperature at the seven meteorological stations. The horizontal bars represent the average daily range in potential temperature at each site. The solid line is a linear fit (between 2 and 3 km) to the average potential temperature profile (dashed line) derived from radiosonde soundings. The average wind vector at each site is depicted in Fig. 2. At the high-elevation sites 6 and 7 where the surface is relatively steep the winds are from the southeast, which is in the downslope direction with a slight deflection to the left with respect to the orientation of the large-scale slope ( 135 ). At the lower sites the winds blow from the east-northeast (except for site 1 where the direction of the wind is largely dictated by the topography of the Scharffenbergbotnen valley), which may be attributed to the fact that the easterly boundary layer winds are forced to flow around the high-elevation area of the Sivorgfjella. Moreover, the surface inclination is less, which increases the influence of Coriolis force relative to that of the katabatic force, allowing a greater angle between slope direction and wind vector. Average winds are indeed strongest on the high-elevation sites 6 and 7, at which the near-surface cooling and hence the katabatic wind forcing are strongest. The strongest gusts and the highest maximum wind speeds occur over blue ice because of their irregular topographical surroundings, which cause tunneling and intermittent vortex shedding. Wind directional constancy (derived from half-hourly values), defined as the ratio of mean vector wind to mean wind speed, was quite low over blue ice at sites 1 and 2, indicating that the wind was relatively variable (Table 2). Directional constancy is highest at the relatively undisturbed high-altitude sites 6 and 7, although not as high as over the other sloped surfaces of the east Antarctic ice sheet where average values over 0.9 have been recorded (Kodama et al. 1989; King 1989). The value of the wind directional constancy at the various measuring locations correlates well with the observed average temperature deficit as shown in Fig. 5. This indicate that boundary layer flow at the snow sites is controlled mainly by the katabatic force inducing unidirectional winds, and that at the blue ice sites other (local) forcings influence the flow as well. Compared to the summer , temperatures at sites 1 and 3 were about 1 C higher, relative humidities were about the same, and the average wind speed was somewhat lower (Bintanja and van den Broeke 1995a). The temperature difference between sites 1 and 3 is about 0.5 C smaller than in Synoptic observations

9 2356 JOURNAL OF APPLIED METEOROLOGY of cloud amount and type exhibit a variability indicative of coastal influence, with transient cyclones occasionally producing overcast conditions and some precipitation. The average cloud amount was 0.38, not significantly higher than in (0.35). Clouds influence the surface flow, increasing incoming longwave radiation and thereby reducing surface cooling by as much as 80 W m 2. I will show in section 5 that the strongest katabatic winds occur in clear-sky situations. To obtain a picture of the temporal changes in the upper-air variables, I used the radiosonde sounding profiles taken at Svea to construct time-height plots of the various variables. Figure 6a shows the variation of potential temperature. Clearly, isotherms roughly follow the pattern of the surface temperature variations: a warm period in the first part of January at all levels, followed by a colder period of about 20 days and, at the end of the measuring period, a significant rise in temperatures. In fact, in the end of January temperatures above 5 km were nearly as high as in early January, whereas surface temperatures remained much lower. The reason for these differences can be found by inspecting the zonal and meridional wind components (Figs. 6b and 6c). Winds are generally from the (north)east at lower levels, and from the (south)west at higher levels (except for periods with strong easterly geostrophic winds, when winds are easterly throughout the column). Upper-air westerlies are associated with the inflow of relatively warm and moist air, whereas low-level easterlies are identified with the outflow of cold interior air (Schwerdtfeger 1984; King and Turner 1997). When the synoptic pressure gradient forces westerly winds at lower levels, such as at the beginning of January, relatively warm air can penetrate to the surface. The westerly geostrophic forcing then opposes the westward-directed katabatic forcing, resulting in weak westerly boundary layer winds. Also, the presence of a nearby high pressure area causes low amounts of cloud. These circumstances allow further warming by solar radiation, especially at surfaces not covered by snow, for example, blue ice and rock. Later on, westerlies did not penetrate to such low levels, which explains why warming was then confined mainly to high altitudes. Easterly geostrophic winds, caused by the presence of low pressure synoptic systems near the coast in the vicinity of the circumpolar trough, usually coincide with fierce easterly winds such as the stormy period of 8 10 January. The katabatic winds in the boundary layer are significantly enhanced by the synoptic pressure gradient. The interaction between a strong synoptic pressure gradient and the katabatic force is known to produce strong boundary layer winds (Murphy and Simmonds 1993). Interestingly, the geostrophic wind changed only marginally with height, illustrating the predominantly barotropic nature of the event, in accordance with the events studied by Murphy and Simmonds (1993). The storm of 8 10 January was a widespread event, since Halley and Neumayer stations also reported gale-force boundary layer and upper-air winds. ECMWF-analyzed surface pressure data (not shown) exhibit a deep trough (minimum under 960 hpa) in the Weddell Sea region at 40 W, producing strong northeasterly geostrophic winds in western Dronning Maud Land. Jonsson (1995) was able to link the strongest wind events in the Scharffenbergbotnen valley (near site 1) during winter 1988 to cyclones passing close to the coast, which thereby enhanced significantly the low-level katabatic winds. The relatively cold period in the middle of the measuring period can be attributed to the frequent occurrence of cold easterly geostrophic winds in that period. Figure 7 shows average upper-air profiles at Svea and at the coastal Halley and Neumayer stations. The troposphere is stably stratified with an average potential temperature lapse rate of 5 K km 1. Interestingly, temperatures at Svea are only slightly lower (maximum 2 C) than at Halley and Neumayer despite the inland location of Svea. The static stabilities are about the same for the three locations except in lowest layers where the stability at Svea is reduced as a result of the stronger shortwave heating over nearby blue ice areas and exposed rocks. The reduced stability of the surface layer over blue ice areas and the associated upward sensible heat fluxes warm the boundary layer from below (Bintanja and van den Broeke 1995a). Relative (and also specific) humidities are slightly lower at Svea (not shown) because of its inland location, where the relatively dry winds from the inland plateau occur more frequently and where advection of moist air from the coast occurs less frequently. The average wind at Svea turns clockwise from northeast in the boundary layer to southwest in the upper troposphere, equivalent to the form of the wind profiles at Halley and Neumayer. A distinct wind maximum is present at about m above the surface, a probable signature of the katabatic wind system. However, the boundary layer wind speeds are probably erroneous, as explained above, and the wind speed maximum will probably be at lower altitudes. A detailed comparison of the boundary layer wind profiles measured by the radiosonde and the cable-balloon system indicates that the wind speed maximum was roughly m lower than indicated by the radiosonde measurements due to the problems described in section 2. Van den Broeke and Bintanja (1995a) analyzed cable-balloon profiles from the field experiment and found no wind maximum below 500 m above the surface. The apparent high altitude of the wind maximum at Svea may be caused by its topographic setting, as the irregular topography of the mountainous region of the Heimefrontfjella increases vertical mixing, enabling strong katabatically forced winds to penetrate upward. An alternative explanation is the development of barrier winds along the Heimefrontfjella escarpment. 4. Geostrophic and thermal winds The meridional and zonal components of the geostrophic and thermal wind were evaluated with the help

10 DECEMBER 2000 BINTANJA 2357 FIG. 6. Time height plots of (a) potential temperature ( C), (b) zonal wind speed (m s 1 ), and (c) meridional wind speed (m s 1 ) during the measuring period as derived from radiosonde soundings. Easterly and northerly winds (gray) are taken to be positive.

11 2358 JOURNAL OF APPLIED METEOROLOGY of radiosonde data from the coastal station Halley (450 km to the west of Svea) and Neumayer (480 km to the north) (see Fig. 1). Above the boundary layer, the flow is in approximate geostrophic balance and the geostrophic wind can be evaluated from 1 p fug y and (2) 1 p f g, x (3) where u g and g are the zonal and meridional components of the wind (taken to be positive in the westward and southward direction, respectively), f is the Coriolis parameter, is the air density, and p is air pressure. The zonal component of the geostrophic wind is shown in Fig. 8. As expected, it agrees well with the observed zonal winds (see Fig. 6b) above the boundary layer. During the stormy period of 8 10 January, easterly synoptic forcing was virtually independent of height, indicating barotropic conditions. The winds were strongest in the boundary layer, though, due to the interaction of the pressure gradient force and the katabatic force near the surface (and possibly by the occurrence of barrier winds). Figure 9 shows average vertical profiles of geostrophic and observed winds. The clockwise wind turning with height is fairly constant and indicates advection of cold air toward the Heimefrontfjella region at all levels on average. It is obvious that whereas the winds in the free atmosphere are clearly geostrophic, the boundary layer winds are not (by definition). Actual wind speeds in the boundary layer are higher than the geostrophic winds (approximately the downward extrapolation from the geostrophic winds aloft). This might indicate that the flow below about 2500 m (1300 m above the surface) is influenced by katabatic forcing at the surface, which enhances its strength and turns it in a more downslope direction. However, it is conceivable that barrier winds induced by the steep orography of the Heimefrontfjella contribute to the low-level wind strengthening (Schwerdtfeger 1984), in particular since the depth of the enhanced wind layer coincides with the height of the barrier. A boundary layer depth of 1300 m is clearly extremely deep for typical Antarctic conditions, because observations and modeling studies generally indicate much smaller summer boundary layer depths, namely m (e.g., Sorbjan et al. 1986; Kodama et al. 1989; Fortuin and Oerlemans 1990). An additional explanation for this is that on average the air flows down from the high plateau Amundsenisen to the lower plateau Ritscherflya over a short horizontal distance, which causes the boundary layer profiles to deviate from their undisturbed forms. Cooling of the air overlying the ice sheet introduces a thermal contrast between the air over the continents and the air over the adjacent ocean. The associated thermal wind relation reads FIG. 7. Average profiles of potential temperature, wind speed, and wind direction at the stations Svea, Neumayer, and Halley during the measuring period as determined from radiosonde data. g g f and (4) z T x ug g f. (5) z T y The mean thermal field (cold air inland, warm air over the ocean) induces a thermal wind parallel to the ice

12 DECEMBER 2000 BINTANJA 2359 FIG. 8. Time height plot of calculated zonal component of geostrophic wind (m s 1 ) derived from the horizontal pressure gradient between Svea and Neumayer. Easterly winds (gray) are taken to be positive. sheet edge and a consequent clockwise wind turning with height. The average meridional and zonal gradients in potential temperature were relatively small (see Fig. 7a), with maximum values of about 0.6 C (100 km) 1 just above the boundary layer. Kodama et al. (1989) found a meridional temperature gradient of 2.5 C (100 km) 1 in Adélie Land during summer. This difference can be attributed to the presence of the Riiser-Larsen ice shelf, which serves to reduce the thermal contrast between the continent and the ocean in western Dronning Maud Land (van den Broeke and Bintanja 1995a). At lower levels, the diabatic heating of the nearby blue ice areas and rocks will counteract the overall cooling of the atmosphere near Svea. Figure 10 shows a time height plot of the potential temperature difference between Svea and Neumayer. Although the atmosphere over Svea is colder most of the time, relatively warm spells appear to occur frequently, reducing the average thermal contrast. By comparing Fig. 10 with Fig. 6, one can link these periods to cooling events in western Dronning Maud Land. Apparently, cooling during such periods is more severe near the coast than inland, which may be attributed to cold-air advection from the Filchner Ronne ice shelf associated with the presence of a high-pressure system in the Weddell Sea region. This relatively cold, negatively buoyant air of the Weddell Sea area will not flow easily onto the continent but will remain confined to the coastal regions. In , van den Broeke and Bintanja (1995a) also found a small thermal contrast between Svea and the coast in the lowest 1000 m during a fine weather period with westerly geostrophic winds. The outflow of cold air from the Weddell Sea region toward Halley and the ice shelves of western Dronning Maud Land during anticyclonic circulations seems to be a frequently occurring phenomenon, which systematically cools the air overlying the ice shelves, reducing the FIG. 9. Average vertical profiles of actual wind, geostrophic wind calculated from pressure gradients, and geostrophic wind calculated from the thermal wind relation (see text for further explanation). Shown are the zonal and meridional wind components and the vector mean wind speed. Easterly and northerly winds are taken to be positive. cross-slope thermal gradient and hence the thermal wind. The associated geostrophic wind components are depicted in Fig. 9. In fact, we used the calculated geo-

13 2360 JOURNAL OF APPLIED METEOROLOGY FIG. 10. Time height plot of the potential temperature difference between Svea and Neumayer ( C). Positive values (gray) indicate warmer conditions at Svea. strophic wind at 5 km and the vertical gradients of the geostrophic wind from (4) and (5) to calculate the geostrophic wind profile down to the surface. This agrees well with the geostrophic wind calculated from the pressure gradients and with the observed winds, which indicates that the observed spatial thermal gradients are in agreement with the vertical clockwise turning of the wind above the boundary layer. The average thermal wind forces northeasterly geostrophic winds in the boundary layer, parallel to the isotherms. Interestingly, the geostrophic wind seems to peak at low levels, which is indicative of a low-level geostrophically forced jet. This is caused mainly by the meridional temperature gradient, which serves to turn the southwesterly winds of the higher troposphere into northeasterly geostrophic winds at lower levels. In Adélie Land, this thermally forced low-level geostrophic jet exists semipermanently and causes very constant boundary layer winds in daytime (Kodama et al. 1989). Further inland, the spatial temperature gradient will decrease and the upper-air westerlies will become less important. The upper-air westerlies carry relatively warm and moist air toward the Antarctic mainland. Relatively cold air is exported through the low-level easterlies. The circulation is closed by continuous subsidence in the middle troposphere, forcing a westerly cyclonal vortex (Parish and Bromwich 1991). Near the coast, summertime westerlies occur quite frequently and penetrate to lower levels (King 1989). This contrasts with the situation near Svea where low-level easterlies dominate the flow. 5. Influence of katabatic forcing Consider again Fig. 5, which depicts the average temperature deficit and its daily range at all seven sites. Clearly, the high-elevation sites 7 and 6 exhibit the strongest inversions, the low-plateau sites 3 and 4 experience intermediately strong temperature deficits, and at the blue ice sites 1, 2, and 5 the average deficit is close to zero. The horizontal bars for each site denote the average daily range in the actual temperature. The strongest nocturnal stabilities are obviously found at sites 6 and 7, whereas weakly stable stratification prevails over blue ice. During daytime, sites 6 and 7 just reach neutral, while unstable conditions occur at the lower sites 3 and 4 and especially over blue ice. Evidently, the importance of the katabatic force exhibits strong diurnal and spatial variations, even when differences in surface slope are disregarded. The importance of the katabatic forcing B relative to the synoptic forcing at the surface can be quantified by evaluating the ratio (King and Turner 1997): B R, (6) f (ug g) where u g and g represent the zonal and meridional components of the geostrophic wind. Here, R represents the ratio of katabatic and synoptic forcing. Obviously, when R k 1 the katabatic force exerts dominant control over the low-level flow, which effectively becomes more or less decoupled from the geostrophically forced upper boundary layer. During times with R 0 the temperature distribution would force (weak) anabatic (upslope) flows. Generally, the surface flow is dominated by the

14 DECEMBER 2000 BINTANJA 2361 FIG. 11. (a) Average diurnal cycle of the ratio of katabatic force to synoptic pressure gradient (R) at sites 1, 3, and 7. Estimated surface inclinations at these sites are 0.001, 0.010, and 0.017, respectively. (b) Average day (1200) and night (0000) radiosonde wind speed profiles over the period 30 Dec 19 Jan (with the strong wind period 8 10 Jan excluded). synoptic-scale pressure gradient when R 1. In evaluating R, we have verified that the synoptic forcing exhibits no diurnal cycle. As a result we can use mean values of geostrophic wind components to evaluate (6). The average diurnal cycles of R at sites 1, 3, and 7 are shown in Fig. 11a. Obviously, the katabatic forcing is largely absent over blue ice at site 1, apart from a brief period during the night. During the greater part of the day the synoptic forcing and other, local forcings appear to dominate the near-surface flow at site 1. Over the snow slopes at sites 3 and 7 the katabatic forcing becomes increasingly important during the night. At site 7, the nocturnal katabatic forcing is more than an order of magnitude larger than the synoptic forcing, indicating complete katabatic control over the surface layer winds. This agrees with the work of Ball (1960), who calculated the near-surface force budget and concluded that under typical Antarctic conditions the katabatic forcing could become much greater than the synoptic forcing. At all snow sites the katabatic forcing diminishes during the day as a result of radiative heating, so the synoptic forcing can dominantly influence the flow as this is the only available forcing mechanism of sufficient strength. Thus, on the slopes of the higher plateau there is a marked diurnal cycle in the controlling mechanisms of the surface flows. As a result, there is a marked diurnal variation in wind direction similar to that described by Kodama et al. (1989). During the night, high values of R cause a flow that is practically downslope, that is, almost pure katabatic flow. In the morning, there is a smooth transition to a synoptically forced circulation with small R values, with winds blowing from the northeast (70 90 ) closely aligned with the prevailing geostrophic wind at that altitude (see Fig. 9). The mean diurnal cycle in wind direction caused by the diurnal cycle in forcing agrees with results of the theoretical approach of Kottmeier (1986), as discussed by van den Broeke and Bintanja (1995). An important question is whether the increase in the actual wind over the geostrophic wind in the lowest 1300 m of the atmosphere as shown in Fig. 9c is caused entirely by katabatic forcing or whether other effects (such as barrier winds) play a role. One way to qualitatively estimate the importance of katabatic forcing is to make use of its marked diurnal variation. We averaged all available daytime (at 1200) and nighttime (at 0000) radiosonde soundings except those in the strong wind period of 8 10 January (during which the synoptically forced winds showed no diurnal variation). Figure 11b shows the resulting average day and night wind profiles. Evidently, at the height of the wind speed maximum nocturnal winds are more than 1 m s 1 stronger than daytime winds. This indicates the effect of nocturnal katabatic forcing even though a low-level jet is still present in daytime, which presumably reflects that occurrence of barrier wind as mentioned above. One should bear in mind, though, that katabatic forcing and the resulting winds lag radiational cooling by about 4 6 h (Kodama et al. 1989), as shown in Fig. 11a and in the next section. Hence, the diurnal variations in katabatically forced winds are underestimated in Fig. 11b by the bad timing of the radiosonde soundings. Nevertheless, on the basis of Fig. 11b and also Fig. 11a we conclude that katabatic forcing dominantly influences the near-surface flow in the Heimefrontfjella region. Other effects, such as barrier wind phenomena, may additionally influence the low level winds, but evidence regarding the extent of such effects is lacking. 6. Influence of synoptic forcing In order to investigate how much control synoptic forcing exerts on the boundary layer flow characteristics, a subdivision has been made between situations with easterly and westerly upper-air winds on the basis of Fig. 6b. High-altitude westerly geostrophic winds occur mainly at the beginning of January and in the last 13 days of the experiment. In total, 18 days could be categorized as having a westerly upper-air circulation,

15 2362 JOURNAL OF APPLIED METEOROLOGY FIG. 12. Average vertical profiles of potential temperature, zonal wind, meridional wind, vector mean wind speed, wind direction, and wind directional constancy during the entire period (mean) and for days with predominant easterly (east) or westerly (west) synoptic forcing. See text for further explanation. with the remaining 19 days exhibiting upper-air easterlies. King (1989), Bintanja and van den Broeke (1995a), and van den Broeke and Bintanja (1995a) made similar (but not equal) subdivisions to study the effect of synoptic forcing on the surface flow. They all found significant differences in meteorological variables between the selected categories, which can be interpreted as indicating the importance of the direction of the synoptic forcing for the climate of western Dronning Maud Land. Figure 12 shows radiosonde profiles averaged over the east and west selection of the measuring period. Clearly, differences in the various variables are significant and thus reflect true differences in flow char-

16 DECEMBER 2000 BINTANJA 2363 acteristics. Free atmospheric temperatures are lower during easterly winds, but the differences are small in the boundary layer. The air above the boundary layer is more stably stratified under westerly forcing. Relative humidity is highest during easterly winds, in particular in the lowest 4 km. This is a reflection of the large amount of clouds during easterly forcing with its depression activity. In both cases, the average zonal wind speed is easterly in the boundary layer. It is obviously largest during easterly synoptic forcing because it is then in the same direction as the topographically induced flow. Even in strong westerly synoptic forcing the boundary layer winds are still from the east, on average, reflecting the fact that the katabatic forcing dominates the boundary layer flow at Svea. The meridional component of the wind turns from north in the boundary layer to south higher up in all cases. The northeasterly winds in the boundary layer turn very sharply with height toward southwesterly winds under westerly forcing and moderately to southeasterly directions during easterly forcing. During westerly synoptic forcing there appears to be a distinct boundary between the topographically induced easterly winds in the boundary layer and the synoptically forced westerlies aloft. Individual profiles often exhibit a sharp transition between the two regimes, as was also observed by King (1989). This reflects the fact that during westerly forcing there are essentially two forcing regimes: a strong, upward increasing synoptic forcing in the free atmosphere and a strong katabatic forcing induced by extensive cooling in the boundary layer. The sharp transition between the two regimes is the cause of the low values of the absolute wind speed and the directional constancy at the boundary of the two regimes, as the boundary layer height will vary for each individual profile. King (1989) found similar wind profiles at Halley under westerly forcing with the boundary at about 1000 m. He further noticed that such strong shear layers can give rise to the formation of gravity waves, which were indeed observed at Halley. The easterly synoptic forcing is generally small, producing weak easterly winds in the free atmosphere. This is also caused by the fact that the easterly synoptic forcing can better be referred to as the nonwesterly synoptic forcing in view of the low values of the directional constancy in the upper troposphere. On the other hand, the westerly synoptic forcing is relatively strong and causes strong westerlies in the upper troposphere. It can be concluded that the westerly zonal circulation typical for the circumpolar region (King and Turner 1997) extends its influence to locations as far inland as the Heimefrontfjella and tends to produce occasional surface westerlies under favorable conditions (weak katabatic forcing). In order to study the influence of the direction of the synoptic forcing on the surface winds, the surface data for site 7 (representative location for the high plateau in the Heimefrontfjella region) have been used to calculate average daily cycles of the various meteorological quantities for the selected periods. These are shown in Fig. 13. An important feature is that during upper-air westerlies nocturnal temperatures are clearly lower. This is caused by predominantly clear skies, enabling the snow surfaces to cool because of a large net outgoing longwave radiation. The stronger nocturnal inversion forces stronger katabatic winds, especially at site 7 where nighttime winds are particularly fierce. As expected, this katabatic wind is directed more in the downslope direction, and is extremely constant with very high directional constancies. The prevailing synoptic pressure gradient in the lower troposphere is too weak to significantly affect the katabatic flow. Paradoxically, westerly geostrophic winds and the associated stronger surface cooling induces the strongest (south)easterly surface winds, at least during the night. This relates directly to the simultaneous occurrence of westerly winds near site 1 over blue ice and more downslope (southeasterly) winds near site 7 during weak westerly synoptic forcing (see Fig. 3d). Occasionally, nighttime wind speeds become very strong (as high as 15 m s 1 ) at site 7 under westerly forcing, while winds at the lower reaches remain calm. These events occurred around 5, 16, and 27 January (see Fig. 3). These strong winds are associated with exceptionally high temperature deficits of 10 C or more, forcing very strong katabatic winds during the night. During these periods of westerly forcing, relatively warm air is advected (see Fig. 6), creating a warm troposphere overlying a significantly colder boundary layer with a strong inversion separating the two air masses. Moreover, strong cooling of the surface air is caused by the net radiation deficit, which further increases the temperature difference with the free atmosphere. The warming events apparently occur only above 2 km MSL, as a result of which the temperature deficit at the lower plateau Ritscherflya remains small. Consequently, during these periods there is a marked difference in the forcing of the surface flow between the high and low areas of the Heimefrontfjella region. These events may be related to the low-level outflow of relatively cold air from the Weddell Sea region under anticyclonic conditions as discussed in section 4. In daytime, heating removes the inversion and the boundary layer wind tends to align with the noneasterly geostrophic wind. These winds are weaker and more variable and cause relatively low directional constancies in daytime. Winds above the boundary layer are weak, which indicates that in the absence of a katabatic force there is no strong dominant control over the surface winds during the day. This contrasts with the observations of Kodama et al. (1989), who found that the daytime geostrophically controlled boundary layer winds were as constant as the nocturnal katabatic winds. The explanation for this difference is that the measurements presented by Kodama et al. (1989) showed a strong meridional temperature gradient and hence a