5.1. Processes in the Antarctic Sea Ice Zone
 The ISPOL measurements by Nicolaus et al.  yielded albedo values ranging from 0.85 to 0.67. Most of the time the difference between their and our values was approximately 0.03 only; it may have arisen from small-scale spatial differences and measurement inaccuracies. In any case, the decrease of albedo during the 34-day period was not enough to trigger any significant positive feedback effect. Therefore the net heat flux and snowmelt did not show an increasing trend during the study period. Also, the results by Brandt et al.  suggest that there is no significant albedo feedback in the Antarctic sea ice zone. They analyzed albedo data from ship-based field experiments in the East Antarctic sea ice zone, and concluded that a representative albedo of snow-covered sea ice is 0.87 in September–November and 0.82 in December–February. Also, based on ship data, Wendler et al.  observed a mean albedo of 0.81 for compact sea ice cover in late December. Although ship-based albedo measurements are not as accurate as those made at an ice station, the numbers are very close to our mean value of 0.83 for December.
 Snow surface albedo is also strongly controlled by precipitation, as snow fall tends to increase albedo [Pirazzini, 2004]. In the Weddell Sea precipitation is less than in other areas at the same latitude because of the Antarctic Peninsula blocking humid westerly winds and depressions on its western side [Tietäväinen and Vihma, 2008]. Air on the east coast of the Peninsula is generally 7 K colder than at similar latitudes and elevations on the west coast. From the point of view of the summer decrease of albedo, the effects of low temperatures and reduced snow fall in the Weddell Sea tend to compensate each other, which partly explains the similarity between our observations and those made at the East Antarctic sea ice zone [Brandt et al., 2005; Wendler et al., 2005].
 In days with surface melt for at least 2 hours, the diurnal cycle of albedo observed at ISPOL was comparable to the seasonal trend through the study period. Our observations on the shape of the cycle are in agreement with results from several locations over the Antarctic continent [Pirazzini, 2004], but we are not aware of any previous studies on the diurnal cycle of surface albedo in the Antarctic sea ice zone. We are neither aware of previous studies on the atmospheric transmissivity for shortwave radiation over the Antarctic sea ice zone. Compared to our December mean of 0.50, Van den Broeke et al.  observed an annual mean value of 0.63 at the Riiser-Larsen Ice Shelf and 0.78 at the Kohnen station 530 km inland from the ice shelf front at the altitude of 2900 m.
 Our results for the ISPOL period indicated an upward sensible heat flux. On the contrary, the prevalence of a downward sensible heat flux over Antarctic sea ice was observed by Andreas et al. [2000, 2004, 2005] in February–May at ISW in the same region as ISPOL, and in four studies in more eastern parts of the Weddell Sea: Kottmeier and Engelbart  in October–November, Launiainen and Vihma  for most of the year, and Vihma et al. [1996, 2002] throughout the year. It seems that the sensible heat flux over Antarctic sea ice is from air to snow for most of the year.
 In summer the direction of the sensible heat flux seems to be variable. Launiainen and Vihma  observed a mean sensible heat flux of 20 W m−2 in December 1990 at 66–67°S, and Vihma et al.  obtained a similar value for December 1996 in the same region. Their study region was located 1000 km east of ISPOL, and this may be one reason for the different flux direction: the central Weddell Sea is more often affected by warm air advection from the north, which is an important factor generating downward sensible heat flux. In the pioneering studies of Andreas  and Andreas and Makshtas , based on ship observations in October–November 1981, the sensible heat flux was directed from air to snow during northerly winds and from snow to air during southerly winds; the difference in the flux magnitude was up to 180 W m−2. Wendler et al.  observed a downward sensible heat flux in the Ross Sea over both broken sea ice (12 W m−2) and compact snow-covered sea ice in the McMurdo Sound (18 W m−2). The observation periods of the work of Wendler et al.  were, however, short: six days over the broken sea ice and four over the compact sea ice. Hence the synoptic-scale situation may have had an essential effect on the results. In any case, the above mentioned results demonstrate that observations on turbulent fluxes from the western Weddell Sea should not be generalized for the entire Antarctic sea ice zone, not even for the entire Weddell Sea.
 On the basis of ISPOL observations and the prevalence of dry, cold katabatic winds, Nicolaus et al.  concluded that in summer turbulent fluxes are predominantly upward over the Antarctic sea ice zone. This may be the case in many regions, but our results indicate that upward heat fluxes can prevail even without this effect: in only 50% of time the air masses observed at ISPOL came from the Antarctic continent, and even among these cases the air mass usually originated from the west or northwest of the Antarctic Peninsula with the flow controlled by the synoptic-scale pressure gradient. On the other hand, despite of the prevailing katabatic winds at the McMurdo Sound, Wendler et al.  observed downward sensible heat flux.
 Our results provide the first extensive quantification on the diurnal cycle in various meteorological quantities affecting the snow metamorphosis over the Antarctic sea ice. We observed diurnal cycles in 15 variables. A simplified picture of the causal chain is that the diurnal cycle in the incoming shortwave radiation is the driving force for the other diurnal cycles, and it most directly affects those in the reflected solar radiation and snow surface temperature. These further affect the outgoing longwave radiation, surface albedo, and the turbulent fluxes of sensible and latent heat. The diurnal cycles in the air temperature as well as the specific and relative humidity follow in this causal chain, where the variables most indirectly affected are the incoming longwave radiation and wind speed. The base height of low clouds is affected by the air temperature and moisture as well as the incoming and reflected shortwave radiation.
 Although the diurnal cycle is a widely studied topic, there have not been many observations reported from the Antarctic sea ice zone. Compared to ISPOL, Wendler et al.  observed a smaller diurnal amplitude for the incoming shortwave radiation (210 W m−2, estimated from their Figure 3) but a larger one for incoming longwave radiation (approximately 15 W m−2). The former was due to higher latitude (78°S). Our drifting buoy data from the central Weddell Sea at 73–65°S in 1990 indicated a diurnal amplitude of 1.4°C in the 2-m air temperature in December [Launiainen et al., 1991], which is slightly larger than the ISPOL result (0.9°C). The buoy data also demonstrated that in addition to season and latitude the diurnal temperature cycle strongly depended on the ice concentration: the more open water around, the smaller was the diurnal cycle. This was also demonstrated by Niros et al. , who compared diurnal cycles over the open and ice-covered Baltic Sea and its snow-covered coasts.
5.2. Comparisons Between the Arctic and Antarctic
 On the basis of ISPOL observations, Willmes et al. [2006, 2007] concluded that there is a fundamental difference in the summer evolution of the snowpack between the Arctic and Antarctic sea ice zones. In the Antarctic, the microwave brightness temperature drops at the onset of snowmelt, although it increases in the Arctic. According to Willmes et al. , the reason for the drop in the Antarctic is the repeated diurnal thawing and refreezing of snow, instead of complete snow wetting as in the Arctic. Our snow surface temperature data support this conclusion.
 Willmes et al. [2006, 2007] and Nicolaus et al.  stressed the differences in snow processes between the Arctic and Antarctic, while Andreas et al.  concluded on the basis of ISW observations that similar processes control the ABL over sea ice in both hemispheres. The statistics on temperature inversions at ISW in February–May were similar to the Arctic: the ABL was virtually always stably stratified with near-surface inversions observed in 96% of the tethersonde soundings. Should we conclude that (1) the processes in the Arctic and Antarctic are basically similar in autumn and winter but different in summer or (2) that the processes are similar in the ABL but different in the snowpack? We try to better understand this issue by comparing the ISPOL observations against two Arctic field experiments: the Surface Heat Budget of the Arctic Ocean (SHEBA) and the Swedish Arctic Ocean Expedition 2001 (AOE-2001).
 In June the SHEBA ice camp was at higher latitudes than ISPOL (76–77°N), but the mean incident shortwave radiation at the top of the atmosphere (504 W m−2) was very close to that of ISPOL (508 W m−2), although the diurnal ranges were somewhat different: 0–1008 W m−2 at ISPOL and 0–800 W m−2 at SHEBA. We calculated from the SHEBA data that in June the monthly mean net shortwave radiation was 73 W m−2, the net longwave radiation −28 W m−2, the latent heat flux −6 W m−2, and the sensible heat flux −2 W m−2 (compare to Figure 20 in Persson et al. ). In ISPOL, the mean net solar radiation was smaller (52 W m−2), the net longwave radiation approximately the same (−30 W m−2), while the fluxes of latent heat (−14 W m−2) and sensible heat (−6 W m−2) had larger magnitudes. In SHEBA, the surface albedo decreased from 0.82 to 0.63 during June, with a mean value of 0.74, while at ISPOL the mean value was 0.83. The mean values of 2-m air temperature were −2°C for ISPOL in December and −0.5°C for SHEBA in June. With colder air and higher albedo, the snowmelt is reduced and a larger portion of the solar heating of the surface is used for evaporation and sensible heat flux to the atmosphere: at ISPOL this portion was 38% compared to 11% at SHEBA, where the solar radiation was mostly used to melt the snow and ice.
 The difference was also seen in the maximum upward sensible heat flux, at ISPOL its magnitude exceeded 50 W m−2, but was only 12 W m−2 in June at SHEBA. Cold air advection is an important factor generating upward sensible heat flux in polar regions [Vihma et al., 2005; Valkonen et al., 2008], but in summer there are not many sources of colder air to flow over the Arctic sea ice, while in the Antarctic the continent provides sources of colder air.
 Although the monthly mean turbulent fluxes were directed upward both at ISPOL in December and SHEBA in June, the partitioning of the incoming solar energy was different. This is illustrated in Figure 11. Although representing a short period only, the results by Wendler et al.  from the compact sea ice cover in the McMurdo Sound are included in Figure 11 to demonstrate the variability in summer conditions in the Antarctic, which may be of spatial or temporal origin. With respect to the net shortwave and net longwave radiation as well as the sensible and latent heat flux, the ISPOL results are closer to SHEBA than the McMurdo Sound results. The residual term Qnet is, however, almost the same at SHEBA and McMurdo Sound, but only approximately 10% of those at ISPOL.
 In SHEBA in June, the incoming shortwave radiation had an amplitude of approximately 180 W m−2 only, while the ISPOL value was 320 W m−2. Hence also other quantities had smaller diurnal cycles at SHEBA: the mean diurnal amplitudes were 0.3°C for the surface temperature, 0.6°C for the air temperature, and 3 W m−2 for the sensible heat flux [Persson et al., 2002, Figures 16 and 17]. The ISPOL values were 1.6°C, 0.9°C and 7 W m−2, respectively.
 In the Arctic, the diurnal cycle in the incoming shortwave radiation at the snow/ice surface is reduced after early May due to the increase in cloudiness toward summer [Persson et al., 2002]. Further, the largest amplitudes of the diurnal cycles of near-surface air temperature and sensible heat flux occur even before – in April. In SHEBA, the air temperature amplitude in July (0.2°C) was already strongly reduced from the June value (0.6°C). This was related to (1) the seasonal evolution of the boundary layer stratification and (2) the onset of melt that limits the diurnal cycle in surface temperature [Persson et al., 2002]. In ISPOL the surface melt was so limited that only in a few cases it notably reduced the diurnal cycle of the surface temperature (the mean cycle peaks just below 0°C; Figure 9), but at SHEBA the diurnal amplitude in June was only 20% of that at ISPOL in December.
 On the basis of AOE-2001 observations, Tjernström [2005, 2007] showed that a diurnal cycle is present over sea ice in summer even very close to the North Pole (most of his data originated from an ice drift station at 88–89°N, from 2 to 21 August 2001). In addition to the incoming shortwave and net radiation, a diurnal cycle was apparent in the wind speed, cloud base, and visibility. Due to the high latitudes, the mean amplitudes observed by Tjernström [2005, 2007] were naturally much smaller than at ISPOL: approximately 25 W m−2 in incoming solar radiation and 5 W m−2 in net radiation. Due to melting and refreezing, the diurnal cycle in the near-surface air temperature remained weak. Tjernström [2005, 2007] observed, however, a significant diurnal cycle in the wind speed with the lowest winds around noon. This is opposite to the ISPOL results. In his case the diurnal cycle in the surface sensible heat flux was qualitatively similar to that at ISPOL with the peak value at noon, but the amplitude was much smaller. Hence there must have been some other factor that has dominated over the stratification effect and thus prevented the generation of maximum winds around noon.
 Tjernström [2005, 2007] observed a consistently high cloud fraction but slightly lower values in the morning and during late night. The cloud base height was on average lowest from 10 to 15 LST, which is approximately in agreement with the ISPOL results (Figure 10a). Tjernström [2005, 2007] did not make separate calculations for low clouds (as in our Figure 10b), and interpreted that the noon minimum in the cloud base height was due to complex interactions in the subcloud and cloud layers involving changes in stratification, cloud top radiative cooling, turbulent mixing, and drizzle. We think that at ISPOL, with a larger diurnal cycle in the surface temperature, the daytime heating and convection were probably enough to explain the noon maximum in the base height of low clouds, although various other processes were naturally also active in the cloud layer.
 In the ice-covered Baltic Sea at the latitudes of 63–64°N in February and early March 1998, Brümmer et al.  observed air temperatures and latent heat fluxes comparable to those at ISPOL, but the sensible heat flux was predominantly from air to ice. This was often due to warm air advection over the sea ice. Its dominating effect was probably related to the small distance from the open sea, which was 0–150 km compared to 220–500 km at ISPOL. Nicolaus et al. (submitted manuscript, 2009) observed stronger snowmelting and superimposed ice formation on ice floes 100 km north of ISPOL, which demonstrates how sensitive the effects of warm air advection are to the fetch over the ice [see also Tisler et al., 2008].