1 JUNE 2005

HUDSON AND BRANDT1673

A Look at the Surface-Based Temperature Inversion on the Antarctic Plateau
STEPHEN R. HUDSON AND RICHARD E. BRANDTDepartment of Atmospheric Sciences, University of Washington, Seattle, Washington(Manuscript received 25 June 2004, in final form 2 November 2004)
ABSTRACT
Data from radiosondes, towers, and a thermistor string are used to characterize the temperature inversion
at two stations: the Amundsen-Scott Station at the South Pole, and the somewhat higher and colder DomeC Station at a lower latitude.
Ten years of temperature data from a 22-m tower at the South Pole are analyzed. The data include 2- and22-m temperatures for the entire period and 13-m temperatures for the last 2 yr. Statistics of the individual
temperatures and the differences among the three levels are presented for summer (December and January) and winter (April–September).
The relationships of temperature and inversion strength in the lowest 22 m with wind speed and downward longwave flux are examined for the winter months. Two preferred regimes, one warming and one
cooling, are found in the temperature versus longwave flux data, but the physical causes of these regimes
have not been determined. The minimum temperatures and the maximum inversions tend to occur not withcalm winds, but with winds of 3–5 m s⫺1, likely due to the inversion wind. This inversion wind also explains
why the near-surface winds at South Pole blow almost exclusively from the northeast quadrant.
Temperature data from the surface to 2 m above the surface from South Pole in the winter of 2001 arepresented, showing that the steepest temperature gradient in the entire atmosphere is in the lowest 20 cm.
The median difference between the temperatures at 2 m and the surface is over 1.0 K in winter; under clear
skies this difference increases to about 1.3 K.
Monthly mean temperature profiles of the lowest 30 km of the atmosphere over South Pole are presented,
based on 10 yr of radiosonde data. These profiles show large variations in lower-stratospheric temperatures,and in the strength and depth of the surface-based inversion.
The near-destruction of a strong inversion at South Pole during 7 h on 8 September 1992 is examinedusing a thermal-conductivity model of the snowpack, driven by the measured downward longwave flux. The
downward flux increased when a cloud moved over the station, and it seems that this increase in radiation
alone can explain the magnitude and timing of the warming near the surface.
Temperature data from the 2003/04 and 2004/05 summers at Dome C Station are presented to show theeffects of the diurnal cycle of solar elevation over the Antarctic Plateau. These data include surface
temperature and 2- and 30-m air temperatures, as well as radiosonde air temperatures. They show thatstrong inversions, averaging 10 K between the surface and 30 m, develop quickly at night when the sun is
low in the sky, but are destroyed during the middle of the day. The diurnal temperature range at the surface
was 13 K, but only 3 K at 30 m.

1. Introduction
It has been known for decades that a surface-based
temperature inversion exists over the Antarctic Plateau
and that its existence is not only a major feature of thisregion’s climate, but also a driver of some aspects of its
climate, including near-surface winds and atmospheric

Corresponding author address: Stephen R. Hudson, Department of Atmospheric Sciences, University of Washington, Box351640, Seattle, WA 98195-1640.
E-mail: hudson@atmos.washington.edu© 2005 American Meteorological Society

JCLI3360

radiation. A significant inversion exists at the surface of
the Antarctic Plateau in every month but December
and January (e.g., Schwerdtfeger 1970a).
The inversion exists because the snow-surface emissivity is greater than the atmospheric emissivity. Whenthe energy absorbed at the surface from solar radiationis small, this inequality in emissivities results in a temperature inversion.
Given that the sum of sensible, latent, and subsurface
heat fluxes provides only about 10%–15% of the energy the surface emits in winter (Schlatter 1972; Carroll
1982) and that the surface temperature is nearly con-

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stant during the 6-month winter (Schwerdtfeger 1984),
we can see that the upward longwave emission by thesurface must be nearly balanced by the downward longwave radiation reaching the surface from the atmosphere. This implies that, in winter,
0.85 ␧s␴T s4 ⬇ ␧a␴T 4a,

共1兲

where ␧s ⬇ 1.0 is the snow-surface emissivity, ␧a is theeffective emissivity of the atmosphere, ␴ ⫽ 5.6696 ⫻10⫺8 W m⫺2 K⫺4 is the Stefan–Boltzmann constant, Ts
is the surface temperature, and Ta is the effective temperature of the atmosphere (i.e., at levels where most of
the emission originates); the factor 0.85 comes from thefact that about 15% of the energy emitted by the surface is replenished by nonradiative mechanisms. Therefore, in winter, (Ts /Ta) ⬇ (␧a/0.85)0.25, so the surface
temperature will have to be lower than the air temperature so long as the emissivity of the atmosphere is lessthan 0.85, a criterion violated only under optically thickclouds. For a typical wintertime inversion, with Ts ⫽213 K and Ta ⫽ 233 K, Eq. (1) implies ␧a ⬇ 0.6.
The inversion is of interest to researchers in a variety
of fields. Waddington and Morse (1994) and Van Lipzig
et al. (2002) showed that the inversion can affect icecore records, and therefore must be considered forproper retrieval of paleoclimate from ice cores. The
ability of inversion layers to resist vertical mixing affects atmospheric boundary layer chemistry and snow
chemistry (Harder et al. 2000; Davis et al. 2001). One of
the rapidly growing fields of research on the Antarctic
Plateau is ground-based astronomy, which can benefit
from the high elevation and cold, dry atmosphere.
However, Marks (2002) and Travouillon et al. (2003)
showed that the presence of strong temperature gradients in the inversion layer can hinder atmospheric “seeing,” while its shallow depth allows for good use ofadaptive optics. Significant temperature gradients inthe lowest 2 m of the atmosphere associated with theinversion mean that satellite retrievals of surface temperature in Antarctica will be lower than the “surface”
air temperature measured at a station, which is usually
measured around 2 m above the surface.
Many studies of the inversion have tried to examine
it through radiosonde data. Among others, these haveresulted in an analysis of the relationship between nearsurface winds and the inversion by Dalrymple et al.
(1966) and Lettau and Schwerdtfeger (1967) and in a
climatology of inversion strength across the continent
by Phillpot and Zillman (1970), which showed that inversions over South Pole average about 20 K in winter,and those over the highest parts of the plateau averageover 25 K. Given the very rapid increase of temperature

VOLUME 18with height in these inversions (Schwerdtfeger 1984)
these radiosondes may not adequately capture the details of the inversion near the surface (Mahesh et al.
1997).
More recently, Connolley (1996) compared the climatology produced by Phillpot and Zillman with that
produced by the Met Office Unified Model, a global
climate model. Connolley’s results are similar to thoseof Phillpot and Zillman, but show a dependence of inversion strength on terrain slope, which was not included by Phillpot and Zillman. This approach is usefulin that it can help to fill in the many data voids that exist
over large portions of Antarctica. On the other hand,
we are a long way from being able to replace observations with model results.
Towers have been used to study the inversion as well;these have the benefit of allowing for more accurate
measurements to be made in the lowest portion of theatmosphere, where the inversion is most pronounced.
Perhaps the best-known tower-based study of the Antarctic Plateau boundary layer was conducted at Plateau
Station, where temperatures and winds were measured
at numerous levels on a 32-m tower (Riordan 1977).
Dalrymple et al. (1966) also presented data collected on
an 8-m tower at South Pole in 1958.
This paper presents data collected at South Pole fromradiosondes and from a 22-m tower over a 10-yr period.
For most of the period the ground-based data include
only 2- and 22-m temperatures, along with 10-m windand downward infrared flux, but some interesting features of the inversion can still be examined with thislimited dataset. Data from a string of thermistors between the surface and 2 m that was operated during thewinter of 2001 are used to examine the inversion very
near the surface at South Pole. The larger-scale characteristics of the inversion are captured in the radiosonde data. Additionally, some comparisons are madeto tower-based and radiosonde data collected in thesummers of 2003/04 and 2004/05 at Dome C, where asignificant diurnal cycle is observed.
The aim of this paper is to present an overview of thetemperature profile over the Antarctic Plateau at different scales, with particular focus on the structure of
the surface-based temperature inversion and on someof the factors that influence this inversion. The use of 10
yr of both radiosonde and tower-based data from South
Pole allows for analyses of the temperature profile thatcome closer to an actual climatology than previous
works. The longer period of record also allows for morerobust analyses of the relationships between the inversion and the winds and downward longwave radiation,
which are believed to influence the inversion. The datafrom the surface to 2 m, while less climatologically ro-1 JUNE 2005

HUDSON AND BRANDTbust, allow for the presentation of the character of thetemperature profile at these very lowest levels.
The limited data from Dome C, in addition to being
the first analysis of the inversion in that area of theplateau, give us a sense of how important the diurnal
cycle is at lower-latitude sites on the Antarctic Plateau.
While the character of the inversion over most of theplateau is likely similar to that at South Pole during
winter months, these data will show that during months
with significant sunlight the inversion at more northerly
locations can differ dramatically from what is seen at
South Pole.
This analysis should allow researchers in fields thatare affected by the inversion to better understand its
structure, and therefore how it will affect their data. It
also presents a baseline against which models can betested, for their ability to simulate both the mean conditions and the relationships between the inversion andwinds and longwave fluxes.
The next section discusses the data used in this paper,and section 3 presents the temperature and inversion
statistics at South Pole and discusses their relationships
with winds and downward longwave flux. A case studyof the destruction of a strong inversion is presented in
section 4, and section 5 shows the diurnal cycle at Dome
C. A summary is given in section 6.

2. Data
Data were obtained from two operational agencies aswell as from three field projects of the University ofWashington and one of the University of Idaho. The
two locations where data were collected, South Poleand Dome C, are shown on the map in Fig. 1.
Most of the ground-based South Pole data presented
here were collected by the National Oceanographic andAtmospheric Administration’s Climate Monitoring and
Diagnostics Laboratory (CMDL) at Amundsen-Scott
South Pole Station. The air temperature and wind data
are collected on a 22-m tower, and the downwelling
longwave fluxes on the top of the building where
CMDL operates, about 50 m from the tower. Both thetower and the building are in the Clean-Air Sector, inthe normal upwind direction from the rest of the station. The temperatures are measured with aspirated
platinum resistance probes, calibrated and intercompared annually, and with an accuracy of ⫾0.1 K. The
instruments are raised once each summer to maintaintheir nominal heights above the snow surface, because
the snow accumulation in the vicinity of the tower isabout 20 cm yr⫺1. Wind data are collected using an
R. M. Young model 05103 anemometer; wind speeds
have an accuracy of ⫾2% with a 1 m s⫺1 threshold, andthe wind directions are accurate to ⫾5° at 1.5 m s⫺1.

1675

FIG. 1. A map of Antarctica showing the location of the stations
mentioned in the text: South Pole is labeled with an “S,” Plateau
with a “P,” and Dome C with a “C.” Surface elevation is contoured at 1000-m intervals; the contour nearest South Pole andDome C indicates an elevation of 3000 m.

The longwave flux data come from an Eppley PIR pyrgeometer, with an accuracy of ⫾4 W m⫺2.
These data include 1-min-resolution observations of
the air temperature at 2 and 22 m for the period 19
January 1994–31 December 2003, as well as 13-m airtemperatures from 14 January 2002–31 December 2003.
They also include 1-min-resolution observations ofwind speed and direction, measured at 10 m above thesurface, from 19 January 1994–29 June 2003. The downward longwave flux data have 3-min resolution from1994–97, and 1-min resolution from 1 January 1998–29
June 2003. Hourly temperature and wind data collectedduring earlier years (1987–93) were not used becausesome periods in this dataset showed obvious biases, andthe data often did not compare well with those from theSouth Pole Meteorological Office.
Two other smaller sets of ground-based data fromSouth Pole are also used. In section 3c temperature
data from thermistors located at the surface and 0.2,
0.5, 1, and 2 m above the surface are presented. These
data were collected just inside the Clean-Air Sector
between 21 March and 21 September 2001 as part of theSouth Pole Atmospheric Radiation and Cloud Lidar
Experiment (Walden et al. 2001). The five YSI 44033
thermistors were intercalibrated in a water–ice bathand in a dry ice and methanol mixture before the experiment and agreed to within ⫾0.05 K. Data were col-

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lected from all five thermistors every 5 min, and theirheights were adjusted every 3 days during the winter.
The 11-m air temperature and downward longwave
flux shown in section 4 were collected using an Eppley
pyrgeometer with an attached thermistor at that height(Walden 1995).
The South Pole radiosonde data were collected by
the South Pole Meteorological Office, currently operated by Raytheon Polar Services Corporation. The data
were provided by the Antarctic Meteorological Research Center. Soundings from 1994–2003 were used.
During the summer operational season, October–February, two soundings are normally conducted each day,at 0000 and 1200 UTC; during most of the rest of theyear only the 0000 UTC sounding is made. These data
were collected with Atmospheric Instrumentation Research Incorporated (AIR) model 4A radiosondes from1994–96, with AIR model 5A sondes from 1997–July
2001, and with Vaisala RS80 and RS90 sondes from
August 2001–December 2003. In addition to rejecting
individual points containing values outside the range
observed at South Pole (T ⱕ ⫺110°C or T ⱖ ⫹10°C,
where T is the temperature at any height), a quality
control check was done in which each sounding wasbriefly viewed as a temperature–height profile. Of the4989 soundings obtained, 208 were removed because ofproblems found during this quality check, most because
they were missing more than 150 m of data at the beginning of the sounding, or because they reported baddata in the lowest part of the sounding. In another 218
soundings there were either large data gaps or bad dataat an elevated level, and all of the data above the problem level were removed from the analysis. Finally, in a
total of 41 soundings taken in June–September of 1997
and 1998 (the first two winters with the AIR 5A system) temperatures below ⫺90°C were recorded as
missing. In these soundings, all data above the level
where the temperature first dropped to ⫺90°C were
removed. In all, these three quality checks affected
about 9% of the soundings and these were distributed
fairly evenly throughout all 12 months.
The Dome C data include radiosonde temperatures,
fixed 2- and 30-m air temperatures and snow-surface
temperature. The air temperatures were measured withYSI 44033 thermistors in ventilated shields. These thermistors have a published accuracy of ⫾0.1 K, but when
the two thermistors were compared side by side at the
beginning of the season they agreed to within 0.01 K in
an ice bath, and to within 0.03 K outside at ⫺21°C. The
surface temperature was measured with a Heitronics
KT19.82 infrared thermometer that looked at the snow
at an angle of about 45°. The instrument measures thebrightness temperature of the snow by looking in the

VOLUME 18

8–14-␮m band. Since the emissivity of the snow isslightly less than 1.0, the actual temperature is about
0.5 K higher than the brightness temperature measured
at this angle (Dozier and Warren 1982). Dealing withthis small error is preferable to dealing with the potentially much larger radiation-induced errors associatedwith trying to measure the snow-surface temperature insitu during sunlit times (Brandt and Warren 1993). Radiosonde data at Dome C were collected using Vaisala
RS80 and RS90 sondes. Two sondes were launched onmost days from mid-December 2003–January 2004. The
launches were generally timed to correspond to certain
satellite overpasses and were most often performed
around 1600 and 2300 local standard time (LST; 0800
and 1500 UTC).

3. The temperature profile at South Polea. Temperature and inversion statistics between 2
and 22 m
It is worthwhile to examine the annual temperature
cycle, shown in Fig. 2, before looking at seasonal statistics. This figure was made using the 10-yr record of2-m air temperature (T2) that is used in the rest of thissection. The familiar seasonal cycle over the Antarctic
Plateau is apparent, with a short, peaked summer, rapidly changing transition seasons, and a long, coreless
winter. This term “coreless winter,” from the German
kernlose, was, according to Wexler (1958), coined by
Middendorfs and Hann in the early twentieth century
and is used to describe the Antarctic winter, during
which the average temperature is nearly constant, withno well-defined core of minimum temperatures. The
standard deviation curve shows that the winter temperatures are much more variable than those in summer, because in winter the temperature can be affectedby clouds that may be much warmer than the surface,by horizontal advection of air of different temperature,by changing wind speeds that affect the vertical mixingof the inversion layer, and by slight vertical motions
due to convergence or divergence of the surface windsthat move colder air up or warmer air down (Schwerdtfeger 1977). The contrast of day-to-day variability between winter and summer is illustrated in a 1-yr timeseries shown by Warren (1996). Given the shape of themean temperature curve, the seasons will be defined asfollows: summer is December and January; autumn is
February and March; winter is April–September; springis October and November (Schwerdtfeger 1970a).
Figure 3 shows a typical temperature sounding for aclear sky in winter. The data above 660 hPa are from aroutine RS80 sounding performed by the South Pole
Meteorological Office, and those below that level are1 JUNE 2005

HUDSON AND BRANDT

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FIG. 2. Mean (thick line) and std dev (thin line) of the 2-m air temperature for each day of the year
at South Pole Station (1994–2003). The calculations were done on all the observations of T2 from the
CMDL data for the 5 days centered on each day.

from a tethered RS80 sounding from the South Pole
Atmospheric Radiation and Cloud Lidar Experiment(Walden et al. 2001). There is a rapid warming withheight very near the surface, with a nearly isothermal
layer above that from about 250–900 m. Above that the
profile transitions to a more normal lapse of temperature with height. Focusing on Fig. 3b, it is possible tosee just how rapid the warming is near the surface. By
just 40 m above the surface the temperature is 15 K
higher than at the surface, and by 100 m the temperature has risen by 20 K of the total 23-K increase. This
concentration of the inversion near the surface is whatmakes it possible to examine some of its features froma 22-m tower, which in this case would span an increasein temperature of more than 5 K.
The distributions and summary statistics for summer
temperatures and inversions are shown in Fig. 4. There
are separate distributions for differences between eachpair of temperatures: 22 m (T22), 13 m (T13), and T2.
Recall that the data for T13 span a much shorter time
period than the other data.
From Fig. 4 it is apparent that summer temperatures
tend to lie within a fairly small range at all three levels.
Over 98% of the T2 data lie in an 18-K range between⫺38° and ⫺20°C. Just comparing the T22 and T2 data
provides some evidence that there is often an inversion
across this layer in summer; every percentile, as well as

the maximum and minimum, is lower in the T2 data,suggesting a tendency for the air to be colder at 2 m thanat 22 m. It is also interesting that the difference between
these two distributions is larger at low temperatures,
where the first percentile is 1.3 K lower for T2 than T22,
than at high temperatures, where the 99th percentile isonly 0.3 K lower for T2. This difference illustrates thefact that the inversion is weaker during warmer periods.
The distribution of the difference between T22 and T2
confirms that, even in summer, inversions occur morethan half the time in this near-surface layer. The inversions are often weaker than 1 K over the 20 m, but cansometimes exceed 5 K. Lapses are observed in thislayer about 26% of the time. These sometimes exceed
the dry-adiabatic lapse rate, but rarely exceed 0.5 K.
Interestingly, the 2–13-m layer rarely experiences lapse
conditions, while the 13–22-m layer has lapse conditions almost 50% of the time. While inversions are almost always present in the lower layer, the strongest
inversions in the lower layer are no stronger than thestrongest inversions in the upper layer.
The analogous winter data are presented in Fig. 5.
Obviously the temperatures in winter are lower thanthose in summer, but they are also much more variable,with the middle 98% of T2 data spanning 34.7 K, nearly
twice the value for summer. Also, the effect of the inversion on the T2 and T22 distributions is more apparent

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lower layer than the upper, but the extreme inversions
are similar in both layers.

b. Relationships of T2 and T22 with downward
longwave flux and winds

FIG. 3. Temperature profile measured at South Pole Station on
25 Sep 2001. Data above 660 hPa are from a routine radiosounding with an RS80; those below 660 hPa are from a tethered sounding with an RS80. (a) The full tropospheric sounding is shown, and(b) the lowest 500 m are enlarged. The surface pressure was 674
hPa.

in these data; the first percentile of T2 is 3.3 K lower
than that for T22, and the 99th percentile is 1 K lower.
The T2 distribution is seen to be skewed positively, illustrating the ease with which the surface can bewarmed in the winter—for instance, by the appearanceof a cloud at the top of the inversion—compared to theslower process of cooling radiatively, which requires
extended periods of clear skies.
The distributions of temperature differences shownin Fig. 5 show that inversions are more common andcan also get much stronger in winter than in summer.
Between 2- and 22-m lapse conditions are observed lessthan 8% of the time, and the median difference is aninversion of 1.7 K. In this 20-m layer inversions sometimes exceed 10 K, and have exceeded 20 K. As insummer, lapse conditions occur much more frequently
in the upper layer (13–22 m) than in the lower layer,
where they are nearly nonexistent. This difference results in a median inversion strength that is larger in the

Using data from CMDL for wind and downwelling
longwave flux, this section will show how the temperature and inversion are affected by these two factors.
Both are expected to influence the inversion strengthand surface temperature. Stronger winds will lead toincreased turbulence, mixing warmer air down and decreasing vertical gradients. Increased downward infrared flux, which will occur under clouds or even at timesof significant blowing snow (Mahesh et al. 2003; Yamanouchi and Kawaguchi 1984), will cause the surface
to warm, or at least reduce its cooling rate. Although
the reasons that wind and longwave flux affect the temperature and inversion are easy to see, the resulting
relationships are not as simple as might be expected.
Only winter data will be shown in this section since thatis when the inversion is most pronounced and mostvariable.
The relationship between T2 and downward infrared
flux is shown in Fig. 6. The dark line shows the medianT2 as a function of downward longwave flux, and thelighter lines show the 10th and 90th percentiles. The
most obvious aspect of this plot is that T2 increases withincreasing flux, as expected. There is also a decrease invariability for fluxes below about 80 W m⫺2. One very
interesting feature of Fig. 6 is that T2 does not seem toincrease in the same fashion across the whole range of
observed fluxes. The slope of the median curve decreases around 90 W m⫺2, then begins to increase again
by 150 W m⫺2. This indicates that T2 is less dependenton downward longwave flux in this middle range offluxes.
To explore this dependence further, the data were
filtered to keep only those observations for which, overthe period from 5 h before to 1 h after the time of theobservation, the difference between the maximum andminimum observed flux was less than 8% of the mean
flux for that time. The idea was to examine only thosedata that represent near-equilibrium conditions. Thesefiltered data are shown in Fig. 7, as a scatterplot and aline showing the median as a function of infrared flux.
Here it is clear that in the filtered data, representing
times when conditions were not changing much, there
are two preferred regimes in this temperature-flux
space. At lower fluxes T2 is significantly higher than theatmospheric brightness temperature, indicating that thesurface is cooling radiatively. At higher fluxes the surface is in radiative equilibrium or warming radiatively.
Conditions apparently do not often remain constant at

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HUDSON AND BRANDT

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FIG. 4. Distributions and statistics of air temperatures at (a) 22 m (T22), (b) 13 m (T13), and (c) 2 m (T2), and thedifferences (d) T22 ⫺ T2, (e) T22 ⫺ T13, and (f) T13 ⫺ T2 for summer months. The statistics include the number ofobservations (N), the mean and std dev (Std) of all observations, the minimum (Min) and maximum (Max)observed values, and various percentiles of the data.

intermediate fluxes, between about 110 and 140 Wm⫺2—that is, in the region with smaller slope in Fig. 6.
A lack of reliable cloud observations in winter madeit difficult to better characterize this relationship, which
is likely related to clouds. From a preliminary analysis
using 3 months of cloud lidar data from winter 2001, it
appeared that data in the high-flux region were observed under low clouds; however, many other observations with low clouds fell in the lower-flux group.

Perhaps future improved cloud observations, includingheight and optical depth, will help explain the patternseen in Fig. 7.
Next, the relationship between winter inversion
strength (T22 ⫺ T2) and downward longwave flux isshown in Fig. 8, using the full dataset. Here again, theobvious feature is what is expected: inversion strength
decreases with increasing infrared flux. Also, the variability of inversion strength tends to decrease with in-

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FIG. 5. As in Fig. 4, but for winter months. Note that the abscissas are on fixed scales for comparison (the same
scales used in Fig. 4), but the ordinates vary between plots to maximize readability.

creasing flux. The interesting part of Fig. 8 is that themedian inversion strength is nearly constant at about 3
K for fluxes less than 80 W m⫺2, then begins to decrease
rapidly, and asymptotes to zero by about 150 W m⫺2.
This decrease begins around 80 W m⫺2, or an atmospheric brightness temperature of about ⫺79°C, which
is near the lowest temperature reached by the surface.
This indicates that, as long as the surface is coolingradiatively, the median inversion strength is about 3 K,
regardless of the magnitude of the cooling. The difficulty in creating lapse conditions in winter can be seenby the fact that the median asymptotes to zero, rather
than dropping below zero for the highest fluxes.
Figure 9 shows how T2 is related to wind speed. The
reason to expect a relationship between temperature
and wind speed is that stronger winds will lead to morewarm air mixing down from higher in the inversion.
The actual relationship does not appear to be so simple.
The minimum air temperatures tend to occur withwinds of 3–5 m s⫺1, not with calm winds. A possible
cause of the location of this temperature minimum will
be presented later in this section.

1 JUNE 2005

HUDSON AND BRANDTFIG. 6. Median 2-m air temperature (dark line) as a function of
downward infrared flux (or corresponding brightness temperatureof the sky) in winter; lighter lines are the 10th and 90th percentiles. The dashed curve shows where T2 would be equal to thebrightness temperature of the sky (actual radiative equilibrium
between the surface and the sky would occur slightly to the left ofthis curve because the surface is generally colder than the air at 2m, and because the surface has an emissivity slightly less than 1).
The lower panel shows the number of observations in each bin ofwidth 1 W m⫺2.

First, let us examine how inversion strength is related
to wind speed. Again, the expectation is that stronger
inversions can develop with weaker winds because of
reduced vertical mixing. Figure 10 shows that this relationship is also more complicated than that. The maximum in the median inversion strength occurs at windspeeds between 3 and 5 m s⫺1, as did the minimum intemperature. Data from Plateau Station suggest a similar pattern (Riordan 1977), but the 1-yr data record inthat paper meant the results were climatologically limited, and Riordan suggested that inversion strength issimilar at all wind speeds less than 6 m s⫺1. These data,with their longer period of record, suggest that this is areal feature of the Antarctic Plateau boundary layer.
The 90th percentile curve (i.e., the strength of strong
inversions for a given wind speed) has two maxima, one
near 3 m s⫺1 and another with very light winds, supporting the idea that very light winds are indeed conducive to the development of strong inversions.
Given that it is unlikely that winds of 3–5 m s⫺1
would promote low temperatures or strong inversions,

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FIG. 7. A scatterplot of filtered, winter T2 data (gray area) andthe median (solid line) of these data as a function of downward
infrared flux (or corresponding brightness temperature of thesky). The dashed curve shows where T2 would be equal to thebrightness temperature of the sky. The lower panel shows thenumber of observations in each bin of width 1 W m⫺2.

the causality is probably in the opposite direction. It
seems that these relationships could be caused by the“inversion wind,” in which the presence of an inversion
over gently sloping terrain forces a wind through a local
thermal wind. This can be investigated using the modelof Mahrt and Schwerdtfeger (1970), who presented asolution for the problem of steady-state flow over uniformly sloping terrain with an exponential temperature
profile, assuming a coefficient of eddy diffusivity that isconstant with height.
As presented in appendix 4 of Schwerdtfeger (1984),
this model specifies the two components of the wind in
the boundary layer as functions of height for given conditions as
u共z兲 ⫽ 共B ⫹ ␷g兲e⫺D sinD ⫺ 共C ⫹ ug兲e⫺D cosD
⫹ Ce⫺zⲐH ⫹ ug,

共2兲

␷共z兲 ⫽ ⫺共B ⫹ ␷g兲e⫺D cosD ⫺ 共C ⫹ ug兲e⫺D sinD
⫹ Be⫺zⲐH ⫹ ␷g,

共3兲

where u(z) and ␷(z) are the eastward and northward
components of the wind at height z, ug and ␷g are theeastward and northward components of the geostrophic

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FIG. 8. Median inversion strength (dark line) in the 2- to 22-m
layer as a function of downward infrared flux (or corresponding
brightness temperature of the sky) in winter; lighter lines are the10th and 90th percentiles. The lower panel shows the number ofobservations in each bin of width 1 W m⫺2.

wind at the top of the inversion, H is the e-folding depthof the inversion (50 m in the example given in the next
paragraph), and
B⫽ ⫺

C⫽
D⫽

H4 f 2␷T ⫹ H2 | f | KuT
K2 ⫹ H4 f 2

H2 | f | K␷T ⫺ H4 f 2uT
K2 ⫹ H4 f 2

冉 冊

0.5
|f|
z
,
2K

,

,

共4兲

共5兲
共6兲

in which f is the Coriolis parameter, K is the coefficientof eddy diffusivity, and uT ⫽ (g/f )(⌬T/T)(⳵h/⳵y) and␷T ⫽ (⫺g/f )(⌬T/T)(⳵h/⳵x) are the thermal wind components at the surface, in which g is the acceleration
due to gravity, ⌬T is the total inversion strength, T is
the temperature at the top of the inversion, and (⳵h/⳵x)and (⳵h/⳵y) are the components of the slope in the eastward and northward directions.
Using this model we can test whether the inversion
alone at South Pole can create a wind of about 4 m s⫺1.
To do this the temperature profile was assumed to beT(z) ⫽ 235–25e⫺z/50, where z is the height above the

VOLUME 18

FIG. 9. Median 2-m temperature (dark line) as a function ofwind speed in winter; lighter lines are the 10th and 90th percentiles. The lower panel shows the number of observations in eachbin of width 0.1 m s⫺1.

surface in meters, and T(z) is the temperature in
Kelvin at z. This gives a surface temperature of ⫺63°C,
and a temperature at the top of the inversion of ⫺38°C.
To isolate the effect of the inversion on the wind, it was
assumed that there was no large-scale pressure gradient(the geostrophic wind at the top of the inversion waszero). The surface slope was set to 2.25 ⫻ 10⫺3, with afall line directed toward 285°. This slope was the average slope over the area from 20 km west to 80 km eastof the pole, and from 20 km south to 80 km north of thepole [all directions around South Pole are given on thelocal grid, where “north” (000°) is parallel to theGreenwich meridian in the direction of Greenwich], as
determined from a 1-km-resolution digital elevation
model (Liu et al. 2001). This area was chosen to emphasize the slope in the area that is usually upwind ofSouth Pole. The final parameter for the model, theeddy diffusivity coefficient, K, is not well known, so
results with two values are compared. Assuming an
eddy diffusivity of K ⫽ 0.1 m2 s⫺1, the value assumed by
Mahrt and Schwerdtfeger (1970), we get a 10-m windof 3.4 m s⫺1 from 067°. From estimates made using
Monin–Obukhov similarity theory (Garratt 1992) it
seemed that a value close to 0.05 m2 s⫺1 might be more
appropriate for these conditions, in which case the resulting wind at 10 m is 5.0 m s⫺1 from 061°. Both results
support the idea that the inversion wind can explain the1 JUNE 2005

1683

HUDSON AND BRANDTFIG. 10. Median inversion strength (dark line) in the 2- to 22-m
layer as a function of wind speed in winter; lighter lines are the10th and 90th percentiles. The lower panel shows the number ofobservations in each bin of width 0.1 m s⫺1.

location of the maximum in inversion strength and theminimum in T2.
In essence, while the development and maintenanceof a strong inversion is favored by light winds, the existence of the inversion can create a wind. This suggeststhat near the tops of the domes of the plateau, where
the surface slope is approximately zero and an inversion therefore cannot create a wind, stronger inversions
may be possible under weak synoptic pressure gradients.
Figure 11 shows that this inversion wind model also
can explain why winds near the surface at South Poleblow almost exclusively from the northeast quadrant.
This figure shows that, for the conditions given in theprevious paragraph and an eddy diffusivity coefficientof 0.1 m2 s⫺1, a geostrophic wind of 5 m s⫺1 from any
direction at the top of the inversion results in 10-m
winds that come from grid directions between 038° and
096°. While Neff (1999) clearly shows that this inversion wind is not the only factor driving the winds at
South Pole, it does seem that the inversion will often
help to ensure that winds will blow from the northeast
quadrant under a variety of synoptic-scale influences.
Finally, it is interesting to examine the relationshipsthat exist between these observations and wind direction. Table 1 shows the annual and seasonal vectormean winds, along with those for very warm and cold

FIG. 11. Winds generated by the inversion under various synoptic conditions. The crosses indicate end points of geostrophic
wind vectors at the top of the inversion. Each is connected by aline to a circle that indicates the end point of the resulting 10-m
wind calculated with the inversion wind model using the corresponding geostrophic wind to specify the synoptic pressure gradient. For example, with a geostrophic wind at the top of the inversion of 5 m s⫺1 from the south (indicated by the cross at the topof the plot), the inversion wind model predicts a 10-m wind of justover 2 m s⫺1 from the east (indicated by the circle that is attachedto the cross at the top). Radial grids are at 2 and 4 m s⫺1. The
downslope direction is 285°.

periods, and times with very strong inversions. Overall,
the resultant wind is from the northeast, at about 4
m s⫺1, weaker in the summer and stronger in the winter. There is a tendency for spring and summer winds tohave more of a northerly component compared to thosein autumn and winter.
Winds at times of the strongest inversions have more
TABLE 1. Vector mean winds and directional constancy (vector
mean wind speed divided by scalar mean wind speed) at South
Pole for various subsets of the data.

All data
SummerAutumnWinterSpring
Winter:T22 ⫺ T2 ⬎ 15 K
T2 ⬍ ⫺75°C
T2 ⬎ ⫺35°C
Summer:T2 ⬍ ⫺38°C
T2 ⬎ ⫺18°C

Speed (m s⫺1)

Direction (°)

Constancy

4.2
3.0
4.1
4.6
4.4

40
30
46
43
27

0.75
0.69
0.77
0.76
0.76

3.2
3.8
7.0

76
87
349

0.86
0.92
0.78

3.0
8.9

92
358

0.94
0.92

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JOURNAL OF CLIMATE

FIG. 12. (a) The median 2-m air temperature and (b) medianinversion strength between 2 and 22 m as functions of wind direction. The frequency of winds from each direction is shown in(c). Data are for the entire year and were grouped into 15° bins
centered on 000°, 015°, . . . , 345°.

of an easterly component, and are slightly weaker thanthe overall average. In both summer and winter thelowest temperatures occur with winds from the east,
while the highest occur with winds from the north tonorth-northwest, as pointed out by Hogan et al. (1979).
Another way to consider this relationship between
temperature and wind direction is to look at the mediantemperature and inversion strength as a function ofwind direction, shown in Fig. 12, for the entire year.
This figure shows the predominance of winds from thenortheast quadrant discussed above. Here we again seethat the lowest temperatures occur with winds out of
the east, which are winds blowing down slope from thehigher, colder parts of the plateau. However, in contrast to the resultant winds in Table 1, we now see thehighest temperatures occurring with winds from thewest, and that winds from the south through the northwest are usually accompanied by very high temperatures, but are also extremely rare.

VOLUME 18

The height of the thermistors was adjusted approximately every 3 days to ensure the lowest thermistor was
right at the surface. In the time between adjustments
drifting or falling snow sometimes buried the surface
thermistor, and drifting occasionally caused it to beslightly above the surface. Since the temporal variability of temperature below the snow surface is less thanthat in the near-surface air, we determine whether thesurface thermistor is buried by comparing the standard
deviation of its time series with that of the 0.2-m thermistor. If, for the 5-h period centered on an observation,
the standard deviation of the temperature measured by
the surface thermistor exceeded the standard deviationof the temperature measured by the 0.2-m thermistor
and all of the data are available from that 5-h period
then the surface thermistor is assumed to be at or above
the surface, and the observation is used in this analysis.
It was unusual for the surface thermistor to be above
the surface, and when it did happen it was usually
within 0.05 m, so no attempt was made to remove theseobservations.
Additional quality control of the data included removing observations in which the temperature difference between the surface and 2 m differed in sign fromthe difference between 0.2 and 2 m, and removing observations in which the temperature difference between
the surface and 2 m showed an inversion of more than
4 K or a lapse of more than 1 K. Inspection of the few
periods outside of this range indicated that those data
were likely in error. These two criteria each resulted in
the removal of 2.5% of the data.
Figure 13 shows the median difference between thec. Inversion statistics between the surface and 2 m
Data from a string of thermistors maintained at
South Pole in the winter of 2001 are presented here toextend the description of the inversion to the lowest
layer of the atmosphere, between the surface and 2 mabove the surface.

FIG. 13. Median temperature difference relative to 2 m. Data
are from South Pole during the 2001 polar night. Separate profiles
are shown for the overall median (All Sky, solid line) and theclear-sky median (dashed line).

1 JUNE 2005

HUDSON AND BRANDT

temperature at each height and that at 2 m, for thesunless period, 21 March–21 September 2001. Two median profiles are shown: the solid line is for all-sky conditions; the dashed line is for times with clear skies.
Downward infrared flux data from an upward-looking
pyrgeometer were used to determine when skies wereclear, using the method described by Town et al. (2005),
which determines a clear-sky threshold for each monthby modeling the downward infrared flux from a clear
sky with the Line-By-Line Radiative Transfer Modelusing the average atmospheric conditions for thatmonth.
It turned out that the subset of data that included
only those observations when the surface thermistor
was not buried oversampled nonclear skies. To correct
for this, the all-sky data shown in this section include allof the clear-sky data when the surface thermistor was
not buried and a random sample of the non-clear-sky
data when the surface thermistor was not buried. The
size of the random sample was chosen so that the fraction of clear-sky data was 50%, which is what it was in
the full dataset, including times when the surface
thermistor was buried.
As expected, Fig. 13 shows that the median temperature differences under clear skies were considerablylarger than those for data including all-sky conditions.
At the surface, temperatures were slightly more than
1.0 K lower than at 2 m in the all-sky median, and wereabout 1.3 K lower in the clear-sky median. The largest
temperature gradient occurs in the lowest layer, so thatthe median temperature difference between 0.2 m andthe surface exceeds 0.8 K under clear skies. Above thislayer the gradient weakens, but under clear skies it isstill more than 0.25 K m⫺1 between 0.2 and 2 m.
Histograms of the temperature difference between 2m and the surface are shown in Fig. 14 to give a senseof the variability observed. Figure 14a includes all-sky
conditions, while Fig. 14b includes only the clear-sky
observations. In both cases the temperature difference
is most often between 0 and 2 K. Some lapses wereobserved in this layer, but never under clear skies.

d. The tropospheric and lower-stratospherictemperature profile
Radiosonde data from South Pole from 1994–2003
were used to create monthly mean temperature profiles
from the surface to 30 km above mean sea level (MSL).
For the lowest 55 m (2835–2890 m MSL) data from thetower rather than the sonde were used to avoid erroneously warm readings resulting from the preparationof the sondes inside a heated building (Mahesh et al.
1997). This height was usually reached about 10 s after
launch, which is approximately twice the e-folding re-

1685

FIG. 14. Relative frequencies of observed temperature differences between 2 m and the surface during the 2001 polar night.
(a) The data from all-sky conditions and (b) those from clear-sky
conditions.

sponse time of the temperature sensors (Hudson et al.
2004), eliminating most of the effect of the heatedbuilding. The 2- and 22-m temperatures from the tower
at the time of launch were inserted into the lowest partof each sounding and a linear interpolation was madefrom the tower temperature at 22 m to the radiosondetemperature at 55 m. To correct for the response timeof the temperature sensor on the radiosondes, the measured temperatures were shifted back in time by 6 s,
assigning each temperature measurement to the heightof the sonde 6 s earlier. Mahesh et al. (1997) found that
this method was as effective at correcting profiles as amore complicated deconvolution method. The datafrom each sounding were then linearly interpolated to a
fixed set of heights: every 10 m from 2890 to 3000 m
MSL, every 20 m from 3020 to 3500 m MSL, every 50 mfrom 3550 to 4500 m MSL, and every 100 m from 4600
to 30 000 m MSL. Then every interpolated sounding
from each month was used to produce the plots in Figs.
15–17. In these monthly mean profiles the effect of the1686

JOURNAL OF CLIMATE

response time correction is largest in the middle of theinversion layer, where it can increase mean temperatures by 2 K in winter months (since the reported temperature is still influenced by the much colder air justbelow the sonde’s position); above the inversion its effect was generally less than 0.25 K.
Each of these plots shows the monthly mean air temperature as a function of height as a solid line; they also
include dashed lines showing the 10th and 90th percentiles of air temperature as a function of height. Three
profiles are presented for each month. All show thesame data, but the latter two show successively moredetail in the lower atmosphere.
Figure 15 shows the soundings from the surface to 30
km MSL. In these we can focus on the progression of
the temperature in the upper troposphere and lower
stratosphere. During January and February the temperature profiles resemble those at lower latitudes, witha well-defined tropopause around 8–9 km and significant warming in the stratosphere. As the sun sets in
March the stratosphere begins to cool, rapidly above
about 15 km, so that by May the area between 15 and
30 km is colder than any part of the troposphere. The
lower stratosphere does not cool as quickly, so a tropopause is still apparent in May, and, to some extent, in
June.
Before looking at the rest of the year it is worth
mentioning that in winter very few of the balloons
make it through the extremely cold air in the stratosphere. As a result, the winter profiles above about
10–15 km are based on relatively few observations.
Moving on, by July and August there is virtually nosign of a temperature tropopause and the air around 20
km has cooled to about ⫺90°C, a temperature notreached in the troposphere or stratosphere anywhereoutside Antarctica, and which permits the formation ofpolar stratospheric clouds and enhances ozone depletion (Solomon 1990). By August and September thehighest parts of these soundings are seeing the sun andbegin to warm. The warming continues throughout October and November, when temperatures at 30 km arehigher than in any part of the troposphere in any
month. However, a temperature tropopause is notclearly reformed until December. Comparing December and January we see that the lower stratosphere
continues to warm, while the upper parts of thesesoundings are already cooling.
It is necessary to explain the unusual features seen in
the September profile in Fig. 15. From all of the September data only 98 of the 314 soundings continue
above 15 km. Of those, eight were from late September
2002. In that year the stratosphere warmed unusually
early, so that temperatures in those eight soundings

VOLUME 18

from about 10–15 km and up were much higher than in
any other September sounding. Between the 0000 UTC
soundings on 20 and 23 September the temperaturerose by about 4 K at 10 km, by about 15 K at 15 km, andby almost 42 K at 21.5 km. The presence of these eight
warm soundings raises the mean enough so it is greaterthan the 90th percentile in places. Six of these eightsoundings continued above 30 km, and the other two
reached 28 km, while the total number of soundings
available had dropped to 73 by 22 km. By that point,
the 90th percentile jumps to the warm group of soundings, which explains the discontinuity in that line. The
fact that soundings are more likely to reach higher levels at times of higher temperatures indicates that mean
profiles could be biased warm at these heights whenever there is a lot of variability in temperature.
Figure 16 shows the profiles from the surface to 10
km MSL. These also show the loss of the temperature
tropopause between April and June. It is also apparent
that the temperature structure in the free tropospheredoes not vary much during the year. The mean temperature at 6 km varies by less than 12 K during theyear, and by only about 3 K during the winter (April–
September).
The temperature structure of the lowest kilometer of
the atmosphere is shown in Fig. 17. This shows thedetails of the full inversion, as well as some of the isothermal layer above the inversion. The profiles for thetwo summer months show only a very slight, shallow
inversion, with nearly isothermal conditions in the lowest kilometer. In February a significant inversion ispresent, as the 2-m temperature drops below ⫺40°C.
The near-surface temperature continues cooling inMarch, but then remains nearly constant during thecoreless winter; the 2-m temperature drops by just 1.3 K
between April and September. In October and November the inversion weakens dramatically as the air near
the surface warms much more than the air above it.
The top of the inversion occurs around 600 m in thesemean profiles except in summer, when it is closer to 300
m. Temperatures at the top of the inversion drop by lessthan 2 K between April and August, despite average
winter radiational-cooling rates at this height measured
to be about 4 K day⫺1 (Schwerdtfeger 1970b). Schwerdtfeger attributed this great reduction in cooling at the
top of the inversion to subsidence and, to a lesser extent, warm-air advection.
Table 2 shows the monthly mean temperature gradient between 2 and 100 m, and inversion strength anddepth calculated from these mean profiles. The average
inversion strength calculated from the individual
soundings is also shown, and is 0.5–1 K larger than the
inversion in the monthly mean profile because of the1 JUNE 2005

HUDSON AND BRANDT

1687

FIG. 15. The monthly mean (solid line) atmospheric temperature profiles from 2 m above the surface to 30 km MSL over South Pole.
The dashed lines show the 10th and 90th percentiles of temperature at each height. Radiosonde data begin 55 m above the surface (2890m), and temperatures from a tower are used at 2 and 22 m. Data from 1994 to 2003 were used.

1688

JOURNAL OF CLIMATEVOLUME 18

FIG. 16. Same as Fig. 15, but only the lowest 10 km.

variability in the height of the top of the inversion. The
fact that the difference between these two inversion
strengths is small suggests that the shape of the inversion is very stable. The strength and persistence of thewinter inversion is clear; on average, the temperature at
100 m is about 12.5 K higher than at 2 m in September.
The last two columns in Table 2 show the values of
inversion strength and depth calculated by Dalrymple

1 JUNE 2005

HUDSON AND BRANDT

1689

FIG. 17. Same as Fig. 15, but only the lowest 1 km above the surface.

(1966) from individual soundings from South Pole from
November 1958–October 1961. These values suggestthat there has been no large change in the inversion
structure at South Pole during the last four decades.

The consistently higher height of the top of the inversion reported by Dalrymple is more likely to be due toimprovements in radiosonde response time than to realchanges in the inversion structure.

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JOURNAL OF CLIMATEVOLUME 18

TABLE 2. Monthly mean temperature gradient between 2 and 100 m above the surface (⳵T/⳵z), inversion strength between 2 m andthe warmest point in the troposphere calculated from the monthly mean sounding (⌬Tmean) and from each individual sounding(⌬Tsound), inversion depth (⌬Z ) calculated from the monthly mean sounding. All of these values were determined from 10 yr of data.
Also shown for comparison are the values presented in Dalrymple (1966, his Table 30) for data from Nov 1958 to Oct 1961 (⌬TDal and⌬ZDal).
Month(⳵T/⳵z) (K m⫺1)

⌬Tmean (K)⌬Tsound (K)⌬Z (m)⌬TDal (K)

⌬ZDal (m)

Jan
Feb
Mar
Apr
May
Jun
Jul
Aug
Sep
Oct
Nov
Dec

0.018
0.054
0.111
0.118
0.113
0.116
0.113
0.114
0.128
0.085
0.030
0.007

2.3
9.8
20.0
22.2
22.0
22.6
21.5
21.0
22.9
16.1
6.6
1.0

3.1
10.6
20.7
22.8
22.6
23.3
22.1
21.6
23.6
16.8
7.3
2.0

265
525
665
665
605
625
585
605
605
585
505
245

3.4
10.6
20.0
23.5
21.1
21.8
20.0
22.0
20.5
15.5
7.7
2.3

357
609
685
706
657
644
641
638
669
596
546
319

It is interesting to compare these results with the
Arctic. Looking at monthly mean soundings from Barrow, Alaska, and Eureka, Nunavut, Walden et al.
(1996) found that the inversion depth was around 1000m at these sites in winter, and the strength was about
10–12 K. At Barrow and Barter Island, Alaska Kahl
(1990), found similar winter inversion strengths, andmedian depths between 600 and 900 m by examining
the strengths and depths from all individual soundings.
Also, Serreze et al. (1992) found that the median inversion depth at the Soviet drifting stations in the Arctic
Ocean in winter was around 1200 m, and the medianstrength was around 12 K. In all these studies the polarnight inversions in the Arctic were found to be much
weaker and at least somewhat deeper than what is observed at South Pole.

4. Case study of inversion destruction
The large day-to-day variability of T2 in winter (Fig.
2) is largely due to the ability of clouds and wind to
rapidly destroy the inversion. Here we examine a casestudy. On 8 September 1992 a very strong inversion wasquickly eroded as T2 increased when a cloud moved
overhead. Temperature data from this day are shown in
Fig. 18; T2 and T22 are hourly data from CMDL, andT11 is the 11-m air temperature measured on a pyrgeometer. The inversion strength (T22 ⫺ T2) was verylarge for several hours and peaked at 19 K; this extreme
inversion was responsible for some optical phenomena
observed visually at that time (S. Warren 2004, personalcommunication): looming of the horizon, a ducting mirage, and the Novaya Zemlya Solar Mirage (Greenler

FIG. 18. Air temperatures at 2, 11, and 22 m, and 10-m wind speed (V10) at South Pole on 8 Sep 1992.

1 JUNE 2005

HUDSON AND BRANDT1980). In the 7 h after 0700 UTC, T2 increased by 15.7
K, while the temperature at the two higher levels remained nearly constant. Warren’s visual observations
indicated that a cloud moved over the station around0720 UTC, and the measured downwelling longwave
radiation increased substantially between 0700 and
0800 UTC. This figure also shows that during the warming the winds were very light, less than 2.1 m s⫺1, suggesting that they likely played only a small role in thenear-surface warming.
Given the light winds and observation of cloud cover,it seems likely that this warming was driven mainly by
the increased downwelling longwave radiation. To see
if the increase in downwelling radiation could, by itself,cause a temperature increase of this magnitude in thistime period, a simple thermal-conductivity model of thesnowpack was used to predict the surface temperature.
This one-dimensional model included the upper 60 cm
of snow, with 1-mm vertical resolution. The temperature change at each layer due to convergence of verticalheat flux was calculated using centered differencing of
the following equation:
k ⭸2T
⭸T
⫽
,
⭸t
␳c ⭸z2

factor of 2 in either direction has negligible impact on
the results of the model. The temperature at a depth of
60 cm was held constant (the measured change during
the 11 h simulated was only 0.1 K). The model was
initialized at 0400 UTC, and was run, with 0.25-s time
steps, until 1500 UTC.
The results of this model are shown as the solid blackline in Fig. 19, along with the downward longwave flux
used to force the model, and the measured T2. Unfortunately, there are no measurements of snow-surface
temperature for this time, so we can only compare thecalculated surface temperature with the measured 2-m
air temperature. This simple model indicates that sucha warming as was observed can be accounted for by theincreased radiation reaching the surface, with no otherfactors contributing. Adding a sensible heat flux, calculated with Monin–Obukhov similarity theory (Garratt1992), reduced the warming slightly, but did not have alarge effect on the overall picture produced by themodel.

5. Diurnal cycle at a lower-latitude site
共7兲

where T is temperature, t is time, k is the effective
thermal conductivity of the snow, ␳ ⫽ 345 kg m⫺3 is theapproximate snow density measured at South Pole in1992 by Brandt and Warren (1997), c ⫽ 2100 J kg⫺1K⫺1 is the heat capacity of ice, and z is depth. Thisequation assumes that k can be taken as constant overa 2-mm layer. An additional term,⭸T Fd ⫺ ␧␴T4
⫽
,
⭸t
␳c␦z

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共8兲

was added to the surface temperature change to account for radiational heating and cooling, where ␦z ⫽0.001 m is the thickness of the layer; Fd is the observed
downward longwave flux; ␧ is the snow-surface emissivity, which was set to 1; and ␴ is the Stefan–Boltzmann constant.
The profile of thermal conductivity was taken fromBrandt and Warren (1997), and their snow temperature
data for this day at depths of 20, 30, 40, and 60 cm wereused to determine the initial temperature profile. The
initial surface temperature was assumed to be 3 K lower
than T2 at that time to account for the continuation of
the inversion to the surface; while this is a reasonablevalue for a strong inversion according to the thermistor
data presented in section 3c, the actual value of thisdifference is unknown; however, varying this value by a

The diurnal cycle is absent at the South Pole but issignificant elsewhere on the plateau. The inversion wasstudied during the summers of 2003/04 and 2004/05 at
Dome C Station (75°S, 123°E). During these periods
the sun was always above the horizon, but its elevation
angle ranged from 3° to 38°, and varied by about 30°
during the course of each individual day.
Figure 20 shows the half-hourly mean temperaturesof the surface, and of the air at 2 and 30 m. The data
used to create this figure and Fig. 21 were collectedfrom 22 December 2004 to 31 January 2005. The largest
diurnal cycle is observed at the surface, where the diurnal range is about 13 K, but the effect of this cycle isseen at all three levels. At the surface the maximum
temperature occurs about 2 h after solar noon, whilethe 30-m temperature maximum lags the sun by nearly
4 h. For several hours around midday the surface iswarmer than the overlying air, but at night an inversion
is clearly present. Schwerdtfeger (1970a) showed similar data from Plateau Station for February. In thosedata the air temperature at 0.25 m varied by about 13 K,
while at 32 m it varied by only about 2.5 K. However,unlike at Dome C in January, at Plateau in February
the near-surface air temperatures remained 1–2 K
lower than the 32-m temperature even at midday
The half-hourly mean inversion strength across thethree layers is shown in Fig. 21. Here we see that lapse
conditions occur in all three layers around local noon.
By just after 1500 LST (UTC⫹8) the surface cools be-

1692

JOURNAL OF CLIMATEVOLUME 18

FIG. 19. Results of the snow–temperature model during the warming in Fig. 18 both with and without sensible heat flux. Also
shown is the measured T2 and the downward longwave flux; the latter was used as an input into the model.

low the 2-m air temperature, and around 1700 LST an
inversion begins to form between 2 and 30 m. The inversion develops rapidly and, on average, reaches amaximum of 10 K between the surface and 30 m, about
6 K between 2 and 30 m, and about 4 K just between thesurface and 2 m. On three nights the inversion in thislowest 2-m layer exceeded 10 K. The layer between 2
and 30 m had inversions exceeding 10 K on about one-

third of the nights, with maximum strengths around 13
K. Over the full surface to 30-m layer the inversion
strength frequently approached 15 K, but never exceeded 17 K. The mean inversion strength between 2
and 30 m on a summer night at Dome C is stronger thanalmost any inversion between 2 and 22 m observed at
South Pole in summer.
Some radiosonde data from Dome C are presented in

FIG. 20. Half-hourly mean surface temperature and air temperature at 2 and 30 m at Dome C Station for the 2004/05 summer.
Time is given in local standard time, UTC⫹8

FIG. 21. Half-hourly mean inversion strength across the threeavailable layers at Dome C Station for the 2004/05 summer. Timeis given in local standard time, UTC⫹8

1 JUNE 2005

HUDSON AND BRANDT

1693

FIG. 22. The main panel shows soundings from Dome C taken every 2 h from 1400 LST 24 Jan 2004 to 1200 LST
25 Jan 2004. The surface is at 3266 m, and the lowest point is the 2-m air temperature measured at the time oflaunch. The 0400 LST sounding is not available, and the 0200 LST sounding is missing data from approximately
3350 to 3450 m. The inset panel shows the mean temperature profile from all soundings at Dome C taken in thehour centered on 1600 LST (dashed line) and on 2300 LST (solid line). Twenty soundings taken between mid-Dec
2003 and the end of Jan 2004 went into each of these two mean profiles. All times are given in local standard time,UTC⫹8.

Fig. 22 to show the effect of the diurnal cycle on a
deeper layer of the atmosphere. The main panel showsa series of temperature profiles from radiosondelaunches every 2 h from the early afternoon of 24 January 2004 through noon on 25 January 2004, presenting
one complete diurnal cycle. We see that on this day the
maximum near-surface temperature in the 2-hourly
data, ⫺24°C, was recorded at 1600 LST, and that theminimum, 15 K lower, occurred 8 h later. Here we canagain see the rapid weakening of the diurnal cycle
above the surface; by about 100 m above the surface thediurnal cycle is overwhelmed by other forms of hourly
variability. Significant lapse conditions are apparent in
the first two afternoon soundings, overlain by weakerlapses, with an elevated inversion at about 200 m. By
1800 LST the atmosphere is nearly isothermal up to theelevated inversion. A strong near-surface inversion develops in the lowest 50 m during the evening, but, as thesolar elevation angle begins to increase in the earlymorning hours, the near surface temperature begins torise. By 0600 LST there is a strong lapse in the lowest 25
m, with the inversion continuing to deepen above thatheight. The surface warming then ceases, and the elevated temperature minimum weakens through the latemorning.
The pattern in the main panel is from one day, and is
not meant to represent a mean diurnal cycle, but rather
to show the type of effects we may see in the lower
atmosphere due to the diurnal cycle. There was unusually dense fog on the night shown here, and that fog
dissipated between 0500 and 0800 LST, possibly leading
to the cooling seen between 0600 and 0800 LST. Thisfog may have caused other irregularities in the diurnal
cycle on this particular day as well.
The inset panel in Fig. 22 is an attempt to average out
the day-to-day variability and show more nearly theaverage effect of the diurnal cycle. The dashed line isthe average of all 20 profiles taken in the hour centered

1694

JOURNAL OF CLIMATEon 1600 LST from mid-December through the end ofJanuary; the solid line is the same, but for those near
2300 LST. These are the two 1-h periods with the mostsonde launches. In the mean profiles the effect of thediurnal cycle does not extend more than about 150 m
above the surface. This plot shows that the large effectof the diurnal cycle does not extend much above thetower data presented above.

6. Conclusions
The South Pole temperature data from 1994 to 2003
show the usual seasonal cycle of the Antarctic Plateau,with a short, peaked summer and a 6-month, coreless
winter. Two-meter air temperatures in summer are usually between ⫺38° and ⫺20°C, while in winter they aremuch more variable, usually between ⫺73° and ⫺38°C.
The lowest temperatures at 22 m are higher than thoseat 2 m because of the temperature inversion. However,
the highest temperatures are similar at both levels because the inversion is generally much weaker on very
warm days.
Distributions of temperatures and temperature differences in the 2- to 22-m layer show no surprising results but provide an analysis of conditions in this largely
unstudied layer with a period of record at the top andbottom that is long enough to eliminate much of theinfluence of interannual variability. In summer an inversion is observed between 2 and 22 m nearly 75% of
the time, but temperature differences across this layer
are usually less than 1 K. Lapse conditions appear to bemuch more common between 13 and 22 m than between 2 and 13 m, or over the entire layer. The 2- to
22-m layer has isothermal or inversion conditions morethan 90% of the time in winter, and these inversions cangreatly exceed 10 K. The 13- to 22-m layer in winter isagain much more likely to have lapse conditions thanthe lower layer, although both layers can have equallysteep inversions.
Analyses of the variations of temperature and inversion strength with winds and downward longwave flux
yielded some surprising and previously unreported results. The 2-m air temperature in winter increases withincreasing downward longwave flux, but not uniformly.
There seem to be two preferred regimes for these datain near-steady-state conditions, one in which the surface is radiatively warming and one in which it is radiatively cooling. This nonuniform and possibly regimelike relationship has not been discussed before and thephysical mechanisms causing it have not been explained. The clear relationship between downward
longwave flux and clouds suggests that better cloud observations could help explain this behavior.

VOLUME 18

The inversion strength does not steadily decrease
with increasing flux, but is nearly constant for values of
downward longwave flux less than about 80 W m⫺2, at
which time the surface is nearly always radiatively cooling, and then quickly decreases as the flux continues toincrease and asymptotes to zero for fluxes greater than
150 W m⫺2. This is the first long-term study of thisrelationship and it is interesting to see that the correlation of these two values is limited to the 80–150 Wm⫺2 range of fluxes.
The relationships of wintertime temperature and inversion strength with wind speed show some features
not previously identified. The minimum temperatures
tend to occur not with calm conditions, but with windsof about 3–5 m s⫺1, as do the strongest median inversions. An analysis of the wind resulting from the inversion over sloped terrain showed that the inversion can
induce winds of this magnitude at South Pole, and suggests that this may be why the minimum temperaturesand maximum inversions occur with nonzero windspeeds.
Minimum temperatures accompany winds that are
blowing downslope from the higher, colder parts of theplateau. The highest temperatures occur with upslope
winds, but similar temperatures are seen for a variety ofunusual wind directions, which are likely caused by synoptic systems, and accompanied by clouds.
The inversion continues below 2 m, right down to thesurface. In winter the median difference between thetemperature at 2 m and the surface is 1.0 K, and thisdifference increases to 1.3 K when looking only at clear
skies. The largest temperature gradient is between thesurface and 20 cm, but it remains large up to 2 m.
Monthly mean temperature profiles above South
Pole show large annual variations in temperature in thelower stratosphere, and the disappearance of the temperature tropopause during the polar night. They also
show relatively small temperature variations in the freetroposphere, and that the inversion over South Pole ismuch stronger and somewhat shallower than those over
and around the Arctic Ocean.
A thermal-conductivity model for the snowpack,forced by downward longwave radiation, shows that theincrease in downward longwave flux due to the appearance of a cloud was sufficient to explain a rapid warming event at South Pole. Light winds at the time likely
reduced the role that sensible heat fluxes could play.
These results support the widely held theory that theappearance of a warm cloud over the Antarctic Plateau
is enough to quickly warm the surface and dramaticallyweaken the inversion.
Finally, a look at summer temperatures at Dome C
Station showed the large effect of the diurnal cycle of1 JUNE 2005

HUDSON AND BRANDT

solar elevation and provided our first look at the temperature profile at this location. The inversion was destroyed during the middle of the day, and lapse conditions existed between the surface and overlying air, but
at night the inversion quickly reformed. This inversion
became much stronger than those seen at South Pole insummer, with an average inversion of 10 K forming
after midnight between the surface and 30 m. Radiosonde data show that the effect of the diurnal cycle islimited to the lowest 150 m.
Acknowledgments. Many of the data presented inthis paper were provided by NOAA’s Climate Monitoring and Diagnostics Laboratory, and we are especially grateful to Thomas Mefford and Ellsworth Dutton for providing access to CMDL’s data. South Pole
radiosonde data were provided by the Antarctic Meteorological Research Center in Madison, Wisconsin.
Paulene Roberts initiated the analyses of the tower
temperatures. We also wish to thank Thomas Grenfell,
Von Walden, Bradley Halter, Lance Roth, Robert
Stone, and Vito Vitale for their help with the collectionof the Dome C data; Delphine Six for helping erect the
32-meter tower at Dome C; and Stephen Warren, Von
Walden, and Michael Town for their many helpful suggestions along the way. Stephen Warren also gave
many useful comments on the first draft of this paper.
This research was supported by NSF Grants OPP-9726676, OPP-02-30466, and OPP-00-03826.
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