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    NCEP reanalysis for the period 1968–97. Global and hemispheric daily area-averaged (a) surface pressure, (b) surface pressure owing to vertically integrated water vapor, and (c) surface pressure owing to dry air (hPa). Note that NH refers to the Northern Hemisphere, SH to the Southern Hemisphere, and G to the globe

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    (a) Distribution of NH cold season dry atmospheric mass fall events as a function of duration. (b) Composite hemispheric dry air surface pressure anomaly for the 25 cold season events of 6–10-day duration shown in (a) with respect to the composite onset time of 0 (T0). The concepts of onset time, event duration and magnitude, discussed in section 2b, are illustrated in (b).

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    Time traces of area-averaged dry air surface pressure anomalies for the NH (thick solid), the SH (dotted), and the globe (thin solid) for 41-day periods centered on the onset time (0) for each of the 25 Northern Hemisphere cold season dry atmospheric mass fall events. Units on the ordinate are hPa, and the two vertical lines in each plot denote the onset and termination times for each event. Event numbers (onset times) are shown in the top right (left) of each plot and correspond to those listed in Table 1. The convention for the onset times is month/day/year (two digit)

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    Composite anomalies of SLP, with a contour interval of 1 hPa to a magnitude of 3 hPa, and every 2 hPa for larger magnitudes: (a) T−5, (b) T−3, (c) T−1, (d) T+1, (e) T+3, (f) T+5, and (g) T+7. Light (dark) shading denotes statistical significance at the 95% (99%) levels according to a two-sided Student's t test. Inset shows plots of the NH (solid) and SH (dash) dry air surface pressure anomalies (hPa) as a function of time from onset (denoted by vertical line). Thick dots denote the time of the horizontal SLP anomaly plot

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    Composite SLP anomalies for the 10 NH cold season dry atmospheric mass fall events between 1979 and 1993: (a), (a′) T−5; (b), (b′) T−3; (c), (c′) T−1; (d), (d′) T+1; (e), (e′) T+3; (f), (f′) T+5; and (g), (g′) T+7. Panels (a)–(g) are derived from the NCEP reanalysis, while (a′)–(g′) are derived from the ECWMF reanalysis. The contour interval is 2 hPa, with negative values dashed

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    (a) Composite SLP anomaly difference, with a contour interval of 1 hPa to a magnitude of 4 hPa, and every 2 hPa for larger magnitudes. For each event the SLP anomaly difference is calculated as the difference in SLP between the times of local max and local min in anomalous NH dry atmospheric mass for each event (see Fig. 3). Prior to calculating the SLP anomaly difference maps, the SLP was temporally filtered, retaining periods between 6 and 30 days. Light (dark) shading indicates the statistical significance at the 95% (99%) level based upon a two-sided Student's t test. The three boxes refer to the northern, southern, and South China Sea areas discussed in section 4. (b) Latitude–time plot of the zonally averaged (0°–360°) composite SLP anomalies (hPa), weighted by the cosine of latitude. Time given on the ordinate is with respect to the onset time, denoted as T0. Shading shows regions of positive SLP anomalies. No filtering was applied in (b)

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    Composite Arctic Oscillation (AO) index for the 25 events of NH dry atmospheric mass fall. The index values represent the projection of the leading standardized principal component of NH SLP (north of 20°N) onto the daily SLP anomalies (see Thompson and Wallace 2000). The values are dimensionless and were obtained online from http://www.colostate.edu/ao/Data/. Time on the abscissa covers the 41-day period centered on the composite onset time (T0)

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    (a) Time series of composite sea level pressure anomalies (hPa), area-averaged over 40°–75°N, 85°–145°E (northern box shown in Fig. 6a; thick solid), and area-averaged over 20°–30°N, 105°–120°E (southern box shown in Fig. 6a; dash). (b) Time series of composite 1000–850-hPa thickness anomalies (m) averaged over the southern box. (c) Time series of composite surface meridional wind anomalies (m s−1) averaged over the South China Sea (5°–10°N, 110°–115°E; see Fig. 6a). Time on the abscissa is for the 41-day period centered on the composite onset time (T0)

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Significant Events of Interhemispheric Atmospheric Mass Exchange: Composite Structure and Evolution

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  • 1 Department of Atmospheric and Oceanic Sciences, McGill University, Montreal, Quebec, Canada
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Abstract

The various modes of atmospheric mass redistribution characterize the principal variations of the general circulation of the atmosphere. Interhemispheric exchanges of atmospheric mass occur with considerable regularity on subseasonal time scales. Observational evidence from previous studies indicates that anomalous and persistent regional atmospheric mass distributions (e.g., atmospheric blocking) may often be related to interhemispheric atmospheric mass exchange.

Using the National Centers for Environmental Prediction (NCEP)–National Center for Atmospheric Research (NCAR) reanalysis surface pressure, significant events when the Northern Hemisphere (NH) loses dry atmospheric mass on subseasonal time scales during the boreal winter from 1968 to 1997 are identified. A total of 25 events is found, with a preferred time scale of 9 days from the time of maximum to minimum NH dry atmospheric mass. The linear correlation coefficient between the dry atmospheric mass anomalies for the NH and Southern Hemisphere (SH) is −0.91 for the 25 events, indicating very strong interhemispheric compensation and increasing confidence in the suitability of the NCEP–NCAR reanalysis dataset for the study of interhemispheric dry atmospheric mass exchange.

Positive sea level pressure anomalies are found over northern Eurasia, the North Pacific, and the North Atlantic prior to the onset of the composite NH dry atmospheric mass collapse event. Over northern Eurasia the building of the Siberian high is found to be a statistically significant precursor to the events. The breakdown of NH dry atmospheric mass occurs in association with the decay of the positive atmospheric mass anomaly in the North Pacific as a cyclone deepens explosively in the Gulf of Alaska. Pressure surges over Southeast Asia and North America, associated with statistically significant positive atmospheric mass anomalies, are mechanisms that act to channel the atmospheric mass equatorward out of the NH extratropics on a rapid time scale (∼4 days). The dry atmospheric mass increase in the SH is manifested as enhanced surface ridging over the South Pacific and south Indian Oceans, two noted regions of atmospheric blocking.

Current affiliation: RSIS Climate Prediction Center, NCEP/NWS/NOAA, Camp Springs, Maryland

Corresponding author address: Dr. Marco L. Carrera, RSIS/Climate Prediction Center, NCEP/NWS/NOAA, 5200 Auth Road, Room 605, Camp Springs, MD 20746. Email: Marco.Carrera@noaa.gov

Abstract

The various modes of atmospheric mass redistribution characterize the principal variations of the general circulation of the atmosphere. Interhemispheric exchanges of atmospheric mass occur with considerable regularity on subseasonal time scales. Observational evidence from previous studies indicates that anomalous and persistent regional atmospheric mass distributions (e.g., atmospheric blocking) may often be related to interhemispheric atmospheric mass exchange.

Using the National Centers for Environmental Prediction (NCEP)–National Center for Atmospheric Research (NCAR) reanalysis surface pressure, significant events when the Northern Hemisphere (NH) loses dry atmospheric mass on subseasonal time scales during the boreal winter from 1968 to 1997 are identified. A total of 25 events is found, with a preferred time scale of 9 days from the time of maximum to minimum NH dry atmospheric mass. The linear correlation coefficient between the dry atmospheric mass anomalies for the NH and Southern Hemisphere (SH) is −0.91 for the 25 events, indicating very strong interhemispheric compensation and increasing confidence in the suitability of the NCEP–NCAR reanalysis dataset for the study of interhemispheric dry atmospheric mass exchange.

Positive sea level pressure anomalies are found over northern Eurasia, the North Pacific, and the North Atlantic prior to the onset of the composite NH dry atmospheric mass collapse event. Over northern Eurasia the building of the Siberian high is found to be a statistically significant precursor to the events. The breakdown of NH dry atmospheric mass occurs in association with the decay of the positive atmospheric mass anomaly in the North Pacific as a cyclone deepens explosively in the Gulf of Alaska. Pressure surges over Southeast Asia and North America, associated with statistically significant positive atmospheric mass anomalies, are mechanisms that act to channel the atmospheric mass equatorward out of the NH extratropics on a rapid time scale (∼4 days). The dry atmospheric mass increase in the SH is manifested as enhanced surface ridging over the South Pacific and south Indian Oceans, two noted regions of atmospheric blocking.

Current affiliation: RSIS Climate Prediction Center, NCEP/NWS/NOAA, Camp Springs, Maryland

Corresponding author address: Dr. Marco L. Carrera, RSIS/Climate Prediction Center, NCEP/NWS/NOAA, 5200 Auth Road, Room 605, Camp Springs, MD 20746. Email: Marco.Carrera@noaa.gov

1. Introduction

The total mass of the atmosphere, as measured by globally averaged surface pressure, is not constant. The annual cycle is approximately 0.5 hPa, with the global atmospheric mass highest during the Southern Hemisphere (SH) winter (Trenberth 1981; Trenberth et al. 1987; Trenberth and Guillemot 1994; Chen et al. 1997b; Hoinka 1998). This annual cycle is a robust feature that results from the surface pressure owing to vertically integrated water vapor, a highly variable constituent of the atmosphere, possessing an annual mean value of approximately 2.5 hPa or 25 mm (Trenberth et al. 1987; Randel et al. 1996; Chen et al. 1996, 1997b). The variable nature of the surface pressure owing to vertically integrated water vapor also gives rise to interannual variations in global atmospheric mass associated with, for example, ENSO (Trenberth et al. 1987; Chen et al. 1997a).

Although the determination of the mass of the atmosphere is important, it is the spatial distribution of the atmospheric mass and the various modes of atmospheric mass redistribution that characterize the principal variations of the general circulation of the atmosphere (Lorenz 1951; Christy and Trenberth 1985; Christy et al. 1989). Atmospheric mass is redistributed on a variety of time scales, ranging from decadal and interannual (Trenberth et al. 1987; Trenberth and Guillemot 1994; Chen et al. 1997a) to seasonal (Van den Dool and Saha 1993; Chen et al. 1997b) and subseasonal (Holl et al. 1988; Christy et al. 1989). The redistribution can take place within the same hemisphere (intrahemispheric), or there can be an exchange of atmospheric mass between the hemispheres (interhemispheric).

a. Seasonal redistribution of atmospheric mass

Lorenz (1951) investigated both regular (seasonal) and irregular sea level pressure (SLP) variations within the Northern Hemispheric (NH). The regular SLP variations involved the movement or shift of atmospheric mass between two homogeneous zones, one located near 65°N and the other near 35°N. More recent studies have confirmed this intrahemispheric mode of atmospheric mass exchange (Trenberth and Paolino 1981; Christy and Trenberth 1985). The irregular SLP variation involved transports across the equator, whereby atmospheric mass was exchanged between the SH and the high latitudes of the NH. Lorenz postulated that the mechanism was most probably linked to shifts of atmospheric mass between the SH and the NH tropical regions and then some intrahemispheric exchange to the northern latitudes thereafter. The exact mechanism was unclear, which is in need of further work.

Gordon (1953) showed that the intrahemispheric seasonal redistribution of atmospheric mass within the NH was dominated by exchanges between land and oceanic regions. The exchanges were most pronounced over the Eurasian continent at the time of equinoxes, with the North American continent playing a less active role. Van den Dool and Saha (1993) emphasized these thermally induced atmospheric mass exchanges using an atmospheric general circulation model, while Chen et al. (1997b) discussed the importance of this seasonal mode of exchange using a combination of model and observed data.

Only recently, with the advent of reliable global-scale atmospheric datasets, have studies focused upon the redistribution of atmospheric mass between the two hemispheres (Trenberth 1981; Christy and Trenberth 1985). The constraint of global dry atmospheric mass conservation implies that anomalous departures of the mean atmospheric mass of dry air in the NH are compensated for in the SH. On an annual time scale, Trenberth (1981) calculated an annual exchange of 0.7 × 1016 kg, or approximately 2.7 hPa. Maximum interhemispheric flow takes place during the transition seasons, southward during Northern (Southern) Hemisphere spring (fall) and northward during Northern (Southern) Hemisphere fall (spring) (Christy et al. 1989).

The mechanisms involved in the seasonal interhemispheric exchange of atmospheric mass are well understood and can be explained on the basis of the annual cycle in solar heating (Trenberth 1981; Van den Dool and Saha 1993; Chen et al. 1997b). During the NH summer, the enhanced solar heating causes the atmosphere to expand, while the enhanced cooling in the SH causes the atmosphere to contract, setting up a pressure gradient at upper levels with flow directed from north to south. At low levels, the heating generates low pressure in the NH, while the cooling generates high pressure in the SH, creating a flow in the opposite direction from upper levels. This thermally directed meridional cell transports a net atmospheric mass to the winter hemisphere owing to the flow at upper levels exceeding that at lower levels.

b. Subseasonal interhemispheric atmospheric mass redistribution

Christy et al. (1989) examined 7 yr of daily global surface pressure analyses from the European Centre for Medium-Range Weather Forecasts (ECMWF) and found considerable variance in the fluctuations of NH atmospheric mass on subseasonal time scales. The authors noted significant peaks in the power spectrum at 40 and 64 days, but also found significant values at periods shorter than 20 days. Holl et al. (1988) found that interhemispheric atmospheric mass exchanges, on the order of several hPa, occurred on a time scale of days to weeks.

Studies of the exchanges of angular momentum between the solid earth and the atmosphere involve polar motions, which refer to oscillations about the equatorial axes of the earth, unlike the more commonly studied oscillations about the axial component. More specifically, polar motions refer to the directional change of the rotation axis with respect to an earth-fixed reference system and have magnitudes of several meters at the earth's surface. Salstein and Rosen (1989) and Nastula and Salstein (1999) examined the regional contributions to high-frequency fluctuations in polar motions, termed “rapid polar motions.” These studies were motivated by the findings from geodetic time series analyses of polar motions that indicated fluctuating signals on time scales of weeks. The latitudinal redistribution of atmospheric mass was found to play an important role in exciting these rapid polar motions, which are on the order of 60 cm. Nastula and Salstein (1999) noted the significance of the atmospheric mass fluctuations over the Eurasian region (51.5°–90°N, 30°–120°E, and 33.7°–51.1°N, 30°W–120°E) in exciting rapid polar motions.

Christy et al. (1989) showed that during events of mean hemispheric atmospheric mass anomalies, the atmospheric mass tended to build up or evacuate from localized regions. They pointed to the importance of cross-equatorial atmospheric mass fluxes as a means by which these localized extreme anomalies could attain their high magnitudes. Observational evidence for the simultaneous occurrence of localized extreme anomalies and interhemispheric atmospheric mass exchange comes from a study by Trenberth (1986). Two major blocking events occurred over the South Pacific Ocean east of New Zealand near 50°S, during July and August of 1979. Between July and August, the hemispheric averaged SLP anomaly for the SH rose from −0.55 to 0.62 hPa. Trenberth (1986) showed that the rise in atmospheric mass was concentrated in the region of 50°S directly associated with the blocking events.

The purpose of this study is to characterize the time evolution of the large-scale circulation associated with extreme events of dry atmospheric mass loss from the NH on subseasonal time scales during the boreal winter. Persistent positive atmospheric mass anomalies are found to preferentially occur over Siberia, the North Pacific, and the North Atlantic during the boreal winter (Dole and Gordon 1983; Dole 1986; Higgins and Mo 1997). When these anomalies decay there is the potential for large interhemispheric atmospheric mass exchanges and, hence, for this reason we focus upon the loss of atmospheric mass from the NH. Using this characterization, we hope to isolate the dominant physical mechanism(s) that both directly and indirectly force these large dry atmospheric mass collapse events.

The studies of Trenberth (1981), Trenberth et al. (1987), Holl et al. (1988), Christy et al. (1989), Van den Dool and Saha (1993), and Chen et al. (1997b) have confirmed that interhemispheric atmospheric mass exchanges do occur with regularity throughout all seasons. The study of Christy et al. (1989) was important in documenting the atmospheric mass distributions at times when the hemispheric atmospheric mass was a local extrema. Their study did not address the physical mechanisms or dynamical processes that create the hemispheric extrema in atmospheric mass. Furthermore, an understanding of the time scale involved in the buildup or evacuation of the atmospheric mass from the hemisphere was lacking. Attempts to examine the redistribution of atmospheric mass have focused largely upon either zonal averages and or lower-frequency time scales (Lorenz 1951; Hsu and Wallace 1976; Christy and Trenberth 1985; Trenberth and Christy 1985; von Storch 2000; Baldwin 2001).

To accomplish these earlier stated objectives we employ a composite study approach. We outline the datasets used in this study and describe the methodology designed to isolate significant events of dry atmospheric mass loss from the NH in section 2. In section 3, we examine the time evolution of the composite NH dry atmospheric mass fall event, focusing upon SLP anomalies. Additionally, we document the localized regions that undergo significant atmospheric mass increases and decreases in both hemispheres during the composite event. In section 4, we focus upon the phenomenon of Southeast Asian pressure surges as a physical mechanism involved in the large and rapid collapses of dry atmospheric mass from the NH. Finally, in section 5 we provide a summary and conclusions.

2. Data and methodology

a. Data

We employ two sets of reanalysis data, the National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis (Kalnay et al. 1996; hereafter referred to as the NCEP reanalysis), and the ECMWF Re-Analysis (ERA; Gibson et al. 1997). Both the NCEP reanalysis and ERA make use of frozen state-of-the-art global data assimilation systems that are important for climatological studies. Additional data observations from a variety of sources, including rawinsonde profiles, surface marine reports from the Comprehensive Ocean–Atmospheric Data Set (COADS), aircraft observations, surface land reports and satellite cloud drift winds, and satellite soundings from TOVS (the TIROS Operational Vertical Sounder), not available in real time, were included in both reanalyses (Kalnay et al. 1996; Gibson et al. 1997).

A subset of the NCEP reanalysis data, extending over the 30-yr period from 1968 to 1997, is the primary dataset utilized to identify significant events of NH dry atmospheric mass fall, and to document the large-scale circulation that accompanies these events. We begin our study in 1968 owing to a problem with the NCEP reanalysis surface pressure and mean sea level pressure data (encoding error) prior to 1968. The complete ERA data extend over a shorter time period, 1979 to 1993, and will be used to examine the reproducibility of the results found with the NCEP reanalysis. Agreement between the two reanalyses will increase the reliability of our findings, especially over data-sparse regions such as the SH oceanic regions.

For both the NCEP reanalysis and ERA we use the daily averaged pressure-level archive data, which have been interpolated from the original model grids, to 2.5° × 2.5° global grids. The specific variables we use from the NCEP reanalysis are SLP, surface pressure, precipitable water, surface meridional wind, and the 1000- and 700-hPa geopotential height. From the ERA, we utilize SLP. Surface pressure values are especially relevant to this study. Trenberth et al. (1987) have shown that global mean areally weighted surface pressure is directly related to the global atmospheric mass, and can be thought of as one of the better known quantities in the atmosphere. Surface pressure is classified as a “B” variable using the NCEP reanalysis terminology (Kalnay et al. 1996). This class of variable is strongly influenced by the assimilating model, with the first-guess field successively modified or corrected by the inclusion of available hourly surface pressure observations (W. Ebisuzaki 2000, personal communication). Hence, surface pressure values are highly dependent on the model orography.

Figure 1a shows the interannual variability of atmospheric mass in the NCEP reanalysis, as represented by area-averaged daily surface pressures (Ps) for the NH, the SH, and the globe (G) for the 30-yr period 1968–97. The most striking feature is the presence of a vigorous annual cycle for both hemispheres and the globe. For both the NH and SH, the maximum (minimum) Ps occur during their cold (warm) seasons, respectively. The mean Ps for the globe, the NH, and the SH are 984.97, 982.21, and 987.70 hPa, respectively.

These values are similar to the long-term values calculated by Trenberth (1981) of 984.68, 982.84, and 987.16 hPa for the globe, the NH, and the SH, respectively. Trenberth utilized a combination of data sources made available through NCAR. The mean values calculated from the ERA for the 15-yr period 1979–93, are 984.52, 981.88, and 987.16 hPa for the globe, the NH, and the SH, respectively (Hoinka 1998). The ERA values were calculated on a higher-resolution 1.125° T106 (∼125 km) resolution grid.

The lower Ps values found in the ERA, when compared with NCEP, could be due in part to the greater mean orographic heights in the ERA. The mean orographic heights for the NCEP reanalysis on the 2.5° grid, for the globe, the NH, and the SH are 237.3, 283.2, and 191.3 m, respectively. The corresponding values for the ERA, on the high-resolution 1.125° grid, are 238.9, 285.6, and 192.2 m, respectively (Hoinka 1998).

The middle curve in Fig. 1a shows a time trace of Ps for the globe. The mass of the global atmosphere is not constant, but rather exhibits an annual cycle, being highest in the NH warm season. The range of atmospheric mass in the mean annual cycle for the NCEP reanalysis is 0.49 hPa, which is similar to that found by Trenberth (1981) of 0.47 hPa, but somewhat larger than the 0.40 hPa found for the ERA (Hoinka 1998). Recall that the global mass of dry air is constant and any fluctuations in the total mass of the atmosphere arise owing to variations in the water vapor content or errors in the analysis.

Figure 1b shows a time trace of surface pressure owing to vertically integrated water vapor (Pw) from the NCEP reanalysis. The precipitable water is calculated from the model data in sigma coordinates, vertically integrating from the model surface to 10 hPa (W. Ebisuzaki 1998, personal communication). The global curve, given in dark solid (Fig. 1b), exhibits a robust annual cycle with maximum (minimum) values in the NH summer (winter) months. The peaks in the global Pw curve are in phase with the peaks in the NH Pw curve and can be explained on the basis of the Clausius–Clapeyron relation (Trenberth 1981; Chen et al. 1996). The larger landmass area, or greater continentality, in the NH results in higher summer temperatures and higher Pw. Poleward of 40°, summer temperatures in the NH are 5°–30°C warmer than in the SH, during its respective summer (Trenberth 1981). The greater continentality in the NH also explains the larger annual range in hemispheric Pw (1.20 hPa) when compared with the SH (0.79 hPa).

Owing to the conservation of dry air mass, the mass of the global atmosphere will be greatest when Pw for the globe is greatest, and hence the in-phase relationship between global Ps and global Pw. The range of global Pw in the mean annual cycle from the NCEP reanalysis is 0.27 hPa. Trenberth (1981) calculated a mean annual cycle of 0.39 hPa, while Trenberth and Guillemot (1994) calculated a value of 0.37 hPa from globally analyzed data from ECMWF for a shorter period between 1985 and 1993. The corresponding annual range for the ERA is 0.25 hPa for the period 1979–93 (Hoinka 1998).

Figure 1c shows time traces of surface pressure owing to dry air (Pd), obtained by subtracting the appropriate curves in Fig. 1b from those in Fig. 1a. One can see from the global dry airmass curve, labeled G, that dry atmospheric mass is not strictly conserved in the NCEP reanalysis. This fact alone is not too surprising, Trenberth (1991) showed that the ECMWF analysis, archived on pressure surfaces, also fails to conserve dry atmospheric mass, while Chen et al. (1997b) found that for the reanalysis data generated by version I of the Goddard Earth Observing System (GEOS-I) Data Assimilation System (DAS; Schubert et al. 1993) dry atmospheric mass was not conserved.

In the NCEP reanalysis the global Pd has an annual range of 0.23 hPa. For comparison, the ERA possesses an annual range of 0.14 hPa (Hoinka 1998), while Trenberth (1981) calculated a range of 0.16 hPa. We argue that the greater annual cycle in global Pd in NCEP results, in part, from the dry bias in global Pw. The mean global Pw in NCEP for the 30-yr period 1968–97 is 2.34 hPa, which is low when compared with the long-term value of Trenberth (1981) of 2.48 hPa, and that from the ERA of 2.37 hPa (Hoinka 1998). Additionally the mean annual cycle in global Pw of 0.27 hPa in NCEP is significantly lower than that of Trenberth (1981) of 0.39 hPa. Given that the mean annual cycles of global Ps are similar between NCEP and that from Trenberth (1981), the greater mean annual cycle in global Pd in NCEP results from the reduced amplitude in global Pw. Trenberth and Guillemot (1998) found a similar dry bias in NCEP Pw. The mean annual cycles in Pw for the NCEP and ERA reanalyses are similar (0.27 versus 0.25 hPa), however the smaller mean annual cycle of global Ps in the ERA (0.40 hPa) when compared with NCEP (0.49 hPa) can explain the reduced amplitude of global Pd in the ERA.

b. Methodology

In this section we outline the procedure used to isolate significant events of dry atmospheric mass loss from the NH. Using the NCEP reanalysis, we calculated the daily NH Pd for the 30-yr period 1968–97 and subsequently removed the mean and the first three harmonics of the 30-yr mean annual cycle. The resultant time series can be thought of as the anomalous subseasonal variations of NH dry air mass, and will be referred to as simply the anomalous dry air mass for the NH.

Because we are interested in times when the NH is losing dry air mass, we consider the rate of change of the anomalous dry air mass using a forward difference scheme. When the rate of change of anomalous dry air mass first becomes negative we count the number of days until the rate of change becomes positive again and label this as an event. The procedure is similar to the threshold-crossing procedure employed by Dole and Gordon (1983) and is depicted schematically in Fig. 2b. The onset time for the given event (i.e., time when the rate of change of anomalous dry air mass first becomes negative), denoted T0, is given by the thin vertical line at time 0 in Fig. 2b. The event duration is defined as the total number of days between the local maximum and local minimum in the NH dry airmass anomalies (two thick vertical lines in Fig. 2b). Finally, we define a total anomalous dry airmass change, denoted as the event magnitude, as the difference in anomalous NH dry air mass between the local maximum and local minimum (see Fig. 2b). For each event we can define a duration (days) and a magnitude (hPa).

A significant event of NH dry atmospheric mass fall is defined as that subclass of events whose magnitude exceeds the 95th percentile (1.23 hPa). This criterion narrows the sample size considerably, but is necessary if we wish to isolate events of sufficient magnitude to impact the large-scale circulation. Greater than 65% of the significant events occurred during the NH cold season, defined as including the months of October–March, indicating a temporal preference for the more intense events to occur during the boreal winter. This fact, combined with the strong seasonality in the distribution of hemispheric-averaged surface pressure, and the dynamical forcing being seasonally dependent, we consider only those events with onset times during the NH cold season. To ensure that we are examining isolated, independent events, we require at least 15 days between the end of one event and the onset of the next.

A plot of the number of significant NH cold season dry atmospheric mass fall events as a function of duration is given in Fig. 2a. The event durations range from a minimum of 4 days (3 events) to a maximum of 17 days (1 event). The most frequent duration is 9 days (12 events), which provides the motivation to focus our study on those events of 6–10-day duration. Examining the entire spectrum of events from 4 to 17 days could pose potential problems as different dynamical processes may be duration-dependent. The result is a robust sample of 25 events. The onset dates for each event, and the total NH anomalous dry air surface pressure fall (event magnitude), are given in Table 1.

The composite NH cold season dry atmospheric mass fall event is shown in Fig. 2b. The time plotted on the abscissa is with reference to the onset time, T0, defined above. The large drop in NH anomalous Pd is evident beginning at T−1. The onset of the drop occurs on day T−1, instead of day T0, owing to the onset time being defined as the first day when the rate of change of anomalous Pd is negative, based upon a forward difference scheme. Hence on day T0, the dry atmospheric mass has already been falling for 1 day. Also plotted in Fig. 2b is the composite SH anomalous Pd, which displays a significant rise in anomalous dry air mass after day T−1, suggesting strong interhemispheric dry atmospheric mass compensation.

For the present study, events are chosen based upon a rate of change of dry atmospheric mass for the NH. The atmospheric mass of the NH is calculated directly rather than inferred as a residual based upon the difference between the global dry air mass and that of the SH. The conservation of dry air mass has not been invoked in the choice of our cases. Nonetheless we must demonstrate that the signal of our events, the fall of NH dry air mass, is not reduced to noise by the nonconservation of global dry atmospheric mass.

Figure 3 shows time traces of the daily anomalies of the NH and SH Pd, along with the daily anomalies of global Pd, for 41-day periods centered on the onset time for each of the 25 events listed in Table 1. The curves for the global Pd are not constant. However, their magnitudes do not mask the signal of the NH Pd for each event. The standard deviation of the anomalies of NH Pd (0.68 hPa) is approximately 5 times larger than that associated with global Pd (0.14 hPa) for the 41-day periods. Additionally, the high-frequency fluctuations in NH Pd are strongly reflected in the time series for the SH Pd, with a linear correlation coefficient of −0.91. Clearly, the dry airmass changes in the NH are reflected strongly in the SH and we feel confident that we can continue using the NCEP reanalysis data and the existing methodology.

3. Composite NH dry atmospheric mass collapse event

a. SLP anomalies

We document in this section the time evolution of the composite NH dry atmospheric mass fall event. In particular, we attempt to isolate physical mechanism(s) that are responsible for the large dry atmospheric mass evacuations from the NH on such a rapid time scale. Additionally, we document the localized regions that undergo significant atmospheric mass increases and decreases in both hemispheres during the composite event. Throughout the discussion the focus is global in scale.

Figure 4 depicts the time evolution of the NCEP composite SLP anomalies from 5 days prior to the onset of the composite NH dry atmospheric mass fall event (T−5) to 7 days following the onset (T+7). An anomaly is calculated for each individual event based upon a weighted monthly climatology (30 yr; 1968–97) centered on the onset time for the given event. For example, the climatological background field for event 1 (25 November 1968; see Table 1) is constructed by appropriately weighting the 30-yr monthly means of November and December. In addition, for each time-lagged composite, the statistical significance of the respective anomalies is determined using a two-sided Student's t test (Hogg and Tanis 1988), assuming a null hypothesis of zero anomaly. The statistical significance gives the probability that the composite mean anomaly is statistically different from zero.

Surface pressure is a more accurate measure of atmospheric mass (Trenberth 1981; Van den Dool and Saha 1993), when compared with SLP, owing to the necessary addition or subtraction of atmospheric mass when deriving SLPs. However, the SLP field is more representative of the circulation, and hence we show the SLP field in Fig. 4. The time-lagged composites for the conserved quantity of dry air surface pressure anomalies are very similar to those for SLP and hence are not shown. To facilitate the discussion, we focus upon three key stages: (i) the buildup of atmospheric mass in the NH prior to onset, (ii) the breakdown or collapse of atmospheric mass in the NH during the event, and (iii) the buildup of atmospheric mass in the SH toward the end of the event.

Beginning at T−5 (Fig. 4a, inset) the dry atmospheric mass for the NH rises steadily to a maximum value at T−1 (Fig. 4c, inset). This rise results largely from the growth of positive SLP anomalies over northern Eurasia, the North Pacific, and the North Atlantic Ocean regions (Figs. 4a–c). At T−1 (Fig. 4c) an extensive, zonally elongated band of positive SLP anomalies characterizes the latitudes poleward of 45°N, extending from northern Eurasia eastward to the North Atlantic, with values locally in excess of 8 hPa. Over northern Eurasia the positive SLP anomalies are related to the building of the Siberian high. Christy et al. (1989) showed that atmospheric mass anomalies in these regions, namely northern Eurasia, the North Pacific, and North Atlantic were positively related to the atmospheric mass anomaly for the NH during extreme events of anomalous hemispheric atmospheric mass. Note additionally the clear separation of the positive atmospheric mass anomalies at higher latitudes from the negative anomalies at lower latitudes over the North Pacific, the North Atlantic, and North America (Fig. 4c).

For the SH between T−5 (Fig. 4a, inset) and T−1 (Fig. 4c, inset) the steady fall of dry atmospheric mass is associated with a zonally extensive area of negative SLP anomalies poleward of 65°S (Figs. 4a–4c) and the weakening of two positive SLP anomaly centers south of Africa and over southeast South America. In South America we see evidence of anticyclonic ridging over the Andes mountains (Figs. 4a,b) followed by lee troughing (Fig. 4c) as the anticyclone propagates northeastward to the Brazilian coast (Lupo et al. 2001).

After the onset time T0 (not shown), the extensive zonally elongated band of positive SLP anomalies breaks down in the NH. The breakup occurs in conjunction with a deepening negative SLP anomaly in the Gulf of Alaska region of the North Pacific at T+1 (Fig. 4d). The negative SLP anomaly associated with the explosive cyclogenesis has a subtropical origin (Fig. 4b) and is not statistically significant until T+3 (Fig. 4e). Examination of the synoptic maps associated with the individual events reveals that the composite suffers from the effects of smearing as the timing of the explosive cyclogenesis in the northeast Pacific varies from event to event, typically occurring within 2 days of the onset date. Statistically significant positive SLP anomalies over both North America and Southeast Asia are seen to propagate equatorward to the east of the major orographic barriers.

The anomalous hemispheric dry air mass continues to fall in the NH, while it continues to rise in the SH from T+1 to T+3 (Figs. 4d,e). Two important features of note are the continued equatorward propagation of statistically significant positive atmospheric mass anomalies over both North America and Southeast Asia, and the deepening and spatial expansion of the negative SLP anomaly in the Gulf of Alaska. Colucci and Davenport (1987) examined events of rapid surface anticyclogenesis over western North America and found that the majority of the events were part of a downstream development involving upstream explosive cyclogenesis in the Pacific Ocean. Downstream of the explosive cyclogenesis, a 500-hPa ridge intensifies and forces a cold air outbreak over North America. The authors further noted that in some cases, the amplification of the trough associated with the cold air outbreak over North America, can induce explosive cyclogenesis downstream over the North Atlantic. At T+1 (Fig. 4d) and T+3 (Fig. 4e) one can see clearly the simultaneous explosive cyclogenesis in the Gulf of Alaska and the anticyclonic outbreak over North America.

Between T+3 (Fig. 4e) and T+5 (Fig. 4f) two regions of interhemispheric interaction are seen in the Tropics along the equator. Near the date line in the central Pacific, a statistically significant positive SLP anomaly extends across the equator, seen as a southward extension of the positive atmospheric mass anomaly extending southeastward from the Asian continent. Westward near Indonesia at T+5 (Fig. 4f) positive SLP anomalies extend southward from approximately 30°N to the extratropical SH between 90° and 120°E. The positive atmospheric mass anomaly over Southeast Asia, present at T+3 (Fig. 4e) has split into two, with one region propagating northeast along the coast of Japan, the other region remaining over China. Comparing the SLP anomaly distributions at T+3 (Fig. 4e) and T+5 (Fig. 4f), we see evidence of substantial ridging equatorward to the west of Australia, which appears to converge with the southward surge of atmospheric mass from Southeast Asia to create a region of positive atmospheric mass along the equator to the west of 120°E (Fig. 4f). Love (1985a) and Kiladis et al. (1994) have shown that the convergence of pressure surges from both hemispheres in the Indian Ocean–tropical western Pacific sector is favorable for the occurrence of westerly wind bursts owing to the enhanced zonal pressure gradient created.

During the Asian winter monsoon, the region of Southeast Asia is affected by frequent cold surges emanating from higher latitudes (Zhang et al. 1997; Compo et al. 1999), and is a preferred region for interhemispheric interaction (Williams 1981; Johnson et al. 1987; Kiladis et al. 1994; Suppiah and Wu 1998). For example, Chu and Park (1984) documented a maximum lower-tropospheric southward atmospheric mass flux between the longitudes of 100° and 130°E of 4.3 × 109 kg s−1 across 30°N and 13 × 109 kg s−1 across the equator, for a cold surge event occurring between 9 and 13 December 1978.

At day T+7 (Fig. 4g) the anomalous atmospheric mass in the SH has continued to rise and is most pronounced to the east of New Zealand in the central South Pacific. In the south Indian Ocean the low-latitude positive SLP anomaly center has retrogressed southwestward to a position near 60°E. Christy et al. (1989) noted a tendency for the atmospheric mass anomaly in the region of the central South Pacific to be positively related to the hemispheric anomaly during times of positive atmospheric mass in the SH summer. Trenberth and Mo (1985) and Sinclair (1996) found a preference for atmospheric blocking in the region of the South Pacific near New Zealand, while Trenberth and Mo (1985) discussed a secondary maximum in blocking frequency in the south Indian Ocean. Our results indicate that when the SH is undergoing an increase in dry atmospheric mass, the atmospheric mass preferentially builds up in localized regions commonly associated with atmospheric blocking.

b. SLP anomaly comparison: NCEP and ERA

A comparison of the time-lagged composites of SLP anomalies between NCEP and the ERA for the 10 common events that occurred between 1979 and 1993 (Table 1) is shown in Fig. 5. Agreement between the two reanalyses will provide increased confidence in our findings, especially over the data-sparse SH oceans. One noteworthy problem with the NCEP reanalysis that affects the surface pressure fields over the SH, from 1979 to 1992, is related to the Australian surface pressure bogus data (PAOBs; Garreaud and Wallace 1998). The PAOBs data were incorrectly assimilated (shifted 180° longitude) into the reanalysis. Sensitivity tests were performed at the Climate Prediction Center to determine the impact of this error. It was found that the SH region poleward of 40° was most affected, with the impacts diminishing as the spatial scale increases from synoptic to global. The PAOBs problem does not exist in the ERA.

The SLP anomalies for both the NCEP and ERA are calculated in a similar manner as described in section 3a, except that the weighted monthly climatologies for both NCEP and ERA are calculated from their respective 15-yr climatologies from 1979 to 1993. The time evolution of the large-scale features are qualitatively very similar between NCEP and ERA. Both reanalyses capture the large region of increasing positive SLP anomalies over the North Pacific and the rapid buildup of positive atmospheric mass anomalies over northern Eurasia prior to the onset (Figs. 5a,a′, 5b,b′, 5c,c′). Commencing at T−1 (Fig. 5c,c′), the surging of atmospheric mass equatorward over Southeast Asia and North America in conjunction with the deepening negative SLP anomaly in the North Pacific is captured by both reanalyses.

Nearing the end of the composite event (Figs. 5f,f′, 5g,g′) there is good agreement between both reanalyses in capturing the positive atmospheric mass anomalies in the south Indian and South Pacific Oceans. A closer examination of the SH reveals that the largest discrepancies between the reanalyses are found in the very high southern latitudes. The negative SLP anomalies in the NCEP reanalysis are consistently more negative to the west of the dateline over Antarctica. In contrast, the localized anomalies, both positive and negative, in the midlatitudes of the SH are more intense in the ERA.

c. Localized regions of atmospheric mass change

Summarizing the collapse of atmospheric mass for the composite NH event, we show in Fig. 6a the average, over all 25 events, of the SLP anomaly difference between the times of local maximum and local minimum of anomalous NH dry atmospheric mass for each event. Statistical significance calculated based upon a two-sided Student's t test is also shown. Note that the SLP anomaly fields, prior to calculating the individual event difference maps, were bandpass filtered, retaining periods between 6 and 30 days, using a Lanczos filter with 121 weights (Duchon 1979). Negative (positive) values denote regions that have undergone an atmospheric mass decrease (increase) throughout the composite event.

Virtually the entire area poleward of 45°N has undergone a loss of atmospheric mass. In particular, three distinct regions in the mid-to-high latitudes of the NH have experienced an atmospheric mass evacuation: (i) the North Atlantic to the south of Greenland, (ii) the North Pacific, extending southeastward to western North America, and (iii) the northern Asian continent. These regions are favored locations for low-frequency persistent anomalies during the NH winter months (Dole and Gordon 1983; Dole 1986; Higgins and Mo 1997). The finding that significant events of NH dry atmospheric mass decrease are manifested regionally as evacuations of atmospheric mass from regions of persistent anomalies is important, as these events may signal a change in circulation regime.

Also evident in Fig. 6a are statistically significant regions of atmospheric mass increase located in the central Pacific to the east of the dateline between 20° and 40°N, and the subtropical North Atlantic off the west coast of Africa. This mode of intrahemispheric redistribution of atmospheric mass, where the pressures in the NH high latitudes are negatively correlated with those in the subtropics, has been noted by several authors (Lorenz 1951; Trenberth and Paolino 1981; Christy and Trenberth 1985; Christy et al. 1989). Our results are consistent with these findings.

Atmospheric mass preferentially accumulates in two regions of the SH, the South Pacific Ocean to the east of New Zealand and the south Indian Ocean (Fig. 6a). These two regions were previously noted in our discussion of Fig. 4, but stand out more clearly in Fig. 6a. An additional region of statistically significant atmospheric mass increase is found in the South Atlantic to the south and east of South America. In the subtropical South Atlantic and south Indian Oceans, the regions of atmospheric mass buildup extend northward toward the equator. The South Pacific region to the south and east of New Zealand, and to a lesser degree, the south Indian Ocean are preferred regions for atmospheric blocking during austral summer (Trenberth and Mo 1985; Sinclair 1996). The dry atmospheric mass increase for the SH is manifested as a regional increase of atmospheric mass in noted areas of persistent anomalies.

Figure 6b presents a latitude–time plot of the zonally averaged (0°–360°) SLP anomalies, weighted by the cosine of the latitude, for the composite NH dry atmospheric mass fall event. Approaching T0 (onset time), the negative SLP anomalies in the SH increase in magnitude and expand equatorward, while in the NH, a buildup of positive SLP anomalies is seen, with much of the atmospheric mass accumulating in the midlatitudes, centered near 50°N. Recall the zonally extensive, statistically significant, elongated band of positive SLP anomalies poleward of 45°N in the NH at T−1 (Fig. 4c). At T−1, the anomalous dry atmospheric mass in the NH (SH) is at a local maximum (minimum) (Fig. 2b).

The composite interhemispheric dry atmospheric mass exchange event occurs in conjunction with an equatorward surge of positive SLP anomalies in the NH, between T+1 and T+4. This equatorward surge results from the combined effects of pressure surges over Southeast Asia and North America, and is depicted as the abrupt transition from positive to negative atmospheric mass anomalies in the mid- to high latitudes of the NH near T+2.

A remarkably similar atmospheric mass transition in the zonally averaged SLP anomalies of the SH occurs approximately 1 or 2 days after that of the NH. Much of the more pronounced changes in atmospheric mass in the SH occur at mid-to-high latitudes, poleward of 30°S. At approximately T+7 and T+8 days, the SH (NH) dry atmospheric mass anomaly reaches a local maximum (minimum) (Fig. 2b). Symmetric about the NH dry atmospheric mass fall event one can clearly see the quadrapole, with positive (negative) SLP anomalies in the NH (SH) midlatitudes prior to the composite event, and negative (positive) SLP anomalies in the NH (SH) midlatitudes after the composite event.

Recently, the meridional redistribution of atmospheric mass on subseasonal time scales has been viewed in the context of a hemispheric-wide oscillation termed the Arctic Oscillation (AO; Thompson and Wallace 1998) or the Northern Hemisphere annular mode (NAM; Thompson and Wallace 2000). Thompson and Wallace (1998) defined the AO as the leading principal component of the wintertime monthly mean SLP poleward of 20°N. The AO index values are calculated by projecting the leading standardized principal component time series of monthly mean SLP north of 20°N onto the daily SLP anomalies (these can be found on D. Thompson's Web site at www.atmos.colostate.edu/ao/Data/). The large atmospheric mass evacuation from the mid- to high latitudes of the NH in conjunction with a rapid transition from positive to negative SLP anomalies seen in our composite event (Fig. 6b) is suggestive of a relationship to the AO or NAM. Baldwin (2001) showed that fluctuations in the NAM index were accompanied by interhemispheric atmospheric mass exchanges. A plot of the AO index for the composite NH dry atmospheric mass fall event is shown in Fig. 7 for the 41-day period centered on the composite onset time. The AO index values were taken from D. Thompson's annular modes Web site.

As the onset time (T0) is approached, the composite AO index falls to a local minimum (Fig. 7) consistent with the building of the positive atmospheric mass anomalies at high latitudes in the NH (Fig. 6b). Subsequently, throughout the composite event, the AO index rises to a local maximum (T+6; Fig. 7) associated with the evacuation of the positive atmospheric mass anomalies from the high latitudes of the NH (Fig. 6b). Figure 7 must be viewed with caution as plots of the time evolution of the AO index for the 25 individual events revealed case-to-case variability. Nonetheless, pressure surges occurring over both Southeast Asia and North America are mechanisms involved in the evacuation of atmospheric mass out of the NH polar regions into the midlatitude and subtropical regions associated with a transition from the low index to the high-index phase of the AO.

4. Possible physical mechanism on subseasonal time scales

Our findings from section 3a showed a prominent feature of the composite NH dry atmospheric mass fall event was the rapid buildup of atmospheric mass over the northern Eurasian landmass prior to onset. Subsequently, the positive SLP anomalies were seen to surge equatorward in conjunction with the large, dry atmospheric mass evacuation from the NH. Composites involving the 25 events were also performed with reference to the time of maximum depletion of dry atmospheric mass for the NH, following the methodology of Wintels and Gyakum (2000). The pressure surge signature over Southeast Asia was again prominent, indicating that the presence or importance of this mechanism was not sensitive to the compositing procedure. Hence we feel that this phenomenon is a robust feature associated with the cold season dry atmospheric mass depletion from the NH.

The influences of Southeast Asian pressure surges are not only regional. Numerous studies have alluded to the planetary-scale impacts of cold surges (Lau et al. 1983; Lau and Li 1984; Lau and Lau 1984; Meehl et al. 1996; Iskenderian and Salstein 1998; Compo et al. 1999; Weickmann et al. 2000; Garreaud 2001). Simultaneous pressure surges occurring to the east of the Himalayas in Asia and to the east of the Rocky Mountains in North America can act as an important stochastic forcing for global atmospheric angular momentum changes through their effects on global mountain torques (Weickmann et al. 2000). A study by Iskenderian and Salstein (1998) provided further evidence for the significant role of pressure surges in forcing high-frequency fluctuations in global atmospheric angular momentum.

Cold pressure surges move rapidly, spreading equatorward from the extratropics to the subtropics over approximately a 4-day span (Garreaud 2001). The Siberian high is found to weaken considerably as atmospheric mass is expelled southward (Ding and Krishnamurti 1987). Numerous studies have documented the enhanced interhemispheric interactions and associated SH tropical and extratropical responses that result from the direct and indirect effects of the pressure surges (Williams 1981; Davidson et al. 1984; Love 1985a,b; Johnson et al. 1987; Kiladis et al. 1994; Suppiah and Wu 1998). The strengthened SH response appears to result from an overall enhancement of the local East Asian Hadley circulation and outbreak of convection in the Tropics following the pressure surges (Chang and Lau 1980; Lau et al. 1983; Meehl et al. 1996; Compo et al. 1999).

The onset of the Australian summer monsoon (ASM) has been linked to the occurrence of cold surges from Southeast Asia. Davidson et al. (1984) documented a link between divergent northerly surges emanating from subtropical anticyclones over Southeast Asia and the onset of the ASM in December 1978. A 12-yr composite study by Suppiah and Wu (1998) found that cold surges typically preceded the onset of the ASM by 5–10 days.

The extratropical flow in the SH is particularly sensitive to anomalous tropical convection over the equatorial western Pacific (Chen et al. 1989; Hurrell and Vincent 1990; Berbery and Nogués-Paegle 1993; Tyrell et al. 1996; Matthews et al. 1996; Mo and Higgins 1998; Renwick and Revell 1999). Kiladis and Weickmann (1992) found that when convection peaked in the equatorial trough north of Australia (15°–5°S, 140°–160°E), a prominent wave train was established after 6 days, with a pronounced upper-tropospheric ridge situated over the central South Pacific (see their Fig. 6b). Knutson and Weickmann (1987) found significant upper-level ridge development in the South Pacific when convection, associated with the Madden–Julian oscillation (Madden and Julian 1994), was enhanced over the equatorial trough north of Australia. Tyrrell et al. (1996) noted strong extratropical wave responses in the SH within 5 days of the period of maximum tropical convection in the western tropical Pacific during the Tropical Ocean Global Atmosphere Coupled Ocean–Atmosphere Response Experiment (TOGA COARE). Renwick and Revell (1999) showed that an extratropical Rossby wave train, induced by anomalous divergence associated with tropical outgoing longwave radiation anomalies in the equatorial western Pacific, can lead to the buildup of a blocking ridge (i.e., buildup of atmospheric mass) over the South Pacific within 10 days.

The surging of atmospheric mass southward over Southeast Asia and the decrease in the lower-tropospheric temperatures over China combined with an increase in the northerly component of the near-surface wind field to the south of China are three prominent characteristics of cold surges over Southeast Asia (Compo et al. 1999). Figure 6a shows three regions that were chosen to capture these prominent characteristics. An area situated over northern Eurasia (40°–75°N, 85°–145°E), hereafter referred to as the northern box, is chosen as it represents the statistically significant area that undergoes a large atmospheric mass evacuation during the composite event. A second area located to the south over southern China (20°–30°N, 105°–120°E), hereafter referred to as the southern box, is chosen as it captures the surging of atmospheric mass southward. Slingo (1998) utilized this same southern box in her study of tropical–extratropical interactions associated with cold surges over Southeast Asia. Finally, a third box located over the South China Sea (5°–10°N, 110°–115°E) is designed to capture the enhancement of the low-level northerly winds associated with the pressure surges.

Figure 8a depicts the time series of the composite SLP anomalies, averaged over the northern box (solid) and southern box (dash), for the 41-day period centered on the onset time for the composite event. A dramatic rise in area-averaged anomalous SLP beginning at T−5 and peaking at T0 is seen in the northern box. Subsequently, the atmospheric mass falls rapidly to a minimum at day T+9. The overall magnitude of the rise is approximately 6 hPa, while the subsequent fall is of the order of 8 hPa. The time series for the southern box also demonstrates a sharp rise in anomalous SLP peaking at T+2, two days after the northern box, and then falling out to T+9. The rise and subsequent fall in anomalous SLP for the southern box are not as pronounced as for the northern box, the rise being approximately 4 hPa, while the fall is on the order of 4.5 hPa. The results in Fig. 8a are not sensitive to the use of SLP, with the results for dry air surface pressures very similar (not shown).

Based upon Fig. 8a, we developed a methodology to determine the subset of the 25 events that possess a Southeast Asian pressure surge. A figure similar to Fig. 8a was constructed for each of the 25 individual events. For a case to be considered to possess a Southeast Asian pressure surge, the following three criterion had to be met: (i) there must be a local maximum in the anomalous SLP, averaged over the northern box, which occurs within 3 days of the event onset; (ii) there must be a local maximum in the anomalous SLP, averaged over the southern box, which occurs no more than 3 days after the peak in the northern box; and (iii) there must be an overall decrease in the anomalous SLP, averaged over the northern box, from event onset to event end. A total of 16 events satisfied these three criteria and referring back to Table 1, the subset of cases consist of events 1, 3, 5, 6, 8, 10, 11, 12, 16, 17, 19, 20, 22, 23, 24, and 25. The fact that close to two-thirds of the events possessed a Southeast Asian pressure surge, following the application of our methodology, is consistent with the hypothesis of the importance of this physical mechanism in relation to the depletion of dry atmospheric mass from the NH during boreal winter.

A time series of the composite 1000–850-hPa thickness anomalies, averaged over the southern box, for the 41-day period centered on the onset time is shown in Fig. 8b. One can clearly see a large drop in the lower-tropospheric mean temperature that coincides with the significant drop in anomalous SLP over northern Eurasia (Fig. 8a). The northerly winds in the South China Sea gradually increase in strength after the onset time and peak 6 days after the onset (Fig. 8c).

5. Summary and conclusions

This study examined the phenomenon of interhemispheric dry atmospheric mass exchange. Specifically, we restricted our investigation to times when the NH loses dry atmospheric mass on subseasonal time scales during the boreal cold season. The research was motivated, in part, by the finding that the atmospheric mass of the NH departed significantly from its mean value on subseasonal time scales (Trenberth 1986; Trenberth et al. 1987; Holl et al. 1988; Christy et al. 1989), and that this departure was manifested in terms of localized extremes of atmospheric mass (Christy et al. 1989). Nonetheless, these studies did not address the physical mechanism(s) or time scales involved in the buildup or evacuation of the atmospheric mass from the hemisphere.

Daily interhemispheric dry atmospheric mass exchanges were examined over a 30-yr period from 1968 to 1997. Such a long-term study is unprecedented and allowed us to address the temporal distribution of NH dry atmospheric mass collapse events with a greater degree of confidence. As a comparison, the study period of Christy et al. (1989) extended from 1980 to 1985. Among the significant events, as defined in section 2b, more than two-thirds of the events occurred during the boreal winter season. One possible explanation is that during the boreal winter the building and strengthening of the Siberian high can create significant positive atmospheric mass anomalies over northern Eurasia (Ding and Krishnamurti 1987; Zhang et al. 1997), which as we saw in section 3 (Fig. 4) is a precursor to the large NH dry atmospheric mass collapse events.

Prior to the onset of the composite NH dry atmospheric mass fall event, the atmospheric mass increase in the NH is manifested as localized positive atmospheric mass anomalies over the western North Atlantic, the central North Pacific, and over northern Eurasia (Figs. 4a–4c). These regions are frequently associated with persistent circulation anomalies (Dole and Gordon 1983). Over northern Eurasia the increase in the positive atmospheric mass anomaly is impressive and is associated with the building of the Siberian high (Figs. 4a–4c). In the SH, the decrease in dry atmospheric mass prior to the onset is associated with an extensive region of negative SLP anomalies poleward of 60° (Figs. 4a–4c).

The breakdown or collapse of the NH dry atmospheric mass is a multiple-phenomenon event. Explosive cyclogenesis in the region of the Gulf of Alaska is seen to be important in breaking up the large positive atmospheric mass anomaly in the central North Pacific (Figs. 4b–4f). Examination of the individual events indicates that explosive cyclogenesis typically occurs within 2 days of the onset and can be traced back several days to the subtropical North Pacific (Figs. 4b,c). Pressure surges over Southeast Asia and North America, associated with statistically significant positive atmospheric mass anomalies, act to channel the atmospheric mass equatorward out of the NH extratropics on a rapid time scale (∼4 days; Figs. 4c–e). The typical time scale for cold pressure surges (∼4 days; Garreaud 2001) is consistent with both the observed rapid equatorward evacuation of atmospheric mass from the higher latitudes, and the preferred time scale of 9 days (see section 2b) for the duration of the collapse events (Fig. 2a).

At the end of the composite event, the atmospheric mass increase in the SH is manifested as enhanced extratropical surface ridging over the South Pacific and south Indian Ocean regions (Fig. 4g). These two regions are noted for their frequent occurrence of atmospheric blocking (Trenberth and Mo 1985; Sinclair 1996). Our findings suggest that when the SH undergoes an increase in dry atmospheric mass, the atmospheric mass shows a preference to accumulate in preferred regions for atmospheric blocking. A comparison of the NCEP and ERA composite SLP anomalies for the 10 common events between 1979 and 1993 (Table 1; Fig. 5) was undertaken to validate the structures. It is noteworthy that the Australian surface pressure bogus data (PAOBs) problem (see section 3b) was not present in the ERA. Differences in the anomaly magnitudes were found, however, the large-scale features were qualitatively similar, increasing our confidence in the findings especially over the data-sparse regions of the SH.

Section 4 considered the role of Southeast Asian pressure surges as a physical mechanism directly related to the large collapses of dry atmospheric mass from the NH. The findings from Fig. 6a showed a dramatic evacuation of atmospheric mass from northern Eurasia during the composite event. Prominent characteristics of cold surges, namely the surging of atmospheric mass southward over Southeast Asia, the decrease in the lower-tropospheric temperatures over China, and the increase in the northerly component of the near-surface wind field to the south of China, were strongly present in the composite event. Close to two-thirds of the events were defined to possess a Southeast Asian pressure surge based upon our methodology derived from Fig. 8a. The presence of statistically significant positive SLP anomalies extending from 30°N to the extratropics of the SH between 90° and 120°E during the composite NH dry atmospheric mass collapse event (Fig. 4f) suggests a preferred zone of interhemispheric interaction associated with the pressure surge.

Efforts to examine the redistribution of atmospheric mass have focused largely upon zonal averages and or lower-frequency time scales (Lorenz 1951; Hsu and Wallace 1976; Christy and Trenberth 1985; Trenberth and Christy 1985; von Storch 2000; Baldwin 2001). We have shown in this study that pressure surges occurring over both North America and Southeast Asia are physical mechanisms forcing the atmospheric mass southward out of the higher latitudes on a rapid time scale during the NH dry atmospheric mass collapse events.

Future work will examine the direct and indirect roles of Southeast Asian pressure surges for a representative case study of dry atmospheric mass loss from the NH. Specially we will show how the pressure surge is linked to a strengthening of the local East Asian Hadley circulation in association with the onset of anomalous convection in the SH Tropics. The redistribution of atmospheric mass in the upper-tropospheric divergent outflow, emanating from this anomalous convection, will be examined through calculations of the vertically integrated dry airmass flux and potential vorticity.

This study has focused upon the loss of atmospheric mass from the NH. One could perform a similar analysis for events of atmospheric mass gain for the NH. It would be interesting to compare the time evolution of the SLP anomalies, similar to Fig. 4, to determine just how linear the two composites are. Are the same geographical areas involved in the gain and loss of atmospheric mass from the NH?

Acknowledgments

We appreciate helpful discussions with Dr. Klaus Weickmann, Dr. Gilbert P. Compo, and Dr. George Kiladis. Professor Gary Lackmann and Dr. Compo provided assistance with the Lanczos filter. The NCEP reanalysis data was provided by NOAA-CIRES Climate Diagnostics Center, Boulder, Colorado, from their Web site at http://www.cdc.noaa.gov and we acknowledge their support. The ECMWF reanalysis data was acquired from the Mass Storage System at the National Center for Atmospheric Research (NCAR). This study benefited from the comments and suggestions of two anonymous reviewers.

Financial support for this research was provided by the Atmospheric Environment Service of the government of Canada through a postgraduate fellowship for one year, and also by the Natural Science and Engineering Research Council of Canada (NSERC) through a two-year postgraduate fellowship. Subsequent support was provided by grants from NSERC and from the Meteorological Service of Canada and an NSERC network research grant on the Mackenzie GEWEX Project. Finally I would like to thank María Belén Herrero for her support and encouragement.

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Fig. 1.
Fig. 1.

NCEP reanalysis for the period 1968–97. Global and hemispheric daily area-averaged (a) surface pressure, (b) surface pressure owing to vertically integrated water vapor, and (c) surface pressure owing to dry air (hPa). Note that NH refers to the Northern Hemisphere, SH to the Southern Hemisphere, and G to the globe

Citation: Journal of Climate 16, 24; 10.1175/1520-0442(2003)016<4061:SEOIAM>2.0.CO;2

Fig. 2.
Fig. 2.

(a) Distribution of NH cold season dry atmospheric mass fall events as a function of duration. (b) Composite hemispheric dry air surface pressure anomaly for the 25 cold season events of 6–10-day duration shown in (a) with respect to the composite onset time of 0 (T0). The concepts of onset time, event duration and magnitude, discussed in section 2b, are illustrated in (b).

Citation: Journal of Climate 16, 24; 10.1175/1520-0442(2003)016<4061:SEOIAM>2.0.CO;2

Fig. 3.
Fig. 3.

Time traces of area-averaged dry air surface pressure anomalies for the NH (thick solid), the SH (dotted), and the globe (thin solid) for 41-day periods centered on the onset time (0) for each of the 25 Northern Hemisphere cold season dry atmospheric mass fall events. Units on the ordinate are hPa, and the two vertical lines in each plot denote the onset and termination times for each event. Event numbers (onset times) are shown in the top right (left) of each plot and correspond to those listed in Table 1. The convention for the onset times is month/day/year (two digit)

Citation: Journal of Climate 16, 24; 10.1175/1520-0442(2003)016<4061:SEOIAM>2.0.CO;2

Fig. 4.
Fig. 4.

Composite anomalies of SLP, with a contour interval of 1 hPa to a magnitude of 3 hPa, and every 2 hPa for larger magnitudes: (a) T−5, (b) T−3, (c) T−1, (d) T+1, (e) T+3, (f) T+5, and (g) T+7. Light (dark) shading denotes statistical significance at the 95% (99%) levels according to a two-sided Student's t test. Inset shows plots of the NH (solid) and SH (dash) dry air surface pressure anomalies (hPa) as a function of time from onset (denoted by vertical line). Thick dots denote the time of the horizontal SLP anomaly plot

Citation: Journal of Climate 16, 24; 10.1175/1520-0442(2003)016<4061:SEOIAM>2.0.CO;2

Fig. 5.
Fig. 5.

Composite SLP anomalies for the 10 NH cold season dry atmospheric mass fall events between 1979 and 1993: (a), (a′) T−5; (b), (b′) T−3; (c), (c′) T−1; (d), (d′) T+1; (e), (e′) T+3; (f), (f′) T+5; and (g), (g′) T+7. Panels (a)–(g) are derived from the NCEP reanalysis, while (a′)–(g′) are derived from the ECWMF reanalysis. The contour interval is 2 hPa, with negative values dashed

Citation: Journal of Climate 16, 24; 10.1175/1520-0442(2003)016<4061:SEOIAM>2.0.CO;2

Fig. 6.
Fig. 6.

(a) Composite SLP anomaly difference, with a contour interval of 1 hPa to a magnitude of 4 hPa, and every 2 hPa for larger magnitudes. For each event the SLP anomaly difference is calculated as the difference in SLP between the times of local max and local min in anomalous NH dry atmospheric mass for each event (see Fig. 3). Prior to calculating the SLP anomaly difference maps, the SLP was temporally filtered, retaining periods between 6 and 30 days. Light (dark) shading indicates the statistical significance at the 95% (99%) level based upon a two-sided Student's t test. The three boxes refer to the northern, southern, and South China Sea areas discussed in section 4. (b) Latitude–time plot of the zonally averaged (0°–360°) composite SLP anomalies (hPa), weighted by the cosine of latitude. Time given on the ordinate is with respect to the onset time, denoted as T0. Shading shows regions of positive SLP anomalies. No filtering was applied in (b)

Citation: Journal of Climate 16, 24; 10.1175/1520-0442(2003)016<4061:SEOIAM>2.0.CO;2

Fig. 7.
Fig. 7.

Composite Arctic Oscillation (AO) index for the 25 events of NH dry atmospheric mass fall. The index values represent the projection of the leading standardized principal component of NH SLP (north of 20°N) onto the daily SLP anomalies (see Thompson and Wallace 2000). The values are dimensionless and were obtained online from http://www.colostate.edu/ao/Data/. Time on the abscissa covers the 41-day period centered on the composite onset time (T0)

Citation: Journal of Climate 16, 24; 10.1175/1520-0442(2003)016<4061:SEOIAM>2.0.CO;2

Fig. 8.
Fig. 8.

(a) Time series of composite sea level pressure anomalies (hPa), area-averaged over 40°–75°N, 85°–145°E (northern box shown in Fig. 6a; thick solid), and area-averaged over 20°–30°N, 105°–120°E (southern box shown in Fig. 6a; dash). (b) Time series of composite 1000–850-hPa thickness anomalies (m) averaged over the southern box. (c) Time series of composite surface meridional wind anomalies (m s−1) averaged over the South China Sea (5°–10°N, 110°–115°E; see Fig. 6a). Time on the abscissa is for the 41-day period centered on the composite onset time (T0)

Citation: Journal of Climate 16, 24; 10.1175/1520-0442(2003)016<4061:SEOIAM>2.0.CO;2

Table 1.

Summary of Northern Hemisphere cold season dry atmospheric mass fall events. The event duration and magnitudes (denoted Δ Pd) defined in section 2b are shown in the last two columns

Table 1.
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