1. Introduction
Little time had passed after aircraft pilots’ early confrontations with jet streams when meteorologists started to attach high importance to these flow features. Much of the rationale for studying jet stream variability is based on its dynamical significance in synoptic-scale cyclogenesis (Sutcliffe 1939, 1947; Nakamura 1992; Baehr et al. 1999; Wernli et al. 2002) and other phenomena dynamically linked to jet variability such as atmospheric blocking and the propagation and dispersion of Rossby waves (Wallace et al. 1988; Nakamura and Wallace 1990; Schwierz et al. 2004). A review by Chang et al. (2002) largely focuses on physical processes vital to storm-track dynamics.
The particular attention the jet stream has received in central,1 South, and East Asia has a number of reasons. First, the multitude of processes relevant to jet stream dynamics makes its study an intriguing and challenging task. In the Tibetan Plateau region, the upper-level flow responds to mechanical and thermodynamical forcing by the Asian continent and the planet’s largest topographic barriers (Academica Sinica 1958a,b; Webster et al. 1998; Yanai et al. 1992; Wu and Liu 2003; Duan and Wu 2005) as well as to the strong diabatic heating associated with the precipitation during the Asian summer monsoon (Rodwell and Hoskins 1996; Liu and Yanai 2001; Randel and Park 2006) or in the eastern Indian Ocean during boreal winter (Barlow et al. 2002, 2005, 2007). Moreover, the region is located downstream of the exit region of the North Africa/Arabian jet and at the entrance region of the East Asian jet. Such regions are of particular interest from a number of perspectives (see the introduction of Barlow et al. 2005 for a discussion of this point). Second, the interaction of the flow with the topography has a pronounced bearing on both regional and hemispheric scales. For example, Nigam and Lindzen (1989) have shown how modest variations in the latitudinal position of the jet over the Himalayas modulate the amount of stationary wave flux reaching mid- and high latitudes. Such processes are of particular relevance because of the importance of orographic gravity wave drag for the extratropical circulation Palmer et al. (1986), and since secondary orographic interactions play a decisive role in the generation of extratropical circulation anomalies in response to anomalies in tropical heating (DeWeaver and Nigam 1995; Nigam and De Weaver 1998). Regionally, the effects of the topography include splitting and confluencing of the westerlies, damping of large eastward-moving troughs, sheltering in so-called dead-water regions, and thermal effects (Academica Sinica 1958a). Third, changes in the jet position can be indicative of large-scale seasonal changes in the prevailing flow regime. For example, it was shown (Yeh et al. 1959; Li et al. 2005; Li and Pan 2006) that a northward jump of the westerly jet over East Asia in early May typically precedes the onset of the summer monsoon over the South China Sea. In the regions both upstream and downstream of the Tibetan Plateau, it has long been recognized that the seasonal meridional translation of the jet stream is a good index for the determination of the natural synoptic season (Bugayev et al. 1957; Academica Sinica 1957; Chanysheva et al. 1995).
The results of early work on the general circulation over East Asia carried out by Chinese researchers are summarized in a series of papers in the English language (Academica Sinica 1957, 1958a,b). These studies are based on observations from a comparatively dense network of sounding and pilot balloon (pibal) stations and include analyses of upper-level fields in the Tibetan Plateau region. Examples of the seasonal evolution of the upper-level flow in the Northern Hemisphere at a wide range of longitudes can be found in Yeh et al. (1959). On the hemispheric scale, an early analysis of the upper-level flow has been provided by Sadler (1975). It has been used in the work of Newton (2004) for the examination of cyclone track and jet variability over the eastern hemisphere including the climatological seasonal cycle of the jet stream latitude in the Tibetan Plateau region at 80°E. The analysis by Blackmon et al. (1977) includes the wind at 500 hPa and the geopotential height at 300 hPa. The bandpassed (2.5–6 days) fields of this analysis were used to identify the storm tracks in the Northern Hemisphere, which were found to be located downstream and somewhat poleward of the jet axes.
Only few studies have analyzed the seasonal cycle and interannual variability of the jet stream in the Tibetan Plateau region by means of modern reanalysis data. Nakamura (1992) used the National Meteorological Center’s (NMC) operational analysis to study the relationship between the jet intensity and baroclinic wave activity over the North Atlantic and North Pacific. They found baroclinic wave activity to be negatively correlated with jet wind speed if this speed exceeds a threshold value of ∼45 m s−1 and related this to the observed midwinter suppression of baroclinic wave activity over the North Pacific. Monthly data from the National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis (Kalnay et al. 1996) were used by Kuang and Zhang (2005) for the investigation of the subtropical westerly jet over East Asia. They examined the seasonal cycle of the jet and how it is related to the meridional temperature gradient in the upper troposphere. Koch et al. (2006) developed a global jet climatology at high temporal resolution on the basis of 6-hourly fields from the 15-yr European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA-15; Gibson et al. 1999). Working with such high temporal resolution is in line with a previous argument of Shapiro and Keyser (1990), who noted that seasonally or zonally averaged atmospheric cross sections depict the subtropical jet, but cannot be used for the study of the more transient polar and arctic jet streams.
The objective of this paper is to follow up on the above studies by focusing on (i) the interannual variability of the westerly jet stream in the Tibetan Plateau region, (ii) the seasonal cycle of the jet location with particular emphasis on the transitions between the north and south sides of the plateau in spring and autumn, and (iii) the covariation of jet location and intensity with the tropospheric temperature fields in order to better understand the previously analyzed climatology. We not only take account of monthly mean fields but also consider the more transient jet characteristics. To this end, we derive a climatology based on jet stream occurrences following an approach similar to that by Koch et al. (2006).
The structure of the paper is as follows: we continue in section 2 with a description of the data used and of how we derive the occurrence-based jet climatology. Subsequently, the climatology is presented in section 3. Section 4 deals with the relationship between the jet and the atmospheric temperature. Thereafter, a short analysis of the covariability between the jet stream location and surface precipitation is presented in section 5. We conclude with a summary and discussion of our results.
2. Data and methods
Currently, the jet identification has been carried out in a spatial domain covering 16.75°–58.25°N, 42.5°–220.5°E and 500–100 hPa at time intervals of 6 h during 1958–2001. In principle, the dataset can be extended to the hemispheric or global scale. The identification successfully captures the major jet axes, as can be verified by means of Fig. 1 and the animation that is provided as supplementary material (available at the Journals Online Web site: http://dx.doi.org/10.1175/2009JCLI2625.s1). We do not distinguish several types of jet occurrences as done by Koch et al. (2006), but our data would allow for an a posteriori classification of this kind.
In addition to the dataset based on ERA-40 wind fields, we derived a completely analogous dataset of jet occurrences from the horizontal wind fields of a regional climate model simulation with the Climate High-Resolution Model (CHRM). Several versions of the model have been validated and used in a wide range of applications over European (Lüthi et al. 1996; Vidale et al. 2003; Schär et al. 2004; Hohenegger and Vidale 2005; Seneviratne et al. 2006; Fischer et al. 2007) and Asian domains (Fukutome et al. 1999; Schiemann et al. 2008). The CHRM simulation relies on ERA-40 data for the specification of its initial and boundary conditions. The horizontal resolution is 0.5° and the simulation spans the time period 1958–2001. Further details of the model setup can be found in Schiemann et al. (2008). Since the jet climatology derived from the numerically downscaled data is very similar to that obtained by means of the ERA-40 wind fields, we chose to mostly present results from the ERA-40 climatology. Observations, in particular radiosonde data, from within the domain of interest are not used as input to the CHRM. The good agreement between the two datasets seems to suggest that the limited availability of radiosonde data in the Tibetan Plateau region is not crucial for the results obtained herein.
Apart from ERA-40 data, we use the monthly all-India rainfall time series as a proxy for the intensity of the Indian summer monsoon (Parthasarathy et al. 1994). These data are provided by the India Meteorological Department and the Indian Institute of Tropical Meteorology for the time period 1871–2004.
3. Jet climatology
a. Monthly mean climatologies
An overview of the monthly mean wind distribution derived from ERA-40 data is provided in Figs. 2 and 3. Since the December and February climatologies are very similar to that of January, and the August climatology strongly resembles that for July, we can give a fairly complete picture of the seasonal cycle by showing climatologies corresponding to nine months.
The horizontal location of the jet on the Northern Hemisphere is shown in Fig. 2 in terms of the horizontal wind speed and the geopotential height at 200 hPa. In winter, the jet axis forms a typical circumpolar spiral-like structure and there is a pronounced maximum of horizontal wind speed in the western Pacific and additional maxima to the east of the North American continent and over the Arabian Peninsula. The range of the jet latitudes is marked by the two end points of the spiral over the North Atlantic at about 15° and 50°N. In spring, the jet intensity weakens and, while the distinction of the two spiral end points over the Atlantic becomes less clear than in winter, two such end points appear over the Pacific. The summer jet, in turn, has a structure similar to that in winter, but it is significantly weaker and shifted to higher latitudes: the spiral end points over southern Europe and the North Atlantic are now located at about 37° and 52°N. The strongest summer jet occurs over Asia and stretches from the eastern Mediterranean to the northern flanks of the Tibetan Plateau. In September, the jet starts to gradually intensify and to recede southward until it adopts its winter structure in December. Figure 2f also shows the easterly Asian jet that is characteristic for the Asian summer monsoon season.
Figure 3 describes the vertical structure of the jet in the Tibetan Plateau region by means of cross sections of horizontal wind speed and potential temperature at 85°E. Throughout the year, the jet axis is located at approximately 200 hPa. In winter, the jet is found at the southern edge of the Tibetan Plateau at about 28°N. While the main jet axis remains at that position until April/May, a strong weakening as well as a northward dispersion of the region of strongest horizontal wind is observed in spring (Figs. 3b,c). During May, the climatological jet axis moves northward and its mean location is over the Tibetan Plateau. May is also the month of the weakest jet intensity. The monthly mean horizontal wind during May is ∼30 m s−1 as compared to a speed of ∼55 m s−1 in January. In June, a well-defined jet reappears to the north of the Tibetan Plateau. The maximum jet latitude of about 42°N is attained in July and August. A gradual recession to winter conditions starts in September. In autumn, the jet is generally somewhat stronger and better defined than in spring.
b. Interannual variability and occurrence-based climatologies
In this section, we consider the occurrence-based jet climatologies derived from the 6-hourly fields of both ERA-40 and the CHRM to complement the results based on monthly means presented in the previous subsection. To this end, not only occurrence count climatologies (Figs. 4 and 5) but also instantaneous jet occurrence maps (see Fig. 1 and the supplementary material) are taken into account. The left-hand side panels in Figs. 4 and 5 show (i) the horizontal distribution of total occurrence counts in 1958–2001 based on 6-hourly ERA-40 fields and (ii) boxplot statistics2 of the jet axis locations obtained from the ERA-40 monthly mean horizontal wind. The right-hand panels show the meridional distribution of the jet counts that was obtained by aggregating all counts between 80° and 90°E (the region marked by the two solid vertical lines in the panels on the left). These distributions are shown for both the ERA-40 and CHRM jet climatologies.
In winter, a pronounced maximum in the jet counts is located at about 28°N and tilts northward to the east of the Tibetan Plateau. This corresponds well to the wintertime jet location as given by the monthly mean climatologies in Figs. 2a and 3a. In addition, however, the occurrence-based climatology captures a minimum of jet counts over the north of the Tibetan Plateau at about 36°N and a weaker secondary maximum of counts to the north of the plateau roughly between 40° and 45°N. These features are—at best—only partially captured by the monthly mean fields and can be interpreted as follows: in a highly simplified picture, two principal flow situations can be distinguished in winter. Either there is a single well-pronounced jet south of the Tibetan Plateau (see Fig. 1a) or the flow is split such that jet occurrences are identified both to the north and to the south of the plateau as in Fig. 1b.
The secondary northern maximum becomes more pronounced in spring, while the primary maximum south of the Tibetan Plateau weakens. The distributions in Figs. 4d,f,h show clearly that spring jet counts are identified over a wide range of latitudes, which is consistent with the increasingly disperse jet structure in Figs. 3b,c. The supplementary material and Fig. 1c give an impression of what the upper-level flow typically looks like in spring: in contrast to the wintertime situation, single well-defined jets that extend zonally over the entire study region are hardly observed. Instead, highly transient jetlike strips are found to move across the domain. In May, the meridional dispersion of the jet is so strong that the corresponding distribution of jet counts is almost uniform north of about 30°N in 80°–90°E.
This changes quite drastically in June. By this time, the jet has moved to the north of the Tibetan Plateau with its meridional distribution peaking sharply at the northern edge of major orography. A well-defined jet can be observed during the entire summer season (e.g., Figs. 1d,e). To the east of the Tibetan Plateau, the axis of maximum jet counts is slightly tilted southward. From June to July, there is a minor northward shift of the jet directly north of the plateau (a bit more than 1° in 80°–90°E). To the west and east of it, however, these shifts are quite substantial and amount to ∼5° over the central Asian plains as well as over eastern China and the Korean Peninsula. During the summer monsoon season, no jet events are counted south of ∼35°N.
The southward recession of the jet during fall is not completely symmetric to its northward transition in spring: in October, the meridional dispersion of the jet counts is smaller than in spring, and well-pronounced jets typically occur over the Tibetan Plateau during that time (e.g., Fig. 1f). By December, the jet has moved back to its winter position.
The meridional spread of the jet counts in Figs. 4 and 5 corresponds well to the interannual variability of the jet monthly mean locations. To give an impression of the interannual variability, box plots have been added to the left-hand side panels. They correspond to the variability of the monthly mean jet latitude in each year determined by the latitudinal position of the jet as represented by the monthly mean wind speed fields discussed in section 3a. Furthermore, the standard deviation of interannual jet count variability can be read off from the distributions on the right-hand side panels. It turns out, however, that the mean meridional dispersion of total jet counts gives a fairly good impression of the interannual variability. It is comparatively small in winter and summer and higher during the transition seasons, especially during spring over the Tibetan Plateau.
c. Meridional transition in the Tibetan Plateau region
There is a strong variability in the meridional position of the jet at synoptic time scales, in particular during its northward and southward transitions in spring and fall (see the animation provided as supplementary material for illustration). This complicates an objective and quantitative characterization of these transitions, and we present here a pragmatic approach to the problem. First, jet counts of the original 6-hourly climatology were aggregated to a dataset of daily jet latitudes in the (western) Tibetan Plateau region (80°–90°E) by computing the median latitude of all jet occurrences within these longitudes on each day. These daily jet latitudes are displayed for three different years in Fig. 6 (open circles and dotted lines), showing considerable variability of the jet at a wide range of time scales.
Figure 6 shows filtered time series with the synoptic variability removed (bold black lines). These were calculated by means of a low-pass filter with a cutoff period of 30 days following the procedure outlined in section 17.5.7 of von Storch and Zwiers (1999). The horizontal dashed line at 36°N marks the latitude of the jet count minimum in spring (see Fig. 4) and roughly separates jet locations corresponding to summer and winter regimes.
The springtime northward transitions differ considerably from one year to another. In 2000, there is a single well-defined northward transition of the jet at the end of April. In most years, however, it is not possible to determine a sharp and monotonic transition based on the low-passed time series of the jet latitude. In 1998, for example, the jet moved back south of the 36°N parallel after its first northward transition. This also happened in 1999 and, moreover, the jet latitude increased only very slowly in a period of four months from March till June. Finally, from the synoptic variability observed it can be concluded that, even though rapid latitudinal transitions, similar to the jumps at 105°–120°E (Li et al. 2005; Li and Pan 2006), do occur in the Tibetan Plateau region, it appears that these transitions are not necessarily associated with persistent changes in the large-scale circulation.
Northward and southward transition dates defined as the dates when the filtered time series crosses the 36°N parallel are summarized in Fig. 7 and in Table 1. The median northward transition date is 28 April, and the median southward transition date is 12 October. The southward transition in autumn is typically more gradual than in spring, yet a unique southward transition date (autumn) can still be defined in most years. Statistically, this transition occurs within a rather narrow time window. In contrast, the variability of the northward transition dates (spring) is considerably more pronounced. Indeed, northward excursions of the jet can occur at any time during the cold season. Despite the northward migration of the midlatitude jet during the last 20 yr documented by Archer and Caldeira (2008), this analysis does not point to a trend in the jet transition dates, neither over the complete ERA-40 period nor when periods before and after the assimilation of satellite data in 1978 are considered separately.
4. Temperature distribution changes and the jet
The thermal wind relation [Eq. (2)] can be used to estimate the change in the jet speed that would be expected because of the seasonal meridional translation of the jet axis alone (supposing other factors remained the same). Since the thermal wind is a geostrophic wind difference, it is inversely proportional to the Coriolis parameter f, just like the geostrophic wind itself. At 85°E, the ratio between summer and winter Coriolis acceleration at the jet latitude is fs/fw = sin(ϕs)/sin(ϕw) ≈ sin(42°)/sin(28°) ≈ 1.43. On the other hand, the observed ratio of the maximum jet speeds in the Tibetan Plateau region (see Fig. 3) is |vw|/|vs| ≈ 1.55; that is, most of the summer decrease in jet speed is due to the increase in Coriolis acceleration at higher latitudes, while the thermal gradient changes by a considerably smaller fraction. This is not typical for other longitudes. For example, analogous estimates at 10°E yield fs/fw ≈ 1.49 and |vw|/|vs| ≈ 2.01. Thus, in this case both the northward translation and the decrease in the meridional temperature gradient contribute almost equally to the summer reduction in jet speed, whereas, over the Tibetan Plateau, the comparatively strong meridional temperature gradient is responsible for maintaining a correspondingly strong summer jet (see Fig. 2).
Several dynamical factors potentially influence the strength of the jet. First, a topographic barrier suppresses baroclinic activity and the meridional exchange of air masses. Because of the large horizontal extent and the approximately zonal alignment of the Tibetan Plateau, this inhibition can be expected to be particularly strong and to affect both the mid- and upper troposphere (Pierrehumbert 1985). This would help to maintain a stronger meridional temperature gradient and a correspondingly strong jet (over Asia, the intensity of the westerly jet is predominantly associated with the baroclinicity in the upper troposphere; see also Kuang and Zhang 2005). Second, the conservation of zonal angular momentum in the atmospheric layer capped from above by strong static stability (see Fig. 2) would cause faster westerly flow in that layer above high topography. Both these effects are qualitatively consistent with the anomalous strength of the jet. Yet, this cannot fully explain the observations since there may be other relevant influences such as the flow-opposing pressure gradient torque and gravity wave drag, and in particular because none of the above effects appears to account for the peculiar seasonal cycle in the jet intensity.
In the following, we use the thermal wind relationship to better understand the evolution of the jet location and strength during spring and early summer. Figures 8a,c,e show the climatologies of the jet for a given month, while the right-hand panels (Figs. 8b,d,f) show the evolution of the horizontal wind speed and tropospheric temperature from a given month to the next. During winter, tropospheric temperatures hardly change and, accordingly, the jet stagnates at the southern flank of the Tibetan Plateau (Figs. 8a,b). This is followed by a rapid warming of the troposphere in spring (Figs. 8c,d). The decisive characteristic of this springtime warming is that it is strongest over the Tibetan Plateau and to the north of it, that is, north of the jet location. Therefore, the meridional temperature gradient decreases in magnitude around the latitude of the jet location and beneath the jet level. This, in turn, is consistent with the meridional dispersal and reduction in jet speed described earlier.
The tropospheric warming in early summer does not resemble the spring warming (Figs. 8e,f). From May to June, the strongest warming is observed in the upper troposphere (300–200 hPa) at the northern edge of the Tibetan Plateau, approximately underneath the jet. Consequently, this warming acts to increase the magnitude of the meridional temperature gradient just north of the jet location, which is consistent with the northward shift from May to June and the reappearance of a rather well-defined and comparatively strong jet. A complementary picture of this is obtained by Zhang et al. (2006), who do not consider the seasonal evolution of the jet at a fixed longitude but follow the zonal and meridional changes of the location of the strongest jet cores over the Asian continent. They show that the appearance of a strong jet core north of the Tibetan Plateau in June is associated with the intensification of diabatic heating over the plateau.
Given the timing of this warming and its qualitative difference from the springtime warming, it is interesting to discuss it in conjunction with the contemporaneously onsetting South Asian monsoon. While the monsoon is a consequence of land–sea temperature contrasts, condensational heating during monsoon rains is well known to have a pronounced bearing on the large-scale flow and tropospheric temperatures in central and South Asia (Rodwell and Hoskins 1996; Liu and Yanai 2001) and is fundamentally important for attaining a reasonable strength and maintaining the monsoon.
As a matter of fact, the climatological temperature difference between June and May shown in Fig. 8f resembles the difference between tropospheric temperature composites in years of a weak and strong Indian summer monsoon (Schiemann et al. 2007). We reexamine this in Fig. 9, which shows the difference between temperature and zonal wind composites with respect to strong and weak all-India rainfall in June. In agreement with earlier studies (Rodwell and Hoskins 1996; Liu and Yanai 2001), Indian monsoon rainfall is found to be associated with a warming of the upper troposphere over central Asia. In June, the center of this warming signal is located over the Caspian Sea and extends downstream along the northern edge of the Tibetan Plateau. The vertical cross section in Fig. 9a corroborates that the warming signal corresponding to interannual variability in the Indian monsoon is similar to the seasonal warming between May and June over the Tibetan Plateau. The zonal wind composite differences show that a strong Indian monsoon is associated with a more intense anticyclonic flow in the upper troposphere over South Asia corresponding to an intensification and northward displacement of the westerly jet. The covariability between summer monsoon intensity and the upper-tropospheric flow over the Asian continent has also been investigated on intraseasonal time scales. In a recent study, Randel and Park (2006) showed that circulation changes of the upper-tropospheric monsoon anticyclone follow fluctuations in deep convection over South Asia with a time lag of ∼5 days. The typical time scale of these fluctuations is ∼10–20 days and they are associated with active/break cycles in the monsoon circulation. Put together, the above arguments suggest strongly that the intensification and sharpening of the jet in June can be linked to the onsetting summer monsoon rainfall over South Asia.
5. Jet variability and precipitation
Jet streams have a profound impact on the climate and also affect the spatial and temporal distribution of surface precipitation (e.g., Hartmann 1994). In this section, we briefly discuss the covariability of the jet and precipitation over central, South, and East Asia. This is to illustrate the jet’s climatological significance and to motivate future use of the occurrence-based jet climatology in applied studies. A more detailed investigation of the covariability of (i) the midlatitude circulation over western Eurasia, (ii) the location of the jet stream in the Tibetan Plateau region, and (iii) the central Asian surface climate is being carried out (see also Schiemann 2007).
Figure 10a shows jet counts on April days where the daily jet latitude at 80°–90°E (determined as described in section 3c) is north of 38°N, and Fig. 10b shows counts for all April days where the jet is south of 34°N. The composites are computed with respect to all days in 1958–2001. Of all 1290 April days in this period, 667 were assigned to the composite of southerly jet location and 399 to the composite of northerly jet location. The ratio of the precipitation composites computed with respect to these two sets of days is shown in Fig. 10c. On days of northerly jet position, precipitation is clearly higher in the central Asian plains and much more so to the north of the Tibetan Plateau than during days of southerly jet position. At the same time, precipitation is reduced over the central parts of the Tibetan Plateau during days of northerly jet position.
Before drawing conclusions from these composites, the fact that both the jet stream and the precipitation fields exhibit a pronounced seasonal cycle has to be taken into consideration. Since the jet undergoes a northward translation in spring, days of northerly jet location occur more frequently in late than in early April. Thus, one might argue that the composite ratios shown in Fig. 10c are merely a consequence of the seasonal changes in the precipitation distribution. To test this hypothesis, we also computed the climatological precipitation for the first and last halves of April and compared the ratio of these two climatologies (not shown) to the ratio of the composites computed with respect to the jet location. It turns out that the previously described patterns of higher precipitation to the west and north, and lower precipitation over the Tibetan Plateau, cannot be identified in the ratio of composites for early and late April. Thus, these features can be attributed to a covariability between the jet location and precipitation fields that cannot be explained simply in terms of the seasonal cycle in both quantities. This does not hold true for other patterns of covariability shown in Fig. 10c: the plot also suggests that precipitation is higher over northeast and lower over northwest India during days of northerly jet location. Here, however, we find a rather similar pattern in the ratio of late and early April precipitation composites and conclude that these patterns are mostly associated with seasonality.
Figure 10d is similar to Fig. 10c, but here we have computed composites with respect to years where the monthly mean jet latitude (as in section 3b) is low or high. This shows that systematic differences in the precipitation pattern are also associated with variability of the jet location on interannual time scales.
6. Findings summary and discussion
The characterization of the westerly jet in the Tibetan Plateau region in terms of monthly mean horizontal winds can be summarized in the following points:
The westerly jet undergoes a pronounced seasonal cycle. It is located at the southern edge of the Tibetan Plateau from December to April. Its intensity, measured in terms of its monthly mean horizontal wind speed, is strongest in winter and decreases in spring. In May, the climatological jet latitude is over the Tibetan Plateau and at the same time the intensity is smallest and the position of the jet varies largely. In June, the jet reaches the northern edge of the Tibetan Plateau and its intensity increases with respect to May. The northernmost monthly mean jet position (about 42°N) is observed in July/August. In September, the jet starts to recede gradually southward and reaches its wintertime position in December.
In general, a climatologically intense jet is associated with a small interannual variability of the jet location. Thus, the interannual variability is smaller in winter and summer than during the transition seasons. It is particularly high over the Tibetan Plateau during April and May. In summer, it is smaller at the northern flank of the Tibetan Plateau than upstream and downstream of it.
The seasonal cycle of the jet location is closely related to the seasonal cycle of the tropospheric baroclinicity. In spring, the mid- and upper troposphere warms strongly over the Tibetan Plateau and to the north of it. The strongest warming is located north of the jet latitude. This leads to the weakening and northward translation of the jet. From May to June, the strongest warming is observed at the northern edge of the Tibetan Plateau beneath the jet. This acts such as to decrease the baroclinicity south of the jet and to increase it north of the jet. The jet moves northward and its intensity increases. The warming in early summer is similar to the tropospheric temperature differences between years of a strong and weak Indian summer monsoon. This suggests that the reintensification of the jet in summer is related to the onset of the monsoon.
For a more detailed study of the temporal and spatial variability of the jet, an occurrence-based jet climatology has been derived. A simple identification of local maxima in the horizontal wind speed for a given longitude has proven successful in capturing the positions of the major jet axes. We have carried out this kind of identification for central and East Asia as well as much of the Pacific during 1958–2001 at 6-hourly intervals. The analysis of the dataset obtained this way allowed to complement the above results as follows:
During months where the climatological location of the jet is south of the Tibetan Plateau (December–April), split-flow situations are not uncommon. This is reflected in a secondary maximum of jet occurrence counts at the northern edge of the Tibetan Plateau and a minimum over the plateau during this time of year.
Well-defined zonally oriented jets that cover a wide range of longitudes are common in winter and summer but rare during the transition seasons, in particular during Apri/May.
Because of the high variability of the jet on synoptic time scales during spring and autumn, our pragmatic definition of the dates of northward and southward jet transitions over the Tibetan Plateau requires substantial smoothing of the respective time series of jet latitude. Nevertheless, it is normally not possible to sensibly define a unique transition date in spring, while in autumn, when the jet transitions tend to be more gradual, such a definition is possible in most years. The median northward transition date is 28 April; the median southward transition date is 12 October.
A simple composite analysis has shown that there is a systematic covariation between the position of the jet stream in the Tibetan Plateau region and the spatial distribution of precipitation in central Asia: in April, a northerly (southerly) jet position in the Tibetan Plateau region is associated with both higher (lower) precipitation upstream and to the north of the plateau and lower (higher) precipitation over the plateau.
The central results of this study are the description of the jet stream in the Tibetan Plateau region as represented by the ERA-40 reanalysis and the construction of a novel occurrence-based jet stream climatology derived in terms of a set of simple criteria on the wind field. These criteria are easy to implement and the occurrence-based climatology can be used to study the jet variability on a wide range of time scales. The presentation of the jet stream climatology has been accompanied by discussions of the covariability between (i) the jet location/intensity and the tropospheric baroclinicity and (ii) the jet location and the spatial precipitation distribution. These analyses contribute to a better understanding of the behavior of the jet and give an indication of how the occurrence-based climatology may be used in applied studies. Yet, evidently, these analyses are far from exhaustive and a number of research questions for future work have been raised. They concern (i) the further investigation of the link between the monsoon and the jet in summer, in particular the clearer identification of source regions of diabatic heating to which the jet responds; (ii) the study of the nature of the apparent asymmetry between the northward and southward migration of the jet in spring and autumn; and (iii) the investigation of the covariability of the jet location and the precipitation distribution with the upstream midlatitude circulation. Furthermore, a multitude of issues such as (i) the relationship between the jet in the Tibetan Plateau region and the jet and storm track farther downstream over East Asia and the Pacific, (ii) a temporally highly resolved analysis of the covariability of the jet location over the Tibetan Plateau and the propagation of orographically forced stationary waves to higher latitudes and to the stratosphere, and (iii) the detection of trends and decadal variability of the jet location and intensity have not been touched upon in this paper. The jet climatology derived herein is available upon request and it is the hope of the authors that it will prove to be a useful analytical tool in addressing issues of the kind mentioned above.
Acknowledgments
We thank Mathew Barlow, Christina L. Archer, and two anonymous reviewers for comments that helped in improving the manuscript, and Cathy Hohenegger and Olivia Martius for discussion. The funding for this study has been provided by the Swiss National Science Foundation (Grant NF 200021–101957 and NCCR Climate).
REFERENCES
Academica Sinica, 1957: On the general circulation over Eastern Asia (I). Tellus, 9 , 432–446.
Academica Sinica, 1958a: On the general circulation over Eastern Asia (II). Tellus, 10 , 58–75.
Academica Sinica, 1958b: On the general circulation over Eastern Asia (III). Tellus, 10 , 299–312.
Archer, C. L., and K. Caldeira, 2008: Historical trends in the jet streams. Geophys. Res. Lett., 35 , L08803. doi:10.1029/2008GL033614.
Baehr, C., B. Pouponneau, F. Ayrault, and A. Joly, 1999: Dynamical characterization of the FASTEX cyclogenesis cases. Quart. J. Roy. Meteor. Soc., 125 , 3469–3494.
Barlow, M., H. Cullen, and B. Lyon, 2002: Drought in central and southwest Asia: La Niña, the warm pool, and Indian Ocean precipitation. J. Climate, 15 , 697–700.
Barlow, M., M. Wheeler, B. Lyon, and H. Cullen, 2005: Modulation of daily precipitation over southwest Asia by the Madden–Julian oscillation. Mon. Wea. Rev., 133 , 3579–3594.
Barlow, M., A. Hoell, and F. Colby, 2007: Examining the wintertime response to tropical convection over the eastern Indian Ocean by modifying atmospheric heating in a global atmospheric model. Geophys. Res. Lett., 34 , L19702. doi:10.1029/2007GL030043.
Blackmon, M. L., J. M. Wallace, N. C. Lau, and S. L. Mullen, 1977: An observational study of the Northern Hemisphere wintertime circulation. J. Atmos. Sci., 34 , 1040–1053.
Bugayev, V. A., V. A. Giorgio, E. M. Kozik, M. A. Petrosyants, A. J. Pshenitshnij, N. N. Romanov, and O. N. Tshernysheva, 1957: Synoptic Processes of Central Asia (in Russian). Publishing Company of the Academy of Sciences of the Usbek Soviet Republic, 477 pp.
Chang, E. K. M., S. Lee, and K. L. Swanson, 2002: Storm track dynamics. J. Climate, 15 , 2163–2183.
Chanysheva, S. G., O. I. Subbotina, U. V. Petrov, K. T. Egamberdiyev, B. A. Aizensh-tat, and G. N. Leukhina, 1995: Variability of the Central Asian Climate (in Russian). Central Asian Hydrometeorological Research Institute, 215 pp.
Chromow, S. P., 1940: Einführung in die synoptische Wetteranalyse. Springer Verlag, 532 pp.
DeWeaver, E., and S. Nigam, 1995: Influence of mountain ranges on the mid-latitude atmospheric response to El Niño events. Nature, 378 , 706–708.
Duan, A., and G. Wu, 2005: Role of the Tibetan Plateau thermal forcing in the summer climate patterns over subtropical Asia. Climate Dyn., 24 , 793–807. doi:10.1007/S00382-004-0488-8.
Fischer, E. M., S. I. Seneviratne, P. L. Vidale, D. Lüthi, and C. Schär, 2007: Soil moisture–atmosphere interactions during the 2003 European summer heatwave. J. Climate, 20 , 5081–5099.
Fukutome, S., C. Frei, D. Lüthi, and C. Schär, 1999: The interannual variability as a test ground for regional climate simulations over Japan. J. Meteor. Soc. Japan, 77 , 649–672.
Gibson, J. K., P. Kållberg, S. Uppala, A. Hernandez, A. Nomura, and E. Serrano, 1999: ERA-15 description. Re-analysis Project Report Series, Part 1. Tech. Rep., European Centre for Medium-Range Weather Forecasts, 84 pp.
Hartmann, D. L., 1994: Global Physical Climatology. Academic Press, 411 pp.
Hohenegger, C., and P. L. Vidale, 2005: Sensitivity of the European climate to aerosol forcing as simulated with a regional climate model. J. Geophys. Res., 110 , D06201. doi:10.1029/2004JD005335.
Holton, J. R., 2004: An Introduction to Dynamic Meteorology. 4th ed. Elsevier Academic Press, 535 pp.
Hoskins, B. J., and K. I. Hodges, 2002: New perspectives on the Northern Hemisphere winter storm tracks. J. Atmos. Sci., 59 , 1041–1061.
Kalnay, E., and Coauthors, 1996: The NCEP/NCAR 40-Year Reanalysis Project. Bull. Amer. Meteor. Soc., 77 , 437–447.
Koch, P., H. Wernli, and H. Davies, 2006: An event-based jet-stream climatology and typology. Int. J. Climatol., 26 , 283–301. doi:10.1002/joc.l255.
Kuang, X., and Y. Zhang, 2005: Seasonal variation of the East Asian subtropical westerly jet and its association with the heating fields over East Asia. Adv. Atmos. Sci., 22 , 831–840.
Li, C., and J. Pan, 2006: Atmospheric circulation characteristics associated with the onset of Asian summer monsoon. Adv. Atmos. Sci., 23 , 925–939.
Li, C., J-T. Wang, S-Z. Lin, and H-R. Cho, 2005: The relationship between East Asian summer monsoon activity and northward jumps of the upper air westerly jet location. Chinese J. Atmos. Sci., 29 , 1–20.
Liljequist, G. H., 1981: Tor Bergeron. A biography. Pure Appl. Geophys., 119 , 409–442.
Liu, X., and M. Yanai, 2001: Relationship between the Indian monsoon rainfall and the tropospheric temperature over the Eurasian continent. Quart. J. Roy. Meteor. Soc., 127 , 909–937.
Lüthi, D., A. Cress, H. C. Davies, C. Frei, and C. Schär, 1996: Interannual variability and regional climate simulations. Theor. Appl. Climatol., 53 , 185–209.
Nakamura, H., 1992: Midwinter suppression of baroclinic wave activity in the Pacific. J. Atmos. Sci., 49 , 1629–1642.
Nakamura, H., and J. M. Wallace, 1990: Observed changes in baroclinic wave activity during the life cycles of low-frequency circulation anomalies. J. Atmos. Sci., 47 , 1100–1116.
Newton, C. W., 2004: Associations between twice-yearly oscillations of the North Pacific cyclone track and upper-tropospheric circulations over the Eastern Hemisphere. Mon. Wea. Rev., 132 , 348–367.
Nigam, S., and R. S. Lindzen, 1989: The sensitivity of stationary waves to variations in the basic state zonal flow. J. Atmos. Sci., 46 , 1746–1768.
Nigam, S., and E. DeWeaver, 1998: Influence of orography on the extratropical response to El Niño events. J. Climate, 11 , 716–733.
Palmer, T. N., G. J. Shutts, and R. Swinback, 1986: Alleviation of a systematic westerly bias in general circulation and numerical weather prediction models through an orographic gravity wave drag parametrization. Quart. J. Roy. Meteor. Soc., 112 , 1001–1039.
Parthasarathy, B., A. A. Munot, and D. R. Kothawale, 1994: All-India monthly and seasonal rainfall series: 1871–1993. Theor. Appl. Climatol., 49 , 217–224.
Pierrehumbert, R. T., 1985: A theoretical model of orographically modified cyclogenesis. J. Atmos. Sci., 42 , 1244–1258.
Randel, W. J., and M. Park, 2006: Deep convective influence on the Asian summer monsoon anticyclone and associated tracer variability observed with Atmospheric Infrared Sounder (AIRS). J. Geophys. Res., 111 , D12314. doi:10.1029/2005JD006490.
Rodwell, M. J., and B. J. Hoskins, 1996: Monsoons and the dynamics of deserts. Quart. J. Roy. Meteor. Soc., 122 , 1385–1404.
Sadler, J. C., 1975: The upper tropospheric circulation over the global tropics. Tech. Rep. UHMET-75-05, Department of Meteorology, University of Hawaii at Manoa, 35 pp.
Schär, C., L. Vasilina, F. Pertziger, and S. Dirren, 2004: Seasonal runoff forecasting using precipitation from meteorological data assimilation systems. J. Hydrometeor., 5 , 959–973.
Schiemann, R., 2007: Forcing and variability of the hydroclimate in Central Asia. Ph.D. thesis, ETH Zurich, Dissertation ETH 17426, 141 pp.
Schiemann, R., M. G. Glazirina, and C. Schär, 2007: On the relationship between the Indian summer monsoon and river flow in the Aral Sea basin. Geophys. Res. Lett., 34 , L05706. doi:10.1029/2006GL028926.
Schiemann, R., D. Lüthi, P. L. Vidale, and C. Schär, 2008: The precipitation climate of Central Asia—Intercomparison of observational and numerical data sources in a remote semiarid region. Int. J. Climatol., 28 , 295–314. doi:10.1002/joc.l532.
Schwierz, C., S. Dirren, and H. C. Davies, 2004: Forced waves on a zonally aligned jet stream. J. Atmos. Sci., 61 , 73–87.
Seneviratne, S. I., D. Lüthi, M. Litschi, and C. Schär, 2006: Land-atmosphere coupling and climate change in Europe. Nature, 443 , 205–209.
Shapiro, M. A., and D. Keyser, 1990: Fronts, jet streams, and the tropopause. Extratropical Cyclones: The Eric Palmén Memorial Volume, C. W. Newton and E. O. Holopainen, Eds., Amer. Meteor. Soc., 167–191.
Sutcliffe, R. C., 1939: Cyclonic and anticyclonic development. Quart. J. Roy. Meteor. Soc., 65 , 518–524.
Sutcliffe, R. C., 1947: A contribution to the problem of development. Quart. J. Roy. Meteor. Soc., 73 , 370–383.
Tukey, J. W., 1977: Exploratory Data Analysis. Addison-Wesley, 688 pp.
Uppala, S. M., and Coauthors, 2005: The ERA-40 Re-Analysis. Quart. J. Roy. Meteor. Soc., 131 , 2961–3012.
Vidale, P. L., D. Lüthi, C. Frei, S. I. Seneviratne, and C. Schär, 2003: Predictability and uncertainty in a regional climate model. J. Geophys. Res., 108 , 4586. doi:10.1029/2002JD002810.
von Storch, H., and F. W. Zwiers, 1999: Statistical Analysis in Climate Research. Cambridge University Press, 484 pp.
Wallace, J. M., L. Gyo-Ho, and M. L. Blackmon, 1988: Relationship between cyclone tracks, anticyclone tracks and baroclinic waveguides. J. Atmos. Sci., 45 , 439–462.
Webster, P. J., V. O. Magaña, T. N. Palmer, J. Shukla, R. A. Tomas, M. Yanai, and T. Yasunari, 1998: Monsoons: Processes, predictability, and the prospects for prediction. J. Geophys. Res., 103 , 14451–14510.
Wernli, H., S. Dirren, M. A. Liniger, and M. Zillig, 2002: Dynamical aspects of the life cycle of the winter storm Lothar (24–26 December 1999). Quart. J. Roy. Meteor. Soc., 128 , 405–429.
Wu, G., and Y. Liu, 2003: Summertime quadruplet heating pattern in the subtropics and the associated atmospheric circulation. Geophys. Res. Lett., 30 , 1201. doi:10.1029/2002GL016209.
Yanai, M., C. Li, and Z. Song, 1992: Seasonal heating of the Tibetan Plateau and its effects on the evolution of the Asian summer monsoon. J. Meteor. Soc. Japan, 70 , 319–351.
Yeh, T. C., S. Y. Dao, and M. T. Li, 1959: The abrupt change of the circulation over the Northern Hemisphere during June and October. The Atmosphere and the Sea in Motion. The Rossby Memorial Volume, Rockefeller Institute Press, Oxford University Press, 249–267.
Zhang, Y., K. Xueyuan, G. Weidong, and T. Zhou, 2006: Seasonal evolution of the upper-tropospheric westerly jet core over East Asia. Geophys. Res. Lett., 33 , L11708. doi:10.1029/2006GL026377.
Sample snapshots of the upper-tropospheric flow (ERA-40). Colored contours correspond to the maximum horizontal wind speed between 100 and 500 hPa (m s−1). Blue circles show locations where jet occurrences are identified as local maxima in latitude–pressure cross sections of the horizontal wind field. Areas where the topography is above 1600 m are stippled; further contour lines correspond to 2400, 3200, and 4000 m.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Sample snapshots of the upper-tropospheric flow (ERA-40). Colored contours correspond to the maximum horizontal wind speed between 100 and 500 hPa (m s−1). Blue circles show locations where jet occurrences are identified as local maxima in latitude–pressure cross sections of the horizontal wind field. Areas where the topography is above 1600 m are stippled; further contour lines correspond to 2400, 3200, and 4000 m.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Sample snapshots of the upper-tropospheric flow (ERA-40). Colored contours correspond to the maximum horizontal wind speed between 100 and 500 hPa (m s−1). Blue circles show locations where jet occurrences are identified as local maxima in latitude–pressure cross sections of the horizontal wind field. Areas where the topography is above 1600 m are stippled; further contour lines correspond to 2400, 3200, and 4000 m.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Seasonal cycle of the monthly mean wind climatology at 200 hPa (ERA-40). Blue contour lines show the geopotential height (m) and colored shading the horizontal wind (m s−1). Topography above 1600 m is stippled. The December and February climatologies strongly resemble that for January and are omitted. Similarly, the August climatology strongly resembles that for July. Latitude circles correspond to 0°, 20°, 40°, 60°, 80°N.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Seasonal cycle of the monthly mean wind climatology at 200 hPa (ERA-40). Blue contour lines show the geopotential height (m) and colored shading the horizontal wind (m s−1). Topography above 1600 m is stippled. The December and February climatologies strongly resemble that for January and are omitted. Similarly, the August climatology strongly resembles that for July. Latitude circles correspond to 0°, 20°, 40°, 60°, 80°N.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Seasonal cycle of the monthly mean wind climatology at 200 hPa (ERA-40). Blue contour lines show the geopotential height (m) and colored shading the horizontal wind (m s−1). Topography above 1600 m is stippled. The December and February climatologies strongly resemble that for January and are omitted. Similarly, the August climatology strongly resembles that for July. Latitude circles correspond to 0°, 20°, 40°, 60°, 80°N.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Seasonal cycle of the monthly mean wind climatology (ERA-40) in a vertical cross section at 85°E. Gray shading and solid contour lines correspond to the horizontal wind speed (m s−1), dashed lines to potential temperature (K). Topography is stippled.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Seasonal cycle of the monthly mean wind climatology (ERA-40) in a vertical cross section at 85°E. Gray shading and solid contour lines correspond to the horizontal wind speed (m s−1), dashed lines to potential temperature (K). Topography is stippled.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Seasonal cycle of the monthly mean wind climatology (ERA-40) in a vertical cross section at 85°E. Gray shading and solid contour lines correspond to the horizontal wind speed (m s−1), dashed lines to potential temperature (K). Topography is stippled.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
For January and March–May: (left) total jet occurrence counts (shading, from 6-hourly ERA-40 fields) and interannual variability of jet latitude (box plots, from ERA-40 monthly mean horizontal wind); (right) meridional distribution of occurrence counts in 80°–90°E, i.e., within the area marked by the solid vertical lines. Bold solid line and gray shading: mean ± one std dev of interannual variability (6-hourly ERA-40 fields). Bold and thin dashed lines: the same based on occurrence counts from 6-hourly CHRM fields.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
For January and March–May: (left) total jet occurrence counts (shading, from 6-hourly ERA-40 fields) and interannual variability of jet latitude (box plots, from ERA-40 monthly mean horizontal wind); (right) meridional distribution of occurrence counts in 80°–90°E, i.e., within the area marked by the solid vertical lines. Bold solid line and gray shading: mean ± one std dev of interannual variability (6-hourly ERA-40 fields). Bold and thin dashed lines: the same based on occurrence counts from 6-hourly CHRM fields.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
For January and March–May: (left) total jet occurrence counts (shading, from 6-hourly ERA-40 fields) and interannual variability of jet latitude (box plots, from ERA-40 monthly mean horizontal wind); (right) meridional distribution of occurrence counts in 80°–90°E, i.e., within the area marked by the solid vertical lines. Bold solid line and gray shading: mean ± one std dev of interannual variability (6-hourly ERA-40 fields). Bold and thin dashed lines: the same based on occurrence counts from 6-hourly CHRM fields.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
As in Fig. 4, but for June, July, and September–November.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
As in Fig. 4, but for June, July, and September–November.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
As in Fig. 4, but for June, July, and September–November.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Samples of the seasonal cycle of the daily jet latitude at 80°–90°E (open circles) and the jet latitude obtained after low-pass filtering with a cutoff period of 30 days (solid line) for (a) 1998, (b) 1999, and (c) 2000.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Samples of the seasonal cycle of the daily jet latitude at 80°–90°E (open circles) and the jet latitude obtained after low-pass filtering with a cutoff period of 30 days (solid line) for (a) 1998, (b) 1999, and (c) 2000.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Samples of the seasonal cycle of the daily jet latitude at 80°–90°E (open circles) and the jet latitude obtained after low-pass filtering with a cutoff period of 30 days (solid line) for (a) 1998, (b) 1999, and (c) 2000.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Dates of spring transitions of the jet to the north of the Tibetan Plateau (open circles) and autumn transitions to the south (filled circles). The horizontal line and shaded bands indicate boxplot statistics of the transition dates. Tick marks on the ordinate are shown every 5 days.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Dates of spring transitions of the jet to the north of the Tibetan Plateau (open circles) and autumn transitions to the south (filled circles). The horizontal line and shaded bands indicate boxplot statistics of the transition dates. Tick marks on the ordinate are shown every 5 days.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Dates of spring transitions of the jet to the north of the Tibetan Plateau (open circles) and autumn transitions to the south (filled circles). The horizontal line and shaded bands indicate boxplot statistics of the transition dates. Tick marks on the ordinate are shown every 5 days.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
(a),(c),(e) Seasonal evolution of the jet stream and associated temperature distribution (ERA-40, monthly mean fields). Horizontal wind (m s−1; solid contour lines and gray shading) and temperature (K; dashed lines). (b),(d),(f) Difference between two consecutive months in temperature (K; colored shading) and in the zonal wind component (m s−1; black contour lines, dashed lines correspond to negative differences).
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
(a),(c),(e) Seasonal evolution of the jet stream and associated temperature distribution (ERA-40, monthly mean fields). Horizontal wind (m s−1; solid contour lines and gray shading) and temperature (K; dashed lines). (b),(d),(f) Difference between two consecutive months in temperature (K; colored shading) and in the zonal wind component (m s−1; black contour lines, dashed lines correspond to negative differences).
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
(a),(c),(e) Seasonal evolution of the jet stream and associated temperature distribution (ERA-40, monthly mean fields). Horizontal wind (m s−1; solid contour lines and gray shading) and temperature (K; dashed lines). (b),(d),(f) Difference between two consecutive months in temperature (K; colored shading) and in the zonal wind component (m s−1; black contour lines, dashed lines correspond to negative differences).
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Differences between tercile (quantiles at probabilities 1/3 and 2/3) composites of temperature (K; colored shading) and zonal wind (m s−1; black contour lines, dashed lines correspond to negative differences) with respect to strong and weak monsoons for June: (a) vertical cross section at 85°E; (b) horizontal section at 200 hPa. Based on monthly mean ERA-40 fields.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Differences between tercile (quantiles at probabilities 1/3 and 2/3) composites of temperature (K; colored shading) and zonal wind (m s−1; black contour lines, dashed lines correspond to negative differences) with respect to strong and weak monsoons for June: (a) vertical cross section at 85°E; (b) horizontal section at 200 hPa. Based on monthly mean ERA-40 fields.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Differences between tercile (quantiles at probabilities 1/3 and 2/3) composites of temperature (K; colored shading) and zonal wind (m s−1; black contour lines, dashed lines correspond to negative differences) with respect to strong and weak monsoons for June: (a) vertical cross section at 85°E; (b) horizontal section at 200 hPa. Based on monthly mean ERA-40 fields.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Jet counts for April days where the jet is located (a) north and (b) south of the Tibetan Plateau (per unit time; 80°–90°E). (c) Ratio between ERA-40 precipitation composites with respect to these sets of days. (d) Same as (c) but with respect to months of northerly/southerly jet position derived from the monthly mean wind field in each year.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Jet counts for April days where the jet is located (a) north and (b) south of the Tibetan Plateau (per unit time; 80°–90°E). (c) Ratio between ERA-40 precipitation composites with respect to these sets of days. (d) Same as (c) but with respect to months of northerly/southerly jet position derived from the monthly mean wind field in each year.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Jet counts for April days where the jet is located (a) north and (b) south of the Tibetan Plateau (per unit time; 80°–90°E). (c) Ratio between ERA-40 precipitation composites with respect to these sets of days. (d) Same as (c) but with respect to months of northerly/southerly jet position derived from the monthly mean wind field in each year.
Citation: Journal of Climate 22, 11; 10.1175/2008JCLI2625.1
Dates of first (FTN, FTS) and last (LTN, LTS) jet transition to the north (FTN, LTN) and south (FTS, LTS) of the Tibetan Plateau.
An illustration of this can be found in the introduction to Bugayev et al. (1957), where the authors emphasize the “notion of the planetary upper-level frontal zone as the fundamental feature of the atmosphere, largely determining the evolution of synoptic processes and most suitable for the classification of microsynoptic situations” as one of the guiding principles in the elaboration of their monograph. This statement can be ascribed to the influence of Tor Bergeron on Soviet meteorology (see Liljequist 1981) following his visits to the Meteorological Office in Moscow in 1930 and 1932 and the ensuing dissemination of his ideas by the book of Chromow (1940).
The whiskers extend to the most extreme data point not more than 1.5 times the interquartile range greater/smaller than the hinges. The hinges, that is, the margins of the box, show the lower and upper quartiles, and the horizontal line within the box corresponds to the median. Outliers beyond the whiskers are not shown. For details see Tukey (1977), sections 2B–E.
* Supplemental information related to this paper is available at the Journals Online Web site: http://dx.doi.org/10.1175/2009JCLI2625.s1.