1. Introduction
The possible atmospheric responses to sea surface temperature (SST) anomalies in the tropical east and west Pacific were explored by Palmer and Mansfield (1984) with the atmospheric general circulation model (GCM) of the U.K. Met Office. They showed that a short-wave train can emanate from the tropical west Pacific and propagate downstream along the North Pacific rim and across North America reaching the Gulf of Mexico. Later, based on the monthly/seasonal summary of weather and climate issued by the National Meteorological Center (NMC; currently the National Centers for Environmental Prediction), Palmer and Owen (1986) inferred that the exceptionally cold weather over the United States in January 1985 and the winter of 1976/77 was related to enhanced convective activity over the tropical west Pacific. Using simulations by the Met Office GCM, Palmer and Owen (1986) were able to demonstrate this link. Recently, Moore et al. (2000) correlated annually averaged precipitation at Mount Logan (located in Canada about 100 km from the Gulf of Alaska) with other regions. Statistically significant correlations emerged over the Mackenzie River basin in Canada and Japan. This precipitation correlation pattern is relatively consistent with the corresponding correlation pattern of sea level pressure. These studies seem to suggest that the climate systems on both sides of the North Pacific may be linked by a teleconnection mechanism.
Exploring the propagation pattern of intraseasonal/subseasonal modes in terms of lagged correlation, Kiladis and Weickmann (1992) and Hsu and Lin (1992) showed that a wavelike structure propagates downstream along a route from North America, to the North Atlantic, to southern Europe, and to the North Pacific. With these observations of the wave propagation path, Hoskins and Ambrizzi (1993) developed a theoretical waveguide of Rossby waves (with zonal wavenumber 5) along the Asian jet from North Africa to the western North Pacific. Hoskins and Ambrizzi adopted a forced bartotropic model to substantiate their theoretical argument. With the climatological 300-mb flow field and a forcing at the North Africa–Asian jet entrance, they generated a Rossby-wave train (of roughly wavenumber 5) that propagated along the Asian jet stream waveguide. Can this waveguide provide a link for the climate systems between both sides of the North Pacific? Earlier, Branstator (1985) tested the response of a forced barotropic model to forcings at various locations. A wave train trapped along the Asian subtropical jet emerged from the response to a forcing in south Asia as inferred from previous observational studies.
Interannually, the tropical Pacific SSTs exhibit an east–west seesaw (Fig. 1) in concert with the El Niño–Southern Oscillation (ENSO) cycle (Rasmusson and Carpenter 1982). Following the ENSO activity, it was demonstrated observationally (Wallace and Gutzler 1981; Horel and Wallace 1981), analytically (Hoskins and Karoly 1981), and numerically (Blackmon et al. 1983; Geisler et al. 1985) that a wave train of the Pacific–North American (PNA) teleconnection pattern emanates from the SST anomaly center in the tropical central Pacific and propagates to North America. As reported by review articles presented in Anderson et al. (1998), this PNA teleconnection pattern has constituted a major theme of the climate variability research prior to and after the Tropical Ocean and Global Atmosphere (TOGA) program. In view of observational evidences and theoretical suggestions of previous studies, one may question whether the warm/cold SST anomalies in the western tropical Pacific (WTP; seesawing with the SST anomalies in the east-central Pacific) induce a short-wave teleconnection pattern along the Asian jet stream in accordance with the ENSO cycle. If this is possible, the interannual variation of the North American climate system is likely affected by both teleconnection wave patterns.
According to the teleconnection wave theory (Hoskins and Karoly 1981), ultralong waves can penetrate to high latitudes, but short waves are trapped in lower latitudes. A Fourier scale separation is introduced in section 2 to identify the aforementioned North Pacific short-wave train and to depict this wave train's three-dimensional structure. Because this wave train is newly identified, its basic dynamics are formulated in section 3 following a streamfunction budget analysis. It is inferred in section 2 that the short-wave train is likely induced by perturbing stationary short waves with a forcing. The latent heat released by the anomalous rainfall generated by SST anomalies and cold surge vortices over the Philippine Sea may contribute to this forcing. An analysis will be conducted in section 4 to illustrate this possible forcing. The hypothesized forcing of the short-wave train is substantiated by simulations of this wave train by version 1 of the NCAR Community Climate Model (Williamson et al. 1987) with a tropical forcing imposed on the western Pacific. A summary and suggestions for future efforts are provided in section 5.
Finally, data from three sources are used in this study: National Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis data (Kalney et al. 1996), the outgoing longwave radiation data of National Oceanic and Atmospheric Administration, and precipitation estimates generated by Susskind et al. (1997) at the Goddard Space Flight Center with the orbiting satellite observations. The first two datasets have a 2.5° × 2.5° horizontal resolution, while the third one is 1° × 1°. Details of the NCEP–NCAR reanalysis and Susskind et al.'s precipitation, refer to the aforementioned documents.
2. North Pacific short-wave train
As inferred from OLR anomalies (reported in Climate Diagnostics Bulletin), rainfall/cumulus convection is enhanced (diminished) in the central equatorial Pacific during the warm (cold) ENSO event (Rasmusson and Carpenter 1982). Correspondingly, there should be anomalous convergent (divergent) flow toward the central Pacific basin to maintain the enhanced (diminished) rainfall/cumulus convection. Following Rasmusson (1991; his Fig. 10.11), this interannual climate change of rainfall and divergent circulation will be illustrated through the winter of 1986/87 (warm ENSO episode) and 1988/89 (cold ENSO episode). Let us denote Δχ (200 mb) and ΔP in Figs. 2a and 2c as seasonally averaged winter departures of 200-mb velocity potential [χ (200 mb)] and rainfall (P) from their three-winter (1986–89) mean values. The intensification (weakening) of upper-tropospheric outflow from the equatorial Pacific and the accompanying rainfall increase (decrease) in this region during the warm (cold) ENSO phase is reflected by a spatial reversal of the planetary-scale Δχ (200 mb) and the Pacific basin ΔP patterns. However, with a careful inspection of Figs. 2a and 2c, one can easily find that embedded in the planetary-scale Δχ (200 mb) pattern are small-scale disturbances along the North Pacific rim.
Only planetary-scale Rossby waves with wavenumbers smaller than or equal to 3 are able to penetrate to the high-latitude region north of 60°N and to form the PNA-like wave train (Hoskins and Karoly 1981). We thus adopt a Fourier scale-separation procedure to divide the Δχ (200 mb) fields into two wave regimes: long [wavenumber 1–3, denoted by (·)L], and short [wavenumber 4–25, denoted by (·)S]. Shown in Figs. 2b and 2d are ΔχS(200 mb) of the short-wave regime; during the two winters this short-wave train emerges along the North Pacific rim propagating from southeast Asia to North America (indicated by a thick dotted line). Can this short-wave train be responsible for the teleconnection between the convective activity/rainfall in the western Pacific and North America, as suggested by Palmer and Owen (1986) and Moore et al. (2000)? Some light may be shed on this question by contrasting the ΔP and ΔχS(200 mb) patterns. In principle, positive (negative) ΔχS cells corresponding to anomalous convergent (divergent) centers should be coincident with negative (positive) ΔP values. As shown in Figs. 2b and 2d, positive and negative cells of ΔχS(200 mb) match well with negative and positive ΔP values, respectively, along this wave train. Despite the fact that ΔP and ΔχS(200 mb) are derived from independent data sources, the ΔχS(200 mb) cells along the North Pacific short-wave train correlate to a good extent with ΔP anomalies. The broad region of large-value ΔP upstream in the western equatorial Pacific seems to support Palmer and Owen's argument; severe winter weather in North America may be remotely affected by the anomalous convective activity in the western equatorial Pacific.
The teleconnection PNA pattern elucidated with the Rossby-wave dynamics is often portrayed in terms of streamfunction. This variable was used by Palmer and Owen (1986) to illustrate the possible teleconnection link between both sides of the North Pacific. Is there a North Pacific short-wave train in the streamfunction corresponding to the ΔχS(200 mb) one? This question may be simply answered with the streamfunction departures of the winters of 1986/87 and 1988/89 in the short-wave regime. Actually, the answer is positive, although it is not shown. In order to obtain an overview of the streamfunction short-wave train, we shall analyze all warm and cold ENSO events identified in Fig. 1 for the past 20 winters to see if that short-wave train is common to these events. For reference, the composite Δψ (200 mb) anomalies [Δ(·) will be used hereafter as seasonally averaged winter departure of variable (·) from its long-term mean value] for all cold and warm ENSO events are displayed in the appendix (Fig. A1). As shown by numerous previous studies (e.g., Hoerling et al. 1997), the PNA teleconnection pattern emanating from the central equatorial Pacific stands out clearly in composite Δψ (200 mb) charts of both warm and cold ENSO phases. For the short-wave regime, ψS(200 mb) (stationary short waves), composite ΔψS(200 mb, cold) and ΔψS(200 mb, warm) are shown in Fig. 3. Like ΔχS(200 mb) anomalies shown in Figs. 2b and 2d, a well-organized short-wave train of ΔψS(200 mb) emerges along the North Pacific rim. An arching short-wave train appears in ΔψS(200 mb) (cold) anomalies (Fig. 3b), while two short-wave trains (one along 20°N and the other along 50°N) emerge from ΔψS(200 mb) (warm) anomalies (Fig. 3c). The phase reversal of the two latter wave trains is just like that of the ψS(200 mb) field in Fig. 3a. Compared to the composite 200-mb eddy streamfunction departures ΔψT(200 mb) during cold and warm phases (Fig. A1), amplitudes of ΔψS(200 mb) over the North Pacific–North America region can reach 30%–100% of ΔψT(200 mb). Evidently, the short-wave train is significant in comparison with the PNA pattern.
The possible effect of the anomalous SSTs over the central equatorial Pacific on the winter climate of North America is linked through the PNA teleconnection pattern. Since an interannual seesaw exists between ΔSST (WTP) and ΔSST (Niño-3.4) in accordance with the ENSO cycle, it is likely that the North Pacific ΔψS (200 mb) short-wave train is induced by the ΔSST (WTP) anomalies. If this argument can be substantiated, this short-wave train bears climatic significance to North America. As shown by the schematic diagram of these two wave trains (Fig. 4), the winter climate system over this continent may be influenced by the competing impacts from the North Pacific short-wave train and the PNA teleconnection pattern. In essence, the interannual variation of the North American winter climate system may be affected remotely by the SST anomalies over both the western and east-central equatorial Pacific, instead of solely the latter region as has been widely accepted over the past two decades.
As the North Pacific short-wave train is documented for the first time, some further analysis of its structure would be of use to the future search for its dynamics. The stationary ultralong waves in winter are baroclinic north of 30°N (Lau 1979), while the PNA teleconnection wave train in middle-high latitudes is barotropic (Philander 1990). Can the same contrast of vertical structure be applied to the short-wave regime? Characterized by the westward tilt in its vertical structure, stationary short waves, ψS(45°N) shown in Fig. 5a, are baroclinic. In contrast, the North Pacific short-wave train depicted by ΔψS(200 mb) and ΔψS(850 mb, not shown) bear a strong resemblance in their spatial structure. The structure resemblance is reflected by the vertically uniform structure of composite ΔψS(45°N) (cold) in Fig. 5b and of composite ΔψS(50°N) (warm) in Fig. 5c. Evidently, the North Pacific ΔψS short-wave train exhibits a barotropic character. Another interesting feature of the North Pacific short-wave train is revealed from the spatial structure comparison between ΔχS and ΔψS. In spite of the barotropic nature of ΔψS and a vertical phase change of ΔχS (not shown), a quarter-phase shift exists between short-wave trains of composite ΔψS and ΔχS.1 With these structure constraints, the simple budget analysis of ΔψS in section 3 will shed some light on the basic dynamics of the North Pacific short-wave train.
A possible waveguide along the North African–Asian jet for the zonal wavenumber-5 Rossby wave was suggested by Hoskins and Ambrizzi (1993). As shown in Fig. 3a, stationary waves in the short-wave regime are dominated by the wavenumber-5 component along this jet stream. Likewise, the North Pacific composite ΔψS(200 mb) short-wave train associated with cold and warm ENSO events (Figs. 3b,c) also exhibits a predominantly wavenumber-5 pattern that is almost spatially coincident with stationary short waves. The PNA teleconnection wave train emanating from the SST anomalies in the central equatorial Pacific does not have any spatial relationship with stationary waves. In contrast, the coherent spatial patterns between the ψS(200 mb) and composite ΔψS(200 mb) short-wave trains strongly suggests that the latter is a result of the amplification (weakening) of the former by some forcing. Recall that a spatially quadrature relationship exists between the North Pacific composite ΔχS(200 mb) and ΔψS(200 mb) short-wave train, and the ΔP anomalies in the western equatorial Pacific coincide with a composite ΔχS(200 mb) cell of opposite polarity in the upstream of the short-wave train. These two factors suggest that the ΔSST(WTP) anomalies are the likely forcing of the North Pacific short-wave train. An effort to substantiate this inference will be made in the next section.
3. Basic dynamics
A heuristic argument concerning the dynamics satisfying the two structural properties of the North Pacific short-wave train may be derived from these two simplified ΔψS budget equations. As revealed from the ΔψS budget (Fig. 7) over the region between south Asia and North America, one can find
The ΔψS(200 mb) cells (stippled areas) are spatially in quadrature with their corresponding
ΔψSA ΔψSA2 An opposite polarity exists between
ΔψSχ1 ΔψSA ΔψSχ1 ΔψSA2 As inferred from Eq. (1),
ΔψSχ1
Next, the vertically uniform ΔψS structure of the North Pacific short-wave train (Fig. 5b,c) is maintained by various dynamical processes in the ΔψS budget:
Both
ΔψSA1 ΔψSA2 ΔψSA1 ΔψSA2 ΔψSA1 ΔψSA2 ΔψSA As inferred from the contrast between
ΔψSχ1 ΔψSχ1 ΔψSχ1 The opposite polarity between
ΔψSA ΔψSχ1
According to Hess's (1979) estimate, Uz − CR at 45°N for wavenumber 5 is about 13.1 m s−1. Apparently, coefficients associated with Δψk's in these two equations are positive. Therefore, the exponential factor of e−iπ/2 in Eqs. (7) and (8) indicates that Δψk and Δχk are spatially in quadrature in the upper and lower troposphere.
4. Possible forcing
The forcing of the PNA teleconnection pattern is the latent heat released by rainfall/cumulus convection over the central equatorial Pacific where the SST anomalies associated with the ENSO cycle reside (e.g., Philander 1990; Anderson et al. 1998). In view of the east–west seesaw of the equatorial Pacific ΔSST anomalies (Fig. 1), can the North Pacific short-wave train be induced by the ΔSST (WTP) anomalies? This conjecture has been suggested by Branstator (1985) based on the response of a forced barotropic model to a south Asian forcing. As shown in Fig. 2, the ΔχS cell over the western Pacific coincides with ΔP anomalies. This correlation may not be accidental, but strongly suggests that these ΔχS cells are a direct atmospheric response to the latent heat released by the western Pacific ΔP anomalies. If this argument is plausible, the next question may be raised; what may be responsible for the generation of ΔP anomalies? Cheang (1977, 1987) pointed out that more than a half of the annual rainfall in Malaysia is contributed by the northeast winter monsoon primarily through cold surge vortices. Conceivably, the southeast Asian rainfall during the northern winter is not only affected by the regional ΔSST anomalies, but also by the tropical midlatitude interaction through cold surges.
a. Cold surge vortices
After the Winter Monsoon Experiment (WMONEX; Greenfield and Krishnamurti 1979), numerous efforts were devoted to uncover the planetary-scale interaction between the Tropics and midlatitudes over the western Pacific (as reviewed by Lau and Chang 1987). In contrast, the cold surge vortex was almost neglected in the post-WMONEX research. We shall compile the climatology of cold surge vortices to test our suggestion for the possible forcing generated by the interannual variation in the cold surge vortex activity. The vortex is identified by the following criteria:
A closed vortex is clearly identified by the 925-mb streamline chart.
Low-value outgoing longwave radiation (OLR)/large-value precipitation [GOES Precipitation Index (GPI) whenever available] appears with the closed vortex.
The identified closed vortex exhibits a clear link to the East Asian cold-air outflow/the northeast Asian cold-air outbreak indicated by the 925-mb streamline chart.
The activity of cold surge vortices in tropical southeast Asia and the tropical western Pacific is part of the Asian winter monsoon. What may be the relationship between this activity and the ENSO cycle? During the cold ENSO phase, the cold-air outflow from east Asia and the tropical cyclonic flow intensity (Fig. 10b) form a favorable low-level environment by which the occurrence frequency of cold surge vortices is enhanced over the three regions. In contrast, the reversed situation occurs during the warm ENSO phase (Fig. 10c). To be more quantitative, the histogram of total cold surge vortex occurrence (Nf) combined over the three regions is shown in Fig. 11a. By contrasting Nf and ΔSST (Niño-3.4) (Fig. 1), an interesting feature of Nf stands out; the occurrence frequency of cold surge vortices increases (decreases), that is, Nf becomes larger (smaller), during the cold (warm) ENSO winter. Furthermore, an interesting inverse relationship exists between Nf and frequency of all cold surges (N) in each winter reaching equatorial south and southeast Asia (Fig. 11b). The comparison between ΔSST (WTP) (Fig. 1) and the cold surge occurrence frequency reveals that cold surge vortices are much easier (more difficult) to form and develop when SSTs become warmer (colder) over the south/southeast Asian region. Apparently, the occurrence frequency of cold surge vortices is more sensitive to ΔSST (WTP) anomalies than to the cold surge frequency reaching tropical south/southeast Asia. Since cold surge vortices are a major winter rain producer in this region, more (less) rainfall is therefore expected during the cold (warm) ENSO phase.
b. Precipitation
Since the cold surge vortex is a major rain producer in southeast Asia during the Asian winter monsoon season, the interannual variation in the western tropical Pacific rainfall is not only a response to the SST (WTP) anomalies, but also a result of the interannual variation in Nf. Our analysis of the cold surge activity in tropical south-southeast Asia shows that the occurrence frequency (Nf) of cold surge vortex increases (decreases) during the cold (warm) ENSO phase. If the latent heat released by rainfall/cumulus convection over the tropical western Pacific plays an important role in inducing/maintaining the North Pacific short-wave train, the rainfall generated by cold surge vortices may form a vital part of this forcing.
Because of the lack of daily precipitation observations over the Pacific Ocean, the precipitation (P) estimates generated by Susskind et al. (1997) with the infrared data observed by orbiting satellites for 1985–97 were analyzed to provide a preliminary indication of the interannual variation in the western tropical Pacific rainfall. The result is then verified with the OLR analysis. Composite seasonally averaged precipitation anomalies (ΔPTs) during cold and warm ENSO winters are displayed respectively in Figs. 12b and 12e. Values of ΔPT ≥ 100 mm season−1 (≤−100 mm season−1) are heavily (lightly) stippled. As reported by the Climate Diagnostics Bulletin, major ΔPT anomalies appear over the western tropical Pacific and the southern part of the South China Sea. The composite seasonally averaged rainfall anomalies (ΔPυ) contributed by cold surge vortices over the three designated regions during both cold (Fig. 12a) and warm (Fig. 12d) ENSO winters are relatively consistent with the corresponding ΔPT distribution. The ratios of ΔPυ/ΔPT (Figs. 12c,f) indicate that the cold surge vortex contribution to ΔPT is particularly significant over southeast Asia and the Philippine Sea.
A more quantitative measurement of the cold surge vortex contributing to ΔPT may be obtained from daily mean histograms shown in Fig. 13; various quantities of rainfall are displayed in the left column, while those of OLR are in the right column. During the cold (warm) ENSO winter, Pυ and PT are larger (smaller) than their long-term seasonally averaged values. The opposite is true for OLR [because OLR values become smaller (larger) during rain/active convection (clear sky)]. The ratio Pυ/PT was computed by the following manner; (Pυ × cold surge vortex days)/(PT × 90 days) over the three designated regions. Arkin and Ardanuy (1989) suggested that ΔOLR (≡235 W m−2 − OLR) may be a reasonable rainfall proxy over the active convection region in the Tropics. With this proxy, the ΔOLRυ/ΔOLRT ratio was computed (Fig. 13f) within the same fashion as the Pυ/PT ratio (Fig. 13c). These two ratios reveal that cold surge vortices contribute in average somewhat more than a half of the winter monsoon rainfall over the three regions. Surprisingly, this contribution undergoes a pronounced interannual variation. As shown in Figs. 13c and 13f, cold surge vortices bring about 80% of winter monsoon rain to the three designated regions during cold ENSO winters, but only about 40% during warm ENSO winters. In light of this distinct contrast in the cold surge vortex contribution to the western tropical Pacific rainfall between the two ENSO extreme winters, the effect of these vortices on the forcing of the North Pacific short-wave train (presumably) becomes obvious.
c. Numerical simulations
Our effort has been so far directed by the following hypothesis: the latent heating generated from the ΔP anomalies over the western tropical Pacific can perturb the stationary short waves (Fig. 3a) to form the North Pacific short-wave train associated with the ENSO event. According to the basic dynamics of this short-wave train formulated in section 3, the forcing of this short-wave train may be reflected by its vorticity source, that is, f∇·VS (or f∇2χS), planetary vortex stretching by local convergence/divergence. Thus, a link between ΔP anomalies and ∇χS should be established. Actually, this link may be illustrated by the anomalous local Hadley circulation. For illustration, this circulation in the short-wave regime, Δ(−ω, υD)S (135°E) corresponding to ΔχS, is shown in Fig. 14 for both cold and warm ENSO winters. During the cold ENSO winter, an upward branch of the short-wave local Hadley circulation is supported by the positive ΔP anomalies in the western tropical Pacific, while a downward branch is centered around 40°N. This midlatitude downward branch is coupled with a positive ΔχS(200 mb) cell (corresponding to a convergence center) that forms a negative
As pointed out previously, Hoskins and Ambrizzi (1993) suggested that the North African–Asian jet may form a waveguide of zonal wavenumber 5. Imposing a forcing at the entrance of this jet in North Africa, they were able to generate a short-wave train propagating along this jet. In contrast, Branstator (1985) showed that a southeast Asia–North Pacific wave train is possibly generated by a south Asian forcing. The mechanism of generating the North Pacific short-wave train suggested in this study differs from Hoskins and Ambrizzi's, but coincides with Branstator's. One may question whether the mechanism proposed in the present study is plausible. In order to test this mechanism, a forcing with a radius of 103 km (corresponding to a divergence or convergence center) was imposed every 15° longitude along the equator in a January perpetual experiment with version 1 of the Community Climate Model developed at the National Center for Atmospheric Research (Williamson et al. 1987). The short-wave train can only be induced by the forcing located between 90° and 150°E. Although it is not our intent to simulate in a precise manner the North Pacific short-wave train, this wave train clearly emerges from the streamfunction departures between the forced experiment and the control experiment (not shown) of the January perpetual condition. For illustration, shown in Fig. 15 are only the eddy component ΔψE(300 mb) generated by a forcing corresponding to divergence (Fig. 15a) and convergence (Fig. 15b) located at 0°, 135°E. The resemblance between the simulated ΔψE(300 mb) wave train and the ΔψS(200 mb) North Pacific short-wave train in Figs. 3b and 3c strongly supports the hypothesized mechanisms of this wave train.
5. Concluding remarks
Interannual variation in equatorial Pacific SSTs and rainfall exhibit a seesaw between the central and western Pacific. The PNA teleconnection pattern emanating from the former region has formed the foundation of climate variability research in the past two decades (reviewed by Philander 1990; Anderson et al. 1998). Can the SST anomalies in the latter region in some way affect the weather and climate system over North America? Evidence accumulated from observations, numerical simulations, and analytic studies strongly suggest such a possibility. However, the basic question is how the rainfall/convective activity in the western Pacific and severe weather/anomalous climate over North America are linked. The teleconnection wave theory (Hoskins and Karoly 1981) showed that only ultralong waves can penetrate to high latitudes, while short waves are trapped in lower latitudes. Based upon this theoretical ground, a Fourier scale separation procedure was introduced to divide the anomalous circulation of the ENSO winter phase into two wave regimes; long wave (wavenumbers 1–3) and short wave (wavenumbers 4–15). The PNA teleconnection pattern is basically formed by the long-wave regime. In contrast, a well-organized wave train in the short-wave regime propagating along the North Pacific rim links the western subtropical Pacific and North America.
Since the North Pacific short-wave train is newly identified, several aspects of its basic characteristics were examined. The major findings are summarized as follows:
a. Structure
During the warm (cold) ENSO winter, the polarity of this short-wave train is almost coincident with (opposite to) that of the winter stationary short waves. This special feature of the short-wave train suggests that it may be formed by some forcing to amplify (weaken) stationary short waves. For this short-wave train, its velocity potential and streamfunction anomalies ΔχS and ΔψS are spatially in quadrature. Because of this relationship, even the short-wave train, which exhibits a vertically uniform structure, does not warrant the barotropic dynamics.
b. Basic dynamics
The formation of the PNA teleconnection wave pattern can be illustrated in terms of the basic dynamics of the stationary barotropic Rossby wave by Hoskins and Karoly (1981). However, following the dynamic constraints of the short-wave train's structure, we find by using the streamfunction budget analysis that the basic dynamics of this short-wave train may be depicted by a Sverdrup balance in the lower troposphere, and by a balance between total vorticity advection and planetary vortex stretching in the upper troposphere. This short-wave train's dynamics explains its vertically uniform structure.
c. Possible forcing
The rainfall anomalies over the western tropical Pacific (WTP) may be enhanced (suppressed) during the cold (warm) ENSO winter when the WTP SSTs become warmer (cooler) and the occurrence frequency of cold surge vortices reduces (increases). Consequently, the latent heat released by rainfall anomalies may intensify (weaken) the local Hadley circulation during the cold (warm) ENSO winter to perturb stationary short waves. To test this hypothesized forcing, a tropical divergent (convergent) center over the tropical western Pacific was imposed on the January perpetual experiment of NCAR CCM1. A short-wave train across the North Pacific resembling the observed short-wave train was simulated.
Results reported in this study are preliminary. Future research of this newly identified North Pacific short-wave train is urged to continue in order to better understand its role in the winter climate system. Several potential studies originating from the current one are suggested:
During the TOGA decade (Anderson et al. 1998), efforts of climate simulations almost always concentrated on the central equatorial Pacific forcing and the atmospheric response in the downstream over North America. No attention has been paid to the possible impact on the North American climate system from the western tropical Pacific SST anomalies through the North Pacific short-wave train. In the past two decades, the proper treatment of the SST anomalies over the extreme western tropical Pacific and the proper simulation of the east Asian cold surges and the southeast Asian cold surge vortices were not actually high priorities in climate simulations. In view of the climate significance of this short-wave train, these concerns should be considered in the future development of the global climate model.
After the WMONEX, the winter Asian monsoon research was primarily focused on the planetary-scale interaction between the east Asian jet and the local Hadley circulation maintained by tropical rainfall/convection related to cold surges (Lau and Chang 1987). Since cold surge vortices are a major rain producer in tropical southeast Asia, the possible impact of midlatitude circulation on the interannual variation of the southeast and south Asian climate may be reflected by that in the cold surge vortex activity. This tropical midlatitude interaction should be addressed in the future under the Climate Variability and Prediction Program (CLIVAR) Asian–Australian monsoon component.
The PNA teleconnection pattern affects not only the weather and climate systems over North America, but also the storm track (Lau 1988) and blocking (Renwick and Wallace 1996) in the North Pacific. As shown by the schematic diagram (Fig. 4) of the PNA teleconnection pattern and the North Pacific short-wave train, what are the competing (cooperating) effects of these two wave trains on synoptic activities over the North Pacific and North America? The NCEP–NCAR reanalysis data (Kalney et al. 1996) and many long-term climate simulations of the global climate model may give us an opportunity to explore the distinguishable weather and climate impacts of these two wave trains.
Acknowledgments
This study is supported by the NASA/NSIPP Grant NAG58293 and the NSF Grant ATM-9906454. We would like to thank Dr. Shu-Ping Weng, Jin-ho Yoon, Paul Tsay, and Peter Hsieh for assistance with data processing and graphics. Special thanks go to Dr. Grant Branstator for offering Fig. 15 to support our study. Typing support provided by Mrs. Reatha Diedrichs and Judy Huang is highly appreciated. Suggestions and comments offered by reviewers were helpful in improving the presentation of this paper.
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APPENDIX
Composite ΔΨT(200 mb)
Composite charts of the 200-mb eddy streamfunction departures [ΔΨT(200 mb)] from the their long-term winter-mean values for cold and warm ENSO events are constructed, respectively, with events identified in terms of the ΔSST (Niño-3.4) index shown in Fig. A1. The Pacific–North America (PNA) patterns indicated by the ΔΨT(200 mb) anomalies are opposite in their polarity during cold (Fig. A1a) and warm (Fig. A1b) ENSO events.
The x–t diagram of tropical ΔSST anomalies averaged over latitudinal zones specified at the top of the figure. Time series of ΔSST(WTP) anomalies averaged over the region (5°–15°N, 120°–160°E) is at the left side, and ΔSST(Niño-3.4) anomalies is at right. Dashed lines indicate that ΔSST(WTP) = ±0.25°C at left and ΔSST(Niño-3.4) = ±0.5°C at right. Letters W and C represent warm and cold sea surface, respectively. Contour intervals of the ΔSST x–t diagram are 0.25°C between 120° and 160°E and 0.5°C between 160°E and 90°W
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Departures of 200-mb velocity potential and precipitation estimate (Susskind et al. 1997), Δ[χ(200 mb), P], from the three winter (1986–89) mean values: (a) Δ[χ(200 mb), P] (1988/89), (b) Δ[χS(200 mb), P] (1988/89), (c) Δ[χ(200 mb), P] (1986/87), and (d) Δ[χS(200 mb), P] (1986/87). Here, (·)S denotes (·) in the short-wave regime. Contour intervals of Δχ(200 mb) and ΔχS(200 mb) are 5 × 105 and 2 × 105 m2 s−1. Values of ΔP ≥ 1 mm day−1 (≤−1 mm day−1) are heavily (lightly) stippled.
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
The 200-mb streamfunction in the short-wave regime: (a) winter stationary short waves ψS(200 mb), (b) composite departures of 200-mb streamfunction from ψS(200 mb) for cold ENSO winters, and (c) same as (b) except for warm ENSO winters. Contour intervals of ψS(200 mb) and ΔψS(200 mb) are 6 × 105 m2 s−1 and 3 × 105 m2 s−1, respectively
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Schematic diagram of the North Pacific short-wave train (indicated by thick solid/dashed lines) and the PNA teleconnection pattern (indicated by thin solid/dashed lines) for the cold ENSO winter. The scalloped region represents the TWP rainfall/convection anomalies, while the central eastern tropical Pacific cold water with ΔSST ≤ −0.5°C is denoted by the hatched region
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Longitude–height cross sections of ψS and ΔψS at different latitudes: (a) ψS(45°N), (b) ΔψS(45°N) (cold), and (c) ΔψS(50°N) (warm). The contour interval of ΔψS(45°N) in (a)–(c) is 5 × 105 m2 s−1.
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Departures of streamfunction tendencies from their long-term winter-mean values: (a)
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Same as in Fig. 6, except for (a)
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Composite longitude–height cross sections of (a)
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Same as in Fig. 8, except for (a)
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
The occurrence frequency (Nf) of cold surge vortices in southeast and south Asia superimposed on the 925-mb streamline: (a) climatology of 1979–99, (b) composite departures for all cold ENSO winters, and (c) composite departures for all warm ENSO winters. The three encircled regions in (a) are designated for more quantitative analysis
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
(a) Histograms of cold surges' vortices combined over the three designated regions in Fig. 10a, and (b) cases of east Asian cold surges (N) identified with the 925-mb streamline charts. Both Nf and N in cold and warm ENSO winters are indicated by dark and gray histograms, respectively
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Composite rainfall departures (ΔP) from their long-term winter-mean values during (a)–(c) cold and (d)–(f) warm ENSO winters for total rainfall (PT) and contribution from cold surge vortices (Pυ) and the ratio ΔPυ/ΔPT
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Histograms of (a)–(c) rainfall (P) and (d)–(f) OLR of every winter averaged over three regions defined in Fig. 10a for three different categories: total (PT, OLRT), contribution from cold surge vortices (Pυ, OLRυ), and ratios Pυ/PT and OLRv/OLRT. Note that ΔOLR ≡ 235 W m−2 −OLR is used as a rainfall proxy. Histograms of warm and cold ENSO winters are denoted by dark and gray color, respectively
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Composite departures of the local Hadley circulation at 135°E in the short-wave regime during (a) cold ENSO winter, and (b) warm ENSO winter from their long-term winter-mean values. Upward and downward motions are denoted by heavily and lightly stippled areas, respectively
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Departures of the simulated eddy streamfunction from the Jan perpetual control experiment: (a) with a divergent forcing center at (0°, 135°E) and (b) with a convergent forcing interval (0°, 135°E). The contour value is 5 × 105 m2 s−1 for both (a) and (b)
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
Fig. A1. Composite departures of 200-mb eddy streamfunction from its long-term winter-mean values for all cold and warm ENSO winters identified in Fig. 1: (a) ΔψT(200 mb) during cold ENSO winter, and (b) ΔψT(200 mb) during warm ENSO winter. The contour interval of ΔψT(200 mb) is 106 m2 s−1 and positive values of this variable are stippled
Citation: Journal of Climate 15, 17; 10.1175/1520-0442(2002)015<2359:ANPSWT>2.0.CO;2
In order to avoid redundancy, ΔχS anomalies are not shown. Because of the relationship between χ and ψχ1[= ∇−2(−f∇−2χ)] shown in Eq. (1) of section 3, the quarter phase shift between ΔψS and ΔχS can be inferred from Figs. 7b and 7d later.
