teleconnections from tropics to northern extratropics ...zhuowang/...symmetric 300-hpa summer...

Teleconnections from Tropics to Northern Extratropics through a Southerly Conveyor

ZHUO WANG*

Department of Meteorology, University of Hawaii at Manoa, Honolulu, Hawaii

C.-P. CHANG

Department of Meteorology, Naval Postgraduate School, Monterey, California

BIN WANG

Department of Meteorology, University of Hawaii at Manoa, Honolulu, Hawaii

FEI-FEI JIN

Department of Meteorology, The Florida State University, Tallahassee, Florida

(Manuscript received 27 October 2004, in final form 4 May 2005)

ABSTRACT

Rossby wave propagation theory predicts that Rossby waves in a tropical easterly flow cannot escapefrom the Tropics to the extratropics. Here the authors show that a southerly flow component in the basicstate (a southerly conveyor) may transfer a Rossby wave source northward; thus, a forcing embedded in thedeep tropical easterlies may excite a Rossby wave response in the extratropical westerlies. It is shown thatthe southerly conveyor determines the location of the effective Rossby wave source and that the extratrop-ical response is relatively insensitive to the location of the tropical forcing, provided that the tropicalresponse can reach the southerly conveyor. A stronger southerly flow favors a stronger extratropicalresponse, and the spatial structure of the extratropical response is determined by the extratropical westerlybasic flows.

1. Introduction

Tropical–extratropical teleconnections have been asubject of numerous studies since 1980s. Rossby wavedispersion is one of the well-known mechanisms to ex-plain the development and maintenance of teleconnec-tion patterns. Using a simple model, Hoskins andKaroly (1981) showed that Rossby wave dispersion in aslowly varying basic flow closely resembles the ob-served Pacific–North American (PNA) pattern.

On a � plane, the dispersion relation for stationaryRossby waves with a basic flow u � u(y) and � � 0 in ameridionally bounded channel can be written as

c � u ��*

k2 � l2� 0, �1�

where c is the phase speed of stationary Rossby waves,�* is the effective � parameter (the meridional gradientof the absolute vorticity), and k and l are zonal andmeridional wavenumbers, respectively. Since generally�* � 0, Eq. (1) indicates that stationary Rossby wavescan only exist in westerly mean flows. However, ob-served tropical forcings are typically embedded in east-erly mean flows, and there is some observational evi-dence of teleconnections in the form of stationaryRossby waves between the tropical forcing and the ex-tratropical circulation anomalies (Horel and Wallace1981; Nitta 1987; Wang et al. 2001). How the forcingwithin the deep tropical easterlies excites and maintainsextratropical stationary waves poses a challenge to theaforementioned stationary wave theory.

Several theories have been proposed to address this

* Current affiliation: Department of Meteorology, Naval Post-graduate School, Monterey, California.

Corresponding author address: Dr. Zhuo Wang, Dept. of Me-teorology, Naval Postgraduate School, Monterey, CA 93943.E-mail: [email protected]

NOVEMBER 2005 W A N G E T A L . 4057

© 2005 American Meteorological Society

JAS3600

issue. Because the basic-state zonal flow varies longitu-dinally, regional westerly flows may be present at somelongitudinal locations even though the zonal mean ofthe basic flow is easterly. These “windows” of westerlyflows may act as ducts and allow propagation of Rossbywaves with zonal wavelength less than the zonal scale ofthe westerly duct (Hoskins et al. 1977; Simmons 1982;Webster and Holton 1982; Branstator 1983; Karoly1983). Webster and Holton (1982) suggested that west-erly ducts play an important role in interhemisphereteleconnections.

Sardeshmukh and Hoskins (1988) suggested the im-portance of vorticity advection by the divergent flow.Even if the convective heating anomalies and the asso-ciated upper-level divergence are within the tropicaleasterlies, the divergent flow may induce an effectiveRossby wave source in the subtropical westerlies. Largevalues of an effective Rossby wave source are associ-ated with the strong vorticity gradient in the vicinity ofthe East Asian jet and are relatively insensitive to theprecise location of the tropical forcing.

An equatorial heating may directly excite equatori-ally trapped baroclinic Rossby waves but is inefficientin exciting barotropic-type responses directly. Lim andChang (1986) found that vertical shear, differentialdamping, and the planetary boundary layer enable anenergy transfer from the heating-forced baroclinicwaves to deep barotropic modes, which allows a tele-connection between the Tropics and the extratropics.Using a two-level model, Wang and Xie (1996) demon-strated that the presence of the mean-flow verticalshear couples the barotropic and baroclinic modes sothat the barotropic mode is a Rossby wave driven by asource term that arises from the forcing of the baro-clinic mode. Furthermore, they showed that, while thebaroclinic Rossby and Kelvin modes remain trapped tothe equatorial region, the barotropic Rossby mode hasits maximum amplitude in the extratropics. Thus, anequatorial heating can generate a barotropic Rossbywave that emanates from the equatorial region to theextratropics through an additional source term that isproportional to the mean-flow vertical shear and baro-clinic mode forcing. These studies shed light on theteleconnection between the heating-induced equatori-ally trapped baroclinic Rossby waves and the midlati-tude barotropic Rossby wave response.

Opsteegh (1982) and Schneider and Watterson(1984) found that the inclusion of a zonally symmetricmeridional circulation allows for Rossby wave propa-gation in the direction of the meridional flow. The zon-ally symmetric meridional flow is an analog of the Had-ley circulation and is related to the divergent compo-

nent of the basic flow, which is usually not considered inbarotropic models. The effects of the Hadley circula-tion on Rossby wave propagation and reflection werealso explored by some other studies (Watterson andSchneider 1987; Held and Phillips 1990; Esler et al.2000; Magnusdottir and Walker 2000; Walker and Mag-nusdottir 2002). However, it is unclear how a zonallyvarying meridional basic flow affects the tropical–extratropical teleconnection.

It is the purpose of this study to investigate the roleof a zonally varying meridional basic flow on the tropi-cal–extratropical teleconnection. The sensitivity ofthe extratropical response to the zonal scale and inten-sity of the meridional basic flow is studied as well. Anondivergent barotropic model is used, and the zonalmean of the meridional basic flow is set to zero, whichmakes this study different from Opsteegh’s (1982) andSchneider and Watterson’s (1984) studies. The advec-tion of vorticity by the divergent flow, which was em-phasized by Sardeshmukh and Hoskins (1988), is alsoexcluded here. A brief description of the nondivergentbarotropic model is given in section 2. The barotropicmodel simulations are presented in section 3. The roleof a regional southerly flow is discussed in section 4,and section 5 summarizes the results and discusses thelimitations of this study.

2. The barotropic model

A linear nondivergent barotropic model is used toexplore the role of meridional basic flow on the tropi-cal–extratropical teleconnection. The model is linear-ized relative to the 300-hPa time–mean streamfunction.It has a T42 spatial resolution and a linear damping rateof (15 day)�1. The steady-state response of the modelcan be obtained by direct matrix inversion. A time in-tegration of about 15 days can also produce a quasi-steady-state response. The barotropic model can bewritten as

��2��

�t� J��, �2�� J��, �2��

� ��

�x� S � �� K�4��2��, �2�

where a prime denotes a perturbation and an overbarrefers to a time–mean basic state. The second term onthe right-hand side represents the linear damping andfourth-order diffusion. The first term on the right-handside is an elliptical vorticity forcing, which takes theform

4058 J O U R N A L O F T H E A T M O S P H E R I C S C I E N C E S VOLUME 62

S � �S0{1 � ��x � x0�2

xd2 �

�y � y0�2

yd2 �1�2}, for

��x � x0�2xd

2 ��y � y0�

2

yd2 �1�2 � 1

0, otherwise.

�3�

An example of a forcing centered at (10°N, 60°E) isshown in Fig. 1b with S0 � 1 10

�10, xd � 12°, andyd � 7°.

The basic state streamfunction takes the form

� � �zm � ALx exp�� x � xc�2Lx

2 � exp�� y � yc�2Ly2 �,�4�

where the first term on the right-hand side of Eq. (4),

zm, is the zonal mean of the 300-hPa summer climato-logical mean streamfunction, which is derived from theNational Centers for Environmental Prediction–National Center for Atmospheric Research (NCEP–NCAR) reanalysis. The second term on the right-handside of Eq. (4) is the deviation of the streamfunctionfrom the zonal mean, where A is the amplitude of the

deviation, Lx and Ly are the zonal and meridionallength scales, and (xc, yc) defines the center of the de-viation. The second term on the right-hand side intro-duces a meridional basic flow and weak zonal variationsof the zonal basic flow, as defined by the followingequations:

u � ��

�y

� uzm �2ALx�y � yc�

Ly2 exp�� x � xc�2Lx2 �

exp�� y � yc�2Ly

2 � �5a�and

��

�x

� �2A�x � xc�

Lxexp�� x � xc�2

Lx2 �

exp�� y � yc�2Ly

2 �. �5b�

FIG. 1. (a) Zonal basic flow calculated from the zonally symmetric 300-hPa summer clima-tological mean streamfunction (contour interval 5 m s�1). The thick solid lines represent thezero contours. (b) The steady eddy streamfunction response (with zonal mean removed) ofthe barotropic model to a divergence forcing (as shown by the shading). The units of thestreamfunction are 106 m2 s�1 and the units of the forcing are 10�10 s�2.


For most of the experiments in this study, the secondterm on the right-hand side of Eq. (5a) is small com-pared to uzm, and u can be approximated by its zonallysymmetric part, uzm (Fig. 1a). Note that

max�� x � xc � Lx�2, y � yc�� max�� 2A exp��0.5� �6�

and

�x � xc � Lx, y � yc� � 2A exp��1.0�. �7�

The intensity of � is controlled by A and the spatialscale of � may be tuned by changing Lx. Thus, the sen-sitivity of the model response to the meridional basicflow may be examined by adjusting these two param-eters. By setting xc � 60°E, yc � 20°N, and Ly � 20° inthe following experiments, we introduce a high cen-tered over Iran with a 1–2 m s�1 southerly (northerly)flow to the west (east), as in Fig. 2a. The meridionalflows connect the easterly and westerly basic flows, withthe maxima around the critical line (u � 0). The highmimics the upper-level anticyclone over South Asia.Observations show that the southerly flow to the westof the anticyclone over the Red Sea has a magnitude of

2–3 m s�1 near the zero line of the zonal flow (figurenot shown).

3. Barotropic model simulation

In this section, a suite of numerical experiments isdesigned to explore the role of the meridional basicflow and the sensitivity of the extratropical response tothe meridional basic flow and the tropical forcing (pa-rameters listed in Table 1). The steady-state responseof the model is solved by direct matrix inversion.

a. Inclusion of meridional flow

To facilitate the sensitivity tests, two control runs areconducted first. In experiment Ctl-0, we use the zonallysymmetric 300-hPa summer climatological meanstreamfunction as the basic state, that is, � zm in Eq.(4). As shown in Fig. 1a, this corresponds to an axiallysymmetric zonal flow with no meridional flow. Theeasterly basic flow extends to 21°N. A vorticity forcingis placed at (10°N, 60°E), deep within the easterly meanflow (shading in Fig. 1b). As expected, the extratropicalresponse is negligible in both hemispheres, and in theTropics the response is dominated by the Rossby waveresponse west of the forcing. Because the zonal mean is

FIG. 2. (a) A departure is added to the zonal mean streamfunction, which introducesmeridional basic flows. (b) The steady eddy response of the barotropic model is indicated bythin black contours and the forcing is indicated by the thick black contour. The units of themeridional flow in (a) are m s�1 and the units of the streamfunction in (b) are 106 m2 s�1.


removed from the response field, there is an anticy-clonic response to the east of the forcing.

Next, meridional flows are introduced to the basicstate (experiment Ctl-1, see Table 1). As shown in Fig.2a, a southerly flow is present west of 60°E and a north-erly flow east of 60°E. Although the westerly and east-erly basic flows are both slightly enhanced near 60°E,these zonal derivations are one order of magnitude

smaller than the zonal mean (figures not shown). As inexperiment Ctl-0, the vorticity forcing is placed at(10°N, 60°E). The striking feature of the response inFig. 2b is a wave train originating from the region of thesoutherly basic flow. This wave train follows an almostgreat circle path over Asia. The magnitude of the ex-tratropical response over Asia is comparable to that ofthe tropical response. As in experiment Ctl-0 (Fig. 1b),the extratropical response in the Southern Hemisphereis negligible.

b. Sensitivity to intensity of the meridional flow

In this subsection, we test the sensitivity of the ex-tratropical response to the intensity of the meridionalbasic flow. We reduce the intensity of the meridionalflow by taking A � A0/4 in experiment Vamp-1 (Figs.3a,b) and A � A0/2 in experiment Vamp-2 (Figs. 3c,d).The zonal scales of the meridional flow (Lx) in bothruns are the same as in experiment Ctl-1. As the inten-sity of the meridional basic flow is reduced, the extra-tropical response is weakened significantly. Compari-son with Fig. 2 indicates that the intensity of the extra-tropical response is approximately linearly proportionalto the strength of the southerly flow while the spatialstructure of the response pattern remains almost thesame. The zonal and meridional wavelengths of the re-sponses decrease slightly in Figs. 3b and 3d. This islikely due to the slight weakening of the zonal basicflow with the decrease of the amplitude A, as shown inEq. (5a).

TABLE 1. Barotropic model experimental design. A perturba-tion is added to the zonal mean streamfunction: � (A0/Lx)exp�� (x � xc)

2/Lx2] exp�� (y � yc)

2/Ly2], where xc, yc, and Ly are

fixed and S0 � 1 1010. The vorticity forcing takes the form S �

S0 f [(x�x0)/xd].

Expts A0 Lx X0 Xd

Control runs Ctl-0 0 60°E 12°Ctl-1 8 105 25° 60°E 12°

Sensitivity to Vintensity

Vamp-1 2 105 25° 60°E 12°Vamp-2 4 105 25° 60°E 12°

Sensitivity to V scale Vscl-1 8 105 12.5° 60°E 12°Vscl-2 8 105 50° 60°E 12°Vscl-3 16 105 12.5° 60°E 12°Vscl-4 4 105 50° 60°E 12°

Sensitivity to forcingscale

Dscl-1 8 105 25° 60°E 6°Dscl-2 8 105 25° 60°E 18°

Sensitivity to forcinglocation

E10 8 105 25° 70°E 12°E20 8 105 25° 80°E 12°E30 8 105 25° 90°E 12°E40 8 105 25° 100°E 12°

FIG. 3. Same as in Fig. 2 except that in (a), (b) experiment Vamp-1 the amplitude of departure from the zonal mean is one-fourthof that in Fig. 2 (A � A0/4) and in (c), (d) experiment Vamp-2 the amplitude of departure from zonal mean is half that in Fig. 2(A � A0/2). The contour levels for the meridional basic flow in (a) and (c) are �0.3, �0.9, �1.5, . . . m s

�1.


c. Sensitivity to the spatial scale of the meridionalflow

Webster and Holton (1982) suggested that the spatialscale of the westerly duct is important for Rossby wavepropagation. When the zonal scale of the westerlies islarger than the zonal wavelength of the perturbations,Rossby waves may propagate through the westerlyduct; otherwise, Rossby waves are blocked at the criti-cal latitude (u � 0). In Fig. 2, the zonal scale of thesoutherly flow is about 60°, which is much smaller thanthe zonal wavelength of the Rossby wave. It suggeststhat the zonal scale of the meridional flow does notconstrain the wavelength of Rossby waves. The sensi-tivity of the extratropical response to the zonal scale ofthe southerly basic flow will be examined in this sub-section. In experiment Vscl-1 (Figs. 4a,b), the zonalscale of the southerly flow (Lx) is reduced to 12.5°,which is half of that in experiment Ctl-1 and is thusmuch smaller than the zonal wavelength of the forcedRossby wave. A wave train still emanates from the re-gion of the southerly basic flow and reaches the BeringSea. The extratropical response in Fig. 4b is muchweaker and the wavelength is reduced slightly com-pared to the control run in Fig. 2b. From Eq. (5a), udecreases with the decrease of Lx. Therefore, the wave-length of the stationary Rossby wave also decreaseswith the decrease of Lx. Since the change of the zonalscale of the southerly flow is much larger than thechange of the Rossby wavelength, the wavelength ofthe stationary Rossby wave is decreased due to the

weakening zonal mean flow instead of being con-strained by the zonal scale of the southerly flow.

This conclusion is further supported by experimentVscl-2 (Fig. 4c,d). As the zonal scale of the meridionalbasic flow is enlarged by a factor of 2, the extratropicalwave response is intensified while the wavelength in-creases only slightly. The spatial structure of the extra-tropical wave train pattern remains almost the same asin experiment Ctl-1 (Fig. 2b). Therefore, the spatialscale of the Rossby wave is independent of the zonalscale of the meridional basic flow, which is differentfrom the scale constraint for a westerly duct. Theseresults suggest that it is the Rossby wave source, insteadof the wave form, that is transmitted by the southerlyflow, which we refer to as a southerly conveyor.

Note also that the extratropical responses are weak-ened in Fig. 4b and strengthened in Fig. 4d. Since theadvection of the Rossby wave source depends on theintegral of the southerly flow, we expect the amplitudeof the extratropical response to remain the same if theintegral of the southerly flow is fixed. From Eq. (5b),we have

�x

xc

dx � �x

xc

d�

� ALx�1 � exp�� x � xc�2Lx

2 �� exp�� y � yc�2

Ly2 �. �8�

FIG. 4. Same as in Fig. 2 except that in (a), (b) experiment Vscl-1 the zonal scale of the meridional basic wind is half that in Fig. 2and in (c), (d) experiment Vscl-2 the zonal scale of the meridional basic wind is twice that in Fig. 2. The maximum magnitude A0 is thesame.


Therefore, the integral of the southerly flow may bekept constant by fixing ALx. The meridional basic flowsand the model responses are shown in Fig. 5 for experi-ment Vscl-3 and experiment Vscl-4, in which Lx is var-ied as in experiment Vscl-1 and experiment Vscl-2, re-spectively, but ALx is held to the value in experimentCtl-1. As we expected, the intensity of the extratropicalresponses in experiment Vscl-3 and experiment Vscl-4are similar to that in experiment Ctl-1. However, thestructure of the pattern changes because the structureand intensity of the basic flows are slightly differentfrom those in experiment Ctl-1.

d. Sensitivity to the forcing scale

To determine the sensitivity of the extratropical re-sponse to the spatial scale of the forcing, we reduce thezonal scale of the forcing to 50% of that in Fig. 2. Al-though the response (Fig. 6a) is weaker because thetotal integrated forcing is reduced, the spatial structureof the response remains essentially the same. In Fig. 6c,the response is shown for the case with the zonal scaleof the forcing enlarged by a factor of 2. The response isstronger, but the spatial structure is still the same. Com-parison of these two experiments with experiment Ctl-1(shown in Fig. 6b for comparison) suggests that thescale of response is determined solely by the basic flowand is independent of the forcing scale. However, thestrength of the response depends on the integratedstrength of the forcing.

e. Sensitivity to the forcing location

The previous sections suggest that the southerly basicflow may transmit a Rossby wave source from the east-erly basic flow and excite a stationary Rossby waveresponse in the extratropical westerlies. An extratropi-cal wave train originates from the southerly conveyorregion. It is hypothesized that the extratropical re-sponse is relatively insensitive to the precise location ofthe forcing but sensitive to the location of the southerlyconveyor. In this subsection, we test this hypothesis byexamining the sensitivity of the extratropical responseto the location of the tropical forcing.

The same basic state as in experiment Ctl-1 is used inthe barotropic model, and the forcing intensity andscale are fixed, but the forcing center is shifted from 60°to 100°E in 10° intervals. As shown in Fig. 7, the re-sponse in the tropical easterly region shifts eastwardconsistently with the shift of the forcing, while the ex-tratropical response remains the same spatial structureand is spatially phase locked to the location of thesoutherly conveyor. As the forcing and its local re-sponse moves farther from the southerly conveyor, theextratropical response becomes weaker, which suggeststhat the amplitude of the extratropical response de-pends on the intensity of the vorticity field near thesoutherly conveyor.

4. The role of a southerly conveyor

How does a southerly component of the basic flowenable forcing in the tropical easterlies to excite Rossby

FIG. 5. Same as in Fig. 4 except that both A and Lx are varied so that the integral of the southerly flow (ALx) is fixed (see the textfor details).


waves in the extratropical westerlies? An explanation issought using the linearized vorticity equation in thepresence of a meridional mean flow,

�

�t� u

�

�x� �u ��x � ��y��f � � � � ��y � S � �,

�9�

where an overbar refers to a time–mean basic state, theprimes for the perturbations have been omitted, S is avorticity source, and �� is a linear damping. The firstterm on the right-hand side of Eq. (9) is the anomalousvorticity advection by the meridional basic flow. It in-troduces an additional “source” term that transportsvorticity perturbation downstream in the meridionalflow. The magnitude of this term depends on the mag-

nitude of the meridional basic flow and the meridionalgradient of the vorticity perturbation. The stronger themeridional flow, the larger the extratropical responsewill be, as demonstrated in Figs. 2 and 3. In addition, ameridional flow with a larger zonal scale allows morevorticity advection and, thus, will also induce a largerextratropical response (Fig. 4). The meridional gradientof the vorticity perturbation near the southerly con-veyor plays a similar role. The stronger the vorticitygradient, the larger the extratropical response will be.Therefore, as the forcing moves farther from the south-erly conveyor, the extratropical response is weakened(Fig. 7).

To illustrate how this mechanism works, the differ-ence of the streamfunction responses between experi-ment Ctl-0 and experiment Ctl-1 from day 2 to day 16

FIG. 6. Response of the barotropic model to forcings with different zonal scale: (a) zonalscale is half the control run, (b) control run Ctl-1, and (c) zonal scale is twice the controlrun.


is shown in Fig. 8 (the model is solved by time integra-tion). If we neglect the weak zonal variations of thezonal basic flow and the basic-state vorticity in experi-ment Ctl-1, the only difference between these two runsis the inclusion of the meridional basic flow and thus an

additional source term in experiment Ctl-1, which isshown by the shading in Fig. 8. The anomalous vorticityadvection by the southerly basic flow increases withtime. It is confined to the region of the southerly con-veyor and extends poleward to the mean westerlies.

FIG. 7. Responses of the barotropic model to the forcing at various longitudinal locationsalong 10°N. The forcing is shifted from (a) 60° to (e) 100°E in 10° intervals. The basic statein Ctl-1 is used in the model, and (a) is identical to Fig. 2b.


Before day 6, the response differences are weak. Onday 6, a stationary wave train appears in the region ofthe southerly conveyor. It takes an almost great circleroute over Asia, which is consistent with Hoskins andKaroly’s (1981) description. The wave train reaches aquasi-stationary state over East Asia in 16 days.

A perturbation expansion approach can also showhow the extratropical wave train is generated from analternative point of view. The steady state of Eq. (9)may be written as

�� u ��x��2 � ��x ��y � �* ��x�� V ��y ��2�� S,

�10�

where � � [�]/[u] � O(10�1), � � �V, and �* is meridi-onal gradient of the absolute vorticity; can be ex-panded for different orders of �. Assuming

� � �0 � �i�1

�

�i�i �11�

FIG. 8. Difference of the streamfunction responses (contours) between Ctl-0 and Ctl-1 at (a) day 2, (b) day 4, (c) day 6, (d) day 8,(e) day 12, and (f) day 16. The contour intervals for streamfunction is 3 106 m2 s�1, with the zero contour omitted. The anomalousvorticity advection by meridional basic flow is represented by shadings (units: 1 106 m2 s�1).


Fig 8 live 4/C

and denoting

L � �� u ��x��2 � ��x ��y � �* ��x�,

we then have

L�0 � S, �12a�

L�i�1 � �V�

�y��2�i�, i � 0, 1, 2, . . . , �12b�

Equation (11) does not converge for the meridionalbasic flow used in experiment Ctl-1. To obtain conver-gence, we can rescale V by a factor �(� � 1), and Eq.(10) can be written as

�� u ��x��2 � ��x ��y � �* ��x�

�� V��

�y��2�� S, �13�

where a prime means that the response is for the re-scaled equation and � may be expanded in terms of ��,

�� 0 � �i�1

�

�ii��i. �14�

The equations for �0, �1, and �2 may be written as

L��0 � S, �15a�

L��i�1 � �V�

�y��2��i�, i � 0, 1, 2, . . . �15b�

Recall from the previous experiments that the ampli-tude of the extratropical response is linearly propor-tional to of the magnitude of the southerly conveyor(�), while its spatial structure is insensitive to the mag-nitude of �. Also note that the O(1) equation [Eq.(15a)] is not affected by � or the rescaling. The totalresponse for Eq. (10) can be approximated as

�* � ��0 �1

�i�1

N

i�i��i � �0 � �i�1

N

i�1�i��i. �16�

The advantage of this method is that � may be chosenarbitrarily as long as Eq. (16) converges, and a larger �leads to quicker convergence. If � � 0.7, Eq. (16) con-verges and N � 3 provides a good approximation to theresponse of the Ctl-1 run.

For simplicity, the zonal mean of the 300-hPa clima-tological mean streamfunction is used as the basic statein the operator L [under this approximation, the zonaladvection of the mean relative vorticity by the anoma-lous flow,

ux,

is neglected. Further calculations show that includingthis term will lead to a better approximation but itseffects are small], and then the zeroth order solution ofEq. (15a) is identical to the response in experimentCtl-0 (Fig. 1b). The source term �V�(�20)/�y is calcu-lated and used to force Eq. (15b); �1 is solved andshown in Fig. 9a. The source term extends poleward inthe region of the southerly conveyor, which is similar toFig. 8. The response is no longer confined to the Trop-ics; rather, a wave train originates from the southerlyconveyor and extends to the Bering Sea: �22 and �

33are solved similarly. Compared to Fig. 9a, the sourceterms in Figs. 9b,c are slightly displaced northward,which causes a spatial phase shift in 2 and 3. How-ever, the wavelengths and wave paths of the first, sec-ond- and third-order responses are rather similar be-cause they are determined by the basic flow. As shownin Fig. 9d, the total response, *, is a good approxima-tion to the Ctl-1 response.

The above analysis suggests that the tropical forcingfirst generates a local Rossby wave response, and thenthe southerly conveyor transfers the vorticity perturba-tions northward to the westerly basic flow through ad-vection, which then excites an extratropical Rossbywave train pattern. For this mechanism to be effective,the southerly conveyor needs to connect both the basic-state easterlies and westerlies. Other numerical experi-ments (figures not shown) in which a southerly flow isconfined to the mean easterlies or to the mean wester-lies do not have a teleconnection between the Tropicsand the extratropics.

5. Concluding remarks

The role of meridional basic flow in generating tele-connections between the deep Tropics and the extra-tropics is studied in this paper. In the barotropic wavetheory that considers only zonal mean flows, a forcingembedded in the tropical easterlies does not allowRossby waves to propagate into the extratropics. Inprevious studies, geographically fixed teleconnectionpatterns are usually explained by barotropic instabilityof the basic flow or eddy–low frequency flow interac-tion (Simmons et al. 1983; Lau 1981; Ting and Lau 1993;Hoerling and Ting 1994; Jin et al. 2004, manuscript sub-mitted to J. Atmos. Sci.; Pan 2003). Sardeshmukh andHoskins (1988) suggested that the advection of vorticityby divergent flows can induce a large effective Rossbywave source in the subtropical jet region, which is rela-tively insensitive to the precise location of the tropicalforcing. Although the meridional vorticity advectionwas also included in their model simulation, its role wasnot studied. Another mechanism is proposed in this


FIG. 9. (a) First-order (�1), (b) second-order (��22), (c) third-order (�

2�33) response, (d) the approximate response * � 0 ��i�1

3 �i�1�ii (black contours) and the exact response of Ctl-1 (green contours). The shading represents the vorticity forcing; � is a smallparameter, and � is a rescaling parameter (see the text for details).


Fig 9 live 4/C

paper to explain the geographically fixed teleconnec-tion patterns using the Rossby wave theory. It wasfound that a southerly basic flow (a southerly conveyor)will transfer Rossby wave sources downstream throughadvection so that forcing embedded in the deep tropicaleasterlies may excite a Rossby wave response in theextratropical westerlies. Stronger southerly flows orsoutherly flows with larger zonal scales favor a strongerextratropical response. Since a southerly conveyor de-termines the location of the effective Rossby wavesource in the westerly basic flows, the spatial structureof the extratropical response is relatively insensitive tothe precise location of the tropical forcing as long as thetropical response can reach the southerly conveyorarea. The spatial structure of the extratropical responseis determined by the westerly basic flows and is insen-sitive to the scale of the tropical forcing or the scale ofthe southerly conveyor.

Since a simple model is used in this study, caution isadvisable in the interpretation of the results. For ex-ample, no Kelvin wave response is generated in thenondivergent, barotropic model, and the local responseis dominant by Rossby waves that are confined to thewest of the forcing. In experiment Ctl-1 (Fig. 2), thevorticity field near the northerly flow region is undis-turbed, which is why the extratropical response is neg-ligible in the Southern Hemisphere. When the forcing isshifted to the east of the northerly flow (Fig. 7), theextratropical response in Southern Hemisphere is stillweak mainly due to the weak northerly flow near thecritical line in the Southern Hemisphere, which cannotgenerate a strong Rossby wave source. Additional ex-periments (not shown) show that a northerly conveyormay help to excite extratropical Rossby waves in theSouthern Hemisphere under favorable conditions. Fi-nally, note that divergent flows are excluded in thisbarotropic model. In the real atmosphere, a diabaticforcing in the Tropics will generate a divergent flow,and the vorticity advection by the divergent flow andsoutherly flows may work together and shift the effec-tive Rossby wave source poleward.

Whereas a meridional flow with a dipole structurehave been used in all of these experiments, the conclu-sions do not depend on this special structure of themeridional flow, as indicated in section 4. Even thoughthe experiments based on a simple model illustrate thebasic mechanism, the real atmosphere is more compli-cated. In the real world, westerly mean flows and tran-sient eddies are strong during northern winter, andbarotropic instability and eddy–mean flow interactionhave dominant roles in maintaining many well-knownteleconnection patterns, such as the Pacific–NorthAmerica pattern and the North Atlantic Oscillation

(Blackmon et al. 1984; Lau 1988; Jin et al. 2004, manu-script submitted to J. Atmos. Sci.). When mean westerlyflows and transient eddies are weak and tropical east-erlies penetrate farther north during northern summer,Rossby wave dispersion related to southerly conveyersmay be important. For example, Wang et al. (2005,manuscript submitted to J. Climate) suggested that thismechanism is essential for the generation of a geo-graphically fixed wave train pattern associated with therainfall variations over the U.S. Great Plains duringnorthern summer.

Acknowledgments. We wish to thank Dr. Janet Beckerfor helpful comments and Professor Russ Elsberry forreading the manuscript. This work was supported byNOAA’s Pan American Climate Studies program andNSF’s Climate Dynamics Program under GrantATM03-29531 at the University of Hawaii, and byNOAA under Grant NA01AANRG0011 and NSF un-der Grant ATM-0101135 at the Naval PostgraduateSchool.

REFERENCES

Blackmon, M. L., Y.-H. Lee, and H.-H. Hsu, 1984: Time variationof 500-mb height fluctuations with long, intermediate andshort time scales as deduced from lag-correlation statistics. J.Atmos. Sci., 41, 981–991.

Branstator, G., 1983: Horizontal energy propagation in a barotro-pic atmosphere with meridional and zonal structure. J. At-mos. Sci., 40, 1689–1708.

Esler, J. G., L. M. Polvani, and R. A. Plumb, 2000: The effect ofthe Hadley circulation on the propagation and reflection ofplanetary waves in a simple one-layer model. J. Atmos. Sci.,57, 1536–1556.

Held, I. M., and P. J. Phillips, 1990: A barotropic model of theinteraction between the Hadley cell and a Rossby wave. J.Atmos. Sci., 47, 856–869.

Hoerling, M. P., and M. Ting, 1994: Organization of extratropicaltransients during El Niño. J. Climate, 7, 745–766.

Horel, J. D., and J. M. Wallace, 1981: Planetary-scale atmosphericphenomena associated with the Southern Oscillation. Mon.Wea. Rev., 109, 813–829.

Hoskins, B. J., and D. J. Karoly, 1981: The steady linear responseof a spherical atmosphere to thermal and orographic forcing.J. Atmos. Sci., 38, 1179–1196.

——, A. J. Simmons, and D. G. Andrews, 1977: Energy dispersionin a barotropic atmosphere. Quart. J. Roy. Meteor. Soc., 103,553–567.

Karoly, D. J., 1983: Rossby wave propagation in a barotropic at-mosphere. Dyn. Atmos. Oceans, 7, 111–125.

Lau, N.-C., 1981: A diagnostic study of recurrent meteorologicalanomalies appearing in a 15-year simulation with a GFDLgeneral circulation model. Mon. Wea. Rev., 109, 2287–2311.

——, 1988: Variability of the observed midlatitude storm tracks in


relation to low-frequency changes in the circulation pattern.J. Atmos. Sci., 45, 2718–2743.

Lim, H., and C.-P. Chang, 1986: Generation of internal- and ex-ternal-mode motions from internal heating: Effects of verti-cal shear and damping. J. Atmos. Sci., 43, 948–957.

Magnusdottir, G., and C. Walker, 2000: On the effects of theHadley circulation and westerly equatorial flow on planetary-wave reflection. Quart. J. Roy. Meteor. Soc., 126, 2725–2745.

Nitta, T., 1987: Convective activities in the tropical western Pacificand their impacts on the Northern Hemisphere summer cir-culation. J. Meteor. Soc. Japan, 65, 165–171.

Opsteegh, J. D., 1982: On the importance of the Hadley circula-tion for cross-equatorial propagation of stationary Rossbywaves. On the Theory and Applications of Simple ClimateModels to the Problem of Long-Range Weather Prediction:Collected Papers Presented at KNMI Workshop, R. Mareau,C. J. Kok, and R. J. Haarsma, Eds., Koninklijk NetherlandsMeteorologisch Instituut, 49–57.

Pan, L.-L., 2003: A study of dynamic mechanisms of annularmodes. Ph.D. dissertation, University of Hawaii at Manoa,193 pp.

Sardeshmukh, P., and B. J. Hoskins, 1988: The generation of glob-al rotational flow by steady idealized tropical divergence. J.Atmos. Sci., 45, 1228–1251.

Schneider, E. K., and I. G. Watterson, 1984: Stationary Rossbywave propagation through easterly layers. J. Atmos. Sci., 41,2069–2083.

Simmons, A. J., 1982: The forcing of stationary wave motion bytropical diabatic heating. Quart. J. Roy. Meteor. Soc., 108,503–534.

——, J. M. Wallace, and G. Branstator, 1983: Barotropic wavepropagation and instability, and atmospheric teleconnectionpatterns. J. Atmos. Sci., 40, 1363–1392.

Ting, M., and N.-C. Lau, 1993: A diagnostic and modeling study ofthe monthly mean wintertime anomalies appearing in a 100-year GCM experiment. J. Atmos. Sci., 50, 2845–2867.

Walker, C., and G. Magnusdottir, 2002: Effect of the Hadley cir-culation on the reflection of planetary waves in three-dimensional tropospheric flows. J. Atmos. Sci., 59, 2846–2859.

Wang, B., and X. Xie, 1996: Low-frequency equatorial waves invertically sheared zonal flow. Part I: Stable waves. J. Atmos.Sci., 53, 449–467.

——, R. Wu, and K.-M. Lau, 2001: Interannual variability of theAsian summer monsoon: Contrasts between the Indian andthe western North Pacific–East Asian monsoons. J. Climate,14, 4073–4090.

Watterson, I. G., and E. K. Schneider, 1987: The effect of theHadley circulation on the meridional propagation of station-ary waves. Quart. J. Roy. Meteor. Soc., 113, 779–813.

Webster, P. J., and J. R. Holton, 1982: Cross-equatorial responseto middle-latitude forcing in a zonally varying basic state. J.Atmos. Sci., 39, 722–733.


teleconnections from tropics to northern extratropics ...zhuowang/...symmetric 300-hpa summer...

Documents