Stratospheric ozone and the morphology of the northernhemisphere planetary waveguide

John R. Albers,1 John P. McCormack,2 and Terrence R. Nathan1

Received 11 April 2012; revised 8 November 2012; accepted 14 November 2012; published 31 January 2013.

[1] A middle atmosphere general circulation model is used to examine the effects ofzonally asymmetric ozone (ZAO) on the Northern Hemisphere planetary waveguide(PWG) during winter (December–February). The morphology of the PWG is measured bya refractive index, Eliassen-Palm flux vectors, the latitude of the subtropical zero wind line,and the latitude of the subtropical jet. ZAO causes the PWG to contract meridionally in theupper stratosphere, expand meridionally in the lower stratosphere, and expand verticallyin the upper stratosphere and lower mesosphere. The ZAO-induced changes in the PWGare the result of increased upward and poleward flux of planetary wave activity intothe extratropical stratosphere and lower mesosphere. These changes cause an increasein the Eliassen-Palm flux convergence at high latitudes, which produces a warmer andweaker stratospheric polar vortex and an increase in the frequency of stratosphericsudden warmings. The ability of ZAO to alter the flux of planetary wave activity intothe polar vortex has important implications for accurately modeling wave-modulated andwave-driven phenomena in the middle atmosphere, including the 11-year solar cycle,stratospheric sudden warmings, and the phase of the Northern Hemisphere annular mode.

Citation: Albers, J. R., J. P. McCormack, and T. R. Nathan (2013), Stratospheric ozone and the morphology of thenorthern hemisphere planetary waveguide, J. Geophys. Res. Atmos., 118, 563–576, doi:10.1029/2012JD017937.

1. Introduction

[2] The concept of a planetary waveguide (PWG) was firstproposed by Dickinson [1968] to describe the regions of theNorthern Hemisphere (NH) winter stratosphere where thebackground zonal wind supports the upward propagation ofplanetary waves. Matsuno [1970] expanded on Dickinson’sPWG concept and derived a refractive index for planetarywaves on a sphere and showed that planetary wave propaga-tion is strongest in regions where the refractive index is largeand positive. For a typical NH winter zonal wind distribution(Figure 1), large, positive values of the refractive index occurin the region bounded on the south by the zero wind line[Tung, 1979] and on the north by the southern edge of thepolar vortex [Chapman and Miles, 1981].[3] The zero wind line and the vortex edge do not, however,

provide a complete description of planetary wave propagationwithin the PWG; indeed, there exist robust features of thebackground flowwithin the PWG that exert substantial controlover how much planetary wave activity enters the PWG fromthe troposphere and where that wave activity propagates onceinside the stratosphere. For example, the vertical shear of the

zonal-mean wind near the tropopause exerts significant influ-ence over the amount of planetary wave activity that propa-gates upwards from the troposphere into the lower stratosphere[Chen and Robinson, 1992]. And once planetary waves arrivein the lower stratosphere, the local minimum in the refractiveindex associated with the negative vertical and horizontalgradients of the zonal-mean wind (above and to the north ofthe subtropical jet) splits the PWG into two “channels”through which planetary wave activity may either propagatepoleward and upward into the interior of the stratosphere orequatorward where the waves are confined to the lowerstratosphere [Chapman and Miles, 1981; Huang and Gambo,1982; Li et al., 2006]. For the planetary waves that propagateupward and poleward into the interior of the stratosphere, theheight that planetary waves are able to propagate is largely con-trolled by the ratio of the meridional gradient of potential vor-ticity to the zonal-mean westerly wind [Charney and Drazin,1961; Matsuno, 1970]. In addition, Nigam and Lindzen [1989]showed that modest latitudinal shifts in the location of thesubtropical jet greatly alter the amount of planetary waveactivity that is refracted into the mid-latitude and polar strato-sphere. These results show that the PWG is largely defined byits meridional width (measured by the location of the zero windline), its vertical extent (measured by the refractive index), andthe strength of wave propagation around the subtropical jet andwithin the waveguide (both measured by the refractive index).[4] The structures of the zonal-mean wind and PWG are

determined by a balance between radiative heating due toozone, long-wave radiative cooling, and dynamical heatingdue to planetary waves. Any physical process that alters thisthree-way balance will in turn modulate the PWG. One such

1Department of Land, Air, and Water Resources: Atmospheric ScienceProgram, University of California, Davis, CA, USA.

2Space Science Division, Naval Research Laboratory, Washington,D.C., USA.

Corresponding author: J. R. Albers, Department of Land, Air,and Water Resources: Atmospheric Science Program, Davis, CA, USA.(albersjohn@hotmail.com)

©2012. American Geophysical Union. All Rights Reserved.2169-897X/13/2012JD017937

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JOURNAL OF GEOPHYSICAL RESEARCH: ATMOSPHERES, VOL. 118, 563–576, doi:10.1029/2012JD017937, 2013

process involves the feedbacks between planetary waveactivity and stratospheric ozone. Specifically, as planetarywaves propagate from the troposphere into the stratosphereduring NH winter, they produce large zonal asymmetries(waves) in wind, temperature, and ozone [Gabriel et al.,2007]. Observations show that zonally asymmetric ozone(ZAO) constitutes a significant fraction of the total strato-spheric ozone field [e.g., Wang et al., 2005; Gabriel et al.,2007; Crook et al., 2008]. That fraction reaches 10% duringboreal winter (based on decadal averages) [Crook et al., 2008],15% in the lower stratosphere near ~70�N and ~65�S [Gillettet al., 2009], 50% during the Antarctic stratospheric suddenwarming of 2002 [Wang et al., 2005], and 50% during thebreakup of the Antarctic ozone hole [Crook et al., 2008]. Thephasing between the wind, temperature, and ozone wavesproduce fluxes that modulate both the spatial-temporaldamping of the wave fields and the driving of the zonal-meancirculation [Nathan and Li, 1991; Nathan and Cordero, 2007;Albers and Nathan, 2012].[5] General circulation modeling studies have also shown

that ZAO has a large effect on the zonal-mean temperatureand wind structure of the wintertime middle atmosphere inboth hemispheres [Kirchner and Peters, 2003; Sassi et al.,2005; Gabriel et al., 2007; Brand et al., 2008; Crook et al.,2008; Waugh et al., 2009; Gillett et al., 2009; McCormacket al. 2011]. For example, Crook et al. [2008] examined theeffect of ZAO on the high latitude Southern Hemisphere andfound that including ZAO produces lower stratosphericcooling that is comparable in magnitude to the coolingproduced by the springtime Antarctic ozone hole. In a studythat focused on the Northern Hemisphere, McCormack et al.[2011] carried out an ensemble of GCM simulations over theDecember–March period and found that ZAO produces awarmer and more disturbed stratospheric polar vortex and an

increased incidence of stratospheric sudden warmings(SSWs). Although McCormack et al. showed that ZAOchanges temperatures and winds in the polar stratosphere,the mechanisms involved in the changes were not fullyinvestigated; yet, understanding the connections betweenZAO and the polar stratosphere is vital to producing reliableassessments of human-caused impacts on key features ofthe climate system, including the upward flux of planetarywave activity, the frequency of SSWs, and the phase of theNorthern Hemisphere annular mode. Here we investigatehow the physics associated with ZAO changes the mor-phology of the PWG. The results that we obtain underscorethe importance of accurately accounting for ZAO in globalclimate models.[6] In the following section, we introduce the model and

diagnostics that we use to measure changes in the mor-phology of the PWG. In section 3, we present our results,and in section 4, we discuss the implications of our resultsfor global climate modeling.

2. Model Description and Diagnostics

[7] We investigate the effects of ZAO on planetary wavepropagation during NH winter using the NOGAPS-ALPHAglobal spectral general circulation model [see, for example,Eckermann et al., 2009; McCormack et al., 2011, andreferences therein]. NOGAPS-ALPHA extends from thesurface to ~90 km in height and utilizes 68 hybrid (s-p)vertical levels with triangular truncation at wave number 79.The model includes prognostic equations for O3 and H2O,which are calculated using the photochemical parameteriza-tions of McCormack et al. [2006, 2008]. Shortwave heatingand long-wave cooling rates are computed using prognosticO3 and H2O fields and a fixed profile of CO2. Sea surface

Figure 1. Three-month average (DJF) ZMO3 ensemble mean (a) zonal-mean wind and (b) n2s withEP-flux vectors. In Figure1a, the thick black line denotes the zero wind line. In Figure 1b, the contoursdenote n2s , which has been non-dimensionalized with a2 (contour interval is 4). The EP-flux vectors(arrows) are plotted only in regions where the EP-flux divergence is less than approximately�25� 106 kg s�2

(this threshold eliminates visually distracting vectors that are inconsequential to the physics). The scaling ofthe EP-flux vectors is discussed in section 2.

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temperatures and ice distributions are specified at the modellower boundary using 12-hourly observations from opera-tional global analyses.[8] We analyze 15 pairs of model simulations spanning

early December to late March (121 days). Each pair ofsimulations was initialized using identical profiles of wind,temperature, and chemical constituents acquired from theNOGAPS-ALPHA data assimilation system [Eckermannet al., 2009]. The initialization dates were staggered in timein order to generate an ensemble of 15 independent pairs ofmodel simulations [for details, seeMcCormack et al., 2011].For each pair of simulations, one used fully prognosticozone in the radiative heating and cooling calculations(designated 3DO3), and the other used zonal-mean values ofthe instantaneous prognostic ozone field to evaluate theradiative heating and cooling rates (designated ZMO3). Weisolate the effect of ZAO on the circulation by taking thedifference between 3DO3 and ZMO3 simulations for anygiven variable. When calculated in this way, the differencebetween the 3DO3 and ZMO3 simulations isolates the effectof ZAO-heating but not the effect of changes in zonal-meanozone heating due to wave-ozone transports. Consistent withAlbers and Nathan [2012], however, the heating due to thewave-ozone transports is of secondary importance comparedto the ZAO-heating. This means that the zonal-mean ozonedistribution is similar for the 3DO3 and ZMO3 simulations.[9] Statistical significance of the differences between ensem-

ble means of model variables is assessed using a Student’st-test at the 95% confidence level. As stated in McCormacket al. [2011], the 3DO3 model runs produced four SSWs,while the ZMO3 runs produced only one SSW. Statisticalsignificance calculations were also carried out with all of theruns containing SSWs removed. Doing so did not changethe qualitative nature of our results.[10] We evaluate the morphology of the PWG in terms of

two properties: (i) its shape, measured by its vertical extentand meridional width, and (ii) by the strength and directionof wave propagation within the waveguide itself. To mea-sure these properties, we use a refractive index, Eliassen-Palm flux vectors, the latitude of the subtropical zero windline, and the latitude of the subtropical jet.[11] To diagnose the shape of the PWG and to distinguish

regions of wave propagation versus evanescence, we employthe spherical form of the quasi-geostrophic refractive indexsquared [Andrews et al., 1987]:

n2s f; zð Þ ¼ �qf�u� s

a cos f

� �2

� f

2NH

� �2

; (1)

where

�qf ¼ 2Ω cos fa

� 1

a2�u cos fð Þfcos f

� �f� f 2

rr�uzN2

� �z

(2)

is the zonal-mean potential vorticity gradient, ū is the zonal-mean wind, f is latitude, z is height, s is the spherical integerzonal wave number, N(z) is the buoyancy frequency, f is theCoriolis parameter, H is the mean scale height [=7 km],r [=r0exp(�z/H)] is the standard density in log-pressure coor-dinates, r0 is the sea-level reference density, a is the radius ofthe Earth, Ω is the Earth’s rotation frequency, and subscriptsdenote derivatives with respect to the given variable.

[12] Planetary waves propagate in regions where n2s > 0and are evanescent in regions where n2s < 0. Thus, the ver-tical extent of the PWG is measured by the height wheren2s ¼ 0. The meridional width of the PWG is determined byits northern and southern boundaries. We measure thenorthern boundary of the PWG by the latitude where n2sbecomes negative. The southern boundary of the PWG,however, cannot be measured directly by n2s . This is becausen2s�� �� ! 1 as ū! 0, so that interpretation of n2s near thesubtropical zero wind line becomes problematic. What isimportant here is that near the zero wind line the sign of ūlargely controls the sign of n2s . We therefore use the locationof the zero wind line as a measure of the southern boundaryof the PWG, consistent with previous studies [e.g., Holtonand Tan, 1982]. In addition, we also use the location of thesubtropical jet as a diagnostic to measure wave propagationwithin the PWG. This diagnostic is motivated by Nigam andLindzen [1989] who showed that small shifts (<3� in lati-tude) in the location of the subtropical jet can greatly alterthe guiding of planetary waves from the troposphere into theextratropical stratosphere.[13] When measuring changes in n2s , we examine the rela-

tive importance of the strength, shear, and curvature of thezonal-mean wind. Equation (1) shows that the strength ofwave propagation within the PWG depends on the ratio of thepotential vorticity gradient to the zonal-mean wind, where thepotential vorticity gradient (2) depends on the shear andcurvature of the wind. Equation (1) also shows that as theplanetary-wave wave number s increases, n2s decreases; thisexplains why the propagation of planetary waves into thestratosphere is largely confined to planetary wave numberss = 1–3 [Charney and Drazin, 1961].[14] We measure the strength and direction of planetary

wave propagation within the PWG by using two relateddiagnostics: the Eliassen-Palm flux (EP-flux) and n2s . Forsteady, slowly varying plane waves, the planetary wavegroup velocity is locally parallel to the EP-flux vector[Edmon et al., 1980]. In addition, the EP-flux vector ⇀F iscurved up the gradient of ns and, in particular, is guidedalong ridges of ns [Palmer, 1981; Palmer, 1982; Karoly andHoskins, 1982]. The magnitude of the EP-flux vector ⇀F isrelated to n2s by [Palmer, 1981; Palmer, 1982]:

⇀F�� �� ¼ 1

2exp �z=Hð Þc2

s ns= afð Þ (3)

where⇀F is defined in (4) and (5) below;cs is the amplitude of asteady, conservative, linear wave with zonal wave number s.Regions of larger n2s are associated with larger EP-flux vectors,and the trajectories of the EP-flux vectors are refracted up thegradient of n2s . Thus, the EP-flux vector and n

2s provide a useful

way of visualizing the propagation pattern of planetary waves inthe latitude-height plane. Under quasi-geostrophic conditions,the EP-flux vector [⇀F ¼ F ’ð Þ;F zð Þ� �

] and its divergence aredefined in log-pressure coordinates as [Andrews et al., 1987]

F fð Þ ¼ �ra cosf u′v′� �

(4)

F zð Þ ¼ raf cosf v′θ′Þ= θz�

(5)

r �⇀F ¼ 1

a cosf@

@fF fð Þcosf

� þ @

@zF zð Þ

� (6)

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565

where u and v are the zonal and meridional winds and θ is thepotential temperature; the overbars and primes denote zonal-mean and perturbation quantities, respectively; and the sub-scripts in (5) denote partial differentiation. In accordance with(4) and (5), an upward directed EP-flux vector corresponds toa poleward heat flux; an equatorward directed EP-flux vectorcorresponds to a poleward momentum flux. The divergenceof the EP-flux (6) measures the net planetary wave driving ofthe zonal-mean wind. Specifically, negative values of the EP-flux divergence (i.e., convergence) exert a westward drag onthe mean wind that weakens the winter stratospheric westerlywinds. When plotting the EP-flux vectors and divergence, weadopt the scaling procedure of Butchart et al. [1982], whichensures that the relative size of the EP-flux vector componentsis preserved when plotted in Cartesian coordinates. For thisscaling, the vector components (4) and (5) are multiplied by2pa cosf and (4) is additionally multiplied by c=0.0091; theunits of the rescaled EP-flux divergence are kg s�2.[15] The relationship between n2s and the direction and

magnitude of the EP-flux has been successfully used to describea variety of stratospheric phenomenon including stratosphericsudden warmings [Butchart et al., 1982; Palmer, 1981], decadaltemperature and circulation trends [Hu et al., 2005], annularmode variability [Lorenz and Hartmann, 2002], and thepropagation of planetary waves between the troposphere andstratosphere [Chen and Robinson, 1992]. However, n2s and itsconnection to the EP-flux are built on several simplifyingassumptions that may not be valid under some circumstances.[16] For example, Harnik and Lindzen [2001] have shown

that when the middle stratosphere contains a reflecting sur-face, the traditional n2s (1) may be an unreliable measure ofthe boundaries of the PWG. This is because n2s (1) is a bulkquantity that does not distinguish between propagation in themeridional and vertical directions. In fact, vertical propaga-tion may be squelched, while n2s remains positive. This canoccur when meridional propagation is sufficiently strong.This point was underscored by Harnik and Lindzen whodefined a wave number diagnostic, which, together with alinear model, provided boundaries for meridional and verti-cal wave propagation. Using the diagnostic for specific casesduring Southern Hemisphere winter, they found that when areflecting surface formed in the upper stratosphere, it wasabout 10 km lower than that produced by the traditional n2s .[17] Because the differences between theHarnik and Lindzen

[2001] wave number diagnostic and the traditional n2s diag-nostic appear to be largest when reflecting surfaces form inthe upper stratosphere, we calculated a reflection index fol-lowing Perlwitz and Harnik [2003] for all of our 3DO3 andZMO3model simulations.We found that for the 3DO3 ensemblesimulations, 6 out of 15 Januarys had reflective surfaces inthe upper stratosphere. For the ZMO3 simulations, only 3 outof 15 winters had reflective surfaces (one in December andtwo in January). For the February simulations, there were noreflective surfaces in the upper stratosphere. For this reason,and since February is also the time when our simulationsshow that ZAO has its strongest effect on the model circu-lation (see section 3), we can expect n2s to be a representativemeasure of the morphology of the planetary waveguide.[18] Another potential concern regarding the diagnosis of

wave propagation using (1)–(6) is that the mathematicalrelationship between the EP-flux and n2s formally holds for

slowly varying plane waves and for sufficiently small wavedamping. These conditions may not be met in the realatmosphere due to wave reflection, tunneling, or damping,which may complicate the interpretation of the n2s and EP-flux relationship or invalidate it altogether [Harnik, 2002].[19] The above caveats notwithstanding, our results will

show that the relationship between the EP-flux divergence andn2s explains a significant portion of the wind and PWG vari-ability observed in the model simulations that include ZAO.

3. Results

[20] We compare the morphology of the PWG for theZMO3 and 3DO3 cases for the winter months of December,January, and February (DJF). We follow with a discussiondescribing how ZAO causes the morphology of the PWG toevolve differently during January and February.

3.1. Mean Winter Results(December–January–February)

[21] Figures 1 and 2 show, respectively, the ZMO3 and3DO3 ensemble mean zonal-mean wind and n2s for stationaryplanetary wave s = 1 averaged over DJF; similar results wereobtained for planetary wave s = 2 (not shown). Figures 1 and2 show features of n2s that are common to previous studies ofNH planetary wave propagation [Chapman and Miles, 1981;Chen and Robinson, 1992]. These features include a localminimum in n2s in the middle latitude lower stratosphere(~42�N latitude at 20 km) and a local maximum in n2s be-tween 50� and 60�N latitude and between 25 and 30 km. Thelocal minimum in n2s is due to the large negative gradient inpotential vorticity (2). The large potential vorticity gradientresults from the negative vertical and meridional gradients inzonal-mean wind located just above and to the north of thesubtropical jet (Figure 1a). For waves originating in theextratropical troposphere, the local minimum inn2s in the lowerstratosphere splits the vertical propagation of planetary wavesalong two “channels” [Chapman and Miles, 1981;Huang andGambo, 1982; Li et al., 2006]. The splitting of wave propa-gation along two wave channels is a common feature of theNH PWG during the midwinter to late winter (the time periodof interest in this study) but may be less common during theearly winter months of November–December [compare Shawet al., 2010, Figure 9 bottom left and bottom right]. We shallrefer to the two channels as the southern and northern wavechannels.[22] Along the southern channel of the PWG, the plane-

tary waves are confined to the lower to middle stratospherewhere they are refracted equatorward, eventually dissipatingbefore they reach the upper stratosphere. Along the northernchannel, the waves propagate through the local maximum inn2s associated with the weak westerly winds located slightlynorth of the region between the subtropical and polar nightjets. The region between the jets is centered at ~56�N lati-tude and 25 km in Figure 1a. From the local maximum in n2s ,waves tend to propagate upwards along the southern edge ofthe vertical axis of the polar night jet, where wave propa-gation is enhanced due to the combination of positive ver-tical and meridional shear of the zonal-mean wind. Wavesthat emanate from the sub-polar and polar regions tend tomerge into the guide formed by the northern channel. The

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EP-flux vectors shown in Figures 1b and 2b clearly showplanetary wave propagation along these two channels.[23] Planetary wave activity within the PWG is generally

strongest during midwinter to late winter, when the polarvortex is generally weaker. This is consistent with Charneyand Drazin [1961] who showed that planetary wave propa-gation from the troposphere into the stratosphere is favoredin westerly flows that are not too strong. Because NH zonalasymmetries in ozone are predominantly generated byplanetary wave activity, we also expect the ZAO field to belargest in late winter [Peters and Entzian, 1999; Peters et al.,

2008]. As winter wanes, the solar zenith angle decreases andthe ozone heating becomes stronger. Thus, we can anticipatethat the effects of ZAO, particularly with regard to the PWG,will be strongest in midwinter to late winter. This is indeedthe case, and in the discussion that follows, we focus ourattention on January and February.[24] McCormack et al. [2011], in a preliminary study

examining the effects of ZAO on the NH wintertime circu-lation, found that ZAO has the largest affect on the dynamicsof the NH polar winter stratosphere during January andFebruary. Figures 3a and 3b show the difference in the

Figure 2. As in Figure 1, but for the 3DO3 ensemble mean.

Figure 3. Monthly average zonal-mean wind difference (ΔU=3DO3 minus ZMO3) for (a) January and(b) February. Contours intervals are in units of m s�1. Statistically significant wind changes at the 95%confidence level are shaded in gray.

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ensemble mean zonal-mean wind (3DO3 minus ZMO3) forJanuary and February. Wind differences that are statisticallysignificant above the 95% confidence level are shaded.Overall, the effects of ZAO produce weaker westerly flow.During January (Figure 3a), the largest ensemble mean dif-ferences (~10m/s) are located throughout the tropical stra-topause region. During February (Figure 3b), the zonal winddifferences extend poleward and downward with the largestdifferences (~17m/s) located in the polar region between 40and 50 km. Although the largest change in the wind speedoccurs in February in the NH, changes in n2s and the EP-fluxdivergence begin in January and grow as the winter seasonproceeds. By February, there is a large change in wind speedin the extratropical stratosphere. We next consider Januaryand February in more detail.

3.2. January Results

[25] Figures 4a and 4b show the 3DO3 EP-flux divergenceand EP-flux difference (3DO3 minus ZMO3) contours forJanuary. Also shown are the corresponding EP-flux vectors.Figure 4b shows two important changes in the EP-flux dueto ZAO. First, between 10 and 15 km and between ~10� and50�N latitude, the EP-flux difference vectors are directedpoleward. This corresponds to an increase in equatorwardmomentum flux, which is associated with the weakenedextratropical upper tropospheric and lower stratosphericwesterly flow. Second, between 10 and 30 km and between~50� and 80�N latitude, the EP-flux difference vectors aredirected upwards. This represents an increase in both verticalwave propagation and the meridional heat flux, which isassociated with the weakening of the westerly flow. Thechange in the momentum flux in January continues to grow

as the winter proceeds and will be important for explainingthe equatorward shift in the subtropical jet seen later inFebruary (Figure 3b). The effect of the increased meridionalheat flux on the PWG, however, becomes readily apparentby the end of January.[26] ZAO affects the zonal-mean wind, and by extension

n2s , via two pathways [Albers and Nathan, 2012]. Along onepathway, ozone-flux convergences due to ZAO modify thezonal-mean ozone field. The change to the zonal-meanozone field produces changes in the zonal-mean radiativeheating rate and thus the zonal-mean temperature. Tomaintain thermal wind balance, the changes to the zonal-mean temperature result in changes to the zonal-mean wind.Along the other pathway, ZAO-heating affects the EP-fluxdivergence, which in turn also modulates the zonal-meanwind. However, the difference between the 3DO3 andZMO3 zonal-mean ozone fields is small between Decemberand mid-February, which means that the first pathway playsa minor role in driving the changes to n2s in the 3DO3simulations. As a result, any changes in the zonal-meanwind and n2s during this time period must be due to changesin the EP-flux divergence associated with ZAO-heating.[27] Figure 4b shows that the increase in the EP-flux di-

vergence within the northern channel of the PWG is pri-marily due to an increase in the vertical component of theEP-flux vector (this is confirmed by the near verticality ofthe EP-flux difference vectors in this region). To confirmthat changes in the meridional heat flux precede changes inn2s within the northern channel of the PWG, we next comparethe time evolution of the meridional heat flux and n2s .[28] Figure 5 shows the time series of the ensemble mean

and area-weighted average of the meridional heat flux and n2s.

Figure 4. January ensemble mean for (a) 3DO3 EP-flux convergence with EP-flux vectors and (b) thedifference (3DO3 minus ZMO3) in the EP-flux convergence and EP-flux vectors. In Figure 4a, the whiteregion between 0� and 40�N latitude and below ~15 km represents EP-flux convergence that is greaterthan the 0 to 300 scale used for plotting. In both plots, the units are 1� 106 kg s�2 and the contour intervalis 5� 106 kg s�2. As in Figures 1 and 2, the EP-flux vectors in Figure 4b are only plotted in regions wherethe EP-flux divergence is less than the threshold approximately �25� 106 kg s�2. The scaling of the EP-flux vectors is discussed in section 2; the magnitude of the vectors in Figures 4a and 4b are consistent withthe EP-flux convergence scale shown alongside each plot.

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The time series spans December 1 to February 28, and thearea-weighted average is between 50� and 75�N at ~27 km inheight. This height and latitude band lies roughly in thecenter of the northern channel of the PWG. During the timeperiod considered, the heat flux and n2s exhibit qualitativelydifferent behaviors before and after 5 January.[29] Prior to 5 January, the difference between the 3DO3

and ZMO3 heat fluxes (Figure 5a) does not produce a sig-nificance difference in the respective values ofn2s (Figure 5b).In contrast, after 5 January, the heat flux of 3DO3 remainslarger than the heat flux of ZMO3 (the exception is a small2-day period just after 10 January). As a result, n2s for the3DO3 ensembles also become larger than n2s for the ZMO3ensembles during mid to late January. Yet, there are twoimportant differences between the temporal evolution of then2s versus the heat flux. First, the increase in the 3DO3 n

2s lags

the increase in the 3DO3 heat flux by ~5 days. Second, theincrease in the 3DO3n2s remains rather small until the final 5–10 days of January. The time lag between the change in theheat flux and the change in n2s is not unexpected because ittakes time for the ZAO-induced changes in the planetarywaves to be manifest as changes in the zonal-mean wind.Once the ZAO-heating triggers the initial increase in themeridional heat flux, however, a positive feedback cycle isinitiated that ultimately affects the PWG and any subsequentplanetary wave activity. In short, the initial increase in theEP-flux convergence due to ZAO-heating causes changesin the strength of the extratropical westerly flow and themeridional potential vorticity gradient that combine toincrease n2s . The increase in n2s , in turn, causes a furtherincrease in vertical wave propagation that reinforces theinitial increase in the EP-flux convergence. To gain furtherinsight into the increase in vertical wave propagation, wenext consider the spatial evolution of the vertical compo-nent of the EP-flux (F(z)) and the PWG.[30] Figure 6a shows contours of the January ensemble

mean difference (3DO3 minus ZMO3) in F(z) (5). Theincrease in F(z) is primarily concentrated in two plumeslocated within the northern and southern wave channels. The

primary plume originates in the sub-polar and polar latitudescentered at ~60�N latitude, while the second, smaller plumeis centered at ~30�N latitude. The two plumes and the gapbetween them are representative of three features of thePWG discussed in section 3.1; that is, the plumes representan increase in vertical wave propagation along both thenorthern and southern planetary wave channels, while thegap between the plumes of F(z) is likely due to the localminimum in n2s seen in Figures 1b and 2b at ~42�N latitudeand 20 km.[31] To examine the connection between the increase in

F(z) along each of the planetary wave channels and changesin the morphology of the PWG, we computed the change inn2s (3DO3 minus ZMO3) for two periods during January.Guided by the n2s time series in Figure 5, we calculated thechange in n2s between 1 and 25 January and between 25 and31 January. Between 1 and 25 January, there is very littlechange in n2s (not shown); this result is consistent with therelatively small change in the January monthly mean zonal-mean wind shown in Figure 3a and the n2s time series shownin Figure 5b. The final 5 days of January, however, showchanges in n2s that match very well with the increase invertical wave propagation. In particular, the two plumes inF(z) shown in Figure 6a are collocated with notable increasesin n2s along both planetary wave channels (Figure 6b).[32] We next examine how the ZAO-related changes in

the PWG that first appear in late January expand and amplifyto produce the large change in the zonal wind in February.

3.3. February Results

[33] Figure 3b shows that ZAO produces statistically sig-nificant decreases in the zonal-mean westerly flow of 5–11m s�1 in the equatorial upper stratosphere and middlemesosphere, as well as decreases of 3–18m s�1 throughoutthe NH extratropical stratosphere and mesosphere duringFebruary. A third, smaller region of statistically significantwind difference is located near the location of the subtropi-cal jet at ~30�N in the upper troposphere and lower strato-sphere (UTLS); the wind change at this location represents

Figure 5. Time series of ensemble mean, area-weighted average (50�–75�N) at ~27 km for the (a) merid-ional heat flux and (b)n2s between December–February for 3DO3 (solid lines) and ZMO3 (dotted lines). Theheat flux is in units of Km s�1, while n2s has been non-dimensionalized by a2.

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an increase of ~1–3m s�1. The differences in the NH zonal-mean wind show three notable results. First, the decrease inthe wind speed in the tropical upper stratosphere and lowermesosphere causes a northward migration of the tropicaleasterlies and a northward shift of the zero wind line. AsFigure 7 shows, the northward shift in the zero wind linevaries from about 1� to 8� latitude between about 30 km and65 km in height. Second, the decrease in wind speed withinthe extratropics and polar region represents a weaker polarvortex and a warmer stratosphere (see also McCormacket al. [2011, Figure 2]. Third, the increased wind speed inthe subtropical UTLS shifts the subtropical jet equatorwardby about 3�. The zonal wind differences in Figures 3 and 7qualitatively agree with recent studies suggesting a connec-tion between northward shifts in the upper stratospheric zerowind line and an increase in the frequency of SSWs. Forexample, Gray [2003] found that similar to the so-called“Holton-Tan” mechanism in the lower stratosphere [Holtonand Tan, 1982], a northward shift in the upper stratosphericzero wind line produces a narrower PWG that directs moreplanetary wave activity poleward. The increase in planetarywave activity induces more wave drag on the zonal-meancirculation and creates a warmer extratropical and polarstratosphere and an increased incidence of SSWs. Indeed, asour EP-flux diagnostics will show, the combination of anarrower PWG and an increase in wave propagation com-bine to produce an increase in wave drag, the weaker polarvortex shown in Figures 1a and 2a, and an increased inci-dence of major SSWs (four in the 3DO3 ensembles versusone in the ZMO3 ensembles).[34] Figures 8a and 8b compare values of the February

ensemble mean n2s (1) for ZMO3 and 3DO3 for stationaryplanetary wave s = 1. The n2s scale varies from red to blue,which corresponds with large to small n2s , respectively.

Regions where planetary waves are evanescent ( n2s < 0)are denoted by white space. In addition to showing n2s ,Figures 8a and 8b are overlaid with a qualitative outline ofthe PWG where values at either extreme of the n2s scale areexcluded and shaded by the transparent light gray color;the solid black line that outlines the PWG traces the 15 and65 n2s contours on the poleward and equatorward side ofthe PWG, respectively. Outlining the PWG this way greatlyaids in visualizing changes to n2s ; moreover, excluding theextreme values of n2s is a physically reasonable approxima-tion for two reasons. First, approaching the subtropical zero

Figure 6. (a) Contours of the difference (3DO3 minus ZMO3) in ensemble mean vertical component ofthe EP-flux vector for January. Units are 1� 10�9 kg�ms�2 and the contour interval is 5� 10�8 kg�ms�2.(b) Difference in the late January n2s for the period between 25 and 31 January. For visual clarity, the n2sdata were filtered so that the n2s is plotted only in regions where the EP-flux divergence is less than thesame minimum threshold value (approximately �25� 106 kg s�2) used when plotting the EP-flux vectorsin Figures 1b and 2b; see Figure 4a for EP-flux divergence scale.

Figure 7. Location of February Northern Hemisphere zerowind line for 3DO3 (solid line) and ZMO3 (dashed line).

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wind line, n2s !1 as ū! 0. In this limit, n2s becomes anunreliable measure of wave propagation. Second, becauseplanetary waves tend to propagate towards and through largevalues of n2s , changes in n2s within the region of very smalln2s on the poleward side of the PWG play a minor role indetermining where planetary waves propagate. Regardless,choosing different extreme values to exclude from thePWG outline does not change the qualitative nature of ourresults.[35] Figure 8 shows that ZAO increases planetary wave

propagation in three important ways. First, the vertical extentof the PWG increases, which allows the waves to propagate10 to 15 km higher within the mesosphere between 45� and70�N. Second, ZAO causes an increase in the magnitude ofn2s throughout the entire stratosphere and, in particular, anincrease in wave propagation along both the southern andnorthern wave channels. The increase in n2s along the south-ern wave channel is significant beginning at the surfaceand extending upwards to ~35 km between 30� and 40�N.The increase in n2s within the northern wave channel extendsupwards from ~20 km all the way to into the mesosphere near~75 km. And third, ZAO also causes a significant increase inn2s in the polar UTLS between ~60� and 80�N. The increasesin wave propagation described above are each due to differ-ent structural changes in the wind and are embodied bycompeting terms in n2s (1), namely, the strength of the zonal-mean wind versus the meridional potential vorticity gradient.[36] According to the first term on the right-hand side

(RHS) of (1), weaker westerly winds and increases in themeridional potential vorticity gradient together increase n2s .Ozone-induced changes in the potential vorticity gradient, in

turn, are due to changes in the meridional and vertical shearterms of the zonal-mean wind [the second and third terms onthe RHS of (2)]. We compared each of the terms in (2) forthe 3DO3 and ZMO3 sets of ensembles and found thatnearly all of the changes in n2s due to changes in the potentialvorticity gradient are due to the third term on the RHS of (2).Further details can then be exposed by expanding the thirdterm on the RHS of (2), which yields

� f 2

rsrs

�uzN2

� �z

¼ f 2

HN2þ f 2

N4

dN 2

dz

� ��uz � f 2

N2�uzz: (7)

[37] The first term on the RHS of (7) is proportional to thevertical shear of the mean wind, while the second term isproportional to the vertical curvature of the mean wind. Thus,an increase in positive vertical wind shear (or a decrease innegative vertical wind shear) increases the meridionalpotential vorticity gradient (2) and increases n2s (1). Based ona numerical comparison of the two terms on the RHS of (7),we found that the vertical shear term is the dominant term andlargely accounts for changes between the 3DO3 and ZMO3ensembles. Thus, the changes in n2s shown in Figure 8 can beexplained by the ratio ūz/ū.[38] With the exception of the UTLS region along the

southern wave channel, the entire extratropical stratosphereand lower mesosphere is characterized by weakened west-erly winds (Figure 3b), consistent with an increase in n2swithin the PWG shown in Figure 8. The weakened polarvortex in the 3DO3 simulations, however, also affects thevertical wind shear. Figure 2a shows that the location of themaximum wind speed in the vortex core extends from~60 km at ~40�N to ~40 km at ~60�N, so that much of the

Figure 8. Monthly averaged n2s for February: (a) ZMO3 for planetary wave s= 1 and (b) 3DO3 for plan-etary wave s= 1. The n2s scale extends from red to blue, which corresponds with large to small n2s , respec-tively; regions where planetary waves are evanescent (n2s < 0) are denoted by white space. We indicate theshape of the PWG by choosing a specific, though arbitrary, contour for n2s. Irrespective of the choice of thecontour, however, the shape of the PWG remains qualitatively unchanged. The solid black line that out-lines the PWG traces the 15 and 65 n2s contours on the poleward and equatorward side of the PWG, respec-tively. In all plots, n2s has been non-dimensionalized by a2; the contour interval is 4.

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region above ~40–50 km is characterized by negative verti-cal wind shear, while the region below ~40 km is largelycharacterized by positive vertical wind shear. These featuresare clearly seen in Figure 9a, which depicts the change in thevertical wind shear between the 3DO3 and ZMO3 ensem-bles. As a consequence of the weaker polar vortex in the3DO3 simulations, the vertical wind shear along the upperportion of the northern wave channel (between 40� and70�N and above ~40 km) and along the southern wavechannel (between 30� 40�N in the UTLS) becomes weaker(less negative); this contributes positively to the increase inn2s . In contrast, the region along the lower portion of thenorthern wave channel (between ~55� and 85�N and between10 and 40 km) is characterized by a decrease in positivevertical wind shear that contributes negatively to n2s ; thus, theincrease in n2s in the this region is due solely to the weakenedwesterly winds in the 3DO3 simulations. Each of the changesin the magnitude and vertical shear of the zonal-mean windjust described will prove to be important for explaining theincreases in the February EP-flux convergence as discussedin the next section.

3.4. Discussion

[39] We have shown that ZAO-heating causes an increasein the EP-flux beginning in early to mid January. This initialincrease in the EP-flux produces changes in the zonal-meanwind that are manifest as changes in the PWG. The changesin the PWG occur via variations in the location of the sub-tropical jet, shifts in the latitude of the subtropical zero windline, and the strength of wave propagation within the extra-tropical stratosphere. These changes in the EP-flux andPWG represent a positive feedback, where the initial ZAO-induced changes in wave amplification/damping triggerchanges in the PWG that affect subsequent planetary wave

activity entering the stratosphere. As the winter season pro-gresses, the ZAO-induced changes in the PWG becomeincreasingly important.[40] To better understand how ZAO affects the PWG

during late January and February, we consider how theZAO-weakened lower stratospheric zonal winds combinewith the equatorward shift in the subtropical jet. First, wepresent some background.[41] Using a primitive equation model, Nigam and Lindzen

[1989] found that small perturbations (<5m s�1) to thezonal-mean wind could produce equatorward shifts in thelocation of the subtropical jet. Although small, these shiftscould nonetheless substantially enhance the guiding of plane-tary waves from their primary source region (the Himalayas at~27�N) into the polar stratosphere. Similarly, Hu and Tung[2002] showed that the low index phase of the NorthernHemispheric annular mode (which is characterized by anequatorward shift in the NH zonal jet) is associated withweakened zonal-mean winds and increased n2s in the UTLSregion between 60� and 80�N. In addition to changes in n2s ,Limpasuvan and Hartmann [2000] showed that the low indexphase of the Northern annular mode is also characterized byanomalous increases in poleward directed EP-flux vectors; theresulting increase in equatorward momentum fluxes are asso-ciated with a weakened stratospheric westerly winds. Ourresults suggest that ZAO modulates n2s and the EP-flux in asimilar fashion to each of the phenomenon described byNigam and Lindzen, Hu and Tung, and Limpasuvan andHartmann. Indeed, the equatorward shift in the subtropicaljet in the 3DO3 simulations is accompanied by an anomalousincrease in upward directed EP-flux vectors into the strato-sphere and an anomalous increase in poleward directed EP-flux vectors in the UTLS between ~30� and 75�N (Figure 9b),weakened westerly winds (Figure 3b), and increased n2s values

Figure 9. (a) The difference (3DO3 minus ZMO3) in the ensemble mean vertical wind shear for Febru-ary; the contour interval is in units of m s�2. (b) The difference (3DO3 minus ZMO3) in the EP-flux vec-tors for February; the ensemble mean February 3DO3 n2s is plotted in the background for reference. As inFigures 1 and 2, the EP-flux vectors in Figure 9b were filtered so that vectors are only plotted in regionswhere the EP-flux divergence is less than approximately �25� 106 kg s�2. The scaling of the EP-flux vec-tors is discussed in section 2.

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(Figure 8b) throughout the extratropical UTLS. The increasein the anomalous poleward and upward directed EP-flux vec-tors observed in February reinforce the positive feedback cyclebetween the EP-flux convergence and n2s first established inlate January; the net result is an increase in wave propagationthroughout both channels of the PWG. To better understandthe implications of the expansion of the PWG in the verticaland the increase in wave propagation within the PWG,we nextconsider the EP-flux convergence.[42] Figures 10a shows contours of the difference (3DO3

minus ZMO3) of the EP-flux convergence in February. In acontinuation of the EP-flux convergence differences that firstappeared in January, the February difference clearly showstwo plumes of increased convergence along both the north-ern and southern wave channels. There is a 30–50% increasein the EP-flux convergence along the northern wave channelbetween 15 and 20 km, and a 15–100% increase along thesouthern wave channel at the same heights. As in January,the two plumes of increased EP-flux convergence corre-spond with significant increases in vertical wave propaga-tion; this can be verified by examining the increased EP-fluxvectors in Figure 9b and, more specifically, the increase inF(z) shown in Figure 10b. In addition to an increase inmagnitude, the ZAO-induced increase in F(z) during Febru-ary also extends ~10–15 km higher than was observed dur-ing January (compare Figures 6a and 10b); this is consistentwith the 10–15 km increase in height of the February PWGdepicted in Figure 8.[43] The increases in vertical wave propagation along the

northern and southern wave channels, however, correspondto the qualitatively different changes in wind structure de-scribed in section 3.3. The increase in vertical wave propa-gation along the southern wave channel (~30�–40�N) in theUTLS is due to a combination of changes in the verticalshear, meridional shear, and curvature of the zonal-meanwind rather than changes in the magnitude of the wind. Thisis verified in Figure 3b, which shows that the UTLS ischaracterized by an increase in the mean wind that con-tributes negatively to the magnitude of n2s . The dominant

contribution to the increase in n2s and the increased verticalpropagation along the southern wave channel is due to theincrease in positive vertical wind shear depicted in Figure 9abetween ~30� and 40�N. The meridional shear and curvatureterms (not shown) also contribute to the increase in n2s , buttheir contribution is secondary to that of the vertical shearterm. The importance of vertical wind shear along the tro-popause is perhaps not surprising in light of the study byChen and Robinson [1992] who showed that even smallchanges in vertical wind shear near the tropopause mayproduce significant changes (20–40%) in n2s and EP-fluxconvergence in the UTLS. In contrast to the changes alongthe southern wave channel, the increase in vertical wavepropagation along the northern wave channel are due toincreases in n2s associated solely with the weakened westerlywinds; this is verified by noting the decrease in positivevertical wind shear (Figure 9a) that contributes negativelyto n2s .[44] Despite the similarities between the distributions in

the January and February EP-flux convergences (compareFigures 4b and 10a), there are two notable changes betweenthe two months. First, the local minimum in the EP-flux con-vergence along the northern wave channel (centered at ~60�Nand 18 km) has increased slightly from �36 to �32 kg s�2,while the local minimum along the southern wave channel(centered at ~32�N and 16 km) has decreased by ~60% from�27 to�43 kg s�2. This indicates that the local increase in n2sand the equatorward shift in the subtropical jet have combinedto significantly increase the poleward and vertical propagationof planetary waves (Figures 9b and 10b), which results in thelarge increase in the EP-flux convergence along the southernwave channel. Second, while the increased EP-flux conver-gence due to ZAO extended all the way from the middlestratosphere down to the surface between 50� and 70�N duringJanuary (Figure 4b), there is no increase in convergencebetween ~10 and 15 km and between 30� and 70�N duringFebruary (Figure 10a); this difference cannot be explained bythe changes in n2s described earlier. There are several possible

Figure 10. (a) February ensemble mean difference (3DO3 minus ZMO3) in the EP-flux convergence.The units are 1� 106 kg s�2, and the contour interval is 5� 106 kg s�2. (b) Contours of the difference(3DO3 minus ZMO3) in ensemble mean vertical component of the EP-flux vector for February. The unitsare 1� 10�9 kg�ms�2, and the contour interval is 5� 10�8 kg�ms�2.

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explanations for this apparent discrepancy between the increasein the EP-flux convergence and n2s .[45] One possible explanation for the increase in the EP-

flux convergence in the UTLS is associated with the abilityof ZAO to alter the EP-flux convergence through waveamplification/damping. Indeed, it is ZAO-heating that trig-gers the initial increase in the EP-flux convergence and thePWG in mid to late January described in section 3.2. Theozone-modified refractive index derived in Nathan andCordero [2007] describes how ZAO and zonal-mean ozonecombine to alter both planetary wave propagation and wavedamping. Albers and Nathan [2012] subsequently used theozone-modified refractive index to show how changes in wavepropagation and damping due to ZAO act together or in op-position to produce small or large changes in the EP-flux in thelower stratosphere. Although the current experimental setupdoes not make it possible to separate ZAO-related effects dueto wave propagation and wave damping, it is indeed possiblethat the apparent discrepancy between the increase in n2s seenin Figure 8 and decrease in the EP-flux convergence seenbetween ~5 and 15 km and between 30� and 70�N inFigure 10 could be related to ZAO-induced changes in wavedamping and wave propagation that are not taken intoaccount by the inviscid form of n2s (1) employed in this study.This issue will be the subject of future modeling and obser-vational studies.[46] A second possible explanation for the apparent dis-

crepancy between the increase in the EP-flux convergenceand n2s was discussed in section 2. That is, wave tunneling orwave reflection, separately or in combination, may invali-date the relationship between the EP-flux and n2s employedhere [Harnik, 2002]. Yet, a third possibility is that the ZAO-induced changes in the stratospheric circulation have causedchanges in the troposphere that increase the generation ofplanetary wave activity. Indeed, our results show that ZAOaffects the circulation of the troposphere (e.g., Figures 3 and10a). Thus, it is possible that some of the increase in the EP-flux convergence in the UTLS is due to an increase inplanetary wave generation in the troposphere.

4. Summary and Conclusions

[47] Using a middle atmosphere general circulationmodel, we have shown that zonally asymmetric ozone(ZAO) alters the morphology of the PWG in three distinctways. First, the PWG contracts meridionally and expandsvertically within the upper stratosphere and lower meso-sphere. Second, the magnitude of wave propagation increa-ses throughout the interior of the stratospheric waveguidebetween ~30� and 75�N. And third, the subtropical jet shiftsequatorward. In combination, the changes in the shape andstrength of the PWG are associated with increased deceler-ation of the zonal-mean westerly flow, resulting in a weakerand warmer stratospheric polar vortex. These changes cor-relate well with the increased incidence of SSW’s in thepresence of ZAO reported by McCormack et al. [2011].[48] Changes in the PWG have important implications for

modeling SSWs, variations in stratospheric dynamics asso-ciated with the 11-year solar cycle, and the NorthernHemisphere annular mode (NAM). For example, Charltonet al. [2007] investigated the ability of a series of GCMs—

which did not consider the effects of ZAO on the PWG—togenerate an accurate number of SSWs. The results showedthat most GCMs failed to produce enough SSWs whencompared to observations. In the current GCM study—where we explicitly consider the effects of ZAO on thePWG—the model runs with ZAO had a noticeably largerEP-flux emanating from the upper troposophere and lowerstratosphere and produced a larger number of SSWs com-pared to similar runs without ZAO. This increase in the EP-flux correlates well with the increase in n2s in the lowerstratosphere, indicating that changes in the PWG associatedwith ZAO may play an important role in modulating the fluxof wave activity entering the stratosphere from the tropo-sphere, which is known to drive SSWs.[49] SSWs have also been correlated with a northward

shift of tropical easterlies and the zero wind line within theupper stratosphere and lower mesosphere [Gray, 2003;Vineeth et al., 2010]. Gray [2003] hypothesized that becausevertically propagating planetary waves have vertical wave-lengths comparable to the distance between the tropopauseand the upper stratosphere during midwinter to late winter,it is possible that shifts of the zero wind line in the upperstratosphere help guide waves poleward in much the samemanner as the traditional Holton-Tan mechanism in thelower stratosphere. Although our experiments were notdesigned to test Gray et al.’s hypothesis, it is nonethelessinteresting to note the correlation between the northwardshift in the zero wind line and the increase in frequency ofSSWs in our model runs that include ZAO.[50] In addition to SSWs, two aspects of the PWG dis-

cussed in this study—the width and strength—may provide ameans for ZAO to communicate and amplify the 11-yearsolar cycle signal. The relationship between the width of thePWG and solar-modulated ZAO has been shown by Corderoand Nathan [2005] to be an important pathway for commu-nicating the solar signal to the quasi-biennial oscillation (seetheir Figure 1). Because our results show that ZAO plays arole in modulating the upper stratospheric tropical easterlies,it is therefore likely that in a similar manner to the Corderoand Nathan solar-quasi-biennial oscillation mechanism,ZAO may also play a role in modulating the semi-annualoscillation in the tropical upper stratosphere. Because thesemi-annual oscillation plays a dominant role in the locationand timing of shifts in the upper stratospheric zero wind line,modulation of the semi-annual oscillation by ZAO mayprovide an important pathway for solar modulation of thewidth of the PWG. Indeed, this is a topic that deserves furtherattention given the parallel timing of the peak in wind mod-ulation from ZAO in our results and the robust solar signalobserved during midwinter to late winter in the tropical upperstratosphere [e.g., Labitzke and van Loon, 1988; Naito andHirota, 1997]. In addition to solar modulation of the width ofthe waveguide, ZAO also provides a way to change thestrength of wave propagation within the waveguide itself.This scenario is supported by work demonstrating the linkbetween the 11-year solar cycle and changes in wave prop-agation and planetary wave drag in the extratropical strato-sphere [Nathan et al., 2011; Gabriel et al., 2011].[51] Finally, we note that the analysis in this paper has

examined the effects of ZAO from a “bottom-up” perspec-tive. From this perspective, ZAO enhances the amount of

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planetary wave activity that is refracted into the interior ofthe stratosphere from the troposphere below. This phenom-enon can be understood as a positive feedback cycle whereincreases in the EP-flux convergence cause wind changesthat increase n2s within the PWG; the increase in n2s within thePWG in turn leads to further increases in the vertical propa-gation of planetary waves with subsequent increases in theEP-flux convergence. However, changes in the upward fluxof planetary wave activity and consequent variations in localwave-mean flow interaction in the stratosphere can also drivedownward propagating zonal-mean wind anomalies [Plumband Semeniuk, 2003], a “top-down” effect. Such downwardpropagating wind anomalies have been shown to be impor-tant for stratosphere-troposphere communication [Koderaand Kuroda, 2000; Christiansen, 2001; Polvani and Waugh,2004; Perlwitz and Harnik, 2004] and NAM variability [e.g.,Baldwin and Dunkerton, 1999; Limpasuvan and Hartmann,2000]. The model simulations examined in this paper provideevidence that ZAO affects both the downward propagation ofwind anomalies and the phase of the NAM. The downwardpropagation of the wind anomalies is seen in the space-timeevolution of the ZAO-induced wind anomalies depicted inFigure 3 and inMcCormack et al. [2011, Figures 4a–4d]. Theeffect of ZAO-heating on the phase of the NAM is seen in theequatorward shift in the subtropical jet and a weakened polarvortex (compare Figures 3 and 7), which are both commonfeatures of the low index phase of the NAM. The equator-ward shift in the subtropical jet shown in our results is con-sistent with the work of Brand et al. [2008] who found thatincluding ZAO produces a shift in the NAM towards its lowindex phase; this indicates that the relationship between ZAOand shifts in the phase of the NAM is likely a robust result.We are currently investigating the effects of ZAO on down-ward signal propagation and annular mode variability.

[52] Acknowledgments. We thank Dr. Nili Harnik and two anony-mous reviewers for providing helpful comments on the manuscript. JRAand TRN were supported in part by NSF grant ATM-0733698 and byNASA/NRL grant NNH08AI67I. JPM was supported in part by the Officeof Naval Research and in part by NASA Heliophysics Living with a StarTR&T Program award NNH08AI67I.

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