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In this chapter we have investigated the evolution of barotropic instability within a critical layer. Our work differs from that of SWW (1978) and Haynes (1989) in that we force the streamfunction in the critical layer directly (rather than force laterally) using the SWW solution as a model for the critical layer. Another difference is that our numerical model is a dimensional model rather than a non-dimensional model used in Haynes (1989).

We examined the effect of the shear across the critical layer on the critical layer evolution, and more specifically on the development of barotropic instability in the critical layer. We found (consistent with Haynes (1989)) that barotropic instability occurred when the shear was suitably small, Λ ≤ 0.3. As the shear across the critical layer decreased the along-channel wavelength of the instability decreased (wavenumber increased). The mixing efficiency of the critical layer increased due to barotropic instability. We also found that barotropic instability

enhances mixing in the critical layer at early times when the shear is small. However this enhancement is much smaller than the enhancement of mixing due to the resonant growth of the critical layer around Λ = 1₂.

In an extension to previous work we carried out a systematic investigation of the critical layer evolution at finite Rossby deformation length. We set up this investigation so that as far as possible everything was the same as the infinite Rossby deformation length investigation, i.e. topographic forcing and average shear in the critical layer region (same closed streamline pattern). We found that regardless of this the critical layer width increased at small LD. This was

due to the effect of LD on the Rossby wave elasticity. We therefore chose to

estimate the growth of the critical layer width with decreasing LD. We adjusted

the topographic forcing amplitude using a scaling factor from the Rossby wave dispersion relation. The scaling factor chosen was close but overcompensated for the increase in the critical layer width. The scaling factor allowed a study of the effect of decreasingLD on the barotropic instability. We discovered that at small LD, LD less than the natural wavelength of the instability, the wavelength of the

Chapter 3

The Influence of Stratospheric

Potential Vorticity on Baroclinic

Instability

3.1

Introduction

Wavebreaking in the stratosphere drives the Brewer-Dobson circulation and affects the mixing and transport of chemicals out of the tropical pipe (see chapter 5 for more details). In this chapter we discuss how this stratospheric wavebreaking affects the troposphere and mixing at the tropopause and the Earth’s surface.

In recent years there has been much interest in the idea of dynamical cou- pling between the stratosphere and the troposphere. It is widely known that wave motions in the troposphere affect the stratospheric circulation, for example, plan- etary waves which propagate up into the stratosphere and lead to stratospheric * This chapter is based on: Smy, L. A. and Scott, R. K. 2009. The influence of stratospheric potential vorticity on baroclinic instability. Q. J. R. Meteorol. Soc., 135, 1673-1683

sudden warmings. However recent observations now suggest that the stratosphere may have an important influence on tropospheric weather and climate.

Observations of correlations between zonally symmetric anomalies of zonal wind and geopotential height in the stratosphere and troposphere (e.g. Kodera

et al., 1990; Thompson & Wallace, 1998, 2000; Baldwin & Dunkerton, 1999) have prompted this recent research into the dynamical coupling between these two regions of the atmosphere. These correlations are time-lagged and show tropospheric anomalies that persist on sub-seasonal timescales, longer than the corresponding stratospheric anomalies. Consequently, it has been suggested that the stratosphere may have a dynamical influence on the tropospheric circulation, even to the extent that medium-range weather forecasts might be improved by improving the representation of the stratosphere in forecast models (e.g. Scaife & Knight, 2008).

Dynamical links between the stratosphere and troposphere have been sug- gested to exist on both sub-seasonal and longer, inter-annual timescales. In the latter case, for example, the stratosphere has been shown to influence tropo- spheric circulation patterns in comprehensive general circulation models. Hart- mannet al.(2000) investigated stratospheric ozone depletion as a possible cause for the trends observed, since the 1970’s, in the stratospheric and tropospheric annular modes. The effect of stratospheric ozone on solar cycle irradiance and its resulting affect on climate change is examined by Shindellet al.(1999, 2001). This influence of the stratosphere on the troposphere has also been shown in sim- plified primitive equation models (Polvani & Kushner, 2002; Kushner & Polvani, 2004), where thermal perturbations in the stratosphere were found to signifi- cantly affect the tropospheric circulation. This indicates the possible sensitivity of the tropospheric circulation to the details of the stratospheric evolution. The stratosphere is now widely believed to play an important role in climate variabil- ity (e.g. WMO, 2007), although the dynamical processes involved are not well

understood.

On shorter timescales, on the other hand, the strong lag-correlations between the strength of the winter stratospheric polar vortex and sea-level pressure dis- tribution (Baldwin & Dunkerton, 1999, 2001; Thompson et al., 2002; Charlton

et al., 2003) again point to a dynamical coupling that remains, however, far from well-understood. While descending zonal wind anomalies within the stratosphere can be explained in terms of descending critical layers (Matsuno, 1971) , the per- sistence of tropospheric anomalies on longer timescales requires consideration of additional dynamical processes. For example, Song & Robinson (2004) have sug- gested that stratospheric sudden warmings might couple with the troposphere through an eddy-feedback mechanism. Another recent study by Thompsonet al.

(2006) suggested, on the other hand, that the purely balanced response of the troposphere to changes in stratospheric wave drag and thermal heating may be sufficient to explain longer tropospheric correlation timescales.

The main driver of surface weather systems in the troposphere is baroclinic instability. A dynamical link between the stratosphere and troposphere involving the modulation of baroclinic instability by stratospheric zonal wind anomalies has also been considered recently (Wittman et al., 2004, 2007) . In simple baroclinic lifecycle experiments using a general circulation model dynamical core, Wittman

et al. (2004) found that the strength of the stratospheric zonal winds had an influence on the tropospheric evolution, both in terms of the synoptic scale de- velopment and in zonal mean quantities, such as surface geopotential. Building on this and earlier work by Muller (1991) that examined linear growth rates in a one-dimensional Eady-type model, Wittman et al. (2007) further examined the dependence of growth rates on stratospheric shear in a variety of simple and more comprehensive models. In particular, they found that increasing the verti- cal shear above the tropopause (a representation of a strong stratospheric vortex) increased growth rates across a range of zonal wavenumbers.

In this chapter, we again investigate the dynamical link between stratospheric anomalies and baroclinic instability, but here restrict attention to anomalies that consist purely of a change to the stratospheric potential vorticity. This approach is motivated by a recognition of potential vorticity as the principle dynamical quantity governing the slow, balanced motion of the stratosphere. A strong polar vortex is an approximately zonally symmetric distribution of high potential vor- ticity, while a weak polar vortex may be due either to a distribution of low poten- tial vorticity, or, alternatively, to a zonally asymmetric distribution of potential vorticity. The latter scenario is what is typically observed during major strato- spheric sudden warmings, when strong planetary wave breaking redistributes the stratospheric potential vorticity, either through a strong displacement of the vor- tex (in the case of a wave-1 warming) or through a vortex split (in the case of a wave-2 warming; e.g. Charlton & Polvani, 2007). The location of the polar vortex affects the horizontal shear around the tropopause and at the surface (see figure 3.1). Therefore we expect changes to the location of the polar vortex to change the critical layer dynamics at the tropopause and at the surface. The ap- proach adopted here allows us to examine in detail how such a redistribution of the stratospheric potential vorticity may effect the tropospheric evolution. Note also that we are interested in resultant changes to the nonlinear tropospheric evolution during the course of a baroclinic lifecycle rather than in changes to the tropospheric circulation resulting from an instantaneous potential vorticity inversion. The effects of zonal mean perturbations to the stratospheric potential vorticity on the tropospheric circulation as a direct result of potential vorticity inversion were considered by Ambaum & Hoskins (2002); Black (2002), but not their influence on baroclinic instability.

For simplicity, and to avoid the difficulties involved in balancing an asym- metric potential vorticity distribution, we use a quasigeostrophic model on an

instability is arranged in the troposphere through an Eady-type distribution of potential temperature at the surface and tropopause. The case in which the stratospheric potential vorticity is exactly zero is treated as a control (discussed in section 3.3.1). Following this we consider the effect on the instability of zonally symmetric perturbations to the stratospheric potential vorticity (as a represen- tation of a strong vortex, section 3.3.2) and asymmetric perturbations (as a rep- resentation of the vortex following a sudden warming, section 3.3.3). Finally, we briefly consider the sensitivity of our results to details of the tropospheric mean state (section 3.3.4)