Search Results
Abstract
Ensemble experiments of decaying shallow-water turbulence on a rotating sphere are performed to confirm the robustness of the emergence of an equatorial jet. While previous studies have reported that the equatorial jets emerging in shallow-water turbulence are always retrograde, predominance of a prograde jet, although less likely, was also found in the present ensemble experiments. Furthermore, a zonal-mean flow induced by wave–wave interactions was examined using a weak nonlinear model to investigate the acceleration mechanisms of the equatorial jet. The second-order acceleration is induced by the Rossby and mixed Rossby–gravity waves and its mechanisms can be categorized into two types. First, the local meridional wavenumber of a Rossby wave packet propagating toward the equator increases because of meridional variation of the Rossby deformation radius and/or the retrograde zonal-mean flow, resulting in a dissipation of the wave packet in the equatorial region. This mechanism always contributes to retrograde acceleration of an equatorial jet. Another mechanism is derived from the tilting of equatorial waves due to meridional shear of the zonal-mean flow. In this case, zonal-mean flow acceleration contributes to the intensification of a given basic flow.
Abstract
Ensemble experiments of decaying shallow-water turbulence on a rotating sphere are performed to confirm the robustness of the emergence of an equatorial jet. While previous studies have reported that the equatorial jets emerging in shallow-water turbulence are always retrograde, predominance of a prograde jet, although less likely, was also found in the present ensemble experiments. Furthermore, a zonal-mean flow induced by wave–wave interactions was examined using a weak nonlinear model to investigate the acceleration mechanisms of the equatorial jet. The second-order acceleration is induced by the Rossby and mixed Rossby–gravity waves and its mechanisms can be categorized into two types. First, the local meridional wavenumber of a Rossby wave packet propagating toward the equator increases because of meridional variation of the Rossby deformation radius and/or the retrograde zonal-mean flow, resulting in a dissipation of the wave packet in the equatorial region. This mechanism always contributes to retrograde acceleration of an equatorial jet. Another mechanism is derived from the tilting of equatorial waves due to meridional shear of the zonal-mean flow. In this case, zonal-mean flow acceleration contributes to the intensification of a given basic flow.
SST; Kirtman and Schneider (2000) and Barsugli et al. (2005) couple an atmosphere to a thermodynamic ocean model and show that tropical precipitation is sensitive to the mean state in the Tropics. Here we use aquaplanet to mean the full coupling of an atmospheric model to an ocean model, both of which represent dynamics and so can transport properties around the globe. Such calculations have rarely been carried out before. In particular, we are only aware of one published study 1 of a coupled
SST; Kirtman and Schneider (2000) and Barsugli et al. (2005) couple an atmosphere to a thermodynamic ocean model and show that tropical precipitation is sensitive to the mean state in the Tropics. Here we use aquaplanet to mean the full coupling of an atmospheric model to an ocean model, both of which represent dynamics and so can transport properties around the globe. Such calculations have rarely been carried out before. In particular, we are only aware of one published study 1 of a coupled
momentum will converge into these source regions, thus sustaining an eddy-driven jet. In the Indian sector, a wide zone of high baroclinicity is observed at the latitudes of the eddy-driven jet (40°–55°). The maximum baroclinicity shifts within this zone from a poleward to an equatorward location between the positive and negative SAM phases, following the shift of the jet and storm tracks. In the Pacific sector, the mean baroclinicity is higher in the Tropics and much weaker at midlatitudes, in
momentum will converge into these source regions, thus sustaining an eddy-driven jet. In the Indian sector, a wide zone of high baroclinicity is observed at the latitudes of the eddy-driven jet (40°–55°). The maximum baroclinicity shifts within this zone from a poleward to an equatorward location between the positive and negative SAM phases, following the shift of the jet and storm tracks. In the Pacific sector, the mean baroclinicity is higher in the Tropics and much weaker at midlatitudes, in
). However the vertical shear in the zonal winds also increases somewhat, a response that we attribute to the reduction in the eddy kinetic energy and poleward heat flux due to the barotropic governor. The poleward shift during the slow adjustment also displays an equivalent barotropic structure outside of the Tropics, as expected from the response to a latitudinal displacement of the eddy-driven component of the wind field forced by a shift of the upper-level eddy momentum flux convergence ( Robinson
). However the vertical shear in the zonal winds also increases somewhat, a response that we attribute to the reduction in the eddy kinetic energy and poleward heat flux due to the barotropic governor. The poleward shift during the slow adjustment also displays an equivalent barotropic structure outside of the Tropics, as expected from the response to a latitudinal displacement of the eddy-driven component of the wind field forced by a shift of the upper-level eddy momentum flux convergence ( Robinson
, the damping, k a , is almost everywhere set to 40 −1 day −1 , the exception being in the low-level Tropics, where it is increased to 4 −1 day −1 at the surface to establish a more realistic Hadley circulation. The equilibrium profile T eq is determined by a small number of parameters, the most significant for our purposes being the equator-to-pole temperature difference, Δ T eq . Momentum is removed from the model near the surface by a Rayleigh drag of strength k f , which decays linearly
, the damping, k a , is almost everywhere set to 40 −1 day −1 , the exception being in the low-level Tropics, where it is increased to 4 −1 day −1 at the surface to establish a more realistic Hadley circulation. The equilibrium profile T eq is determined by a small number of parameters, the most significant for our purposes being the equator-to-pole temperature difference, Δ T eq . Momentum is removed from the model near the surface by a Rayleigh drag of strength k f , which decays linearly
10°S. And, as before, initiate the zonal flow without maintaining it. Here the response is totally different. Rossby wave radiation is absorbed in the Tropics, preventing any significant penetration into the Southern Hemisphere. This absorption feeds easterly momentum into the Tropics, accelerating it. A shelf of easterly momentum moves toward the wave source ( Fig. 9 ), in just the manner of the quasi-biennial oscillation model of Lindzen and Holton (1968) . Note that this migration and
10°S. And, as before, initiate the zonal flow without maintaining it. Here the response is totally different. Rossby wave radiation is absorbed in the Tropics, preventing any significant penetration into the Southern Hemisphere. This absorption feeds easterly momentum into the Tropics, accelerating it. A shelf of easterly momentum moves toward the wave source ( Fig. 9 ), in just the manner of the quasi-biennial oscillation model of Lindzen and Holton (1968) . Note that this migration and
.g., Garcia 1987 ; Haynes 1998 ; Scott and Haynes 1998 ). In the present case, forcing continually injects energy into the system, which then accumulates in the zonal flow. The flow in midlatitudes is damped by the effect of the relaxation on the streamfunction, but in the Tropics the relation between height and streamfunction is weaker and there is less damping of the flow there. Note that the equatorial jet is both stronger than the midlatitude jets and prograde, two features in common with the zonal
.g., Garcia 1987 ; Haynes 1998 ; Scott and Haynes 1998 ). In the present case, forcing continually injects energy into the system, which then accumulates in the zonal flow. The flow in midlatitudes is damped by the effect of the relaxation on the streamfunction, but in the Tropics the relation between height and streamfunction is weaker and there is less damping of the flow there. Note that the equatorial jet is both stronger than the midlatitude jets and prograde, two features in common with the zonal
is located slightly equatorward of its counterpart, so too does the Southern Hemisphere forcing reach a bit more into the Tropics. Hence, the direct response to trial 6 in the Southern Hemisphere is larger than its Northern Hemisphere counterpart (note particularly the direct streamfunction responses in Fig. 18 ). The annular mode dynamics may be more obscured in the Southern Hemisphere, and the annular mode response is therefore weaker. 6. Summary and discussion A simple general circulation
is located slightly equatorward of its counterpart, so too does the Southern Hemisphere forcing reach a bit more into the Tropics. Hence, the direct response to trial 6 in the Southern Hemisphere is larger than its Northern Hemisphere counterpart (note particularly the direct streamfunction responses in Fig. 18 ). The annular mode dynamics may be more obscured in the Southern Hemisphere, and the annular mode response is therefore weaker. 6. Summary and discussion A simple general circulation
the poleward months there is a distinct upper-tropospheric jet peak at lower latitudes. These winds peak in winter, as evident from U at 36°S shown in Fig. 2a . To remove externally forced variation, the mean winds are removed at each latitude, for each of the 12 months in turn, so leaving anomalies relative to the mean annual cycle. Likewise any linear trend over time is removed. Finally, the variability in zonal means forced from the Tropics by ENSO ( Seager et al. 2003 ) is removed using
the poleward months there is a distinct upper-tropospheric jet peak at lower latitudes. These winds peak in winter, as evident from U at 36°S shown in Fig. 2a . To remove externally forced variation, the mean winds are removed at each latitude, for each of the 12 months in turn, so leaving anomalies relative to the mean annual cycle. Likewise any linear trend over time is removed. Finally, the variability in zonal means forced from the Tropics by ENSO ( Seager et al. 2003 ) is removed using
? To answer this question, we must estimate the likely values of the forcing and damping parameters. Several authors have suggested, in analogy with Earth’s Tropics, that Jovian thunderstorms can transport most of the heat flux through the cloud layer ( Banfield et al. 1998 ; Gierasch et al. 2000 ). If so, then the globally averaged thunderstorm mass flux is Ṁ ∼ F /( c p Δ θ ), where F is the heat flux, c p is specific heat, and Δ θ is the static stability across the layer (or equivalently
? To answer this question, we must estimate the likely values of the forcing and damping parameters. Several authors have suggested, in analogy with Earth’s Tropics, that Jovian thunderstorms can transport most of the heat flux through the cloud layer ( Banfield et al. 1998 ; Gierasch et al. 2000 ). If so, then the globally averaged thunderstorm mass flux is Ṁ ∼ F /( c p Δ θ ), where F is the heat flux, c p is specific heat, and Δ θ is the static stability across the layer (or equivalently
