In these simulations, the zonal asymmetry of the boundary conditions (i.e. the topography) is fixed across the range of climates. Any changes in the amplitude of stationary waves must thus result from changes in the zonal-mean basic state. For this reason we first present the zonal-mean basic state in these simulations. For the most part, the zonal-mean climate is unchanged by the presence of zonal asymmetries. Zonal-mean quantities are mostly unaffected by the presence of zonally asymmetric boundary conditions. We compare zonal-mean quantities with the zonally-symmetric southern hemisphere or a separate zonally-symmetric experiment where relevant.
60 N 30 N EQ 30 S 60 S
0.2 0.4 0.6 0.8
−20 0 20 40 60
60 N 30 N EQ 30 S 60 S
0.2 0.4 0.6 0.8
−20 0 20 40 60
60 N 30 N EQ 30 S 60 S
0.2 0.4 0.6 0.8
−20 0 20 40 60 80 100 a odp = 1.0
b odp = 2.0
c odp = 6.0
Figure 5.1: Zonal-mean zonal wind [u] (shading, contour interval 5 m s−1) and potential temperature θ
(black contours, contours interval 20 K) for (a)α= 1.0, (b)α= 2.0, and (c)α= 6.0 of a reference experiment with no zonally asymmetric forcing. A dark green line shows the tropopause height as determined by the 2 K km−1lapse-rate criterion. All fields are computed from the R45 experiment.
The symmetry of the northern and southern hemispheres shows that these fields are not dependent on the zonally asymmetric forcing.
5.3.1 Jet changes
The jet stream and atmospheric thermal structure are the primary properties of the zonal-mean climate that determine the response of stationary waves to zonally anomalous forcing. We present the zonal-mean zonal wind and potential temperature in Fig. 5.1 for the reference climate and 2 warmer climates. As the climate warms, the atmosphere becomes more stable (as determined by the vertical derivative of potential temperature), meridional temperature gradients get weaker in the lower troposphere but stronger in the upper troposphere, the midlatitude eddy-driven jet disappears in favor of a single subtropical jet, and the height of the tropopause increases. Note that because of the increase in tropopause height, the jet at tropopause level gets stronger with warming despite the decrease in lower-tropospheric meridional temperature gradients. Another notable change is that the atmosphere becomes superrotating, where [u] > 0 at the equator, for the warmest two climates (α= 4.0,6.0). The superrotation is weak enough such that it does not significantly affect the propagation of Rossby waves across the equator, but it may play a role in the dynamics of the equatorial heating experiment (Chapter 6).
An interesting aspect to the zonal wind change in this model is that the midlatitude surface westerlies shift equatorward and polar easterlies strengthen and expand as the climate warms (Fig.
5.2a). The presence of a mountain tends to weaken the zonal-mean surface westerlies, but does not explain the equatorward shift with warming (Fig. 5.2b). Most evidence in comprehensive models and observations indicate a poleward shift of surface westerlies with warming (Kushner et al., 2001;
Yin, 2005; Chen and Held, 2007), though this response has only been probed over a fraction of the range of global-mean surface temperatures explored here. The changes in the idealized model are consistent with the changes in transient-eddy kinetic energy (see 5.3.2). In light of these differences, the orographic forcing results should be interpreted as the response to the particular zonal surface- wind change in this idealized GCM. Generalizing these ideas to climate change in Earth’s real climate system requires an understanding of the differences in zonal surface-wind response.
Global−mean surface temperature (K)
285 290 295 300 305 310 315
EQ 30 N 60 N
−5 0 5
Global−mean surface temperature (K)
285 290 295 300 305 310 315
EQ 30 N 60 N
−2
−1 0 1 2 a
b
m s -1 m s -1
Figure 5.2: Zonal-mean zonal surface wind across a wide range of climates in a) the northern hemisphere of R45 b) the difference between the northern hemisphere and southern hemisphere.
Dashed lines show the extent of the maximum topographic relief in the R45 experiment.
5.3.2 Transient eddy changes
Synoptic eddies affect the forcing of stationary eddies in 2 main ways. First, transient-eddy momen- tum flux convergence sets the strength and position of the surface zonal winds. Second, zonally- anomalous transient-eddy heat and momentum fluxes can either force or damp stationary waves through their influence on the stationary-eddy energy and momentum budgets. If the transient eddies can be thought to act diffusively (cf. Caballero and Langen, 2005), these zonally anoma- lous fluxes should scale with a bulk measure of the transient-eddy activity and the strength of the stationary-eddy temperature or vorticity gradients they are acting on. As a bulk measure of transient-eddy activity, we choose the transient-eddy kinetic energy (tEKE). This is shown in Fig.
5.3a, integrated over the troposphere. The non-monotonicity of tEKE has been explained previously in the same model as resulting from changes in mean available potential energy (MAPE) (O’Gorman and Schneider, 2008a).
The slight poleward shift of full-tropospheric tEKE does not fit with the equatorward shift of surface winds (Fig. 5.2) or the much stronger poleward shift of lower-troposheric tEKE previously reported for this model (Schneider et al., 2010). To understand the reasons for this we split tEKE into lower-tropospheric (belowσ= 0.72) and upper-tropospheric (tropopause level plus two levels above
Global−mean surface temperature (K)
285 290 295 300 305 310 315
EQ 30 N 60 N
0 2e5 4e5 6e5 8e5
Global−mean surface temperature (K)
285 290 295 300 305 310 315
EQ 30 N 60 N
0 1e5 2e5
Global−mean surface temperature (K)
285 290 295 300 305 310 315
EQ 30 N 60 N
0 1e5 2e5 3e5 4e5 a
b
c
J m -2
Figure 5.3: Zonal-mean transient-eddy kinetic energy (tEKE) vertically integrated (a) over the full troposphere (b) from the surface toσ= 0.72, and (c) from 2 levels above the tropopause to 3 levels below the tropopause.
and 3 below) components. The tEKE behaves very differently with climate change for the upper- tropospheric and lower-tropospheric components. While the lower-tropospheric tEKE exhibits the strong poleward shift reported in Schneider et al. (2010) and explained in a dry-model context by Mbengue and Schneider (2013), the upper-tropospheric tEKE shifts equatorward. The poleward shift of lower-tropopsheric tEKE and other measures of storm track activity has been seen in observations and comprehensive models (Fyfe, 2003; Yin, 2005; Bender et al., 2012), but, as far as we are aware, this is the first mention of an equatorward shift of upper-tropospheric tEKE with warming. This result could be unique to this model. The relevance of this is that the surface winds, which follow the upper-tropospheric eddy momentum flux convergence, also shift equatorward (Fig. 5.2). Some aspects of the orographic aspects of the orographic stationary Rossby wave response depend on this equatorward shift with warming, as we will see in the following sections.
90 180 270 60 S
30 S EQ 30 N 60 N
−4e+10
−2e+10 0 2e+10 4e+10
90 180 270
60 S 30 S EQ 30 N 60 N
−4e+10
−2e+10 0 2e+10 4e+10
90 180 270
60 S 30 S EQ 30 N 60 N
−4e+10
−2e+10 0 2e+10 4e+10
90 180 270
60 S 30 S EQ 30 N 60 N
−4e+10
−2e+10 0 2e+10 4e+10
90 180 270
60 S 30 S EQ 30 N 60 N
−4e+10
−2e+10 0 2e+10 4e+10
90 180 270
60 S 30 S EQ 30 N 60 N
−4e+10
−2e+10 0 2e+10 4e+10
α = 1.0
α = 1.0 α = 1.0
α = 3.0 α = 3.0 α = 3.0
R45 R54 R36
kg s -1 kg s -1
Figure 5.4: Barotropic streamfunction response to orographic forcing in the reference simulation (α= 1) and in the 3x optical depth simulation (α= 3) for all three mountain configurations (R45, R54, R36). The black contour is the 1000 m contour of surface height.