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  • View in gallery

    Aquaplanet SST configuration. (a) Background zonally symmetric and (b) zonally asymmetric SST (see appendix A for analytical functions). Contour interval is (a) 5 and (b) 1 K.

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    Aquaplanet and reanalysis circulation climatology. (a) The 925-hPa stationary eddy streamfunction (black, contour interval 2 × 106 m2 s−1, negative values dashed) and transient eddy streamfunction amplitude (orange, contour interval 1 × 106 m2 s−1 starting at 8 × 106 m2 s−1) and (b) zonal-mean streamfunction. (c),(d) As in (a),(b), but for 150 hPa with contour interval 5 × 106 (black) and 1 × 106 m2 s−1 (orange, starting at 18 × 106 m2s−1). Red contour indicates where the subcloud moist entropy is 5860 J kg−1 K−1, which exceeds the supercritical moist entropy for the ZSYM SST (see Emanuel 1995). (e)–(h) As in (a)–(d), but for the ERA-Interim summertime climatology.

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    Response of the Northern Hemisphere to the five different climate change scenarios (see text for description): (a) jet latitude (latitude of maximum 925-hPa zonal wind) and (b) Hadley cell edge (latitude where 925-hPa zonal wind transitions from easterly to westerly). In all panels, black and red symbols indicate the response for the ASYM and ZSYM background SST, respectively. Response to climate change of the latitude of maximum vertically integrated transient eddy: (c) momentum and (d) MSE transport. The response of dry static and latent energy transports in (c) are shown as triangles and diamonds, respectively. Response of subtropically averaged (20°–40°N) stationary eddy amplitude (streamfunction variance ) at (e) 925 and (f) 150 hPa.

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    Response of zonal-mean zonal wind (color shading) and temperature (black contours) to (a),(c) land warming and (b),(d) ocean warming for the (a),(b) ASYM and (c),(d) ZSYM background SST. Contour interval is 0.5 K (black) and 0.5 m s−1 (color). Green contour indicates the position of the tropopause in the control climate and the magenta contour indicates the position of the tropopause in the warmed climate.

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    Response of stationary eddy streamfunction (color shading) at (a),(b) 925 hPa (contour interval 5 × 105 m2 s−1) and (d),(e) 150 hPa (contour interval 1 × 106 m2 s−1) to (a),(d) land and (b),(e) ocean warming relative to control climate streamfunction (black contours as in Figs. 2a,c). Solid and dashed red contours indicate where the subcloud moist entropy is 5860 J kg−1 K−1 for the control (see Fig. 2) and warmed climates, respectively. Response of zonal-mean streamfunction at (c) 925 and (f) 150 hPa for land (brown lines) and ocean (blue lines) warming for ASYM (solid) and ZSYM (dashed) background SSTs.

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    Response to a range of regional climate change perturbations (dots indicate land perturbations and diamonds indicate ocean perturbations) of Asian-averaged 925-hPa (a) streamfunction and (b) vorticity decomposed into changes in moist entropy gradients and (black) and changes in temperature stratification and (red) [see (7) and (8)] and (c) vorticity decomposed into changes in energy input to the atmosphere (black) and changes in GMS (red) [see (9)]. The GMS changes due to enthalpy and latent energy alone are shown in blue. (d) Response of stationary eddy streamfunction variance vs change in vorticity due to moist entropy (black) and energy input (red) at 925 hPa.

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    Response of stationary eddy (a) moist entropy and (b) energy input to the atmosphere to land warming. Contour interval is (a) 2 J kg−1 K−1 and (b) 10 W m−2. (c) Response of 925-hPa zonal-mean stationary eddy moist entropy (black; J kg−1 K−1, multiplied by 5) and energy input (red; W m−2) variance in response to land warming (solid) and ocean warming (dashed). (d) Response of 925-hPa zonal-mean stationary eddy streamfunction variance vs zonal-mean stationary eddy moist entropy (black) and energy input (red) variance. Circles indicate land perturbations and diamonds indicate ocean perturbations.

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    Response of vertically integrated stationary eddy (black lines) (a),(b) momentum and (c),(d) MSE transport to (a),(c) land and (b),(d) ocean warming. The momentum transport is decomposed into squared eddy streamfunction (red line) and meridional wavenumber (blue line) changes [see (10)]. The MSE transport is decomposed into diffusivity (red line) and MSE gradient (blue line) changes [see (11)].

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    Response of vertically integrated stationary eddy (black lines) and transient eddy (blue lines) (a),(b) momentum and (c),(d) MSE transport to (a),(c) land and (b),(d) ocean warming. Dashed lines indicate transport averaged over the Asia–Pacific sector.

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    As in Fig. 7, but for the transient eddy response.

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    As in Figs. 2a–d, but for a 300-K zonally symmetric background SST.

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    Zonal mean of SST perturbations GWARM (red), GWARM for equivalent potential temperature (black), LWARM (brown), OWARM (blue) that mimic surface warming. See Table 1 for simulation description. Response of near-surface (925 hPa) equivalent potential temperature to (b) LWARM + 2 K, (c) LWARM + 4 K, (d) OWARM + 2 K, (e) OWARM + 4 K, and (f) warming of equivalent potential temperature by 2 K. Contour interval in (b),(d),(f) is 1 K and in (c),(e) 2 K.

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    As in Fig. 5, but for the difference between the response with and without cloud and water vapor radiative feedbacks.

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    (a) Near-surface divergence and (b) approximate divergence following (C8) for ASYM background SST. Contour interval is 0.5 × 10−6 s−1.

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    Contributions to approximate divergence [(C8)] from (a) Ekman, (b) Sverdrup, and (c) equatorial convergence for the climatology shown in Fig. C1b. Contour interval is 0.5 × 10−6 s−1.

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Understanding the Links between Subtropical and Extratropical Circulation Responses to Climate Change Using Aquaplanet Model Simulations

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  • 1 Department of the Geophysical Sciences, University of Chicago, Chicago, Illinois
  • 2 Lamont-Doherty Earth Observatory, Columbia University, Palisades, New York
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Abstract

Previous research has shown that subtropical and extratropical circulations are linked seasonally and in response to climate change. In particular, amplification (weakening) of subtropical stationary eddies is linked to a poleward (equatorward) shift of the extratropical circulation in the Northern Hemisphere. Here the mechanisms linking subtropical and extratropical circulation responses to climate change are examined using prescribed sea surface temperature aquaplanet simulations with a subtropical zonal asymmetry that mimics land–ocean contrasts. A poleward circulation shift occurs in response to uniform global warming even in the presence of subtropical stationary eddies. Subtropical stationary eddies exhibit a weak response to global warming; however, regional warming of temperature (or equivalent potential temperature) over land (ocean) increases (decreases) stationary eddy amplitude and shifts the extratropical circulation poleward (equatorward), consistent with comprehensive models. The stationary eddy response to regional warming is connected to regional moist entropy gradient, energy input to the atmosphere, and gross moist stability changes. Stationary eddy amplitude changes directly affect momentum and moist static energy transport following linear wave and mixing length theories. The transport changes do not follow a fixed-diffusivity framework. Extratropical transient eddy transport changes compensate ~70%–90% of the subtropical stationary eddy transport response. This assumes exact subtropical compensation accounts for a large fraction of the meridional shift of the extratropical circulation in response to regional climate change.

Corresponding author address: Tiffany A. Shaw, Department of the Geophysical Sciences, University of Chicago, 5734 South Ellis Ave., Chicago, IL 60637. E-mail: tas1@uchicago.edu

Abstract

Previous research has shown that subtropical and extratropical circulations are linked seasonally and in response to climate change. In particular, amplification (weakening) of subtropical stationary eddies is linked to a poleward (equatorward) shift of the extratropical circulation in the Northern Hemisphere. Here the mechanisms linking subtropical and extratropical circulation responses to climate change are examined using prescribed sea surface temperature aquaplanet simulations with a subtropical zonal asymmetry that mimics land–ocean contrasts. A poleward circulation shift occurs in response to uniform global warming even in the presence of subtropical stationary eddies. Subtropical stationary eddies exhibit a weak response to global warming; however, regional warming of temperature (or equivalent potential temperature) over land (ocean) increases (decreases) stationary eddy amplitude and shifts the extratropical circulation poleward (equatorward), consistent with comprehensive models. The stationary eddy response to regional warming is connected to regional moist entropy gradient, energy input to the atmosphere, and gross moist stability changes. Stationary eddy amplitude changes directly affect momentum and moist static energy transport following linear wave and mixing length theories. The transport changes do not follow a fixed-diffusivity framework. Extratropical transient eddy transport changes compensate ~70%–90% of the subtropical stationary eddy transport response. This assumes exact subtropical compensation accounts for a large fraction of the meridional shift of the extratropical circulation in response to regional climate change.

Corresponding author address: Tiffany A. Shaw, Department of the Geophysical Sciences, University of Chicago, 5734 South Ellis Ave., Chicago, IL 60637. E-mail: tas1@uchicago.edu

1. Introduction

The transition of the atmospheric circulation from winter to summer involves significant changes in subtropical and extratropical climate in the Northern Hemisphere (NH). In particular, the intertropical convergence zone (ITCZ) migrates 20° latitude northward into the subtropics (Gadgil 2003; Schneider et al. 2014), summertime subtropical monsoon cyclones over land and anticyclones over ocean replace wintertime land anticyclones and oceanic cyclones (Wang and Ting 1999; Rodwell and Hoskins 2001), and the extratropical eddy-driven jet stream shifts approximately 10° northward (Shaw 2014). From a dynamical perspective the seasonal transition involves increased stationary zonally asymmetric rotational flow in the lower and upper subtropical troposphere (Wang and Ting 1999; Shaw 2014). Summertime subtropical stationary eddy amplitude is actually larger than its wintertime midlatitude counterpart in the lower troposphere (see Fig. 8 of Wang and Ting 1999), and stationary eddies dominate moisture transport during summertime (Shaw and Pauluis 2012; Shaw 2014), determining the northward extent of the monsoons (Chou and Neelin 2003).

The seasonal growth of summertime subtropical stationary eddies is due to increased diabatic heating according to idealized models (Ting 1994; Wang and Ting 1999; Chen et al. 2001; Rodwell and Hoskins 2001; Shaw 2014). Interactions with topography, stationary nonlinearity, and transient eddy transport are of secondary importance in stationary wave models [see Figs. 15 and 16 in Wang and Ting (1999)]. Topography is mainly important for shaping the spatial structure of near-surface moist entropy (Boos and Kuang 2010) and thus diabatic heating following convective quasi-equilibrium (Emanuel et al. 1994). The strength of oceanic anticyclones is affected by land–sea atmospheric feedbacks (Miyasaka and Nakamura 2005) and coupling with sea surface temperatures (SSTs; Seager et al. 2003).

Summertime subtropical stationary eddies exhibit a baroclinic vertical structure (Chen 2010) consistent with first baroclinic mode diabatic heating and Sverdrup balance (Gill 1980). Wintertime stationary eddies exhibit a barotropic vertical structure consistent with the importance of zonal advection of vorticity by the jet stream (Hoskins and Karoly 1981). In the upper troposphere the divergent circulation driven by diabatic heating is a source of stationary eddies (Sardeshmukh and Hoskins 1988). Stationary eddies can affect synoptic-scale transient eddies by influencing baroclinicity (e.g., Hoskins and Valdes 1990; Kaspi and Schneider 2013) or the barotropic flow (e.g., Manabe and Terpstra 1974; Lee 1995; Harnik and Chang 2004; Brayshaw et al. 2009; Son et al. 2009; Sauliere et al. 2012). Subtropical stationary eddies can also influence the meridional position of the subtropical critical line (Shaw 2014).

The response of the summertime circulation to anthropogenic climate change is of considerable interest because of its role in determining regional moisture transport and precipitation (Xie et al. 2015; Shepherd 2014). Shaw and Voigt (2015) connect robust changes in summertime stationary eddies and dynamical moisture transport to land–sea moist-entropy contrast changes. The role of land–sea moist-entropy contrast changes is revealed using atmospheric general circulation model (AGCM) simulations with prescribed SST. When SSTs are fixed and CO2 is quadrupled, near-surface moist entropy and energy input to the atmosphere increase over land (positive meridional moist entropy gradient change) and stationary eddy amplitude increases. In contrast, when SSTs are warmed uniformly but CO2 is held fixed, near-surface moist entropy and energy input increase over the ocean (negative meridional moist entropy gradient change) and stationary eddy amplitude decreases [see Figs. 1b–d of Shaw and Voigt (2015)]. The negative meridional land–ocean moist-entropy gradient response to SST warming is not dominated by temperature changes: in AGCM simulations land warms more than the ocean in response to SST warming implying a positive gradient change. In coupled climate models the direct radiative effect of CO2 and indirect SST warming compete transiently to produce a weak and nonrobust equilibrium summertime circulation response over the Asia–Pacific region [see Fig. 1a of Shaw and Voigt (2015)].

The subtropical circulation response to climate change is linked to the extratropics. In particular, when subtropical stationary eddy amplitude increases there is a poleward shift of the North Pacific jet stream, whereas when the amplitude decreases the jet shifts equatorward (Shaw and Voigt 2015). The equatorward Pacific jet shift during summertime is very unique in response to increased CO2 (Simpson et al. 2014; Grise and Polvani 2014). A similar connection between subtropical stationary eddy amplitude and extratropical jet stream occurs in response to seasonal insolation: stationary eddy amplitude increases and the jet stream shifts poleward for July minus May, whereas for September minus August wave amplitude decreases and the jet stream shifts equatorward [see Figs. 3 and 4 of Shaw and Voigt (2015)].

Here we focus on the thermodynamic and dynamic mechanisms linking subtropical and extratropical circulation responses to regional climate change during summertime. Our goal is to explain how regional climate change impacts stationary eddy amplitude and how stationary eddy amplitude changes affect the extratropical circulation. Building on previous work that used idealized models to study the zonally symmetric circulation response to global warming (e.g., Lorenz and DeWeaver 2007; Butler et al. 2010; Lu et al. 2014; Mbengue and Schneider 2013; Medeiros et al. 2015; Voigt and Shaw 2015), we investigate the summertime circulation response to warming using aquaplanet simulations with subtropical zonal asymmetries because of their leading order contribution to the summertime circulation. The aquaplanet simulations are described in section 2. Our results are presented in section 3. The conclusions and discussion are summarized in section 4.

2. Methods

As discussed in the introduction, subtropical and extratropical circulation responses to regional climate change (e.g., direct radiative forcing vs indirect SST warming) are linked. Here we examine these links using prescribed SST aquaplanet simulations. While highly idealized, a prescribed SST setup is useful for examining the AGCM responses because regional climate change can be controlled directly via SST perturbations as discussed below.

Prescribed SST aquaplanet models are an important part of the climate model hierarchy (comprehensive moist physics with simplified lower boundary condition; see Blackburn and Hoskins 2013) and have been used to study the jet stream response to SST gradients (Brayshaw et al. 2008), Walker circulation (Neale and Hoskins 2000), NH (Brayshaw et al. 2009) and Southern Hemisphere (SH; Inatsu and Hoskins 2004) wintertime storm tracks, and the response of the extratropical circulation to uniform global warming (Medeiros et al. 2015). Shaw (2014) showed that the seasonal connection between subtropical stationary eddy amplitude and jet position can be understood using aquaplanet simulations. Here we use the aquaplanet protocol to understand the subtropical stationary eddy response to climate change.

The prescribed SST aquaplanet model setup used here follows Shaw and Voigt (2015), which builds on the results of Shaw (2014). We use the MPI-ESM-LR aquaplanet model (Stevens et al. 2013). The background SST involves the sum of zonally symmetric and zonally asymmetric contributions. The zonally symmetric contribution is the “Qobs” SST profile, which is a simple geometric function that mimics the observed zonal-mean SST (Neale and Hoskins 2000). The maximum Qobs SST is shifted to 10°N (Fig. 1a), which mimics the northward shift of insolation during late spring (appendix A). The zonally asymmetric SST contribution mimics subtropical land–ocean heating asymmetries via a subtropical zonally asymmetric perturbation (Fig. 1b). We define “land” as areas enclosed by the black contours in Fig. 1b (labeled Asia and N. America; see appendix A) and “ocean” areas elsewhere. These land areas represent idealized saturated land; that is, they have the same surface properties (same surface flux parameterization) as the ocean areas.

Fig. 1.
Fig. 1.

Aquaplanet SST configuration. (a) Background zonally symmetric and (b) zonally asymmetric SST (see appendix A for analytical functions). Contour interval is (a) 5 and (b) 1 K.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

The stationary eddy streamfunction, transient eddy streamfunction variance, and zonal-mean streamfunction for the ASYM SST (Fig. 1a plus Fig. 1b) climate are shown in Figs. 2a–d. There is a large-amplitude near-surface cyclone over Asia, which mimics the monsoon circulation, and anticyclones over the ocean regions. The circulation in the upper troposphere is out of phase with that at the surface (cf. Figs. 2a and 2c). The circulation over Asia has a larger amplitude than that over North America, in agreement with reanalysis data (Figs. 2e–h; Wang and Ting 1999; Rodwell and Hoskins 2001; Shaw 2014). The extratropical circulation for the ASYM SST exhibits a clear Pacific storm track (orange contours in Figs. 2a,c) but not an Atlantic storm track; hence, our results focus on the Pacific storm track. The maximum near-surface moist entropy (red line in Figs. 2a,c) is in good agreement with reanalysis data (red line in Figs. 2e,g). The value of moist entropy shown in Fig. 2 is supercritical according to zonally symmetric criteria [see (11) in Emanuel (1995)]. Overall the aquaplanet does a reasonable job of capturing the important summertime features (stationary eddy streamfunction, near-surface moist entropy) in reanalysis data. We note that by imposing the SST we are controlling the background near-surface moist entropy, which allows us to test its role in the circulation response to warming. However, we are not taking into account the role of ventilation (Chou and Neelin 2003) and topography (Boos and Kuang 2010; Liu et al. 2007) in shaping the maximum entropy.

Fig. 2.
Fig. 2.

Aquaplanet and reanalysis circulation climatology. (a) The 925-hPa stationary eddy streamfunction (black, contour interval 2 × 106 m2 s−1, negative values dashed) and transient eddy streamfunction amplitude (orange, contour interval 1 × 106 m2 s−1 starting at 8 × 106 m2 s−1) and (b) zonal-mean streamfunction. (c),(d) As in (a),(b), but for 150 hPa with contour interval 5 × 106 (black) and 1 × 106 m2 s−1 (orange, starting at 18 × 106 m2s−1). Red contour indicates where the subcloud moist entropy is 5860 J kg−1 K−1, which exceeds the supercritical moist entropy for the ZSYM SST (see Emanuel 1995). (e)–(h) As in (a)–(d), but for the ERA-Interim summertime climatology.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

We also construct an equivalent zonally symmetric (ZSYM) background SST, which has no zonal asymmetries but has the same zonal-mean SST as ASYM (e.g., zonal mean of Fig. 1a plus Fig. 1b). Note the climate with ZSYM SST has no stationary eddies. Our approach follows Kaspi and Schneider (2013), who considered a zonally asymmetric ocean heat flux convergence perturbation in the NH and an equivalent zonally symmetric ocean heat flux convergence in the SH. The comparison of the circulation response for ASYM and ZSYM allows us to assess the extent to which zonally averaged concepts apply in the presence of zonal asymmetries. We also assess the role of cloud and water vapor radiative feedbacks using the locking technique of Voigt and Shaw (2015). The technique disables radiative feedbacks from clouds and water vapor by locking their radiative properties to those in the control simulation.

Using the aquaplanet setup we run several simulations with different perturbations in order to mimic regional climate change (i.e., land–ocean moist-entropy contrasts). We impose uniform global warming (GWARM) of SST, global warming of equivalent potential temperature,1 land warming (LWARM), and ocean warming (OWARM). Global warming of equivalent potential temperature tests the importance of moist entropy gradients for the circulation response.2 We focus on the simulations that involve regional and global warming by 2 K (Table 1) because they produce regional changes of near-surface equivalent potential temperature on the order of 10 K, in agreement with the response in AGCM (Shaw and Voigt 2015) and CMIP5 (Byrne and O’Gorman 2013) simulations. Perturbations imposed on the ZSYM SST involve only zonal-mean components; that is, the zonal-mean components of land and ocean warming are imposed on top of the ZSYM SST (Fig. A2a). We also assess the impact of direct radiative forcing by quadrupling CO2 while holding SSTs fixed.

Table 1.

Zonally asymmetric sea surface temperature perturbation (K) applied to ASYM background SST (Fig. 1a plus Fig. 1b) to mimic climate change. The perturbations involve global uniform warming (GWARM), global uniform equivalent potential temperature warming (GWARM θe), land warming (LWARM), and ocean warming (OWARM).

Table 1.

To understand the thermodynamic and dynamic mechanisms controlling the circulation response across a range of regional perturbations, we run simulations with perturbations ranging from ±1 to ±4 K over land, ocean, and Asia for a total of 24 simulations. The 4-K LWARM and OWARM perturbations lead to changes of near-surface equivalent potential temperature on the order of 20 K. Thus the 24 simulations involve near-surface equivalent potential temperature changes in the range of ±20 K (Figs. A2b–f).

In the analysis that follows the angle brackets represent a mass-weighted vertical integral, the square brackets represent a zonal average, asterisks represent a deviation from the zonal average, overbars represent a time average, and primes represent a deviation from the time average. Here the eddy-driven jet position is defined as the maximum near-surface (925 hPa) zonal-mean zonal wind and the Hadley cell edge is defined as the latitude where the near-surface zonal wind is zero in the subtropics (transition from tropical easterlies to extratropical westerlies). The jet position and Hadley cell edge are calculated after interpolating the zonal wind data onto a 0.1° grid.

3. Results

a. Response to global versus regional warming

We begin by examining the aquaplanet circulation response in the NH to global warming by 2 K and direct radiative forcing by 4 × CO2. While we examine zonal-mean quantities, the extratropical climate with ASYM SST has a Pacific storm track (Fig. 2a), and thus our results apply primarily to the Asia–Pacific region (0°–240°E). The ASYM and ZSYM climates both exhibit a poleward shift of the jet stream, equal to 1.0° and 0.6°, respectively (Fig. 3a), and robust tropical upper tropospheric warming (not shown) in response to global warming. Statistical tests (t and Kolmogorov–Smirnov tests) applied to daily distributions of jet latitude for the climatology and warmed climate suggest we can reject the null hypothesis at the 5% significance level. The poleward jet shift is consistent with CMIP5 aquaplanet models, which exhibit a poleward shift in response to uniform global SST warming by 4 K [see Fig. 6 of Medeiros et al. (2015)]. The jet stream response to global equivalent potential temperature warming by 2 K is much weaker (Fig. 3a). This suggests moist entropy gradient changes, induced by uniform SST warming via the Clausius–Clapeyron relation (Shaw and Voigt 2016), are important for the circulation response as discussed below. In response to direct radiative forcing by 4 × CO2 the jet shift is weaker than that in response to global warming (Fig. 3a). The jet stream response to uniform warming or direct radiative forcing does not depend significantly on whether the background SST is zonally asymmetric (ASYM) or zonally symmetric (ZSYM).

Fig. 3.
Fig. 3.

Response of the Northern Hemisphere to the five different climate change scenarios (see text for description): (a) jet latitude (latitude of maximum 925-hPa zonal wind) and (b) Hadley cell edge (latitude where 925-hPa zonal wind transitions from easterly to westerly). In all panels, black and red symbols indicate the response for the ASYM and ZSYM background SST, respectively. Response to climate change of the latitude of maximum vertically integrated transient eddy: (c) momentum and (d) MSE transport. The response of dry static and latent energy transports in (c) are shown as triangles and diamonds, respectively. Response of subtropically averaged (20°–40°N) stationary eddy amplitude (streamfunction variance ) at (e) 925 and (f) 150 hPa.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

The jet stream shift in response to uniform warming coincides with a shift of the edge of the Hadley cell. The Hadley cell expands in response to global warming but does not change significantly in response to equivalent potential temperature warming or direct radiative forcing (Fig. 3b). Consistent with the jet response, the response of the Hadley cell edge to uniform warming or direct radiative forcing does not depend significantly on the background SST.

The circulation response to regional warming is not necessarily similar to the response to global warming or direct radiative forcing. When only land regions are warmed the jet shifts poleward but the shift is much larger than that for global warming (Fig. 3a). The magnitude of the poleward shift also depends on the background SST. In particular, there is a 2.2° poleward jet shift for ASYM; in contrast there is only a 0.6° poleward shift for ZSYM (Fig. 3a). Cloud and water vapor radiative feedbacks account for roughly 20% of the jet shift in response to land warming for the ASYM SST (appendix B). More surprisingly the jet shifts equatorward by 1.5° in response to ocean warming for ASYM (Fig. 3a). Cloud and water vapor radiative feedbacks do not significantly impact the jet response to ocean warming (appendix B). The jet also shifts 1.5° equatorward for ZSYM in response to ocean warming, consistent with the impact of subtropical SST gradient changes on zonally symmetric storm tracks (Brayshaw et al. 2008). The Hadley cell edge follows the jet shift for the ZSYM SST; however, the shifts are not the same magnitude for the ASYM SST (Figs. 3a,b). This result suggests subtropical stationary eddies affect the connection between the jet position and Hadley cell edge.

The meridional shifts of the circulation are coupled to shifts of maximum vertically integrated transient eddy momentum and moist static energy (MSE) transports. The maximum vertically integrated transient eddy momentum transport shifts poleward in response to uniform global warming and land warming but shifts equatorward in response to ocean warming (Fig. 3c). The covariation of jet stream and maximum transient eddy momentum transport is expected because transient eddy momentum flux convergence, occurring poleward of the transport maximum, maintains the eddy-driven jet. The maximum vertically integrated transient eddy MSE transport also shifts poleward in response to uniform global warming and land warming but shifts equatorward in response to ocean warming (Fig. 3d). The shift of transient eddy MSE transport is interpreted as a shift of the storm track since MSE transport maximizes in the lower troposphere. The meridional shift of maximum vertically integrated transient eddy MSE transport follows the dry-static energy transport (triangles in Fig. 3d). The maximum transient eddy moisture transport shifts northward in response to all forcings, including ocean warming (diamonds in Fig. 3d).

An important feature distinguishing the circulation response to regional versus global warming is the response of stationary eddy amplitude. There is no stationary eddy amplitude response to global warming, regional warming, or direct radiative forcing for the ZSYM background SST (Figs. 3e,f) by definition. The stationary eddy amplitude response to global warming and direct radiative forcing is small for the ASYM background SST (Figs. 3e,f). In contrast, land warming increases stationary eddy amplitude whereas ocean warming decreases it (Figs. 3e,f). The subtropical stationary eddy amplitude and extratropical circulation responses to regional warming occur in conjunction with changes in subtropical stratification. For the ASYM SST, land warming increases the subtropical upper-tropospheric temperature, consistent with the adjustment to a moist adiabat (black solid lines, Fig. 4a). The subtropical tropopause also shifts upward. In response to ocean warming the lapse rate is destabilized with anomalous cooling aloft (black dashed lines, Fig. 4b) and the suptropical tropopause shifts downward; however, there is warming aloft in the tropics. Subtropical stratification changes are weaker for the ZSYM SST (Figs. 4c,d).

Fig. 4.
Fig. 4.

Response of zonal-mean zonal wind (color shading) and temperature (black contours) to (a),(c) land warming and (b),(d) ocean warming for the (a),(b) ASYM and (c),(d) ZSYM background SST. Contour interval is 0.5 K (black) and 0.5 m s−1 (color). Green contour indicates the position of the tropopause in the control climate and the magenta contour indicates the position of the tropopause in the warmed climate.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

The area of supercritical moist entropy increases over land in response to land warming and over the ocean in response to ocean warming (cf. solid red and dashed lines in Figs. 5a,b). The lower-tropospheric stationary eddy streamfunction response to land and ocean warming is consistent with anomalous cyclonic circulations over the warmed regions and anticyclonic circulations elsewhere (Figs. 5a,b). In the upper troposphere the responses have the opposite phase, consistent with a vertical baroclinic structure (Figs. 5d,e). These aquaplanet responses are consistent with AGCM response to direct radiative forcing and SST warming [see Figs. 1 and 2 in Shaw and Voigt (2015)]. Cloud and water vapor radiative feedbacks amplify the stationary eddy response over the ocean but only amplify the circulation over land in response to ocean warming (Fig. B1). The zonal-mean streamfunction response, which reflects the Hadley circulation, is similar for ASYM and ZSYM in response to ocean warming (blue solid and dashed lines in Figs. 5c,f); however, in response to land warming the streamfunction response for ASYM is much larger than ZSYM (brown solid and dashed lines in Figs. 5c,f).

Fig. 5.
Fig. 5.

Response of stationary eddy streamfunction (color shading) at (a),(b) 925 hPa (contour interval 5 × 105 m2 s−1) and (d),(e) 150 hPa (contour interval 1 × 106 m2 s−1) to (a),(d) land and (b),(e) ocean warming relative to control climate streamfunction (black contours as in Figs. 2a,c). Solid and dashed red contours indicate where the subcloud moist entropy is 5860 J kg−1 K−1 for the control (see Fig. 2) and warmed climates, respectively. Response of zonal-mean streamfunction at (c) 925 and (f) 150 hPa for land (brown lines) and ocean (blue lines) warming for ASYM (solid) and ZSYM (dashed) background SSTs.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

b. Linking stationary eddy amplitude changes to the rotational and divergent flow

The response to land and ocean warming involves large changes in stationary eddy amplitude. What controls the stationary eddy amplitude changes? Summertime stationary eddies are driven by surface convergence and diabatic heating (Wang and Ting 1999; Rodwell and Hoskins 2001). The stationary eddy amplitude response to land and ocean warming can be understood via the connection between thermodynamics and rotational and divergent components of the circulation. Here we make use of two frameworks that connect thermodynamics to the rotational and divergent flow: 1) convective quasi-equilibrium dynamics (Emanuel et al. 1994; Emanuel 1995) and 2) the MSE budget (Neelin and Held 1987).

Convective quasi-equilibrium dynamics assumes the circulation is tied to the region of supercritical near-surface moist entropy and constrains the vertical structure of the atmosphere to be a moist adiabat (Emanuel et al. 1994; Emanuel 1995). This framework is appropriate for summertime when the atmosphere is close to moist adiabatic up to 50°N (Korty and Schneider 2007) and reanalysis data support a quasi-equilibrium view of summertime monsoons (Nie et al. 2010). Following quasi-equilibrium dynamics, the near-surface moist entropy and its gradients are directly linked to the geopotential Φ and the rotational component of the balanced surface flow :
e1
e2
where the subscript s refers to the near surface, ψ is the streamfunction, f is the Coriolis parameter, is the subcloud moist entropy and is its critical value,3 and is the temperature stratification [i.e., the difference between surface and tropopause temperature; see Eqs. (27) and (26) in Emanuel (1995)]. Here we use the thermal tropopause from the World Meteorological Organization. Quasi-equilibrium dynamics also provides an approximate expression for the vertical component of near-surface vorticity:
e3
The MSE framework links the divergent flow (e.g., divergence ) to energy input to the atmosphere —that is, the difference of energy fluxes between the surface and top of the atmosphere and the gross moist stability (GMS, M)4:
e4
[see Eq. (2.12) in Neelin and Held (1987)].
It is well known that the rotational and divergent flow in the boundary layer can be connected via the Gill model (Gill 1980; Held and Hoskins 1985; Neelin 1989; Chiang et al. 2001). The connection can be straightforwardly derived assuming a linear momentum balance between Coriolis, pressure gradient, and frictional forces (see appendix C):
e5
where r is a linear damping parameter. This equation reflects a linear vorticity balance. The terms on the right-hand side of (5) represent Sverdrup balance, Ekman balance, and equatorial convergence via the trade winds. In general, the dominance of the different terms depends on the length scale of the perturbation. Assuming rf then Ekman balance holds if the length scale L < f/β (appendix C). In the control climate, (5) captures near-surface divergence (Fig. C1) with Ekman balance dominating in the subtropics (Fig. C2) such that (5) becomes
e6
In the upper troposphere, the linear balance in (5) reduces to the classic Sverdrup balance; however, nonlinear terms are nonnegligible (Sardeshmukh and Hoskins 1988).
According to quasi-equilibrium dynamics, changes in near-surface streamfunction are linked to changes in near-surface moist entropy or temperature stratification. Likewise changes in vorticity are linked to changes in near-surface moist entropy gradient or temperature stratification:
e7
and
e8
Combining the linear momentum balance in (6) with the MSE framework in (4) implies near-surface vorticity changes are linked to energy input to the atmosphere or GMS changes:
e9
The relationships in (8) and (9) imply that in a region of climatological cyclonic circulation, a cyclonic circulation response (e.g., ), which occurs over land regions in response to land warming (Fig. 5a), may be due to 1) increased supercriticality gradient, 2) increased temperature stratification, 3) increased energy input to the atmosphere, or 4) decreased GMS. In contrast, an anticyclonic circulation response, which occurs over land regions in response to ocean warming (Fig. 5b), involves opposite-signed changes.

To determine the factors that affect the stationary eddy amplitude response to regional warming, we run several aquaplanet simulations with a range of regional surface perturbations (see section 2). We focus on changes around Asia; more specifically we average the response over regions of climatological cyclonic circulation. Across the range of imposed regional climate perturbations, we find the linearization in (7)(9) is a good approximation (cf. x and y axes in Fig. 6). The changes in quasi-equilibrium streamfunction [(7)] around Asia are dominated by changes in near-surface moist entropy (Fig. 6, black). In particular, the cyclonic response to land warming (Fig. 5a) occurs because of increased supercriticality, whereas decreased supercriticality in response to ocean warming leads to an anticyclonic response (Fig. 5b). Changes in temperature stratification have a smaller impact on vorticity changes (Fig. 6a, red).

Fig. 6.
Fig. 6.

Response to a range of regional climate change perturbations (dots indicate land perturbations and diamonds indicate ocean perturbations) of Asian-averaged 925-hPa (a) streamfunction and (b) vorticity decomposed into changes in moist entropy gradients and (black) and changes in temperature stratification and (red) [see (7) and (8)] and (c) vorticity decomposed into changes in energy input to the atmosphere (black) and changes in GMS (red) [see (9)]. The GMS changes due to enthalpy and latent energy alone are shown in blue. (d) Response of stationary eddy streamfunction variance vs change in vorticity due to moist entropy (black) and energy input (red) at 925 hPa.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

Along similar lines near-surface moist entropy gradient changes (Fig. 6b, black) dominate quasi-equilibrium vorticity [(8)] changes around Asia. Vorticity increases (strengthened cyclonic circulation) because of increased supercriticality gradient in response to land warming, whereas vorticity decreases (weakened cyclonic circulation) in response to ocean warming. Vorticity changes due to temperature stratification changes (Fig. 6b, red) are much weaker, consistent with the quasi-equilibrium rotational flow response in AGCMs [see Fig. 2 in Shaw and Voigt (2015)]. These results highlight the importance of land–ocean moist entropy gradient changes. If there is no regional change of supercriticality gradient then there is no change of stationary wave vorticity, which explains why stationary eddies are not sensitive to global warming of equivalent potential temperature (Figs. 3e,f).

Changes in MSE vorticity [(9)] around Asia are associated with changes of energy input to the atmosphere (Fig. 6c, black) and changes in GMS (Fig. 6c, red). Increased energy input due to land warming increases vorticity over land, whereas decreased energy input due to ocean warming decreases vorticity. The GMS decrease in response to ocean warming decreases vorticity over land (red diamonds, bottom-left quadrant in Fig. 6c). The more nuanced GMS changes reflect opposing behavior between different GMS components (cf. Chou et al. 2013)—for example, thermal (enthalpy plus latent energy terms in MSE) stratification changes are positively correlated with vorticity changes (blue in Fig. 6c); however, they are opposed by geopotential stratification changes, which reflect tropopause height changes (Figs. 4a,b).

The changes in near-surface vorticity associated with near-surface moist entropy gradients and energy input around Asia are significantly correlated with zonal-mean stationary eddy amplitude (streamfunction variance) response to regional warming in the lower (Fig. 6d) and upper (not shown) troposphere. The streamfunction and vorticity are related via a squared length scale , , which can be estimated from the slope in Fig. 6d—for example, (black), and (red). The linear relationship between vorticity around Asia and zonal-mean stationary eddy amplitude across a range of regional climate changes is likely set by the zonal width of the Asian land area (Fig. 1b).

A more general interpretation of the stationary eddy response to regional climate change is that regional perturbations project directly onto stationary eddy thermodynamics. For example, in response to land warming the change of stationary eddy moist entropy and energy input reflect increased values over land and decreased values over the ocean (Figs. 7a,b). The opposite occurs in response to ocean warming (not shown). These stationary eddy thermodynamic changes project onto the zonal mean via stationary eddy variance ( and , Fig. 7c) and are correlated with zonal-mean stationary eddy amplitude changes (Fig. 7d), consistent with reanalysis data [see Fig. 2 in Shaw (2014)].

Fig. 7.
Fig. 7.

Response of stationary eddy (a) moist entropy and (b) energy input to the atmosphere to land warming. Contour interval is (a) 2 J kg−1 K−1 and (b) 10 W m−2. (c) Response of 925-hPa zonal-mean stationary eddy moist entropy (black; J kg−1 K−1, multiplied by 5) and energy input (red; W m−2) variance in response to land warming (solid) and ocean warming (dashed). (d) Response of 925-hPa zonal-mean stationary eddy streamfunction variance vs zonal-mean stationary eddy moist entropy (black) and energy input (red) variance. Circles indicate land perturbations and diamonds indicate ocean perturbations.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

c. Changes in transport and shifts of the extratropical circulation

Changes in subtropical stationary eddy amplitude in response to regional warming coincide with large changes in vertically integrated stationary eddy momentum and MSE transport. In particular, vertically integrated stationary eddy momentum and MSE transport increases in response to land warming and decreases in response to ocean warming (Fig. 8, black lines). Stationary eddy amplitude is connected to transport of momentum and MSE via linear wave and mixing length theories. For example, linear wave theory predicts and mixing length theory predicts , where is the stationary eddy zonal wavelength, is the stationary eddy meridional wavelength, is the stationary eddy kinematic diffusivity, and is the near-surface MSE. Following Kushner and Held (1998), the stationary eddy diffusivity is defined kinematically as , where is the standard deviation of the stationary eddy streamfunction and α is a coefficient of proportionality chosen to satisfy the mixing length approximation. The meridional wavenumber is also defined kinematically: , where is fixed by the imposed subtropical SST perturbation. {The meridional wavenumber is also related to (Harnik and Lindzen 2001).} Note that we consider vertically integrated transport; however, momentum and MSE transport by stationary eddies peak in the upper and lower troposphere, respectively, and applying the relationships at individual levels leads to similar conclusions.

Fig. 8.
Fig. 8.

Response of vertically integrated stationary eddy (black lines) (a),(b) momentum and (c),(d) MSE transport to (a),(c) land and (b),(d) ocean warming. The momentum transport is decomposed into squared eddy streamfunction (red line) and meridional wavenumber (blue line) changes [see (10)]. The MSE transport is decomposed into diffusivity (red line) and MSE gradient (blue line) changes [see (11)].

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

According to linear wave and mixing length theories, stationary eddy transport changes can be decomposed as follows:
e10
and
e11
Across the range of imposed regional climate perturbations, we find the linearization is a good approximation (cf. colored and black lines in Figs. 8 and 10). When the decomposition in (10) and (11) is applied to the stationary eddy transport response to land and ocean warming the results show changes in streamfunction and diffusivity dominate the transport response (Fig. 8, red lines). The changes of meridional wavenumber and meridional MSE gradient (Fig. 8, blue lines) are much weaker. Thus, subtropical stationary eddy transport changes do not conform to a fixed-diffusivity framework in agreement with results for tropical stationary eddies (Lee 2014). As discussed previously, the stationary eddy amplitude changes reflect the divergent flow response to regional climate change, which is coupled to the rotational flow [see (6)]. Consistent with the dominance of stationary eddy streamfunction changes, the dynamical component of the MSE and moisture transport responses to regional warming is much larger than the thermodynamic component (cf. Seager et al. 2010).

The subtropical stationary eddy transport changes in response to regional warming coincide with significant changes in vertically integrated transient eddy momentum and MSE transport (blue lines in Fig. 9). In the NH subtropics, transient eddy transport changes are similar in magnitude but opposite in sign to the stationary transport changes suggesting compensation. For example, subtropical stationary eddy transport increases in response to land warming, whereas transient eddy transport decreases and vice versa for ocean warming (cf. solid black and blue lines in Fig. 9).

Fig. 9.
Fig. 9.

Response of vertically integrated stationary eddy (black lines) and transient eddy (blue lines) (a),(b) momentum and (c),(d) MSE transport to (a),(c) land and (b),(d) ocean warming. Dashed lines indicate transport averaged over the Asia–Pacific sector.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

Stationary and transient eddies can interact in many ways. For example, stationary eddies can shift the critical line for transient eddies in the subtropical upper troposphere (Shaw 2014) and stationary eddies can change near-surface MSE gradients that affect transient eddy transport (Kaspi and Schneider 2013). When we apply the linear wave and mixing length decomposition to transient eddy transport changes we find that wavenumber changes , where and are the transient eddy zonal and meridional wavenumbers, are important for momentum transport changes (see blue lines in Figs. 10a,b) and transient eddy diffusivity changes are important for MSE transport changes (see red lines in Figs. 10c,d). This suggests momentum compensation occurs via eddy propagation (wavenumber changes) and MSE compensation occurs via diffusivity (eddy amplitude) changes.

Fig. 10.
Fig. 10.

As in Fig. 7, but for the transient eddy response.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

Here we quantify the degree of transient eddy compensation by comparing the subtropical averaged transport. The subtropical averaged transient eddy momentum transport response to land and ocean warming compensates approximately 70%–90% of the stationary eddy transport response (Table 2). Along similar lines, the transient eddy MSE transport response to land and ocean warming compensates approximately 90% of the subtropical stationary eddy transport response (Table 2). Compensation does not occur at each longitude but does dominate in longitudinal sectors (e.g., the Asia–Pacific sector; dashed black and blue lines in Fig. 9).

Table 2.

Response of vertically integrated subtropically averaged (20°–40°N) stationary eddy () and transient eddy () momentum transport (1 × 1012 m2 Pa) and momentum compensation (mom comp) = 100 × . Response of vertically integrated subtropically averaged (20°–40°N) stationary eddy () and transient eddy () MSE transport (PW) and MSE compensation (MSE comp) = 100 × .

Table 2.

A consequence of the transient eddy momentum and MSE transport changes is a meridional shift of the latitude of maximum transport (Figs. 3c,d; i.e., meridional shifts of the extratropical circulation). The shift of the latitude of maximum transport in response to regional climate change depends on the meridional gradient of transport change at the latitude of maximum transport in the control climate. The poleward shift of the latitude of maximum transient eddy momentum and MSE transports in response to land warming is dominated by the negative transient eddy transport change in the NH subtropics and not by the positive transport change in high latitudes (Tables 3 and 4). Along similar lines, the equatorward shift of the latitude of maximum transient eddy momentum and MSE transports in response to ocean warming is dominated by the positive transport change in the NH subtropics (Tables 3 and 4). Note that high-latitude transport changes are important for the flux convergence response.

Table 3.

Response of vertically integrated transient eddy momentum transport () to land and ocean warming. Shift of latitude of maximum transient eddy momentum transport . Shift of latitude of maximum transient eddy momentum transport due solely to negative transport changes (negative portion of blue curve in Figs. 9a,b) and positive transport changes (positive portion of blue curve in Figs. 9a,b) . Finally, shift of latitude of maximum transient eddy transport momentum assuming subtropical compensation [Eq. (12)] and percent compensation.

Table 3.
Table 4.

As in Table 3, but for the response of vertically integrated transient eddy MSE transport () to land and ocean warming. Shift of latitude of maximum transient eddy MSE transport . Shift of latitude of maximum transient eddy MSE transport due solely to negative tranport changes (negative portion of blue curve in Figs. 9c,d) and positive transport changes (positive portion of blue curve in Figs. 9c,d) . Finally, shift of latitude of maximum transient eddy transport MSE assuming subtropical compensation [Eq. (13)] and percent compensation.

Table 4.

We quantitatively assess the impact of compensation on the shift of the latitude of maximum transient eddy transport as follows. If we assume the transient eddy transport changes in the NH subtropics, responsible for the shift of the latitude of maximum transport, exactly compensate for stationary eddy transport changes in response to regional climate change then we have the following:
e12
and
e13
We apply this assumption in latitudes where the stationary eddy and transient eddy transport are of opposite sign (i.e. NH subtropics). Upon adding the assumed transient eddy transport changes in (12) and (13) to the transient eddy transport from the control climate we then calculate the latitude of maximum transport and infer a shift of the latitude of maximum transient eddy transport in response to regional warming. Assuming compensation quantitatively accounts for approximately 70%–90% of the meridional shift of the latitude of maximum transient eddy transport (Tables 3 and 4). While compensation does a reasonable job of explaining the meridional shifts of the latitude of maximum transport, it does not explain the dipole structure of the transient eddy transport response (blue lines in Fig. 9). This dipole is likely a result of internal eddy–zonal flow feedbacks (e.g., Lorenz and Hartmann 2001; Boer et al. 2001), which are known to impact the extratropical response to warming (Chen et al. 2013; Lu et al. 2014).

4. Conclusions and discussion

The subtropical and extratropical circulation in the NH exhibit significant changes in response to seasonal insolation and recent work has shown that a similar connection exists between the subtropical and extratropical summertime circulation response to direct radiative forcing and indirect SST warming (Shaw and Voigt 2015). Here we examined the mechanisms linking the subtropical and extratropical summertime circulation response to climate change using idealized prescribed SST aquaplanet simulations with a subtropical perturbation that mimics land–ocean contrast. While highly idealized, a prescribed SST setup is useful for examining the circulation responses because regional climate change can be controlled directly via SST perturbations. Here are the conclusions:

  • Global SST warming produces a poleward jet shift and Hadley cell expansion even in the presence of subtropical stationary eddies consistent with zonally symmetric aquaplanet models (Medeiros et al. 2015). Global warming of equivalent potential temperature, which accounts for changes in moisture via the Clausius–Clapeyron relation, and direct radiative forcing produce very weak circulation shifts. Stationary eddies do not exhibit a significant response to global SST warming, global equivalent potential temperature warming, or direct radiative forcing in prescribed SST aquaplanet simulations.
  • Regional warming can produce circulation responses that are opposite to those in response to global warming. For example, ocean warming shifts the jet equatorward whereas land warming shifts the jet poleward in aquaplanet simulations, consistent with the summertime Pacific jet stream response in AGCMs (Shaw and Voigt 2015). This is true for zonally symmetric and zonally asymmetric background SSTs.
  • Regional warming produces large changes in stationary eddy amplitude. Wave amplitude increases (decreases) in response to equivalent potential temperature warming over land (ocean), consistent with comprehensive models. The stationary eddy amplitude response depends on changes in near-surface moist entropy following quasi-equilibrium dynamics (Emanuel 1995) and changes in energy input to the atmosphere and GMS following the MSE budget (Neelin and Held 1987) in the region of climatological convergence. Eddy amplitude increases (decreases) with increased (decreased) near-surface moist entropy or energy input. The circulation weakens in response to ocean warming because of increased GMS, consistent with previous work (Chou and Chen 2010; Chou et al. 2013; Merlis et al. 2013a,b).
  • Latent heating dominates the stationary eddy response to regional warming. Cloud and water vapor radiative feedbacks enhance the poleward shift of the NH jet stream in response to land warming by 20% but do not significantly impact the jet stream shift in response to ocean warming.
  • Subtropical stationary eddy transport changes in response to regional warming do not follow a fixed-diffusivity approximation. In particular, subtropical stationary eddy momentum and MSE transport responses to regional warming are dominated by changes in stationary eddy amplitude (diffusivity), not by changes in moist entropy gradient or changes in meridional wavelength. Thus, when wave amplitude increases (decreases), stationary eddy transport increases (decreases).
  • Transient eddy momentum and MSE transport changes largely compensate for changes in subtropical stationary eddy transport in response to regional warming. In particular, in response to land warming, increased subtropical stationary eddy transport leads to a reduction of subtropical transient eddy transport. Compensation is due to changes in eddy propagation (wavenumber changes) and eddy diffusivity (eddy amplitude). Quantitatively, compensation accounts for ~70%–90% of the transient eddy transport response.
  • Compensation can be used to predict the direction of extratropical circulation shift in response to regional climate change. In particular, the assumption of perfect compensation of stationary eddy transport by transient eddy transport accounts for ~70%–90% of meridional shift of the latitude of maximum transient eddy momentum and MSE transport in prescribed SST aquaplanet simulations. The meridional shift of the latitude of maximum transient eddy transport is connected to the meridional shift of the jet stream and storm track.

We have focused on the interaction between subtropical stationary eddies and transient eddies in response to regional climate change. This interaction occurs during summertime when subtropical monsoon cyclones and oceanic anticyclones are large amplitude. Previous work has found compensation between extratropical stationary and transient eddies during wintertime in response to changes in orography (Manabe and Terpstra 1974; Park et al. 2013) and in response to zonally asymmetric ocean heat flux convergence (Kaspi and Schneider 2013). Compensation assumes small changes of total MSE transport, which has been connected to the weak latitudinal dependence of outgoing longwave radiation (Stone 1978) and can be predicted by energy balance models (Frierson et al. 2007); however, its connection to the momentum budget requires further investigation.

Our results have shown that the circulation response to climate change is similar for equivalent zonally symmetric and zonally asymmetric background climates under certain circumstances. For example, global warming produces a poleward shift with and without subtropical stationary eddies in the background state. Similarly there is an equatorward shift in response to ocean warming irrespective of the background SST. In contrast land warming did not produce equivalent circulation responses for symmetric and asymmetric background SSTs, which suggests zonally averaged concepts do not apply exactly when climate change is localized (small projection onto zonal mean) and generates stationary eddies.

We have assessed the connection between the subtropical and extratropical circulation response to regional climate change using prescribed SST aquaplanet model simulations. The prescribed SST aquaplanet was a useful framework because near-surface moist entropy contrasts, which are important in AGCM simulations (Shaw and Voigt 2015), could be controlled directly. While the connection between the subtropical stationary eddy and extratropical circulation response to regional climate change in the aquaplanet agrees with the AGCM response, there are several limitations worth mentioning. The prescribed SST framework used here assumes a saturated land surface without topography. An assessment of the circulation response with a real land surface, where temperature and saturation (where temperature and moisture responses can be of opposite sign) vary with radiative forcing, is needed. Our results show that stationary eddies respond weakly to direct radiative forcing, which suggests that the response of stationary eddies to direct radiative forcing in AGCM simulation may be dominated by land adjustment. Future work is focused on testing the theoretical relationships developed here in more realistic aquaplanet configurations including those with idealized land and topography and coupling with SSTs (e.g., Xie and Saiko 1999; Brayshaw et al. 2008; Merlis et al. 2013a,b). Previous work has shown that cloud radiative feedbacks and surface heat fluxes significantly impact the circulation response to subtropical zonally asymmetric ocean heat transport perturbations (Shaw et al. 2015).

We have shown that combining the subtropical stationary circulation response to regional climate change predicted from quasi-equilibrium dynamics or the MSE framework with an assumption of transient eddy compensation leads to a prediction of the sign of the extratropical circulation response to regional warming, which is useful for understanding circulation shifts in future climates and potentially also in past climates.

Acknowledgments

TAS and AV are supported by the David and Lucile Packard Foundation. TAS is also supported by NSF Award AGS-1538944 and the Alfred P. Sloan Foundation. TAS thanks O. Pauluis, M. Cane, and I. Held for helpful discussions. We also thank the reviewers whose comments helped to improve the manuscript. The simulations in this paper were completed with resources provided by the University of Chicago Research Computing Center. The model data in this study are available from TAS upon request.

APPENDIX A

Aquaplanet Model Simulation

The aquaplanet model simulations involve prescribed zonally symmetric and zonally asymmetric SST patterns. The zonally symmetric SST (Fig. 1a) follows the Qobs distribution of Neale and Hoskins (2000) with maximum shifted to 10°N:
ea1
where λ is longitude in radians, ϕ is latitude in radians, , and . The zonally asymmetric SST perturbation (Fig. 1b) is given by
ea2
where , , , and . This represents a zonal wavenumber-1.5 perturbation between 0° and 240°E and a zonal wavenumber-3 perturbation between 240° and 360°E. There is no seasonal cycle of insolation (permanent equinox conditions). Each simulation is run with spectral T63 resolution for 10 yr with the last 8 yr used for diagnostics. We used 6-hourly output for calculating the eddy transport.

The circulation response to the zonally asymmetric SST perturbation (Fig. 1b) in the presence of the Qobs background SST (Fig. 1a) is shown in Figs. 2a and 2d. The response in the presence of a globally uniform 300-K SST is shown in Fig. A1. The 300-K uniform background assesses the response in the absence of a meridional SST gradient (e.g., in the absence of a mean meridional circulation and jet stream). It is analogous to a moist Gill model response. The circulation response with and without the background meridional SST gradient (Figs. 2a–d vs Fig. A1) reveals the importance of the mean meridional circulation and jet stream in shaping the circulation over the regions labeled North America and Atlantic. The perturbations mimicking regional climate change are shown in Fig. A2.

Fig. A1.
Fig. A1.

As in Figs. 2a–d, but for a 300-K zonally symmetric background SST.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

Fig. A2.
Fig. A2.

Zonal mean of SST perturbations GWARM (red), GWARM for equivalent potential temperature (black), LWARM (brown), OWARM (blue) that mimic surface warming. See Table 1 for simulation description. Response of near-surface (925 hPa) equivalent potential temperature to (b) LWARM + 2 K, (c) LWARM + 4 K, (d) OWARM + 2 K, (e) OWARM + 4 K, and (f) warming of equivalent potential temperature by 2 K. Contour interval in (b),(d),(f) is 1 K and in (c),(e) 2 K.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

APPENDIX B

Radiative Feedbacks

The impact of water vapor and cloud radiative feedbacks on the circulation response to regional warming can be assessed using the cloud and water vapor locking technique (Voigt and Shaw 2015). The technique disables radiative feedbacks from clouds and water vapor by locking the radiative properties of clouds and water vapor to those in the control simulation. Figure B1 shows the difference of the circulation response to land and ocean warming with and without cloud and water vapor radiative feedbacks. When cloud and water vapor radiative feedbacks are disabled the NH jet stream shifts 1.8° latitude northward in response to land warming as compared to a 2.2° shift when the feedbacks are enabled. The NH jet shifts 1.5° equatorward when cloud and water vapor radiative feedbacks are disabled as compared to a 1.5° shift when the feedbacks are enabled. Thus, cloud and water vapor radiative feedbacks account for ~20% of the northward jet stream shift in response to land warming but do not contribute significantly to the equatorward shift in response to ocean warming.

Fig. B1.
Fig. B1.

As in Fig. 5, but for the difference between the response with and without cloud and water vapor radiative feedbacks.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

APPENDIX C

Linear Boundary Layer Model

Consider a linear boundary layer momentum balance involving Coriolis, pressure gradient, and frictional forces:
ec1
where f is the Coriolis parameter, u is the horizontal velocity (), Φ is the geopotential, and r is the inverse of the frictional time scale. Upon decomposing the flow into rotational and divergent components , where , we obtain
ec2
where we have used . In matrix form (C2) is
ec3
and the solution is
ec4
where
ec5
The components of divergence are as follows:
ec6
and
ec7
After making use of the following relations and we obtain
ec8
where we have ignored higher-order terms in the parameter (r2 + f 2)−1. Note that (C8) reduces to Sverdrup balance when r = 0. At the equator where f = 0 convergence occurs in the presence of easterly (westward) flow consistent with ITCZ dynamics. At the equator there is a singularity in the boundary layer depth that can be alleviated using a stably stratified boundary layer model. In the stratified boundary layer model the tropical boundary layer depth is constant (Schneider and Lindzen 1976), and this underlies the success of models of tropical convergence (Lindzen and Nigam 1987; Back and Bretherton 2009).

Figure C1 shows the approximate divergence in (C8) for the control simulation with background zonally asymmetric SST, where the parameter r (unit day−1) is determined by a linear regression at each grid point. Ekman balance dominates the terms in (C8) (Fig. C2).

Fig. C1.
Fig. C1.

(a) Near-surface divergence and (b) approximate divergence following (C8) for ASYM background SST. Contour interval is 0.5 × 10−6 s−1.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

Fig. C2.
Fig. C2.

Contributions to approximate divergence [(C8)] from (a) Ekman, (b) Sverdrup, and (c) equatorial convergence for the climatology shown in Fig. C1b. Contour interval is 0.5 × 10−6 s−1.

Citation: Journal of Climate 29, 18; 10.1175/JCLI-D-16-0049.1

If we nondimensionalize (C8) using
ec9
then the first and third terms on the right-hand side scale as βL/f. Hence, Ekman balance holds if L < f/β, which is consistent with the warm SST perturbations in Fig. 1b in the subtropics.

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1

Global warming of equivalent potential temperature () is achieved by assuming a near-surface water vapor increase of 7% K−1 following the Clausius–Clapeyron relation. For example, the SST perturbation for uniform global warming of equivalent potential temperature is determined by solving for ΔSST in the following equation: , where is the latent heat of vaporization, is the near-surface specific humidity in the control climate, and is the specific heat at constant pressure.

2

Uniform SST warming does not introduce a near-surface temperature gradient; however, it does introduce a near-surface moisture and thus equivalent potential temperature gradient. The uniform equivalent potential temperature warming experiment assesses the impact of moisture gradient changes.

3
The subcloud entropy is defined at 925 hPa. Determining the zonally varying critical subcloud moist entropy is a difficult task in general; thus, to simplify things we define
eq1
[see Eq. (27) in Emanuel (1995)]. Since the critical subcloud moist entropy is exceeded in regions of cyclonic near-surface vorticity.
4

Here we define the gross moist stability as the difference between the tropopause mt and near surface ms (925 hPa) MSE multiplied by the pressure difference between the near surface and tropopause.

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