1. Introduction
As the climate warms due to increased greenhouse gases in the atmosphere, the atmospheric general circulation is expected to change. Climate model simulations have found a weakening of the tropical atmospheric circulation (Held and Soden 2006; Vecchi and Soden 2007), a poleward expansion of the Hadley cell (Lu et al. 2007), a poleward shift of the tropospheric zonal jets (Kushner et al. 2001; Lorenz and DeWeaver 2007) and the midlatitude storm tracks (Yin 2005), as well as a rise in tropopause height (Kushner et al. 2001; Lorenz and DeWeaver 2007). These circulation changes have also been noticed in observational analyses for recent decades (e.g., Hu and Fu 2007; Chen and Held 2007). Stratospheric ozone depletion in the second half of the twentieth century might dominate over the role of CO2 increase in explaining Southern Hemisphere (SH) trends (Polvani et al. 2011; McLandress et al. 2011), and there is a possible contribution from natural variability in both hemispheres (e.g., Seager and Naik 2012).
Some mechanisms have been proposed to understand the cause for the extratropical circulation response to global warming. Lorenz and DeWeaver (2007) suggested that the midlatitude circulation response is predominantly driven by a rise in tropopause height based on the similarities in extratropical circulation response between a simple dry general circulation model (GCM) when the tropopause height is raised and the global warming simulations of models participating in the Coupled Model Intercomparison Project phase 3 (CMIP3) and assessed by the Intergovernmental Panel on Climate Change Fourth Assessment Report (IPCC AR4). Lu et al. (2008) proposed two possible mechanisms for the zonal mean circulation response to global warming by analyzing the CMIP3–IPCC AR4 models. The first mechanism suggests that the rising tropospheric static stability stabilizes the subtropical jet streams on the poleward flank of the Hadley cell, shifting the Hadley cell, the baroclinic instability zone, and the midlatitude eddies poleward. The second mechanism points to the importance of the increased phase speed of the midlatitude eddies. They suggested that the strengthened midlatitude wind in the upper troposphere and lower stratosphere, as a result of enhanced tropical upper-tropospheric warming and/or stratospheric cooling along the sloped tropopause, accelerates the eastward phase speeds of the midlatitude eddies, shifting the subtropical breaking region and the transient eddy momentum flux convergence and surface westerlies poleward. Butler et al. (2010) prescribed a heating in the tropical troposphere in a simple atmospheric GCM and found similar poleward jet and storm-track displacements as in the CMIP3–IPCC AR4 models, suggesting that the tropical upper-troposphere heating drives the circulation response to climate change. Kidston et al. (2010, 2011) found a robust increase in eddy length scale in the CMIP3–IPCC AR4 models, which is possibly caused by increased static stability in the midlatitudes. They argued that the increase in eddy length scale is a possible cause of the poleward shift of the eddy-driven jets and surface westerlies by reducing the eddy phase speed relative to the mean flow on the poleward flank of the jets and shifting the dissipation and eddy source regions poleward.
In addition, the stratosphere and its coupling with the troposphere have also been found to be important in determining the circulation response in the troposphere to global warming. Sigmond et al. (2004) studied the climate effects of middle-atmospheric and tropospheric CO2 doubling separately using the ECHAM middle atmosphere climate model with prescribed sea surface temperatures (SSTs). They found strengthened Northern Hemisphere (NH) midlatitude tropospheric westerlies as a consequence of a uniform CO2 doubling everywhere in the atmosphere and attributed this mainly to the middle atmosphere CO2 doubling.
The mechanisms mentioned above emphasize the close link between the thermal structure and circulation responses to global warming and suggest the warming in the mid- and upper troposphere and/or the cooling in the stratosphere as possible causes. The stratospheric cooling is caused directly by increased emission due to increased CO2, while the mid- and upper-tropospheric warming in the tropics arises from increased boundary layer temperature and humidity and a shift to a warmer moist adiabatic lapse rate (e.g., Hansen et al. 1984; Held 1993). This explanation for the tropospheric warming is essentially the same as that for the enhanced tropical upper-tropospheric warming during El Niños. However, in contrast to the broad warming response under global warming, the heating in the atmosphere during El Niño events is confined in the tropics and anomalous cooling occurs in the midlatitude troposphere induced by anomalous eddy-driven ascending motion (Seager et al. 2003). Also, the Hadley cell strengthens and narrows, and the tropospheric jets and midlatitude transient eddies shift equatorward in response to El Niños (see also Lu et al. 2008). In contrast, the warming in the mid- and upper troposphere in response to global warming, as simulated by the CMIP3–IPCC AR4 models (e.g., Fig. 10.7 in Meehl et al. 2007a), expands beyond the tropical convective region to about 40°N(S). It is not clear what causes the warming expansion into the extratropics.
In this study, we investigate the transient atmospheric adjustment to an instantaneous doubling of CO2. The response is investigated using the National Center for Atmospheric Research (NCAR) Community Atmosphere Model, version 3 (CAM3), coupled to a Slab Ocean Model (SOM). In contrast to previous studies on the equilibrium response to global warming (e.g., Hansen et al. 1984; Manabe et al. 1990; Meehl and Washington 1996; Shindell et al. 2001; Sigmond et al. 2004; Held and Soden 2006; Meehl et al. 2007a; Lu et al. 2008), our work focuses on the transient evolution that allows an assessment of the sequence of cause and effect in the circulation and thermal structure response prior to the establishment of a quasi-equilibrium state. Since the actual rate of anthropogenic CO2 increase is slow compared to the instantaneous CO2 doubling in our model experiments, the instantaneous CO2 doubling framework may not be strictly comparable to that in the actual response to global warming in every aspect. However, we demonstrate that our simulations in both transient and equilibrium states agree well with that from the CMIP3–IPCC AR4 models in which the CO2 concentration is gradually increased. Therefore, we believe that the transient atmospheric adjustment to instantaneous CO2 doubling provides valuable insight into the actual mechanisms underlying the extratropical tropospheric circulation response to global warming. In the paper, the following questions will be addressed: 1) what gives rise to the broad warming in the mid- and upper troposphere between 40°S and 40°N? and 2) what are the dynamical mechanisms involved in the extratropical circulation response to increased greenhouse gases? First, we describe the model and numerical experiments in section 2. The quasi-equilibrium response in thermal structure and circulation is presented in section 3. Furthermore, section 3 also presents the transient evolution step by step, and in particular the diagnostics of the cause of the broad warming expansion in the extratropical mid- and upper troposphere. Finally, a mechanism of the extratropical tropospheric circulation response to increased CO2 is proposed. Section 4 extends the analysis of the linkage between the eddy-driven vertical motion anomaly and the warming expansion in the subtropical mid- and upper troposphere to 14 CMIP3–IPCC AR4 coupled models. Discussions and conclusions are presented in section 5. In Wu et al. (2012, manuscript submitted to J. Climate; hereafter Part II), we will mainly focus on the transient, sequential, response day by day before the structure of the extratropical tropospheric circulation response is established—in particular, the perturbations in both the stratosphere and the troposphere and their coupling.
2. Model experiments
a. Model description
The NCAR CAM3 is a three-dimensional atmospheric general circulation model (AGCM), which includes the Community Land Model, version 3 (CLM3), an optional Slab Ocean Model, and a thermodynamic sea ice model. There are substantial modifications in the physics and dynamics of CAM3 from the previous version, Community Climate Model version 3 (CCM3), a detailed description of which is in Collins et al. (2006). CAM3 includes options for Eulerian spectral, semi-Lagrangian, and finite-volume formulations of the dynamical equations. The implementation of CAM3 with T85 spectral dynamics is the version used in the Community Climate System Model, version 3 (CCSM3), which is a fully coupled climate model for the CMIP3–IPCC AR4. CAM3 includes revised parameterizations of cloud condensation and precipitation processes as well as for radiative processes and atmospheric aerosols. The changes to the model lead to a more realistic tropical upper-troposphere temperature, a less pronounced double intertropical convergence zone, and an improved simulation of tropical continental precipitation. However, biases remain, such as the underestimation of the tropical variability associated with the Madden–Julian oscillation, the underestimation of the implied oceanic heat transport in the SH, excessive midlatitude westerlies, and surface stress in both hemispheres (Collins et al. 2006; Hurrell et al. 2006; Rasch et al. 2006).
b. Experimental design
A control experiment of CAM3–SOM is run for 140 yr with the CO2 concentration fixed at 355 ppmv. The year-by-year evolution of the global annual mean surface temperature (Ts) is shown in Fig. 1a (gray line) and has an average value of 288.5 K. The model asymptotes toward a quasi-equilibrium state after approximately 20 years (not shown).
(a) The global annual mean Ts for the control experiment for 140 yr (gray lines), 10 of the 100 1CO2 climatological runs (each for 22 yr) (blue lines), and instantaneous 2CO2 runs (each for 22 yr) (red lines). (b) As in (a), except that the time series are shifted to the same starting year (year 1) and last for 22 yr.
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
Using 1 January of each year of the last 100 yr of the control experiment as initial conditions, we generated a 100-member ensemble of single- and doubled-CO2 pair runs (1CO2 and 2CO2, respectively). The 1CO2 run is the same as the control experiment and keeps the CO2 level constant at 355 ppmv and is integrated forward for 22 yr. The 2CO2 experiment is a branch model run lasting for 22 yr as well and doubles the CO2 concentration instantaneously to 710 ppmv at the beginning of the experiment on 1 January. The difference between the 1CO2 run and the instantaneous 2CO2 run provides the atmospheric response to an instantaneous doubling of CO2. The ensemble average across the 100 runs to a large extent removes the model’s internal variability and allows for an assessment of the day-to-day adjustment of the atmospheric general circulation. Several variables, such as zonal and meridional winds, temperature and specific humidity, are output daily for the first 2 yr of the model integration. This methodology has been applied successfully to the study of cause and effect in the tropospheric response to El Niño and tropical Atlantic SST anomalies (Seager et al. 2009, 2010a,b; Harnik et al. 2010).
3. Results
a. Global mean response
Figure 1 shows the year-by-year evolution of the global annual mean Ts for the 1CO2 runs (blue lines) and the 2CO2 runs (red lines) for 10 of the 100 ensemble runs. The global annual mean Ts immediately increases by about 0.5 K in the first year after the doubling of CO2 on 1 January. After about 20 years, the 2CO2 runs reach a quasi-equilibrium state with Ts asymptoting toward an increase of 2.2 K (shown in Figs. 1a and 1b).
The CO2 forcing and the model’s climate sensitivity are also examined in the 2CO2 runs. Following Gregory et al. (2004), a scatterplot of the ensemble mean change in global annual mean Ts and the change in global annual mean net radiative flux at the top of the atmosphere (TOA) for the 22 yr of integration is shown in Fig. 2. The intercept of the regression line provides an estimate for the CO2 forcing at the time of doubling, F2× = 3.33 W m−2, and the slope indicates the climate response parameter α = 1.54 W m−2 K−1. In Gregory and Webb (2008), they found a doubled CO2 forcing of 2.93 ±0.23 W m−2 and a climate feedback parameter of 1.1 W m−2 K−1 for the CCSM3 T85 slab ocean model. The two results generally agree with each other despite different horizontal resolutions.
Scatterplot of the year-to-year change in global annual mean Ts and net radiative flux at TOA for the 22 yr of 2CO2 integration. It provides an estimate for the doubling CO2 forcing F2× = 3.33 W m−2 and the climate sensitivity of about 2.2 K.
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
b. Equilibrium response
As shown in Fig. 1, the 2CO2 simulations reach equilibrium after about 20 years. Figure 3 shows the equilibrium response in zonal mean temperature (T), zonal wind (u), transient eddy momentum flux (
The 2CO2 equilibrium response in (a),(b) 
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
The responses in transient eddy momentum flux and variance of meridional velocity include a prominent poleward and upward shift, especially in the upper troposphere and lower stratosphere. There is also an intensification in 
The response in transient eddies agrees well with the temperature anomaly and the change in linear baroclinic instability in CAM3–SOM. The largest increase in meridional temperature gradient occurs in the midlatitude upper troposphere and lower stratosphere. This is consistent with the strengthened transient eddies in this region. The close linkage between the thermal structure change and the circulation response to increased greenhouse gases has also been found in other studies (e.g., Yin 2005; Wu et al. 2010; O’Gorman 2010; Butler et al. 2010). Because neither daily variables nor monthly covariances in the NCAR CCSM3 coupled model are available for the CMIP3–IPCC AR4 experiments, the transient eddy activity and its future projections in the coupled model cannot be assessed and compared with our results.
c. Transient atmospheric adjustment and thermodynamics
1) Transient response
Figures 4– 6 show the month-by-month evolution of 
The month-by-month transient response in 
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
As in Fig. 4, but for 
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
As in Fig. 4, but for 
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
Figures 7a and 7b show the day-by-day evolution of the zonal mean temperature and zonal wind averaged over 30°–70°N from 1 January 1 to 30 April 30 of year 1 as a function of time and pressure level. The average over 30°–70°S is shown in Figs. 7c and 7d. The response is robust for different choices of latitudinal bands. A 5-day running average has been applied to all the variables. It is noted that the cooling in the stratosphere occurs first in the upper stratosphere and extends to the lower stratosphere in about a month. The substantial warming (0.5 K) in the mid- and upper troposphere takes place in early March. The eastward zonal wind anomaly clearly begins in the upper stratosphere and then gradually moves downward into the lower stratosphere and the troposphere, with the whole process taking about 100 days. The succession of events, first happening in the stratosphere and subsequently in the troposphere, resembles that in observations of subseasonal to seasonal variability (Baldwin and Dunkerton 2001) as well as in the “downward control” theory (Haynes et al. 1991).
The 2CO2 transient day-by-day response in (a),(c) 〈T〉 and (b),(d) 〈u〉 shown as a function of day (from 1 Jan to 30 Apr) and pressure level averaged between (a),(b) 30° and 70°N, and (c),(d) 30° and 70°S. A 5-day running average has been applied for plotting. The contour intervals are 0.25 K (−0.5 K) for positive (negative) values in (a),(c) and 0.25 m s−1 for (b),(d).
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
Figure 8 shows the day-by-day response in 
The 2CO2 transient day-by-day response in (a) 〈T〉, (b) 〈u〉, (c) total, and (d) bandpass-filtered eddy momentum flux convergence (defined in text) as a function of day (from 1 Jan to 30 Apr) and latitude. These variables are averaged from 150 to 500 mb, and a two-dimensional (latitude–pressure level) 1–2–1 smoothing and a 10-day temporal running average are applied. Gray shadings denote the 90% significance level. The contour intervals are (a) 0.25 K, (b) 0.25 m s−1, and (c),(d) 0.1 m s−1 day−1.
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
2) Thermodynamic diagnostics
As mentioned in the introduction, there is a close link and consistency between the tropospheric thermal structure and the circulation response in both the El Niños and the global warming scenario. During the El Niños, there is a confined tropical tropospheric warming and a cooling in the midlatitudes that is associated with an equatorward displacement of the tropospheric jets and midlatitude transient eddies. In contrast, as a result of the CO2 increase, the tropospheric warming extends broadly into the subtropics and the midlatitudes along with a poleward shift of the zonal wind and the storm tracks. Previous studies have demonstrated that the midlatitude cooling in response to the El Niños is driven by eddy-induced ascent anomaly (Seager et al. 2009, 2010a,b; Harnik et al. 2010). Therefore, what is responsible for the extensive tropospheric warming to CO2 increase?
Figure 9 shows the latitude–pressure level plot of the net temperature tendency 
(a) The actual zonal mean temperature tendency 
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
(a) The transient ωeddy (mb day−1) and (b) the actual vertical motion ω (mb day−1) from model output in March of year 1. Equation (3) for eddy-driven vertical motion is not applicable in the deep tropics, and thus regions between 10°S and 10°N are masked out in (a). The colors show the 2CO2 response, while the contours show the climatological results with a contour interval of 5 mb day−1. The positive (negative) values in ωeddy (ω) denote downward (upward) motion.
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
As shown in Fig. 8a, the subtropical warming expansion in the SH takes place in mid-February of year 1, albeit of small magnitude. A similar thermodynamic analysis has been done for February, and it is found that the SH subtropical warming at that time is also driven by enhanced transient eddy momentum flux convergence and resulting anomalous downward vertical motion (not shown).
The heating anomaly in the subtropical mid- and upper troposphere in this model experiment is induced by the dynamical circulation change rather than vice versa. It is the enhanced transient eddy momentum flux convergence in response to increased CO2 that causes anomalous descending motion and adiabatic heating in the subtropical mid- and upper troposphere. The dynamics of the changing transient eddies is closely connected with the response in the stratosphere and coupling between the stratosphere and the troposphere, and this will be further investigated in Part II.
d. Possible dynamical mechanisms
Based on the above diagnostic work, we propose a possible dynamical mechanism for the extratropical circulation response to increased CO2 with the following sequence:
- The CO2 doubling gives rise to a westerly zonal wind anomaly in the stratosphere. The westerly acceleration propagates downward into the lower stratosphere and upper troposphere.
- The westerly acceleration in the lower stratosphere and upper troposphere changes the propagation of baroclinic eddies, leading to enhanced transient eddy momentum flux convergence between 40°N(S) and 60°N(S).
- The increased transient eddy momentum flux convergence drives an anomalous mean meridional circulation in the troposphere as well as a poleward displacement of the tropospheric jets.
- The induced anomalous descending motion in the subtropical mid- and upper troposphere leads to an adiabatic heating anomaly and thus a broad warming expansion beyond the tropical convective region. The subtropical warming allows for adjustment to the thermal wind balance with the poleward-shifted jets.
A schematic figure showing the hypothesized sequence of the dynamical response is shown in Fig. 11. Other mechanisms are also possible. For example, it is expected that the increase in tropopause height could cause an increase in the length scale of transient eddies, which has been associated with a poleward jet shift (Williams 2006). The dynamical mechanisms of the transient adjustment and their cause and effect, explaining all possibilities, will be analyzed in detail in Part II. Here we provide a brief summary of Part II. Part II will emphasize the transient sequential response in both the stratosphere and the troposphere before the circulation change is well established in the extratropical troposphere. Three phases are defined during the period of transient adjustment. A fast, radiatively induced thermal response occurs during January (defined as phase 1) and an easterly anomaly is generated in the NH subpolar stratosphere. Phase 2, which covers February, features a westerly acceleration in the stratosphere, and this is driven dynamically by the intensified momentum flux convergence of planetary stationary eddies. A “downward propagation” of the westerly acceleration from the lower stratosphere to the troposphere is seen in March and April (phase 3), and this is followed by a poleward displacement of the tropospheric midlatitude jets. In this final phase, the transient eddies play an important role in shifting the position of the tropospheric zonal wind.
Summary of the proposed mechanisms causing the tropospheric extratropical circulation response to increased CO2 concentration.
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
4. Eddy-driven vertical motion in CMIP3–IPCC AR4 coupled models
The work so far has demonstrated the importance of the eddy-driven vertical motion in inducing the warming anomaly in the mid- and upper troposphere from our instantaneous CO2 doubling experiments in CAM3–SOM. This section extends the work to an ensemble of CMIP3–IPCC AR4 coupled models (Meehl et al. 2007b) and shows that the above conclusions also apply in these models. Because the CMIP3–IPCC AR4 Special Report on Emissions Scenarios (SRES) A1B experiments are quasi-equilibrium runs and the diabatic heating term is not available in the standard output, we cannot examine the causality sequence or close the zonal mean temperature equation as in the previous section. Instead, we calculate ωeddy from Eq. (3) using 
The 14 IPCC AR4 coupled models and their resolution for the atmospheric component used in this study.
Figure 12 shows the multimodel annual average of 
The late twenty-first-century trend in annual mean (a) 
Citation: Journal of Climate 25, 8; 10.1175/JCLI-D-11-00284.1
5. Discussions and conclusions
We have explored the transient evolution of the atmospheric adjustment to an instantaneous doubling of CO2 concentration. The sequence in the general circulation response in the atmosphere helps reveal the dynamical mechanisms underlying the equilibrium circulation response, for example, the poleward expansion of the Hadley Cell (Lu et al. 2007), and the poleward shift of the tropospheric jets and storm tracks (e.g., Kushner et al. 2001; Yin 2005), as found in CMIP3–IPCC AR4 models. In contrast to previous studies suggesting that the thermal forcing in the tropical upper troposphere drives the tropospheric circulation response (e.g., Butler et al. 2010), our results indicate that the broad warming expansion in the subtropical mid- and upper troposphere is a consequence of the circulation change. Enhanced transient eddy momentum flux convergence in the lower stratosphere and upper troposphere, possibly originating from the stratospheric westerly acceleration, drives an anomalous mean meridional circulation in the troposphere. The induced anomalous descending motion in the subtropical mid- and upper troposphere warms the air adiabatically. Afterward, the subtropical warming and the poleward displacement of the jets and the baroclinic eddies can potentially feedback positively onto each other via a poleward shift in the eddy generation region, leading to a further poleward shift of the jets and the eddies and a further warming expansion in the subtropical troposphere.
It is noted here that this study focuses primarily on the atmospheric transient adjustment process before there is any appreciable SST change. Certainly SST change in a longer time scale may cause further circulation change. Nonetheless, the transient adjustment shown here is similar to the results in CMIP3–IPCC AR4 fully coupled models. As also demonstrated in Lu et al. (2009), the observed widening of the tropics in the latter half of the twentieth century is entirely attributed to the direct effect in the atmosphere of changing trace gas concentrations rather than to SST warming.
Our results also show the sequence of the zonal wind anomaly in the vertical column of the atmosphere, indicating that the poleward displacement of the tropospheric jets follows the subpolar westerly anomaly in the stratosphere. It suggests the importance of the stratosphere, and the coupling between the stratosphere and the troposphere, in regulating the extratropical tropospheric circulation response to increasing CO2. A detailed analysis of the stratospheric response and the stratosphere–troposphere coupling, including how the response “migrates” downward into the troposphere and how the eddies respond step by step, will be further examined in Part II. It is noted here that our study intends to understand the circulation response and the dynamical mechanisms in CMIP3–IPCC AR4–like models, albeit most of the models have poorly resolved stratospheres. Some studies have argued that a well-resolved stratosphere is required to reproduce observed behavior (e.g., Shindell et al. 1999; Sassi et al. 2010). However, Sigmond et al. (2008) suggested that the atmospheric circulation response to CO2 doubling does not necessarily require a well-resolved stratosphere, but rather a realistic simulation of the zonal wind strength in the mid- and high-latitude lower stratosphere. The zonal mean zonal wind in CAM3 agrees with the reanalysis data in this region. The circulation response to a CO2 doubling in both the troposphere and the stratosphere in our results also agrees to a large extent with those from studies that used models with much finer vertical resolution in the middle atmosphere (e.g., Shindell et al. 2001; Sigmond et al. 2004). However, a model lid in the midstratosphere is known to impact the vertical propagation of stationary planetary scale waves during NH winter (Shaw and Perlwitz 2010; Sassi et al. 2010). Assessing the transient and equilibrium responses to CO2 doubling in a model with high vertical resolution and a high model lid height is the subject of future investigation. Finally, because our experiments double the CO2 concentration on 1 January, it would be interesting to change the time of CO2 doubling to see if the model responds differently. A set of experiments with an instantaneous CO2 doubling on 1 July is currently under investigation.
Acknowledgments
The authors greatly appreciate the useful views of two anonymous reviewers and Dr. Joe Kidston. We would also like to thank the Global Decadal Hydroclimate (GloDecH) group at Lamont-Doherty Earth Observatory and Columbia University for their comments and advice, in particular, Prof. Lorenzo Polvani and Dr. Michela Biasutti as well as Dr. Nili Harnik. We also thank Profs. Ngar-Cheung Lau, Geoffrey K. Vallis, and Benjamin Lintner for the helpful discussions. YW was supported by NASA headquarters under the NASA Earth and Space Science Fellowship Program Grant NNX08AU80H. RS, MT, and NN were supported by NOAA Awards NA08OAR4320912 and NA10OAR4320137, and NSF Awards ATM-0804107. TAS was supported by a postdoctoral fellowship from the Natural Sciences and Engineering Research Council of Canada.
REFERENCES
Baldwin, M. P., , and T. J. Dunkerton, 2001: Stratospheric harbingers of anomalous weather regimes. Science, 294, 581–584.
Boville, B. A., , and C. S. Bretherton, 2003: Heating and kinetic energy dissipation in the NCAR Community Atmosphere Model. J. Climate, 16, 3877–3887.
Butler, A. H., , D. W. J. Thompson, , and R. Heikes, 2010: The steady-state atmospheric circulation response to climate change–like thermal forcings in a simple general circulation model. J. Climate, 23, 3474–3496.
Chen, G., , and I. M. Held, 2007: Phase speed spectra and the recent poleward shift of Southern Hemisphere surface westerlies. Geophys. Res. Lett., 34, L21805, doi:10.1029/2007GL031200.
Collins, W. D., and Coauthors, 2004: Description of the NCAR Community Atmosphere Model (CAM3.0). NCAR Tech. Note NCAR/TN-464+STR, 226 pp.
Collins, W. D., and Coauthors, 2006: The formulation and atmospheric simulation of the Community Atmosphere Model Version 3 (CAM3). J. Climate, 19, 2144–2161.
Gregory, J. M., , and M. Webb, 2008: Tropospheric adjustment induces a cloud component in CO2 forcing. J. Climate, 21, 58–71.
Gregory, J. M., and Coauthors, 2004: A new method for diagnosing radiative forcing and climate sensitivity. Geophys. Res. Lett., 31, L03205, doi:10.1029/2003GL018747.
Hansen, J., , A. Lacis, , D. Rind, , G. Russell, , P. Stone, , I. Fung, , R. Ruedy, , and J. Lerner, 1984: Climate sensitivity: Analysis of feedback mechanisms. Climate Processes and Climate Sensitivity, Geophys. Monogr., Vol. 29, Amer. Geophys. Union, 130–163 pp.
Harnik, N., , R. Seager, , N. Naik, , M. Cane, , and M. Ting, 2010: The role of linear wave refraction in the transient eddy–mean flow response to tropical Pacific SST anomalies. Quart. J. Roy. Meteor. Soc., 136, 2132–2146, doi:10.1002/qj.688.
Haynes, P. H., , M. E. McIntyre, , T. G. Shepherd, , C. J. Marks, , and K. P. Shine, 1991: On the downward control of extratropical diabatic circulations by eddy-induced mean zonal forces. J. Atmos. Sci., 48, 651–678.
Held, I. M., 1993: Large-scale dynamics and global warming. Bull. Amer. Meteor. Soc., 74, 228–241.
Held, I. M., , and B. J. Soden, 2006: Robust response of the hydrological cycle to global warming. J. Climate, 19, 5686–5699.
Hu, Y., , and Q. Fu, 2007: Observed poleward expansion of the Hadley circulation since 1979. Atmos. Chem. Phys., 7, 5229–5236.
Hurrell, J. W., , J. J. Hack, , A. S. Phillips, , J. Caron, , and J. Yin, 2006: The dynamical simulation of the Community Atmosphere Model version 3 (CAM3). J. Climate, 19, 2162–2183.
Kidston, J., , S. M. Dean, , J. Renwick, , and G. K. Vallis, 2010: A robust increase in the eddy length scale in the simulation of future climates. Geophys. Res. Lett., 37, L03806, doi:10.1029/2009GL041615.
Kidston, J., , G. K. Vallis, , S. M. Dean, , and J. A. Renwick, 2011: Can the increase in the eddy length scale under global warming cause the poleward shift of the jet streams? J. Climate, 24, 3764–3780.
Kushner, P. J., , I. M. Held, , and T. L. Delworth, 2001: Southern Hemisphere atmospheric circulation response to global warming. J. Climate, 14, 2238–2249.
Lorenz, D. J., , and E. T. DeWeaver, 2007: Tropopause height and zonal wind response to global warming in the IPCC scenario integrations. J. Geophys. Res., 112, D10119, doi:10.1029/2006JD008087.
Lu, J., , G. A. Vecchi, , and T. Reichler, 2007: Expansion of the Hadley cell under global warming. Geophys. Res. Lett., 34, L06805, doi:10.1029/2006GL028443.
Lu, J., , G. Chen, , and D. M. W. Frierson, 2008: Response of the zonal mean atmospheric circulation to El Niño versus global warming. J. Climate, 21, 5835–5851.
Lu, J., , C. Deser, , and T. Reichler, 2009: Cause of the widening of the tropical belt since 1958. Geophys. Res. Lett., 36, L03803, doi:10.1029/2008GL036076.
Manabe, S., , K. Bryan, , and M. J. Spelman, 1990: Transient response to a global ocean–atmosphere model to a doubling of atmosphere carbon dioxide. J. Phys. Oceanogr., 20, 722–749.
McCaa, J. R., , M. Rothstein, , B. E. Eaton, , J. M. Rosinski, , E. Kluzek, , and M. Vertenstein, 2004: User’s guide to the NCAR Community Atmosphere Model (CAM 3.0). NCAR Tech. Rep., 85 pp.
McLandress, C., , T. G. Shepherd, , J. F. Scinocca, , D. A. Plummer, , M. Sigmon, , A. I. Jonsson, , and M. C. Reader, 2011: Separating the dynamical effects of climate change and ozone depletion. Part II: Southern Hemisphere troposphere. J. Climate, 24, 1850–1868.
Meehl, G. A., , and W. M. Washington, 1996: El Niño-like climate change in a model with increased atmospheric CO2 concentrations. Nature, 382, 56–60.
Meehl, G. A., and Coauthors, 2007a: Global climate projections. Climate Change 2007: The Physical Science Basis, S. Solomon et al., Eds., Cambridge University Press, 747–845.
Meehl, G. A., , C. Covey, , T. Delworth, , M. Latif, , B. McAvaney, , J. F. B. Mitchell, , R. J. Stouffer, , and K. E. Taylor, 2007b: WCRP CMIP3 multimodel dataset: A new era in climate change research. Bull. Amer. Meteor. Soc., 88, 1383–1394.
O’Gorman, P. A., 2010: Understanding the varied response of the extratropical storm tracks to climate change. Proc. Natl. Acad. Sci. USA, 107, 19 176–19 180.
Polvani, L., , D. Waugh, , G. Correa, , and S.-W. Son, 2011: Stratospheric ozone depletion: The main driver of twentieth-century atmospheric circulation changes in the Southern Hemisphere. J. Climate, 24, 795–812.
Rasch, P. J., and Coauthors, 2006: A characterization of tropical transient activity in the CAM3 atmospheric hydrologic cycle. J. Climate, 19, 2222–2242.
Reichler, T., , M. Dameris, , and R. Sausen, 2003: Determining the tropopause height from gridded data. Geophys. Res. Lett., 30, 2042, doi:10.1029/2003GL018240.
Sassi, F., , R. R. Garcia, , D. Marsh, , and K. W. Hoppel, 2010: The role of the middle atmosphere in simulations of the troposphere during Northern Hemisphere winter: Differences between high- and low-top models. J. Atmos. Sci., 67, 3048–3064.
Seager, R., , and N. Naik, 2012: A mechanisms-based approach to detecting recent anthropogenic hydroclimate change. J. Climate, 25, 236–261.
Seager, R., , N. Harnik, , Y. Kushnir, , W. Robinson, , and J. Miller, 2003: Mechanisms of hemispherically symmetric climate variability. J. Climate, 16, 2960–2978.
Seager, R., and Coauthors, 2009: Mexican drought: An observational modeling and tree ring study of variability and climate change. Atmósfera, 22, 1–31.
Seager, R., , N. Naik, , W. Baethgen, , A. Robertson, , Y. Kushnir, , J. Nakamura, , and S. Jurburg, 2010a: Tropical oceanic causes of interannual to multidecadal precipitation variability in southeast South America over the past century. J. Climate, 23, 5517–5539.
Seager, R., , N. Naik, , M. Ting, , M. A. Cane, , N. Harnik, , and Y. Kushnir, 2010b: Adjustment of the atmospheric circulation to tropical Pacific SST anomalies: Variability of transient eddy propagation in the Pacific–North America sector. Quart. J. Roy. Meteor. Soc., 136, 277–296, doi:10.1002/qj.588.
Seager, R., , N. Naik, , and G. A. Vecchi, 2010c: Thermodynamic and dynamic mechanisms for large-scale changes in the hydrological cycle in response to global warming. J. Climate, 23, 4651–4668.
Shaw, T. A., , and J. Perlwitz, 2010: The impact of stratospheric model configuration on planetary-scale waves in Northern Hemisphere winter. J. Climate, 23, 3369–3389.
Shindell, D. T., , R. L. Miller, , G. A. Schmidt, , and L. Pandolfo, 1999: Simulation of recent northern winter climate trends by greenhouse-gas forcing. Nature, 399, 452–455.
Shindell, D. T., , G. A. Schmidt, , R. L. Miller, , and D. Rind, 2001: Northern Hemisphere winter climate response to greenhouse gas, ozone, solar, and volcanic forcing. J. Geophys. Res., 106, 7193–7210, doi:10.1029/2000JD900547.
Sigmond, M., , P. C. Siegmund, , E. Manzini, , and H. Kelder, 2004: A simulation of the separate climate effects of middle-atmospheric and tropospheric CO2 doubling. J. Climate, 17, 2352–2367.
Sigmond, M., , J. F. Scinocca, , and P. J. Kushner, 2008: Impact of the stratosphere on tropospheric climate change. Geophys. Res. Lett., 35, L12706, doi:10.1029/2008GL033573.
Vecchi, G. A., , and B. J. Soden, 2007: Global warming and the weakening of the tropical circulation. J. Climate, 20, 4316–4340.
Williams, G. P., 2006: Circulation sensitivity to tropopause height. J. Atmos. Sci., 63, 1954–1961.
Wu, Y., , M. Ting, , R. Seager, , H.-P. Huang, , and M. A. Cane, 2010: Changes in storm tracks and energy transports in a warmer climate simulated by the GFDL CM2.1 model. Climate Dyn., 37, 53–72, doi:10.1007/s00382-010-0776-4.
Yin, J. H., 2005: A consistent poleward shift of the storm tracks in simulations of 21st century climate. Geophys. Res. Lett., 32, L18701, doi:10.1029/2005GL023684.
Without bandpass filtering.
The change in 
In this model experiment, the NH tropospheric jet shift in summer is very weak (in both the transient and equilibrium states), while that in winter is slightly stronger (shown in Figs. 3c and 3d). However, there is not much seasonal dependence for the tropospheric jet displacement in the SH.
The time filter used here is a standard 21-point two-sided bandpass filter that keeps the variability within 2–8 days. It skips the first and last 10 days in the time series of daily eddy momentum flux convergence.
There is a cosϕ term missing in the denominator of Eq. (7) in Seager et al. (2003).
Daily atmosphere data are output to standard levels up to 200 mb.
