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  • View in gallery

    Precipitation differences between experiments with 1 × CO2 and 2 × CO2 concentrations for different convection heights: (a) Ms − 1, (b) Ms + 0, and (c) Ms + 1. The thick dashed curves indicate a convective margin: that is, ω = 0 at 500 hPa. The unit is mm day−1.

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    Scatter plots for fractional changes of hydrological variables and the strength of tropical circulation averaged over convective regions between the last and first 20 yr of the twenty-first century for each of the climate models listed in Table 1: (a) PE vs the column-integrated convergence of the moisture flux −〈 · vq〉, (b) PE vs precipitation, (c) surface temperature vs PE, and (d) PE vs the strength of tropical circulation (vertical velocity at 500 hPa).

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    The anomalous pressure velocity between the last and first 20 yr of the twenty-first century. The anomalies are averaged over convective regions for the 16 CMIP3 models listed in Table 1. Positive values indicate downward motion (reduced) and negative values indicate upward motion (enhanced).

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    Fractional changes of precipitation and water vapor and tropical circulation in the QTCM1 experiments associated with different convection depths. The numbers are normalized by the change of global-mean surface temperature. Depth of convection becomes shallower from left to right. (a) Pressure velocity at 500 hPa, (b) evaporation averaged over convective regions (indicated by thick curves in Fig. 1), and global-mean (c) precipitation and (d) column-integrated water vapor.

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    The anomalous pressure velocity averaged over convective regions (indicated in Fig. 1) between experiments with 1 × CO2 and 2 × CO2 concentrations for different convection heights: Ms − 1, Ms + 0, Ms + 1, Ms + 2, and Ms + 3.

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Depth of Convection and the Weakening of Tropical Circulation in Global Warming

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  • 1 Research Center for Environmental Changes, Academia Sinica, and Department of Atmospheric Sciences, National Taiwan University, Taipei, Taiwan
  • 2 Department of Earth Sciences, National Taiwan Normal University, Taipei, Taiwan
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Abstract

Anthropogenic forcings, such as greenhouse gases and aerosols, are starting to show their influence on the climate, as evidenced by a global warming trend observed in the past century. The weakening of tropical circulation, a consequence of global warming, has also been found in observations and in twenty-first-century climate model simulations. It is a common belief that this weakening of tropical circulation is associated with the fact that global-mean precipitation increases more slowly than water vapor. Here, a new mechanism is proposed for this robust change, which is determined by atmospheric stability associated with the depth of convection. Convection tends to extend higher in a warmer climate because of an uplifting of the tropopause. The higher the convection, the more stable the atmosphere. This leads to a weakening of tropical circulation.

Corresponding author address: Chia Chou, Research Center for Environmental Changes, Academia Sinica, P.O. Box 1-48, Taipei 11529, Taiwan. Email: chiachou@rcec.sinica.edu.tw

Abstract

Anthropogenic forcings, such as greenhouse gases and aerosols, are starting to show their influence on the climate, as evidenced by a global warming trend observed in the past century. The weakening of tropical circulation, a consequence of global warming, has also been found in observations and in twenty-first-century climate model simulations. It is a common belief that this weakening of tropical circulation is associated with the fact that global-mean precipitation increases more slowly than water vapor. Here, a new mechanism is proposed for this robust change, which is determined by atmospheric stability associated with the depth of convection. Convection tends to extend higher in a warmer climate because of an uplifting of the tropopause. The higher the convection, the more stable the atmosphere. This leads to a weakening of tropical circulation.

Corresponding author address: Chia Chou, Research Center for Environmental Changes, Academia Sinica, P.O. Box 1-48, Taipei 11529, Taiwan. Email: chiachou@rcec.sinica.edu.tw

1. Introduction

Under global warming, the global-mean precipitation of general circulation model (GCM) simulations tends to increase at a rate of around 2% K−1, whereas the atmospheric water vapor increase roughly follows the Clausius–Clapeyron expansion of saturated water vapor with a constant relative humidity, which is around 7.5% K−1 (Allen and Ingram 2002; Held and Soden 2006; Stephens and Ellis 2008; Trenberth et al. 2003; Vecchi and Soden 2007; Wentz et al. 2007). Based on a simplified equation P = Mcqs, where P is precipitation, Mc is convective mass flux, and qs is a typical boundary layer mixing ratio, the fractional change of global-mean precipitation is roughly equal to the sum of the changes in convective mass flux and the boundary layer mixing ratio; that is, P′/P = Mc/Mc + qs/qs, where the overbar () is climatology and the prime (′) is the change. Because the water vapor increases much faster than the global-mean precipitation, tropical circulation should be weakened in the future warmer climate (Held and Soden 2006; Knutson and Manabe 1995; Vecchi and Soden 2007); that is, Mc/Mc < 0. The weakening of tropical circulation is very robust among the twenty-first-century GCM simulations and occurs most in a longitudinal direction (Held and Soden 2006; Vecchi and Soden 2007). A weakening of the Walker circulation has already been observed (Vecchi et al. 2006).

What kind of mechanisms could cause this weakening of tropical circulation? Two possible mechanisms have been proposed by Chou et al. (2009, hereafter C09): the upped-ante mechanism and the effect of deepened convection. The upped-ante mechanism is associated with a dry advection from subsidence regions to convective regions. This mechanism usually occurs over margins of convective regions. The effect of deepened convection, on the other hand, increases the gross moist stability and stabilizes the atmosphere. In this study, we would like to further examine the effect of deepened convection, which is commonly found under global warming. The model and data used here are briefly described in section 2. A vertically integrated water vapor budget used to determine possible contributions to precipitation changes is discussed in section 3. The effect of deepened convection has been examined by using an intermediate climate system in section 4, followed by discussion and summary.

2. The model and data

a. The quasi-equilibrium tropical circulation model

To examine the effect of atmospheric stability associated with convection depth on tropical circulation, a coupled ocean–atmosphere–land model of intermediate complexity (Neelin and Zeng 2000; Zeng et al. 2000) with prescribed divergence of ocean heat transport (Q flux) is used. Based on the analytical solutions derived from the Betts–Miller moist convective adjustment scheme (Betts and Miller 1993), typical vertical structures of temperature, moisture, and winds for deep convection are used as leading basis functions for a Galerkin expansion (Neelin and Yu 1994; Yu and Neelin 1994). The atmospheric model constrains the flow by quasi-equilibrium thermodynamic closures and is referred to as the quasi-equilibrium tropical circulation model with a single vertical structure of temperature and moisture for deep convection (QTCM1). Because the basis functions are based on vertical structures associated with convective regions, these regions are expected to be well represented and similar to a GCM with the Betts–Miller moist convective adjustment scheme. Far from convective regions, QTCM1 is a highly truncated Galerkin representation equivalent to a two-layer model. A cloud-radiation scheme (Chou and Neelin 1996; Zeng et al. 2000), which is simplified from the full radiation schemes (Fu and Liou 1993; Harshvardhan et al. 1987), is included. In this scheme, the deep and cirrocumulus–cirrostratus cloud fraction is estimated by an empirical parameterization (Chou and Neelin 1999).

An intermediate land surface model (Zeng et al. 2000) is used to simulate the interaction between the atmosphere and land surface. This model simulates processes such as evapotranspiration and surface hydrology in a single land surface layer for calculating energy and water budgets. Soil moisture is balanced by precipitation, evaporation, surface runoff, and ground runoff. This model does not include snow and sea ice feedback, so the surface warming at higher latitudes is weaker than in most coupled GCM simulations. A slab mixed layer ocean model with a fixed mixed layer depth of 50 m is used. The Q flux is obtained from a prescribed seasonal sea surface temperature (SST) run. In this study, a revised QTCM1 version 2.3 is used. The main improvement in the model physics is the inclusion of a simple ABL, which assumes a steady-state, vertically homogeneous mixed layer with fixed height (Stevens et al. 2002).

b. Vertical structure of vertical velocity

Under quasi-equilibrium convective closures, a typical vertical profile of deep convection is prescribed in QTCM1, so vertical velocity can be written as
i1520-0442-23-11-3019-e1
where ω is vertical velocity and · v1 is divergence induced by baroclinic winds. The vertical profile of vertical velocity Ω(p) is independent of time and horizontal position but depends on convection depth (Neelin and Yu 1994; Yu et al. 1998).
With the prescribed vertical profile of vertical velocity, the vertically integrated vertical advection of moisture and dry static energy can be written as
i1520-0442-23-11-3019-e2
i1520-0442-23-11-3019-e3
where q is specific humidity and the dry static energy is s = T + ϕ, with ϕ being the geopotential. Both moisture q and temperature T are in energy units by absorbing the latent heat per unit mass L and the heat capacity at constant pressure Cp, respectively; 〈·〉 denotes a mass integration through the entire troposphere; Ms is dry static stability,
i1520-0442-23-11-3019-e4
and Mq is gross moisture stratification,
i1520-0442-23-11-3019-e5
Thus,
i1520-0442-23-11-3019-e6
where moist static energy (MSE) is h = s + q and the gross moist stability is M = MsMq (Yu et al. 1998).

c. Experiment design

A set of experiments is designed to study the effect of convection depth. In the control experiment, the current climate with the normal CO2 concentration (i.e., 1 × CO2 and seasonal Q flux, obtained from a fixed SST run with observed seasonal climatology) is used. In the standard global warming experiment (Ms + 0), the same Q flux and a doubled CO2 concentration (i.e., 2 × CO2) are used. The convection height in these two experiments is fixed at 150 hPa. The differences between these two experiments are induced by the increase of the anthropogenic greenhouse gas, CO2. Because convection height is determined beforehand, such as in (1), and the associated effect is precalculated in QTCM1, convection depth can therefore be much more easily changed than in most coupled GCMs. In this study, four experiments with different depths of convection are conducted with the same setting as that of the standard global warming experiment (Ms + 0): that is, 2 × CO2 and the same Q flux. Changes in convection height affect 〈ωps〉 [i.e., Ms in QTCM1; see (3)] but have much less of an impact on 〈ωpq〉 because of little water vapor at upper troposphere. In other words, changes in Ms are roughly equivalent to changes in convection height. Thus, we varied Ms in those four experiments instead of directly changing convection height. The convection heights for the experiment Ms − 1, Ms + 1, Ms + 2, and Ms + 3 are fixed roughly at 155, 145, 141, and 137 hPa, respectively, so one (Ms − 1) is shallower than in the standard global warming experiment (Ms + 0) and three (Ms + 1, Ms + 2, and Ms + 3) are deeper than in the standard run. All experiments are 50-yr runs but averaged only over the last 40 yr. Figure 1 shows precipitation anomalies for the experiments Ms − 1, Ms + 0, and Ms + 1 minus the control experiment. The patterns are similar, but the magnitudes of precipitation anomalies tend to be smaller as convection becomes deeper. The precipitation anomalies are further reduced in the other two experiments, Ms + 2 and Ms + 3, but with similar spatial patterns (not shown). Overall, all five experiments show reasonable changes induced by global warming.

d. Data

The twenty-first-century GCM simulations from the World Climate Research Programme (WCRP) Coupled Model Intercomparison Project phase 3 (CMIP3) multimodel dataset (Table 1) in the A1B scenario are used here. As shown later in the budget analysis (Table 2), the moisture budget is slightly imbalanced in some models (e.g., GFDL_CCM2.1, JP_CCSR3.2H, and FR_IPSL_CM4.1). Causes for this imbalance might be associated with the interpolation of model variables from the original vertical coordinate to the standard pressure coordinate. This could be larger over regions with variations in topography. Thus, saving flux variables from the original coordinate is important in the budget analysis. Another possible cause is related to the transient component. In our calculation, monthly data are used, so the transient terms are excluded. We also notice that the availability of model variables is not consistent in some model. At some grid points, for instance, the horizontal velocity u and υ exist, whereas the moisture q is missing. This creates an inconsistency between the flux terms, such as − · vq, and the advection terms, such as −v∇ · q. Overall, the moisture budget in most models is still balanced quite nicely (Fig. 2).

3. Global water vapor budget

To more precisely estimate the change of precipitation, a column-integrated water vapor budget is used, which can be written as
i1520-0442-23-11-3019-e7
where pressure velocity ω is assumed to be zero at the surface ps and at tropopause pT. For simplicity and consistency, the precipitation (latent heat release) P is in energy units (W m−2); divided by 28, these units become mm day−1. Here, E is evaporation (surface latent heat flux) and v is horizontal velocity. The anomalous column-integrated moisture budget can be further written in an advection form,
i1520-0442-23-11-3019-e8
The nonlinear term −〈ω′∂pq′〉 is neglected here. In a warmer climate, based on (7) and (8), the fractional change of PE should be equal to the change in the column-integrated convergence of moisture flux,
i1520-0442-23-11-3019-e9
The first term on the right-hand side (rhs) of (9), −〈ωpq′〉, is a thermodynamic component, which is mainly associated with the water vapor change. The second term on the rhs, −〈ω′∂pq〉, is a dynamic component, which is related to the change in tropical circulation (C09; Emori and Brown 2005; Held and Soden 2006).

Most precipitation occurs over convective regions, which are roughly defined with ω < 0 (upward motion) at 500 hPa (e.g., indicated by thick dashed curves in Fig. 1), so we focus on the moisture budget (9) averaged over convective regions. Figure 2a shows a rough balance between the fractional changes of PE and −〈 · vq〉, so the moisture in convective regions is conserved in all climate models. Because the horizontal moisture advection and the residual term [the third and fourth terms on the rhs of (9)] are relatively small for most climate model simulations compared to the thermodynamic and dynamic components (Table 2), they are usually neglected in the first-order approximation. Thus, (9) could become similar to the simplified precipitation equation discussed in section 1 (i.e., P′/P = Mc/Mc + qs/qs), if the fractional changes of precipitation and evaporation averaged over convective regions were the same. For global averages, the fractional changes of precipitation and evaporation must be equal because of the balance of the global moisture budget (Held and Soden 2006; Lambert and Webb 2008; Vecchi and Soden 2007). Averaged over convective regions, on the other hand, it is not guaranteed that the precipitation change is equal to the evaporation change. In fact, the evaporation change is usually smaller than the precipitation change (Table 2), making the fractional change of PE larger than that of precipitation in most climate model simulations (Fig. 2b). Unlike the global precipitation change, which is controlled by the global energy balance with its fractional change being much smaller than the Clausius–Clapeyron thermal scaling (Held and Soden 2006; Lambert and Webb 2008; Vecchi and Soden 2007), PE averaged over convective regions does not depend on the global energy balance, so the fractional change of PE could be either larger or smaller than the Clausius–Clapeyron thermal scaling. For the climate model simulations used in this study, however, the fractional change of PE is still consistently smaller than the Clausius–Clapeyron thermal scaling (Fig. 2c), even though some of those changes in PE are much larger than the fractional change of precipitation (Fig. 2b).

Considering the moisture convergence terms [i.e., the right-hand side of (9)], the fractional change of the thermodynamic component is smaller than the value indicated by the Clausius–Clapeyron thermal relation but is still larger than the fractional change of PE (Table 2). Thus, the fractional change of the dynamic component must be negative, as shown in Table 2, and tropical circulation should be weakened in the future warmer climate. In the 16 climate models (Table 1) that we analyzed here, all show a weakening of tropical circulation, which is indicated by pressure velocity at 500 hPa (Fig. 2d). Because the change of PE is not controlled by the global energy balance, the negative fractional change of the dynamic component should not be directly associated with less efficient radiative cooling via precipitation change, as indicated in previous studies (Held and Soden 2006; Vecchi and Soden 2007). In other words, tropical circulation could be strengthened in a warmer climate even though the fractional change of precipitation is much less than the fractional change of water vapor. For instance, the fractional change of PE could be larger than the Clausius–Clapeyron thermal scaling if the change of evaporation was much less than that of precipitation or even negative.

4. Impacts of deepened convection on the strength of tropical circulation

If tropical circulation cannot be determined by different fractional changes between the global-mean precipitation and water vapor, what kind of mechanisms control the change of tropical circulation? In previous studies (Chou and Neelin 2004, hereafter CN04; Chou et al. 2006; C09), several mechanisms for regional tropical precipitation change have been proposed through the further analysis of the MSE budget, which is
i1520-0442-23-11-3019-e10
The net energy flux into the atmosphere is Fnet = FtFs. The net heat flux at the top of the atmosphere (TOA) is Ft = StStRt, and the net heat flux at the surface is Fs = SsSs + RsRsEH. Subscripts s and t on the solar (S and S) and longwave (R and R) radiative terms denote surface and model top, and H is sensible heat flux. Positive Ft and Fs indicate downward heat fluxes. The anomalous MSE budget can then be written as
i1520-0442-23-11-3019-e11
i1520-0442-23-11-3019-e12
Thus, the fractional changes of the column-integrated MSE budget can be written as
i1520-0442-23-11-3019-e13

In (13), the change in the vertical velocity that is associated with tropical circulation 〈ω′∂ph〉 is mainly determined by changes in the term associated with atmospheric stability −〈ωph′〉, the horizontal MSE advection −〈v · (q + T)〉′, and the net energy flux into the troposphere Fnet′. In a warmer climate, the tropospheric vapor water increases and mainly concentrates at lower troposphere, so the atmosphere becomes more unstable. On the other hand, the vertical profile of the warming induced by latent heating should follow the moist adiabatic process; thus, the atmospheric temperature increases more at higher troposphere than at lower troposphere (Karl et al. 2006; Trenberth et al. 2007). This amplification of tropospheric warming can stabilize the atmosphere. Although these two effects tend to cancel each other out, moisture contributes slightly more than temperature (C09). In other words, the atmosphere becomes a little more unstable in a warmer climate when considering only changes in temperature and moisture. This effect has been termed the “rich-get-richer” mechanism (CN04; C09). In a warmer climate, tropical convection also tends to extend higher (Fig. 3) because of the uplifting of the tropopause (Holzer and Boer 2001; Santer et al. 2003; International Ad Hoc Detection and Attribution Group 2005; Lorenz and DeWeaver 2007), which stabilizes the atmosphere because of the increase in its effective static stability (C09). The changes of convection depth are around 2%–3%: that is, deepening for all 16 GCM simulations (Table 2). Convection height here is roughly estimated by assuming a similar MSE at the base and top of convection, so it represents the maximum level of convection (e.g., Yu et al. 1998). Besides the changes associated with atmospheric stability, other effects, such as the MSE advection, also affect tropical circulation. In convective regions, the MSE advection is usually negative because the inward flow at lower troposphere brings relatively dry air from subsidence regions to convective regions; this process is termed the upped-ante mechanism (CN04; C09). This effect often occurs over margins of convective regions. The change in the net energy input into the atmosphere is relatively small, because changes in evaporation and sensible heat tend to balance the radiative cooling when averaging over a large enough domain (Lambert and Webb 2008; Vecchi and Soden 2007). Overall, tropical circulation is strengthened by the rich-get-richer mechanism but weakened by the upped-ante mechanism and the effect of convection depth.

According to the previous discussion, one of the robust mechanisms that can reduce tropical circulation is associated with depth of convection. To examine this effect associated with the change of convection depth, an intermediate climate model coupled with a mixed layer ocean (Neelin and Zeng 2000; Zeng et al. 2000) is used. This model (QTCM1) has been used in studying mechanisms of regional tropical precipitation changes under global warming (Neelin et al. 2003; CN04; Chou et al. 2007).

In a set of experiments, convection height is gradually increased from 155 to ∼137 hPa, which is equivalent to −1.2% to ∼3.3% of changes in convection depth (Table 3), similar to those in the GCM simulations (Table 2), so the atmosphere becomes more stable. The fractional change of the global-mean water vapor varies between 7.0% and 8.2% K−1, which roughly follows the Clausius–Clapeyron scaling. The change of the global-mean precipitation, on the other hand, ranges from 2.5% to 1.5% K−1, which is far below the corresponding changes in water vapor (Fig. 4). According to the argument discussed in previous studies (Held and Soden 2006; Vecchi and Soden 2007), tropical circulation must weaken in all experiments, because the increase in precipitation is slower than in water vapor. On the contrary, tropical circulation is not always weakened in these experiments. In Fig. 4a, the fractional change of vertical velocity ω at 500 hPa, which can be used to roughly estimate the strength of tropical circulation (Vecchi and Soden 2007), is positive when convection is shallower and becomes negative as convection becomes deeper. In other words, tropical circulation is not necessarily weakened, even when the fractional change of the global-mean precipitation is consistently less than that of the global-mean water vapor. The vertical profile of changes in vertical velocity is shown in Fig. 5. The vertical profiles with weakened tropical circulation (Ms + 2 and Ms + 3) are similar to those in Fig. 3. It implies that the changes in vertical velocity shown in Fig. 3 are associated with the deepening of tropical convection. In the MSE budget shown in Table 4, −〈ωph′〉, which is associated with the opposite sign of the change in the effective static stability, is gradually reduced as convection becomes deeper, so the atmosphere becomes more stable. The horizontal MSE advection, which is another mechanism that can suppress tropical circulation (CN04; C09), is negative in all experiments but does not vary too much among these experiments. The anomalous net energy input into the atmosphere Fnet′ increases as convection deepens, so Fnet′ tends to enhance, not suppress, tropical circulation. Thus, the gradual weakening of tropical circulation in these experiments is mainly due to the change in atmospheric stability that is associated with depth of convection.

On the rhs of the moisture budget (9), because the fractional change of water vapor does not vary too much, the thermodynamic component is roughly constant among these QTCM1 experiments, ranging from 5.3% to 6.6% (Table 3). The dynamic component, on the other hand, varies greatly (from 4.2% to −4.6%) because of the variation in the strength of tropical circulation. Both the horizontal moisture advection and the residual term are relatively small. Thus, the fractional change in the convergence of moisture flux becomes smaller as the atmosphere becomes more stable (Table 3). On the lhs of (9), the precipitation averaged over convective regions remains relatively unchanged as the atmosphere becomes more stable, whereas the evaporation becomes larger (Fig. 4). This leads to a smaller change in PE (Table 3). In other words, the weaker (stronger) convergence of moisture flux associated with the weakened (enhanced) tropical circulation compensates for the increase (decrease) of the evaporation over convective regions, so the precipitation change does not vary too much with the change in convection depth. Overall, the fractional change of PE averaged over convective regions varies with convection depth but not fractional changes in either the global-mean precipitation or water vapor, which are controlled by different mechanisms.

5. Discussion and summary

As global warming becomes dominant, many climate changes are starting to show. One of the well-known climate changes is a weakening of tropical circulation, which could affect rainfall amount and intensity both in regional and global scales. Previous studies (Held and Soden 2006; Vecchi and Soden 2007) show that the atmospheric water vapor increases faster than precipitation, so the associated tropical circulation must weaken. However, we demonstrated here that this condition cannot guarantee the weakening of tropical circulation without changes in depth of convection.

We provided a possible mechanism, which is associated with depth of convection, for inducing the weakening of tropical circulation under global warming. Conventionally, vertical motion associated with tropical circulation is determined by atmospheric stability. Atmospheric stability is usually affected by the vertical profile of atmospheric temperature and moisture (Karl et al. 2006; Trenberth et al. 2007), which tends to cancel each other out if the atmosphere is mainly controlled by convection, such as in the tropics (C09). A previous analysis (Yu et al. 1998) shows that atmospheric stability can also be affected by depth of convection. Under global warming, a deepening of convection is consistently found in the CMIP3 dataset. This leads to a weakening of tropical circulation in the CMIP3 global warming simulations. The change in atmospheric stability also modifies the evaporation gradient between convective and subsidence regions, which compensates for the change in the convergence of moisture flux to maintain a similar change in global-mean precipitation. Overall, by examining mechanisms for the weakening of tropical circulation, this study offers an alternative way to detect global warming impacts on tropical circulation.

Acknowledgments

We acknowledge the international modeling groups for providing their data for analysis, the Program for Climate Model Diagnosis and Intercomparison (PCMDI) for collecting and archiving the model data, the JSC/CLIVAR Working Group on Coupled Modelling (WGCM) and their Coupled Model Intercomparison Project (CMIP) and Climate Simulation Panel for organizing the model data analysis activity, and the IPCC WG1 TSU for technical support. The IPCC Data Archive at Lawrence Livermore National Laboratory is supported by the Office of Science, U.S. Department of Energy. The authors thank three anonymous reviewers for their helpful comments for improving the quality of this paper. This work was supported by the National Science Council Grants NSC98-2628-M-001-001 and NSC98-2625-M-492-011.

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Fig. 1.
Fig. 1.

Precipitation differences between experiments with 1 × CO2 and 2 × CO2 concentrations for different convection heights: (a) Ms − 1, (b) Ms + 0, and (c) Ms + 1. The thick dashed curves indicate a convective margin: that is, ω = 0 at 500 hPa. The unit is mm day−1.

Citation: Journal of Climate 23, 11; 10.1175/2010JCLI3383.1

Fig. 2.
Fig. 2.

Scatter plots for fractional changes of hydrological variables and the strength of tropical circulation averaged over convective regions between the last and first 20 yr of the twenty-first century for each of the climate models listed in Table 1: (a) PE vs the column-integrated convergence of the moisture flux −〈 · vq〉, (b) PE vs precipitation, (c) surface temperature vs PE, and (d) PE vs the strength of tropical circulation (vertical velocity at 500 hPa).

Citation: Journal of Climate 23, 11; 10.1175/2010JCLI3383.1

Fig. 3.
Fig. 3.

The anomalous pressure velocity between the last and first 20 yr of the twenty-first century. The anomalies are averaged over convective regions for the 16 CMIP3 models listed in Table 1. Positive values indicate downward motion (reduced) and negative values indicate upward motion (enhanced).

Citation: Journal of Climate 23, 11; 10.1175/2010JCLI3383.1

Fig. 4.
Fig. 4.

Fractional changes of precipitation and water vapor and tropical circulation in the QTCM1 experiments associated with different convection depths. The numbers are normalized by the change of global-mean surface temperature. Depth of convection becomes shallower from left to right. (a) Pressure velocity at 500 hPa, (b) evaporation averaged over convective regions (indicated by thick curves in Fig. 1), and global-mean (c) precipitation and (d) column-integrated water vapor.

Citation: Journal of Climate 23, 11; 10.1175/2010JCLI3383.1

Fig. 5.
Fig. 5.

The anomalous pressure velocity averaged over convective regions (indicated in Fig. 1) between experiments with 1 × CO2 and 2 × CO2 concentrations for different convection heights: Ms − 1, Ms + 0, Ms + 1, Ms + 2, and Ms + 3.

Citation: Journal of Climate 23, 11; 10.1175/2010JCLI3383.1

Table 1.

A list of the 16 coupled atmosphere–ocean climate model simulations in the A1B scenario from the CMIP3 archive.

Table 1.
Table 2.

Fractional changes of the moisture budgets for 16 CMIP3 models on the A1B emission scenario between the last and first 20 yr of the twenty-first century. Rows 1–3 are averaged over the entire globe, and the rest are averaged only over convective (ascending) regions. The values are normalized by the global-mean surface temperature change.

Table 2.
Table 3.

Fractional changes of the QTCM moisture and MSE budgets for different Ms, which indicates the contribution of vertical temperature profile. Rows 1–3 are averaged over the entire globe (80°S–80°N), and the rest are averaged only over convective (ascending) regions. The values are normalized by the global surface temperature change.

Table 3.
Table 4.

Fractional changes of the QTCM MSE budgets for different Ms, which represents different convection depths. From left to right, convection becomes deeper. The values are averaged over convective (ascending) regions and normalized by the global-mean surface temperature change.

Table 4.
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