1. Introduction
Traditionally, monsoon refers to the seasonal reversal of the prevailing surface winds in the tropics (Gadgil 2003). Monsoon regions are distinct from the midlatitude and polar regions by having two seasons, one dry and the other wet. While the winter season is characterized by cool and dry air blowing from the continent, during summer, moist tropical maritime air moves onshore, bringing abundant rainfall to monsoon regions (Webster et al. 1998). The enormous amount of latent heat released by monsoon precipitation accounts for a large fraction of the total diabatic heating in the tropics (Yanai et al. 1973). This latent heating is important for driving mesoscale convective systems as well as large-scale tropical circulations (Gill 1982; Holton 2004).
Figure 1 shows the climatological July geopotential height Z at 150 hPa from the European Centre for Medium-Range Weather Forecasts (ECMWF) interim reanalysis (ERA-Interim) (Dee et al. 2011). During the boreal warm season, the circulation of the upper troposphere and lower stratosphere (UTLS) is dominated by anticyclones over Asia and North America.
TMPA climatological July precipitation rate R (color) and ERA-Interim climatological July geopotential height Z at 150 hPa (contours; interval: 5 dam). The Asian and North American monsoon anticyclones are labeled and indicated by arrows. Four longitude sectors (EH1, WH1, WH2, WH3) of equal size are labeled and outlined in red. The longitudinal-mean precipitation of these sectors is examined in Fig. 6.
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
The Asian monsoon anticyclone (AMA), also known as the South Asian high, is centered at ~30°N, ~70°E and is bounded by the subtropical westerly jet to the north and the tropical easterly jet to the south (Dethof et al. 1999). The onset of the AMA is typically in mid-May. At its peak strength in July, the anticyclone extends over the entire Eastern Hemisphere (EH). The AMA begins weakening in September and typically disappears in October.
The North American monsoon anticyclone (NAMA), also known as the Mexican high, is located near 30°N, 110°W, but the circulation is not as strong and persistent as its Asian counterpart (Dunkerton 1995; Chen 2003). The NAMA emerges in late May off the Pacific coast of Central America, near the eastern Pacific intertropical convergence zone (ITCZ). It then moves northward along the Pacific coast of Mexico until it is centered near northwestern Mexico and the southwestern United States (Douglas et al. 1993). The circulation enters its mature phase in July and August and gradually decays from late September (Vera et al. 2006).
The origin of the upper-tropospheric NAMA circulation, however, is not clear. There are a number of studies on the spatial and temporal variations of precipitation in the North American monsoon (NAM) region, including Mexico and the United States (e.g., Douglas et al. 1993; Adams and Comrie 1997; Barlow et al. 1998; Higgins et al. 1997, 1998, 1999). Higgins et al. (1997, 1998) showed that during July and August the strength of the NAMA is related to the amount of precipitation in Arizona and New Mexico. Stensrud (2013) demonstrated the importance of diabatic heating in simulating the NAMA using a mesoscale numerical model. Chao and Chen (2001) asserted that land–sea contrasts and orography are important for simulating the North American monsoon. Other thermal forcings may also be important. The heaviest North American monsoon precipitation is located along the Sierra Madre Occidental in northwestern Mexico; however, the heaviest precipitation in the western Hemisphere (WH) falls in the equatorial Pacific and Atlantic ITCZs, Central America, and the northern part of South America (Fig. 1). Given its magnitude and proximity, the ITCZ precipitation may contribute to the formation of the NAMA. In addition, heat sources in the Asia monsoon region may play a role through downstream wave effects. Chen et al. (2001) showed that in a linear, quasigeostrophic model the subtropical anticyclones in the lower troposphere over the North Pacific and the North Atlantic are a remote response to the Asian heat sources. Jiang and Lau (2008) showed that intraseasonal variability of the North American monsoon is associated with convective activity in the subtropical western North Pacific via a trans-Pacific wave train.
The fundamental dynamics of the NAMA may be a Matsuno–Gill-type response, but much is not understood about the origin and dynamics of the anticyclone and the contributions of heating from different regions. In this paper, we investigate whether the NAMA is a response to the diabatic heating from the Asian monsoon region, the North American monsoon region, the ITCZ, or a combination thereof. Our approach is through running numerical experiments with a simplified dry GCM dynamical core forced by observed and idealized heating distributions.
2. Data
a. Atmospheric reanalysis
To provide an observational foundation for the numerical experiments, we use 39 years (1979–2017) of the ERA-Interim product (ECMWF 2009). Data files are obtained from the Research Data Archive (RDA) at the National Center for Atmospheric Research (NCAR). The ERA-Interim system has a T255 spectral resolution and the corresponding reduced N128 Gaussian grid has a horizontal grid spacing of ~0.7° × ~0.7° (~80 km × ~80 km). Analyses are available at 6-h intervals. For this study we use data on 37 unevenly spaced pressure levels. The topmost pressure level is 1 hPa. Monthly averages are computed as a simple arithmetic average of the 6-hourly data. Climatological monthly averages are computed in a similar manner from the monthly averages.
b. Precipitation data
To prescribe the horizontal latent heating distribution, which is used to specify diabatic heating rates in the GCM, we use 19 years (1998–2016) of 3-hourly TMPA, version 7, precipitation analyses (Tropical Rainfall Measuring Mission 2011; Huffman and Bolvin 2017). The TMPA has a horizontal grid spacing of 0.25° × 0.25° over the latitude band from 50°S to 50°N. Monthly averages are computed as a simple arithmetic average of the 3-hourly data. Climatological monthly averages are computed in a similar manner from the monthly averages.
3. Simplified GCM
a. Dynamical core
We use the dry Eulerian spectral dynamical core of the NCAR Community Atmosphere Model (CAM), version 5.4, which is the atmospheric component of the Community Earth System Model (CESM). The dynamical core solves the hydrostatic primitive equations formulated in vorticity-divergence form on a hybrid sigma–pressure (σ–p), or η coordinate, using a spectral transform method in the horizontal, a finite difference method in the vertical, and a semi-implicit leapfrog scheme for time integration (Neale et al. 2012). The model has no topography or moisture. A linear harmonic 
b. Idealized physics package
One of the goals of version 2 of the CESM is to support simpler model configurations. As a result, CAM 5.4 offers an idealized physics configuration based on Held and Suarez (1994, hereafter HS94) as an alternative to full physics parameterizations. While the default HS94 simulation resembles the general circulation of the atmosphere in equinoctial conditions, the height of the climatological tropical tropopause differs somewhat from observations (340–350 vs 360–380 K) and the tropical lapse rate in radiative equilibrium follows the dry adiabatic lapse rate rather than the moist one (Tandon et al. 2011). Since this study is focused on the tropical UTLS region, we adopt an HS94-like physics package from Schneider (2004) and Schneider and Walker (2006, hereafter SW06) to better reproduce the tropical circulation. The two Schneider configurations differ slightly in their treatment of the moist adiabatic profile; therefore, we make reference only to the SW06 configuration from here on. Values of selected parameters of the idealized physics package are given in Table 1.
Selected parameters of the idealized physics package described in section 3. Parameters related to the horizontal distribution of thermal forcing are summarized in Tables 2 and 3.
1) Surface drag
A quadratic drag is applied to the horizontal wind within the planetary boundary layer. The depth of the boundary layer extends to 
2) Radiative processes
The temperature field T is relaxed toward the zonally symmetric radiative equilibrium temperature profile 
3) Convective adjustment
The convective heating 
4) Thermal forcing
In this study, the horizontal distribution of the thermal forcing 
For the latter, two idealized horizontal precipitation distributions are being used. The first one is a zonally elongated thermal forcing, which resembles the appearance of the Pacific ITCZ and is specified in (6), where 
The second is a more compact regional forcing given by (7).
c. Experimental design
To evaluate the contribution of latent heating in different parts of the tropics and Northern Hemisphere subtropics to driving the NAMA, we carry out numerical experiments using the simplified GCM. The parameters of the realistic and idealized heating experiments are summarized in Tables 2 and 3, respectively. The control experiments (runs 1a–1c) have zero forcing (Q0 = 0 K day−1). The first set of perturbed experiments (runs 2a–2c) uses realistic heating distributions derived from the TMPA climatological July precipitation rate to investigate the contribution of the global and hemispheric heating to driving the NAMA. The second set of perturbed experiments (runs 3a–3i) further analyzes the contribution of different regions in the Western Hemisphere to the NAMA response. Additional experiments 4–6 examine the response to the idealized heating distributions specified by (6) and (7). The last set of experiment is used to test the model sensitivity to both horizontal resolution (runs 7a–7e) and vertical resolution (runs 7f–7j).
Summary of the numerical experiments with thermal forcing based on the TMPA observations. The second column gives the model resolution. The third through fifth columns give the description, latitude range, and longitude range of the forcing, respectively. The last two columns give the maximum magnitude Q0 and total amount Qtotal of the forcing, respectively. Information for TMPA observations is included for reference. All experiments use the same vertical distribution of thermal forcing (Table 1).
Summary of the numerical experiments with idealized zonally elongated and compact regional thermal forcings. Second column gives the model resolution. Third through seventh columns give the description, center longitude 
Most of the numerical experiments are integrated for 600 days at T42L30 resolution, denoting a triangular truncation of the spherical harmonic series at wavenumber 42 (~2.8° latitude × ~2.8° longitude on a Gaussian grid), and 30 evenly spaced vertical levels between η ≈ 0.9833 and 0.0167. To test the model sensitivity, several experiments are repeated at T85 horizontal resolution (~1.4° × ~1.4°) while other experiments use an L60 vertical resolution (60 evenly spaced vertical levels between η ≈ 0.9917 and 0.0083). The model is initialized with a resting isothermal atmosphere at 250 K. To satisfy the numerical stability criterion, the model time step size is set to 900 (450) s for T42 (T85) experiments. To prevent energy accumulation at small scales, the biharmonic dissipation coefficient is set to 1.17 × 1016 (7.14 × 1014) m4 s−1 for T42 (T85) experiments, with a damping time scale of 12 h at the smallest wavelength (MacVean 1983).
The zonally symmetric basic state temperature profile 
(top) Basic state for all numerical experiments: (a) reference temperature 
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
Runs 1b and 1c, which are integrated at T85L30 and T42L60 resolution, respectively, produce a very similar climatology (not shown). Increasing horizontal resolution leads to a slightly stronger boreal (weaker austral) subtropical jet, a somewhat stronger winter Hadley cell, and a lower tropical temperature near the surface. Increasing vertical resolution results in slightly stronger subtropical jets in both hemispheres, a relatively stronger winter Hadley cell, and a rather lower temperature at the cold-point tropical tropopause.
4. Results and discussions
a. Response to the realistic forcing
1) Model tuning
To determine the major forcing regions of the NAMA, numerical experiments are carried out using realistic global and regional geographical distributions of latent heating derived from the TMPA precipitation data. A top-down approach is used to analyze the impacts of various regions on the upper-tropospheric circulation. The model is first subjectively tuned to produce a realistic zonal-mean state and reasonable representations of the AMA and NAMA when forced by latent heating throughout the tropics and subtropics (run 2a). The time-averaged model response to latent heating within selected regions is then examined.
Because of several significant simplifications in the model, including the absence of topography, land–sea contrast, and temporal variations of the heating, a perfect simulation is not expected. To produce the best possible simulation of the monsoon anticyclones, the principal tuning parameter for the basic state is the pole-to-equator temperature gradient 
Preliminary experiments show that the subtropical jet strengthens when either the pole-to-equator temperature gradient or the total external thermal forcing increases. To include a representative geographical extent of the forcing region, but at the same time not make the jet too strong, we set Δh = 30 K and limit the TMPA-derived forcing to the zone from 5°S to 40°N.
Preliminary experiments also show that the response of the anticyclones is not very sensitive to the vertical extent of the forcing, especially the bottom-level 
Finally, the forcing magnitude parameter 
2) Global heating
The performance of the model when forced by heating in both the Eastern and Western Hemispheres (run 2a) is assessed through comparisons with the ERA-Interim July climatology. Figure 3 shows selected time-mean, zonal-mean fields. CAM temperatures are high near the surface, which is typical for a dry GCM without surface energy fluxes (Figs. 3a and 3b), but the discrepancy becomes smaller in the free troposphere (Fig. 3c). The model produces a reasonable structure for both summer and winter subtropical jets (Figs. 3d and 3e), although the summer jet is somewhat strong compared to observations (31.2 vs 21.8 m s−1). The tropical easterlies are also stronger in the model than in observations, but the difference is modest (Fig. 3f). The difference in jet strength can be attributed to a lack of other mechanisms of heat transport (Baker et al. 2017). The tropopause height in the deep tropics generally agrees well with observations. With a very weak stratospheric lapse rate (Fig. 2c), the polar-night jet is absent in this model.
Latitude–pressure cross sections of selected time-mean, zonal-mean fields. (left) ERA-Interim July climatology. (center) CAM run 2a. (right) Difference (CAM minus ERA-Interim). (a),(b) Temperature T (black) and potential temperature θ (red). (c) Temperature T. (d)–(f) Zonal wind u. The tropopause pressure ptrop (blue) is shown in the left and middle columns. Contour intervals of T for (a) and (b) are 10 K and (c) are 5 K. Contour intervals of θ and u are 10 K and 5 m s−1, respectively. Negative values are dashed, and zero values of T and u are in bold contours. Contours of θ above 450 K are omitted.
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
Figures 4a and 4b shows the observed and simulated geopotential height Z (contours) for the reanalysis and run 2a, respectively; the thermal forcing distribution 
(a) ERA-Interim climatological July geopotential height Z. (b)–(d) Time-mean CAM Z (contours) for runs 2a–2c, respectively, with geographical distributions of TMPA-derived thermal forcing QLH (color). Both fields indicate values at 150 hPa. Contour interval of Z is 5 dam. The horizontal distribution of QLH is specified in Table 2. The model has no topography or land–ocean contrast, and the continents are shown only to provide a geographic reference.
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
Figure 5 compares meridional and zonal vertical sections of the time-mean horizontal wind structure through the centers of the AMA and NAMA. Zonal averages of u and meridional averages of υ are computed over 45° sectors and 10° zones, respectively. Note the shift in longitude for the zonal section of the AMA in Fig. 5d. For the AMA, the subtropical westerlies are stronger in the model, as expected from Fig. 3e, but the difference of strength in the tropical easterlies is small (Figs. 5a and 5c). The meridional wind in the AMA is slightly stronger in the model (Figs. 5b and 5d). The vertical extent of the AMA is between 300 and 50 hPa and agrees well between the observations and simulation. For the NAMA, the simulation satisfactorily reproduces the magnitude and location of the subtropical jet (Figs. 5e and 5g). Contrary to the AMA, the meridional wind in the NAMA is slightly weaker in the model (Figs. 5f and 5h). Also the NAMA does not extend as high as the AMA. Overall the simplified GCM is capable of reproducing the major features of both anticyclones with realistic amplitudes when forced by the observed, time-averaged, geographical distribution of latent heating.
(left) Latitude–pressure cross sections of time-mean, longitudinal-mean zonal velocity u. (right) Longitude–pressure cross sections of time-mean, latitudinal-mean meridional velocity υ. Only cross sections between 500 and 50 hPa are shown. Note that the longitude (latitude) ranges of observations and model used for averaging u (υ) are different but the extents of longitude 
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
The diabatic heating from the physics package in (2) consists of the radiative heating 
3) Hemispheric heating
Experiment 2a provides the basis for analyzing the contributions to the forcing from individual geographical regions. The forcing is first divided into Eastern and Western Hemisphere contributions (Figs. 4c and 4d; runs 2b and 2c). The total heating in run 2a is equal to the sum of the heating in runs 2b and 2c (Table 2). The AMA is slightly weaker when the Western Hemisphere heating is removed (Fig. 4c). A deep trough is still distinct in the Pacific. Similarly, the NAMA is slightly weaker when the Eastern Hemisphere heating is removed (Fig. 4d). In runs 2b and 2c the westerlies are somewhat weaker than in run 2a. These experiments demonstrate that in the model the anticyclones are little affected by remote heat sources. That is, the anticyclones are primarily a response to nearby thermal forcing in their respective hemispheres, and there is no indication that the NAMA is a downstream response to the Asian monsoon circulation.
4) Partitioning the Western Hemisphere heating
Runs 2a–2c show that the NAMA is principally driven by latent heating in the Western Hemisphere. Runs 3a–3i are then used to determine the importance of heating in different longitude sectors and latitude zones within the Western Hemisphere. The heating is first partitioned into three 60° longitude sectors (WH1, WH2, and WH3) as defined in Fig. 1. Figure 6 shows the longitudinal-mean precipitation in these three sectors. One Eastern Hemisphere sector (EH1) is shown for comparison. The maximum WH and EH precipitation rates are similar, but the WH sectors are dominated by the narrow ITCZ precipitation located between 0° and 15°N, while EH1 has a large amount of precipitation between about 10° and 40°N. There is little subtropical precipitation in WH1 and WH3 (less than 1 mm day−1), but there is a substantial amount of precipitation poleward of 15°N in the middle sector (WH2).
TMPA longitudinal-mean climatological July precipitation rate R of sectors EH1, WH1, WH2, and WH3, as defined in Fig. 1.
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
Results for the three WH sectors simulations (runs 3a–3c) are shown in Figs. 7a–c. Note that the latitudinal extent of the heating is limited to be between 5°S and 40°N. The total heating in run 3b (middle sector) comprises ~51% of the total Western Hemisphere heating in run 2c (Table 2), and it has substantially more precipitation in the subtropics than the other two runs. For runs 3a and 3c, which have heating primarily within the Pacific and Atlantic ITCZs, respectively, the subtropical response is weak. Run 3b, however, produces a substantial anticyclone in the subtropics.
As in Fig. 4, but for runs 3a–3c, 3e, and 3h.
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
To investigate the relative importance of the magnitude of the heating and its meridional distribution, each longitude sector is divided at 15°N into tropical and subtropical latitude zones, giving six regions. Experiments 3d–3i are carried out using heating from those individual regions. Runs 3d and 3f (not shown), which include only the tropical part of the heating from runs 3a and 3c, produce weak anticyclones similar to Figs. 7a and 7c. Similarly, runs 3g and 3i, which have only a small amount of heating in the subtropics, produce weak anticyclones (not shown). In the middle sector, however, both runs 3e (tropical heating) and 3h (subtropical heating) produce a subtropical anticyclone (Figs. 7d and 7e), with a much stronger response coming from the latter case. Note that although the response from the subtropical heating is larger, the total heating in the subtropical sector is only ~62% of the tropical sector (Table 2). These two simulations suggest that the latitude of the heating plays a significant role in determining the strength of the anticyclone.
We note that precipitation over Mexico and the southern United States, along with the NAMA response in the UTLS, exhibits a northward propagation during late spring and early summer (Higgins et al. 1999). Run 3e resembles the early stage of the NAMA, with a weak anticyclone at a low latitude, while runs 3b and 3h, which are very similar, can be viewed as the mature stage of the NAMA. A closer look at the geographical distribution of the heating in these simulations suggests that the onset of the North American monsoon precipitation across different regions is intimately related to the northward propagation and development of the NAMA.
b. Response to the idealized forcing
1) Zonally elongated heating
Although all three WH sectors have substantial precipitation in the deep tropics, only the heating in WH2, which also has substantial precipitation in the subtropics, produces a significant subtropical response. Experiments 3e and 3h show that, despite the large magnitude of forcing, the response to the tropical precipitation in WH2 is small, and the anticyclone in run 2c is produced primarily by subtropical heating. The preliminary experiments and these realistic forcing experiments suggest that the magnitude and latitude of the forcing are both important for controlling the strength of anticyclones, which supports the Gill-type response, (1). To explore this question in a simpler context, we use two different idealized horizontal distributions, as described in section 3.
The experiments in this section (group 4) investigate the role of the latitude of the heating in the atmospheric response using idealized zonally elongated heating distributions. The total amount of heating is the same at all latitudes to isolate its effect in (1). The meridional profile of the heating is given by (6), with the meridional extent 
Runs 4a–4e use a global zonally symmetric heating distribution centered at different latitudes. As expected, zonally symmetric forcing does not produce a localized anticyclone, even if the forcing is located in the subtropics, so no results from these experiments are shown here.
Runs 4f–4j restrict the zonally elongated heating to a 180° longitude sector centered at 90°W (
Figure 8 shows the results from experiments 4g and 4i. The quantity displayed is the zonally asymmetric part of the atmospheric response 
Time-mean geopotential height anomaly Z* (contours) and geographical distributions of zonally elongated heating QLH (color) for runs (a) 4g and (b) 4i. Both fields indicate values at 150 hPa. Contour interval of Z* is 2 dam (solid for positive values, dashed for negative values, and bold for zero). Note the change of the color-bar scale from Figs. 4 and 7.
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
2) Compact regional heating
Because narrow and zonally elongated heating in the deep tropics produces little response in the subtropics, we next consider the response to more compact regional heating. For this we use the idealized heating distribution given by (7) with a zonal extent 
Figure 9 shows the eddy response Z* to the regional heating at different latitudes. Forcing at the equator (run 5a), excites weak extratropical responses to the north (Fig. 9a). As the forcing is shifted poleward, the responses in the tropical channel to the west and in the extratropics strengthen somewhat (Fig. 9b). For 
As in Fig. 8, but for runs (a) 5a–(e) 5e with the geographical distributions of compacted regional heating QLH (color). Contour interval of Z* is 1 dam (solid for positive values, dashed for negative values, and bold for zero). Note the change of the color-bar scale from Figs. 4, 7, and 8.
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
Compared to the zonally elongated forcing experiments, a distinct extratropical response appears for a lower 
Experiments 1–6 use T42L30 resolution. It has been reported that increasing the number of vertical levels could improve the simulation of Asian monsoon in the UTLS region in the comprehensive version of the CESM (Wang et al. 2018). To test the sensitivity of the model to changes in both horizontal and vertical resolution, additional runs were carried out with higher horizontal resolution (runs 7a–7e at T85L30) and higher vertical resolution (runs 7f–7j at T42L60). In both cases the results do not change significantly from the T42L30 runs (not shown).
3) Linearity of the response
The linearity of the atmospheric response is evaluated by repeating runs 5a–5e with the heating scaled by factors of 0.5 (runs 6a–6e) and 2 (runs 6f–6j). Details of the numerical experiments are given in Table 3. Figure 10a shows the strength of the anticyclone as a function of the total amount of forcing 
(a) Linear least squares fits of maximum time-mean geopotential height anomaly 
Citation: Journal of the Atmospheric Sciences 76, 7; 10.1175/JAS-D-18-0340.1
Figure 10b shows the strength of the anticyclone as a function of 
5. Conclusions
The origin and dynamics of the upper-tropospheric North American monsoon anticyclone are investigated through numerical experiments with an idealized general circulation model and comparisons with observational analyses. The analysis focuses on the time-averaged response of the atmosphere. The model uses a simplified physics package based on HS94 that relaxes the temperature to a prescribed zonally symmetric basic state. To simulate a more realistic tropical circulation, the physical parameterizations and basic state are modified somewhat from HS94 according to SW06, and an explicit dry convection scheme is included. The model is forced by a steady diabatic heat source based on TMPA precipitation data or prescribed from idealized heating distributions. Significant approximations in the idealized model include the absence of topography, land–sea contrast, and temporal variations of the forcing. These effects are partially compensated by adjusting the magnitude of the thermal forcing. The model uses a constant vertical heating profile throughout the region where heating is applied and ignores heating outside of the tropics and Northern Hemisphere subtropics. Experiments with increased horizontal and vertical resolution (T85L30 and T42L60) show that the simulations are not sensitive to either horizontal or vertical resolution.
The model is first tuned to give a realistic representation of the boreal summer upper-tropospheric circulation, including the AMA, NAMA, and the mid-Pacific trough, using a realistic geographical distribution of diabatic heating based on the observed TMPA time-mean precipitation. To produce a realistic zonal-mean circulation, the latent heating is scaled to 59% of its observed value. This heating distribution is partitioned geographically in various ways, and the response to each individual region is examined. When the heating is turned off in either the Eastern or Western Hemisphere, the anticyclone in that hemisphere disappears, while the anticyclonic circulation in the other hemisphere slightly weakens. This demonstrates that the AMA and NAMA are fundamentally responses to diabatic heating within their respective hemispheres, and the NAMA in particular is not a downstream wave response to Asian monsoonal heating. During the warm season the majority of the precipitation in the Western Hemisphere falls in a narrow zone between about 5° and 15°N within the Atlantic and Pacific ITCZs and across the northern part of South America. The NAMA, however, is primarily a response to heating poleward of 15°N in the longitude sector between 60° and 120°W, which includes the northern part of Central America, Mexico, the southern United States, the Caribbean Sea, and the Gulf of Mexico. This demonstrates the importance of the latitude of the heating to the response.
To explore the atmospheric response to the shape and location of the heating, the model is forced with idealized heating distributions that represent either zonally elongated (ITCZ-like) or compact regional precipitation features. When the zonally elongated heating is placed near the equator, the response in the extratropics is weak. The extratropical response is significant only when the heating is located outside the deep tropics. The response to a compact heating distribution depends on the latitude of the heating as well. An extratropical response is evident when the heating is centered poleward of 10°N, and the extent of the response is more localized compared to the zonally elongated forcing. These experiments support the conclusion that the AMA is stronger than the NAMA is because of both the heavier precipitation in the Eastern Hemisphere and the location of the intense precipitation at higher latitudes over Asia compared to North America. Compared to the zonally elongated heating, the compact heating is shorter in the zonal direction and wider in the meridional direction. The numerical experiments indicate that the longitudinal extent of the AMA is much larger than the NAMA is in part because the zonal extent of the subtropical precipitation in the Eastern Hemisphere is greater than in the Western Hemisphere.
The linearity of the atmospheric response of the idealized forcing experiments is examined by varying the magnitude of the heating. To a good approximation the strength of the anticyclone is linearly proportional to the magnitude of the applied heating, but for a fixed magnitude of heating the response depends nonlinearly on the latitude of the heating.
The model results support the idea that the AMA and NAMA are largely independent of one another, being forced primarily by diabatic heating in their respective hemispheres. That is, both anticyclones are Matsuno–Gill-type responses. The differences in the amplitude and zonal extent of the AMA and NAMA are due in part to the total amount and zonal distribution of heating in each hemisphere, but are primarily a result of the different meridional distributions of heating in the two hemispheres.
Acknowledgments
We are grateful to three anonymous reviewers for their constructive comments that helped improve the manuscript. We thank Andrew Dessler, Ramalingam Saravanan, and Courtney Schumacher for their comments on an earlier version of the manuscript. We thank the CESM working groups at the NCAR Climate and Global Dynamics (CGD) Laboratory for developing the model. We particularly thank Isla Simpson for providing developer’s access of CAM and assistance with running the model. The version of CAM 5.4 in this study was made available through the Simpler Models Initiative as part of the CESM project; this initiative is supported by NCAR under the sponsorship of the National Science Foundation (NSF) and the United States Department of Energy (DOE). This research is funded by NSF Grant AGS–1550611 to Texas A&M University. Portions of this research were carried out with advanced computing resources provided by Texas A&M High Performance Research Computing (HPRC), and we thank Ping Luo for setting up CAM on the Texas A&M Ada system. We thank the ECMWF for producing the ERA-Interim product, which was obtained from the NCAR Research Data Archive. The TMPA data were obtained from the NASA Goddard Earth Sciences (GES) Data and Information Services Center (DISC). Outputs from the model simulations are available upon request.
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