mechanisms controlling the annual cycle of …david/mechanisms.pdf4708 journal of climate volume 17...

4708 VOLUME 17J O U R N A L O F C L I M A T E

q 2004 American Meteorological Society

Mechanisms Controlling the Annual Cycle of Precipitation in the Tropical AtlanticSector in an Atmospheric GCM*

M. BIASUTTI, D. S. BATTISTI, AND E. S. SARACHIK

Department of Atmospheric Sciences, University of Washington, Seattle, Washington

(Manuscript received 29 September 2003, in final form 6 May 2004)

ABSTRACT

A set of AGCM experiments is used to study the annual cycle of precipitation in the region surrounding thetropical Atlantic Ocean. The experiments are designed to reveal the relative importance of insolation over landand the (uncoupled) SST on the annual cycle of precipitation over the tropical Atlantic Ocean, Africa, and thetropical Americas.

SST variations impact the position of the maritime ITCZ by forcing the hydrostatic adjustment of the at-mospheric boundary layer and changes in surface pressure and low-level convergence. The condensation heatingreleased in the ITCZ contributes substantially to the surface circulation and the maintenance of the SST-inducedITCZ anomalies.

The remote influence of SST is felt in equatorial coastal areas and the Sahel. The circulation driven bycondensation heating in the maritime ITCZ extends to the coastal regions, thus communicating the SST signalonshore. Conversely, the Sahel responds to variations in SST through boundary layer processes that do notinvolve the maritime ITCZ. The atmospheric response to changes in subtropical SST is advected inland andforces changes in sea level pressure and low-level convergence across a large part of tropical Africa.

The impact of local insolation on continental precipitation can be explained by balancing net energy input atthe top of the atmospheric column with the export of energy by the divergent circulation that accompaniesconvection. Increased insolation reduces the stability of the atmosphere in the main continental convectioncenters, but not in monsoon regions.

Insolation over land impacts the intensity of the maritime ITCZ via its influence on precipitation in Africaand South America. Reduced land precipitation induces the cooling of the Atlantic upper troposphere and theenhancement of convective available potential energy in the maritime ITCZ.

1. Introduction

Decomposing the observed annual cycle of precipi-tation over the tropical Atlantic, South America, andAfrica into fundamental building blocks is crucial initself as a test of our understanding of basic climateprocesses. Moreover, it can provide powerful tools fordiagnosing the reasons behind the shortcomings of cur-rent models, and guidance in the investigation of inter-annual variability. The latter expectation rests on thefact that the Atlantic interannual variability has longbeen interpreted as a modulation of the annual cycle(e.g., Hastenrath 1984; Mitchell and Wallace 1992) andis strongly affected by the annual cycle of the basic state(Biasutti 2000).

* Joint Institute for the Study of the Atmosphere and OceansContribution Number 1004.

Corresponding author address: Michela Biasutti, Lamont-DohertyEarth Observatory of Columbia University—Oceanography, 61 Route9W, P.O. Box 1000, Palisades, NY 10968-8000.E-mail: [email protected]

This study aims at understanding the processes thatgovern the annual cycle of precipitation over the At-lantic sector as simulated by a global atmospheric gen-eral circulation model (GCM). We identify the processesthat directly force the annual cycle of precipitation asa response to local forcing and boundary conditions(insolation and SST) and the processes that couple thecontinental climate to the climate of the tropical ocean.Specifically, in this paper we discuss how changes inlocal insolation affect precipitation over the tropicalcontinents, how changes in SST (which in our modelexperiments can be prescribed independently from in-solation) affect precipitation in the ITCZ, and howchanges in temperature and rainfall over the ocean affectprecipitation over the continents and vice versa.

A companion study (Biasutti et al. 2005, hereafterBBS2) addresses modifications to the mutual influencesof land and ocean described here that arise from coupledinteractions of the atmosphere and the ocean.

In this study, we focus on the monthly mean, large-scale fields, with the understanding that the large-scalefeatures are the backdrop for rain-bearing synoptic andmesoscale processes. Thus, while recognizing the sin-

15 DECEMBER 2004 4709B I A S U T T I E T A L .

TABLE 1. List of experiments: name, insolation forcing, SST boundary conditions, condensational heating over land, condensationalheating over the oceanic ITCZ region.

Name Insolation SST

Landcondensational

heating

ITCZcondensational

heating Relevant figures

Control (CTL) Climatological Climat. Calculated (5climat.) Calculated(5climat.)

Fig. 1

Varying insolation(PMS)

Climatological Mar Calculated (;climat.) Calculated(;Mar)

Figs. 2, 3, 9a, 10a, c, e

Varying SST (PVE) Vernal equinox Climat. Calculated (;Mar) Calculated(;Climat.)

Figs. 4a, 5a, 6, 7, 8

Perpetual Mar Vernal equinox Mar Calculated (;Mar) Calculated(;Mar)

Figs. 4b, 5b, 9b, 10d, f

Perpetual Mar withSep ITCZ heating

Vernal equinox Mar Calculated (;Climat.) Imposed (5Sep) Figs. 4b, 5b

Perpetual Mar withJun African andSouth Americanheating

Vernal equinox Mar Imposed over SouthAmerica and Africa(5Jun)

Calculated(;climat.)

Figs. 9b, 10b, f

Perpetual Mar withJun South Africanheating

Vernal equinox Mar Imposed over SouthAfrica (5Jun)


Fig. 11a

Perpetual Mar withJun South Ameri-can heating

Vernal equinox Mar Imposed over SouthAmerica (5Jun)


Fig. 11b

gularity of rain events at each location and in each sea-son, we hope to build a conceptual model that can ex-plain the coarse features of the simulated annual cyclein the Atlantic sector and, by extension, in the Tropics.

In its focus on mechanisms, this paper complementsour previous work (Biasutti et al. 2003, hereafter BBS1).BBS1 documented where remote influences were mostimportant in producing the simulated annual cycle ofsurface air temperature and precipitation and also ex-plored issues of nonlinearity in combining the effectsof local and remote forcings.

The paper is organized as follows. In the remainderof this section we present our experimental design anda review of the main results of BBS1. We describe themechanisms responsible for local forcing of continentalconvection in section 2, local forcing of oceanic con-vection in section 3, oceanic forcing of continental con-vection in section 4, and continental forcing of oceanicconvection in section 5. Section 6 summarizes the paper.

a. Experimental design

We employ an atmospheric general circulation model(AGCM) to artificially decompose the annual cycle into‘‘locally generated’’ and ‘‘remotely generated’’ com-ponents by separately imposing SST and insolation asif they were two independent forcings. We allow annualvariations in only one of the forcings at a time (eitherSST or insolation) and prescribe a fixed, arbitrary valuefor the other one. While SST is obviously not indepen-dent of insolation, we can simulate these conditions inour AGCM experiments because SST is prescribed.

The simulations presented in this paper are listed inTable 1, together with a description of the forcings used

in each experiment and a reference to the relevant fig-ures.

We present results from two main experiments: the‘‘varying insolation’’ experiment [the perpetual MarchSST (PMS) simulation of BBS1] and the ‘‘varyingSST’’ experiment [the perpetual vernal equinox (PVE)simulation of BBS1]. The annual cycle in the varyinginsolation experiment—in which insolation varies an-nually while SST is fixed at March value—is the re-sponse to insolation changes over land. Thus this ex-periment indicates where the annual cycle over land isdriven locally and where the annual cycle over the oceanis forced remotely by circulation changes over the con-tinents, subject to the restriction of fixed SST. The vary-ing SST experiment—in which SST varies annuallywhile insolation is fixed at its 21 March value—indicateswhere the annual cycle of precipitation over the oceanis driven by the local SST and where the annual cycleover land is remotely forced by SST changes.

Additionally, we present experiments in which bothinsolation and SST are kept fixed at March values, andthe latent heat released in convection in selected tropicalareas is prescribed to a value corresponding to differentboundary conditions. The prescribed heating is takenfrom either the varying SST experiment or the varyinginsolation experiment and corresponds to SeptemberSST conditions or June insolation conditions, respec-tively. (See the discussion in the following sections fora more complete description of these experiments.) Wecompare these experiments with a control perpetualMarch simulation (perpetual 21 March insolation andfixed March SST) in order to assess the atmosphericresponse to elevated condensation heating when every-thing else is left unchanged. We caution the reader that,


because the heating is specified, these experiments arenot constrained to conserve heat.

We use the (National Center for Atmospheric Re-search (NCAR) Community Climate Model, version 3(CCM3) at the standard resolution (19 levels in the ver-tical and T42 wave truncation, corresponding to a 2.88by 2.88 horizontal grid). All simulations that have anannual cycle in the forcing were run for at least 12 yr;the first 4 yr were discarded, and the analysis was con-ducted on the average of the remaining years. Simula-tions in which all forcings were kept constant were runfor 12 months. All the features discussed in the paperhave been verified to be significant at the 95% level.

It is worth emphasizing that this is a modeling studyand that our results are only valid in the framework ofan atmospheric GCM with prescribed SST. Moreover,some of our results might be specific to CCM3 andmight be influenced by that model’s biases and char-acteristics (Hack et al. 1998; BBS1).

b. The precipitation response to insolation and SSTforcings

Figure 1 summarizes the main results detailed inBBS1. Shown are the first empirical orthogonal function(EOF) and first principal component (PC) of the annualcycle of precipitation in the tropical Atlantic sector inthe control simulation, the varying insolation simula-tion, and the varying SST simulation. The EOF, PC pairin each row of Fig. 1 gives a compact description ofthe main spatial and temporal characteristics of the an-nual cycle of precipitation that is forced in the AGCMby the annual cycle of SST and insolation (Figs. 1a,b),insolation only (Figs. 1c,d), and SST only (Figs. 1e,f).For example, Fig. 1a shows that the main characteristicof the control annual cycle is a north–south movementof the major precipitation centers in both the ocean andland.

The similarity of Figs. 1c and 1a over the continentsindicates that forcing by the local insolation is sufficientto recreate the annual meridional march of precipitationover the continents. Both Figs. 1b and 1d indicate thatmaximum precipitation over the continents occurs dur-ing the summer months. Yet, when only the insolationforcing is applied (Fig. 1d), African precipitation reach-es the farthest north in June, and not in August, as seenin observations and in the control experiment (Fig. 1b);this phase relationship is also apparent in the raw annualcycle and is not an artifact of EOF analysis. The positiveloadings of EOF 1 over the Atlantic ITCZ in Fig. 1cindicate that the insolation-induced annual cycle of con-tinental climate remotely forces an annual cycle in theintensity of the maritime ITCZ; the ITCZ is stronger inconjunction with weaker precipitation in South Americaand South Africa.

The north–south dipole straddling the mean ITCZ po-sition in Fig. 1e indicates that, in the prescribed SSTexperiments, SST determines the meridional march of

the ITCZ. The ITCZ reaches its northernmost positionin September and its southernmost position in March(Fig. 1f), following the highest SST. The same SST-induced north–south displacement is apparent in the pre-cipitation field over the coastal regions in northeastSouth America and Africa, and over central Africa, theSahel, and Sudan.

2. Local control of continental precipitation

Neelin and Held (1987) show how, in the thermallydirect circulations of the Tropics, a net influx of energyin the atmospheric column (Fnet , the net downward ra-diative fluxes at the top of the column plus the upwardradiative and turbulent fluxes at the surface) is balancedby the export of moist static energy by the divergentcirculation. Neelin and Held derive this conclusion froman analysis of the dominant balance in the equations ofenergy and moisture in the precipitating regions of theTropics. They develop their argument in terms of themoist static energy; for consistency with the rest of thepaper, we will recast it in terms of equivalent potentialtemperature.

We describe the energy of a moist air parcel as cpue,where cp is the atmospheric heat capacity at constantpressure and ue is the equivalent potential temperature.The total rate of change in cpue equals the total energyinput,

D c u 5 ] c u 1 = · c u V 1 ] c u v 5 g] F,t p e t p e p e p p e p

where F is the total vertical energy flux, that is, the sumof the radiative, sensible, and latent vertical heat fluxes.If we take the vertical integral over the atmosphericcolumn (indicated with ^ · &), and the time mean, thefirst and third term in the lhs vanish and we obtain ^=· cpueV& 5 Fbottom 2 Ftop.

Neelin and Held now make two major assumptions.First, they assume the vertical structure of convectiveregions to be the first baroclinic mode (thus, the lower-level convergence is equal and opposite to the upper-level convergence). Second, they assume that horizontalgradients of both temperature and moisture are negli-gible in the Tropics and so is the transport by transients.Under these approximations, the equation for the ver-tically integrated mean equivalent potential temperaturesimplifies to

lowerlevel upperlevel lowerlevel bottom top(u 2 u )= · V } F 2 F .e e

In regions of deep convection, the adiabatic cooling andthe diabatic heating nearly balance, so Due 5 2upperu e

ø 0. However, Neelin and Held note that, for theloweru e

time-mean circulation to be thermally direct, Due mustbe positive. Thus, in this simple model, low-level con-vergence and precipitation require Fnet 5 Fbottom 2 Ftop

to be positive. The intuitive interpretation is that a pos-itive import of energy is balanced by the horizontalexport by a thermally direct, divergent circulation. Thus,the Neelin and Held model uses a thermodynamic quan-


FIG. 1. First EOF and PC of the annual cycle of precipitation in (a),(b) the control run, (c),(d) the experiment with varying insolation andSST fixed at Mar value, and (e),(f ) the experiment with varying SST and insolation fixed at Mar value. (a),(c),(e) The contour interval is2 mm day21 per standard deviation, starting with 61. Dashed contours indicate negative values. The shading indicates positive values.

tity to diagnose the occurrence of convection and ther-mally direct circulations.

There is some circularity in this argument. The modelassumes that the atmospheric column can be describedby the first baroclinic mode and thus should only beapplied to describe an atmospheric column in whichconvection is occurring, and Sobel et al. (2004) showthat indeed cloud radiation feedback can be a deter-mining factor. Similarly, the energy input into a columnis determined in part by whether convection is occurringin it. Finally, we remark that while the argument appears

to be strictly thermodynamic, the dynamics of the trop-ical atmosphere are assumed.

Nevertheless, recent studies indicate that Fnet can beused as a zeroth-order indicator of where convection ispossible. For example, Sobel and Bretherton (2000)show that a model based on the balance of Fnet and thedivergent circulation captures the distribution of con-vective regions in the Tropics equatorward of 158, al-though they also show that the addition of moistureadvection improves the simulation. Chou and Neelin(2003) show an overall good correspondence between


FIG. 2. Varying insolation experiment. Jun–Dec precipitation (con-toured, c.i. 5 4 mm day21, starting with 62, negative values dashed)and total energy flux into the atmosphere Fnet (shaded, interval of 50W m22 starting at 625 W m22, dark gray shading indicates negativevalues less than 225 W m22, white indicates value between 625W m22, and light gray shading indicates positive values larger than25 W m22).

regions of positive Fnet and precipitation during Julyboth in observations and in an intermediate complexitymodel. Our GCM simulations also reproduce the hy-pothesized coincidence of rainfall and positive Fnet overmost of the tropical land, but not in the Atlantic ITCZor in land regions affected by the ITCZ (e.g., the Guineacoast).

Over land, where the negligible heat capacity of thesoil makes the energy flux into the surface vanish ontime scales longer than a day, the requirement of a pos-itive Fnet for the development of convection translatesinto the requirement of a downward energy flux at thetop of the atmosphere. Therefore, it is tempting to ex-plain the influence of insolation on the annual cycle ofprecipitation in the tropical continents in terms of Fnet.Figure 2 shows the seasonal change (June–December)in precipitation brought about by insolation only, su-perimposed on the corresponding change in total columnenergy flux (Fnet). It is apparent that, to zeroth order,changes in Fnet and changes in precipitation over landgo together: where the June–December Fnet is negative,so is the precipitation change and vice versa. A notableexception is the Sahara, where dryness and high albedomake Fnet negative all year-around in our simulations(not shown) as well as—to a lesser extent—in obser-vations (Chou and Neelin 2003).

The idealized model outlined above has the merit offraming the question of where convection can occur interms of the energetics of the atmosphere, but says noth-ing about the processes at play. In order to gain someinsight as to how the added energy that comes fromsummer insolation translates into added convection, weneed to look at the vertical structure of the atmospherein more detail. Although deep convection over the con-tinents is strongly modulated by the diurnal cycle andvaries on short temporal and spatial scales, the monthlymean, area-mean vertical profiles of potential temper-ature (u), equivalent potential temperature (ue), and sat-urated equivalent potential temperature (ues) are good

indicators of where deep convection can be sustainedby the large-scale environment. Convection occurswhen the free troposphere is unstable or neutral to moistconvection (]zues # 0) and when a mixed-layer plumecan reach the level of free convection [ues 5ue(mixedlayer) at the level of free convection]. More-over, the seasonal changes in u and ue indicate whetherchanges in temperature or in moisture are responsiblefor changes in convection.

Figure 3 shows how the vertical structure of the at-mosphere over two convective regions in Africa is in-fluenced by insolation; it shows the vertical profiles ofu, ue, and ues in northern Africa (208–408E, 68–128N)and southern Africa (158–358E, 58–208S). It is apparentthat insolation influences the stability of different landregions in different ways. In South Africa, insolationinduces a substantial change in the stability of the at-mosphere. The low-level austral summer warming isresponsible for the erosion of the wintertime inversionat 700 mb. This change in stability is responsible, to-gether with the increase in low-level humidity, for mak-ing the December ‘‘sounding’’ conditionally unstable(the interception between the ues line and the gray linein Fig. 3d indicates that the level of free convection forsurface plumes is at about 750 mb), supporting the De-cember increase in precipitation in southern Africa.

Conversely, the sounding in northern Africa indicatesthat the surface temperature and atmospheric stabilitybarely change in response to insolation changes betweenDecember and June. What makes the North Africasounding conditionally unstable is the increased low-level moisture due to the monsoon flow (see the inso-lation-induced changes in ue in Fig. 3e). Hence, the roleof large-scale circulation is paramount for northern sub-Saharan Africa, while it is somewhat secondary in theCongo basin, farther to the south. As an aside, we notethat, although the raison d’etre for the monsoon is thewarm continent/cold ocean summer contrast, the mon-soon rain falls on land that is no warmer during thesummer than it is in winter. It is a truism, but worthrepeating: insolation drives the African monsoon bywarming up the Sahara, not the regions where the rainends up falling. We will return to the role of the Saharain the dynamics of monsoon circulation in section 4.

3. Local control of oceanic precipitation

Figure 4a shows the difference between Septemberand March precipitation in the varying SST simulation,that is, the response of precipitation to September minusMarch SST difference. In this section, we will focus onthe oceanic portion of the response. In our experiments(which, the reader is reminded, are uncoupled), SSTdetermines the position of the ITCZ: when the belt ofhighest SST migrates to the north, so does the ITCZ.Note that while this simulation is driven by seasonalchanges in SST in the global ocean, the position of theITCZ changes in response to the local SST (this was


FIG. 3. Vertical profiles of potential temperature (u, solid), equivalent potential temperature (ue,dashed), and saturated equivalent potential temperature (ues, dotted) over (left column) northernAfrica (68–128N) and (right column) southern Africa (58–208S) during (top) Jun and (middle)Dec, and (bottom) the Jun–Dec difference in the varying insolation experiment. (a),(d) The grayvertical lines denote the surface ue.


FIG. 4. (top) Sep–Mar precipitation in the varying SST experiment.(bottom) Precipitation anomalies due to Sep–Mar elevated latent heat-ing anomalies in the oceanic ITCZ (perpetual Mar with Sep ITCZheating–perpetual Mar). The contour interval is 4 mm day21, startingwith 62, negative values are dashed, positive values larger than 2mm day21 are shaded.

confirmed by an experiment in which SST was changedindependently in the Atlantic and in the Indo–Pacific,not shown). This result supports the view of the SST–ITCZ coupling first suggested by Lindzen and Nigam(1987) in an uncoupled study of the eddy componentof the tropical circulation. The skeleton of the Lindzenand Nigam mechanism is the following: a positive sur-face temperature perturbation is communicated by tur-bulent mixing throughout the boundary layer, creates anegative low-level pressure anomaly, and thus influ-ences the boundary layer convergence that sustains con-vection. Hastenrath and Greischar (1993) have used thesame argument to explain the coherent decadal vari-ability in SST and rainfall in the tropical Atlantic. (Butnote that, when the SST is allowed to vary, the ITCZcan feedback on the ocean; e.g., Chang et al. 1997).

While it is true that boundary layer processes drivethe seasonal shift of the ITCZ in an uncoupled setting,convective heating in the ITCZ has itself the potentialto thermally drive changes in surface pressure and thussurface wind. The role of elevated condensation heatingin driving surface wind was investigated by the theo-retical study of Wu et al. (2000a,b) and was quantifiedfor the Pacific in a linearized GCM by DeWitt et al.(1996) and for the Atlantic in an intermediate-com-plexity model by Chiang et al. (2001). Recently,McGauley et al. (2004) have estimated from observa-tions a minor but sizable influence of elevated conden-

sation heating on the sea level pressure and surfacewinds in the Pacific ITCZ region.

To identify the role of elevated heating in forcing low-level winds and convergence, we will refer to a pre-scribed-heating experiment. This is another AGCM ex-periment in which both insolation and SST are fixed atMarch values (perpetual March) and in which the three-dimensional heating field was artificially changed. Atevery time step of the simulation, we discarded the sim-ulated latent heating due to convection in the oceansbetween 158S and 258N (i.e., we discard the heatingassociated with the ITCZ for March conditions) andreplaced it with a prescribed value taken from the cli-matology of the varying SST simulation and corre-sponding to September SST (i.e., we substitute the heat-ing associated with the ITCZ for September conditions).Over land and outside the band between 208S and 308Nthe model remains free to determine the latent heatingreleased in convection. The transition between the re-gions forced by prescribed and calculated latent heatingis smoothed over two grid points in every direction. Wewill refer to this simulation as the perpetual March withSeptember ITCZ heating experiment.

Note that the model is free to determine its own pre-cipitation everywhere. Thus, the simulated precipitationand the imposed latent heating are not necessarily con-sistent. Also, not all the heating associated with con-vection is prescribed: for example, radiative heating isleft for the model to calculate. We wish to disclose thatour main reason for constructing this experiment wasto look at the remote response to elevated heating inthe ITCZ (see section 4). Because of the inherent in-consistency between the heating field and the precipi-tation, temperature, and stability fields, strong cautionis advised in interpreting the local response to an im-posed heat source. We will return to this point at theend of this section.

Figure 5b shows the difference in surface winds andthe column-integrated moisture convergence betweenthe perpetual March with September ITCZ heating ex-periment and a control perpetual March experiment.(The total moisture convergence is calculated as pre-cipitation minus evaporation, as the change in the stor-age term is small.) This difference is the response forcedby differences in the elevated heating between the Sep-tember and the March ITCZ. Figure 5a shows the Sep-tember–March difference in wind and moisture con-vergence in the varying SST run, and portrays the re-sponse to differences in SST. Because the anomalies inFig. 5b develop in the absence of SST forcing, com-paring Figs. 5a and 5b gives a rough estimate of theinfluence of the elevated condensation heating upon thelow-level convergence. We interpret Fig. 5 as follows.In regions where Figs. 5a and 5b are substantially dif-ferent, boundary layer forcing primarily determines thelow-level convergence and precipitation, while elevatedheating (i.e., precipitation) can be thought of as a passiveresponse. In regions where Figs. 5a and 5b are similar,


FIG. 5. (top) Sep–Mar surface wind and column-integrated moistureconvergence in the varying SST experiment, calculated as precipi-tation minus evaporation (P 2 E). (bottom) Surface wind and column-integrated moisture convergence (P 2 E) anomalies due to Sep–Marelevated latent heating anomalies in the oceanic ITCZ (perpetual Marwith Sep ITCZ heating–perpetual Mar). The contour interval is 4mm day21, starting with 62, negative values are dashed, positivevalues larger than 2 mm day21 are shaded. Only wind vectors largerthan 2 m s21 are plotted.

the circulation forced by the elevated condensation heat-ing contributes substantially to the maintenance of thesurface winds and low-level convergence initiated bySST. The similarity between Figs. 5a and 5b would sug-gest that the elevated heating plays a substantial role,in particular in the equatorial western Atlantic and east-ern tropical Pacific.

Again, we stress that any quantitative conclusion onthe forcing of convergence by the elevated heating isprecluded in this experimental design, because of theeffect of the imposed heating on stability. The precip-itation response to the September–March difference inelevated heating in the tropical Atlantic (Fig. 4b) isweaker than the response to SST (Fig. 4a). However,one cannot measure the intensity of the response to heat-ing on the basis of how much weaker the response is:a weaker response can be easily explained in terms ofthe effect of the imposed heating on the stability of theatmospheric column. The imposed heating warms theupper troposphere more than the lower troposphere, thusincreasing the stability of the atmospheric column or,equivalently, decreasing the convective available poten-tial energy (not shown). Hence, in the region of imposedheating, deep convection is less efficient than in theoriginal simulation (the simulation from which the heat-ing field itself was taken) and precipitation is weaker

(cf. Figs. 4a and 4b). The heating anomalies are toostrong to be consistent with the precipitation anomaliesthey force. It follows that imposing the convective heat-ing as an external forcing is problematic if one is in-terested in the local response. We suggest that only thequalitative aspects of the local response to the heatingcan be considered robust. The remote response to theelevated heating in the ITCZ is described in the nextsection.

4. Oceanic control of continental precipitation

We have shown in Fig. 1e that annual changes in SSThave the ability to force annual changes in precipitationover the continents, specifically in those coastal regionsof South America and Africa that border the ITCZ, andin a zonal band that spans the whole of Africa at thelatitude of the Sahel. Figure 1f shows that the SST sea-sonal cycle is responsible for the fact that the Africanmonsoon reaches its northernmost position in late (asopposed to early) summer.

There are two possible ways for SST to influencecontinental climate: by directly changing the boundarylayer wind and the temperature and moisture advectedonshore, or by causing anomalies in oceanic precipi-tation (and thus elevated heating) which, in turn, drivefar-reaching circulation anomalies. To elucidate the rel-ative importance of the boundary layer forcing and ofthe elevated heating forcing in inducing precipitationchanges over land, we refer again to the perpetual Marchwith September ITCZ heating experiment introduced inthe previous section and presented in Figs. 4b and 5b.We recall that, in this experiment, SST and insolationare kept fixed at March values, while the elevated heat-ing field in the tropical oceans is taken from the varyingSST experiment and corresponds to September SSTconditions. Because we are now focusing on the remoteresponse, the effect of the imposed heating on localstability is inconsequential and the precipitation differ-ence between the imposed-heating simulation and a con-trol perpetual March simulation (Fig. 4b) is, over land,directly comparable to the September–March precipi-tation difference in the varying SST experiment (Fig.4a).

In the Guiana Highlands and Northeast Brazil, theresponse forced by the ITCZ heating (Fig. 4b) and thatforced by SST (Fig. 4a) are very similar, indicating thatthe influence of the ocean in this region can be explainedmainly in terms of the circulation changes that accom-pany the movement of the ITCZ, in particular the chang-es in low-level convergence.1 This is consistent with the

1 Note that the perpetual March with September ITCZ heating ex-periment underestimates the anomalies in Colombia and Venezuela,but we attribute this to the fact that the forcing in the eastern Pacificand the Gulf of Mexico is not applied at full strength, because inthose regions the forcing is a linear interpolation between the fullforcing (over the open ocean) and zero forcing (over the CentralAmerican continent and the Caribbean Islands).


FIG. 6. (top) Sep–Mar SLP and SAT differences in the varying SST experiment. (bottom)SLP and SAT differences due to Sep–Mar elevated latent heating anomalies in the oceanicITCZ (perpetual Mar with Sep ITCZ heating–perpetual Mar). SLP is contoured (c.i. 5 1 mb),and SAT is shaded (c.i. 5 18C).

fact that most of the SST-induced precipitation changesin coastal South America can be forced by changes inthe Atlantic SST alone. Conversely, over Africa the pre-cipitation response to SST is not reproduced by the per-petual March with September ITCZ heating experiment.The ITCZ heating forces anomalies of the right sign inequatorial central Africa, but their magnitude is far tooweak. More importantly, the ITCZ heating does notforce any coherent anomaly over the Sahel. We concludethat Sahel rainfall is mainly influenced by SST viaboundary layer processes and not by the circulationforced by the ITCZ. The ocean influences the Sahelthrough the advection of maritime boundary layer tem-perature anomalies over northern Africa; the tempera-ture anomalies, in turn, cause the development of anom-alies in sea level pressure (SLP) and in surface windconvergence. Specifically, the SST changes from Marchto September generate a cross-isobaric southerly windthat converges in the Sahel region, where it generatesthe observed enhancement in precipitation.

Figure 6a shows the SLP and surface air temperature(SAT) changes generated by the September–March SSTdifference. The similarity between the changes in SLPand SAT in North Africa supports the interpretation that

the SLP changes are a boundary layer response to theSAT anomalies. Figure 6b shows the SLP and SAT re-sponse to changes in the elevated heating field over thetropical oceans. In response to elevated heating, thereare substantial SLP anomalies over the Sahara, but veryweak SAT anomalies. More importantly, the low-levelwind associated with this pattern of SLP anomalies isnot converging in the Sahel (Fig. 5b).

A good indicator of the location of precipitation isthe mean boundary layer equivalent potential temper-ature (ue), which contains information about both tem-perature and humidity near the earth’s surface: the max-imum precipitation is collocated with the maximum ofue. Figure 7 portrays the effect of SST changes onboundary layer ue, SLP, and precipitation over Africa.The solid (dashed) lines in Figs. a,b portray ue (SLP)averaged over 158W–158E (Fig. 7a, West Africa and theeastern Atlantic) and 158–458E (Fig. 7b, central Africa)as a function of latitude for September and March SSTconditions (in the varying SST simulation). The hatchedvertical bands indicate regions of precipitation greaterthan 5 mm day21. The correspondence between the pre-cipitation and ue maxima is apparent.

During September, the maxima of ue and precipitation


FIG. 7. Sep (thick lines) and Mar (thin lines) latitudinal profiles of boundary layer equivalentpotential temperature (ue, solid) and SLP (dashed) over (a) West Africa (158W–158E), and (b)East Africa (158–458E) in the varying SST experiment. The hatched gray vertical bands indicatelatitudinal bands of precipitation larger than 5 mm day21 in Sep (positive slope) and in Mar(negative slope). (a) The wavy line at the bottom indicates the extent of the oceanic region.

are shifted to the north, in both West and central Africa.To the south of West Africa, the September–March dif-ference in ue is a direct expression of the changes inSST; in the Sahara, the September increase in ue is dueto changes of land surface temperature, and in the Sahelregion it is due to changes in humidity. Similarly, sea-sonal changes in ue in central Africa are due to tem-perature except in the Sahel, where they are due to hu-midity. The profiles of ue show the important role ofchanges in the Sahara in shifting the maximum ue andprecipitation in both western and central Africa.

An additional experiment (not shown), in which wedisabled changes in the elevated condensation heatingover land, confirms that boundary layer processes aloneare sufficient for establishing the bulk of the ue responsein the Sahel that is due to SST changes. At the sametime, it suggests that the feedback provided by the el-evated condensation heating is in part responsible forthe local humidity changes and the northward shift ofprecipitation (the shift is reduced by one grid point whenthe feedback is disabled).

To summarize, the northward shift of precipitation inthe Sahel is a consequence of warm SAT anomaliesdeveloping in the Sahara in response to the March-to-September changes in SST. But, how is the seasonalchange of SST communicated inland across the entireAfrican continent? The differences in the temperatureadvection and subsidence between March and Septem-ber in the varying SST simulation (not shown) indicatethat the anomalies are maintained in equilibrium by hor-izontal advection. The advection of 1) mean temperatureby the anomalous wind, 2) anomalous temperature bythe mean wind, and 3) anomalous temperature by theanomalous wind are all of comparable magnitude. Whatmaintains the equilibrium anomalies is not necessarilywhat created the anomalies in the first place, but wespeculate that advection is indeed responsible for com-municating the warm SAT anomalies forced over theocean in the subtropical North Atlantic in the zonal

direction, and thus for forcing the bulk of the SAT re-sponse in the Sahara.

Many previous studies (Folland et al. 1986; Schneideret al. 2003; Giannini et al. 2003) have linked Africanrainfall interannual and interdecadal variations to SSTforcing from the global ocean. What ocean regions havethe strongest influence on the annual cycle of NorthAfrican rainfall? Figure 8 shows the response of ue inWest and central Africa to September–March changesin SST that are imposed in all ocean basins, only in theAtlantic, and only in the Indo–Pacific. The bold grayline is the ue difference between September and Marchin the varying SST–only experiments (i.e., the differ-ence between the ue profiles in Fig. 7). The remaininglines are difference plots between simulations with vary-ing SST and with fixed SST in one or more ocean basinsand provide an alternative estimate of the response toglobal SST, as well as estimates of the response to SSTin each basin.2 This figure suggests that western Africaresponds predominantly to forcing from the AtlanticSST, while central Africa responds to forcing from allbasins. Note that the two estimates of the ue responseto the global SST derived from the Varying SST–onlysimulation and from the CTL–PMS difference, respec-tively, are qualitatively similar, but quantitatively dif-ferent. This is consistent with the conclusions of BBS1,who noted that insolation and SST interact nonlinearlyin the Sahel–Sudan region.

Finally, we note that over the Sahel, the pattern ofthe precipitation anomalies induced by the September–March SST changes (Fig. 4a) is remarkably uniform in

2 Specifically, the estimate of the response to September–MarchSST anomalies in all basins is given as the difference between acontrol simulation and one with perpetual March SST (CTL–PMS inBBS1), the response to Atlantic SST–only is given as the differencebetween CTL and a Perpetual March Atlantic SST experiment (PMASin BBS1) and the response to the Indo–Pacific SST is given as thePMAS–PMS difference.


FIG. 8. Latitudinal profiles of differences in boundary layer equivalent potential temperatureover (a) West Africa (158W–158E, and (b) East Africa (158–458E) due to the Sep–Mar differencesin SST over the global ocean (bold gray and black lines), over the Atlantic only (dashed line)and the Indo–Pacific only (dot–dashed line). The two estimates of the response to global SST arethe Sep–Mar difference in the varying SST only run (gray line) and the CTL–PMS difference inSep (black line). The hatched gray vertical bands indicate latitudinal bands of precipitation largerthan 5 mm day21 in Sep (positive slope) and in Mar (negative slope). (a) The wavy line at thebottom indicates the extent of the oceanic region.

the zonal direction. Yet, there are some interesting dif-ferences in how SST affects West and central Africa.In central Africa, the SST forcing leaves the latitudinalextent of the precipitation maximum unchanged and thearea of maximum precipitation straddles the equatorwhether SST assumes March or September values (Fig.7b). The anomalous circulation is mostly limited to thelatitude of the Sahel (Fig. 5b). In West Africa, Septem-ber SST conditions both shift the location of the pre-cipitation maximum and reduce its extent to a narrowband at about 158 (Fig. 7a). In this region, the anomalouscirculation is a cross-equatorial, monsoonlike circula-tion (Fig. 5b).

5. Continental control of oceanic precipitation

Figure 9a shows the June-minus-March difference inthe varying insolation simulation, that is, the precipi-tation changes forced by the June–March changes ininsolation over land (Note the similarity between Figs.9a and 1c). Recall that, in this experiment, SST con-ditions are fixed at the March value and the position ofthe ITCZ is anchored by the SST gradient to be nearthe equator throughout the year. The remote effect ofJune–March insolation changes over land is to increasethe precipitation intensity in the region of the ITCZ byabout 6 mm day21, with peak anomalies of 10 mmday21. The location of the ITCZ—when SSTs are pre-scribed—is not affected by changes over land. (BBS2shows that this is not the case when the atmosphere iscoupled to an ocean.)

How is the increase of oceanic precipitation intensityachieved? There are two main possibilities. The first isthat the insolation-induced changes in land surface tem-perature directly affect the sea level pressure field insuch a way as to enhance the intensity of the low-levelconvergence into the ITCZ. The second possibility is

that changes in land temperature force changes in landconvection, which in turn induce anomalies in the sta-bility of the atmosphere over the ocean.

Figure 10 (left column) documents the remote effectof insolation-induced seasonal changes over land on thevertical structure of the atmosphere in the ITCZ region;it shows the vertical profiles of u, ue, and ues averagedover the equatorial Atlantic for the months of June andMarch in the varying insolation simulation. Under Juneand March insolation conditions alike (Figs. 10a and10c), the atmosphere is conditionally unstable and thelevel of free convection (indicated in Fig. 10 by theintersection of ues with the light gray line) is reachedjust below 900 mb. Figure 10e suggests that the heavierprecipitation under June insolation condition over landfollows a cooling of the upper troposphere over theAtlantic, which is, most likely, due to the reduced con-vection over the tropical continents.3

We hypothesize that June–March land surface tem-perature changes do not have a direct effect on oceanicprecipitation; instead they cause a reduction in deepconvection over Africa and South America, which coolsthe tropical free troposphere and makes the equatorialAtlantic more convective. This mechanism is akin tothat explored by Chiang and Sobel (2002) for the At-lantic response to ENSO.

To test this hypothesis, we have performed some ad-ditional experiments with a perpetual March configu-ration of insolation and SST, and a modified elevated

3 Note that Fig. 10e does not indicate an increase in low-levelmoisture, but that would not necessarily be expected: because thebasic state is convective to begin with, any increase in moisturethrough low-level convergence would likely show up in strongerupdrafts and precipitation, and a moistening of the free troposphere,rather than a moistening of the surface. Surface convergence is indeedlarger when precipitation is more intense (not shown), but this cor-respondence shows only consistency, not causality.


FIG. 9. (a) Jun–Mar precipitation and 850-mb wind anomalies inthe varying insolation experiment (precipitation c.i. 5 4 mm day21,starting at 62 negative values are dashed and shaded in dark gray,positive values are solid and shaded in lighter gray; only wind anom-alies greater than 0.75 m s21 are plotted). (b) Precipitation and 850-mb wind difference between the perpetual Mar with Jun Africa andSouth America heating run and the perpetual Mar run [plotting con-vention as in (a)].

heating field over selected areas (Africa only, SouthAmerica only, both Africa and South America, or thewhole tropical belt except the Atlantic Ocean). In theseregions we disregard the elevated heating field com-puted by the model at any time step, and we impose aconstant heating that is taken from the output of thevarying insolation simulation and corresponds to Juneinsolation conditions over land. Any change in the in-tensity of precipitation in the Atlantic ITCZ betweenone of these ‘‘modified heating’’ simulations and a con-trol perpetual March simulation should be interpretedas forced by June–March changes in convection in theselected area.

Figure 9b shows precipitation and 850-mb windchanges forced by the June–March elevated condensa-tion heating difference over both South America and

Africa (perpetual March with June South America andAfrica heating minus perpetual March). We first notethat, as was seen in section 3 in reference to the ITCZ,forcing the atmosphere with a given elevated heatingdoes not induce a consistent precipitation intensity atthe location of the forcing (cf. Figs. 9a and 9b overSouth America and Africa). More importantly, we notethat the elevated condensation heating in the Africanand South American convection centers does indeedforce more intense rainfall in the Atlantic ITCZ, but theresponse in the central and western Atlantic is only athird of that caused by the full continental forcing (cf.Figs. 9a and 9b over the Atlantic). Additional experi-ments in which the elevated condensation heating isimposed over South Africa only or over South Americaonly indicate that the two continental convection centersplay comparable roles in producing the rainfall anom-alies in the ITCZ, but the response to African heatingis more widespread and somewhat stronger (Fig. 11).Moreover, the low-level wind in the Atlantic respondspredominantly to the African forcing.

Finally, an experiment in which the elevated conden-sation heating is imposed over the entire tropical beltexcept the Atlantic Ocean, exhibits a response in theAtlantic ITCZ that is closer to what is seen in Fig. 9ain terms of spatial extent (it reaches the northeast coastof South America) and intensity (it reaches a peak valueof 6 mm day21), but is still too weak (not shown).

Figures 10b and 10d show the profiles of u, ue, andues in the equatorial Atlantic. Both simulations weredone under March SST and insolation conditions butone (Fig. 10b) has March elevated condensation heatingconditions in South America and Africa and the other(Fig. 10d) has June elevated condensation heating im-posed over South America and Africa. It is apparent thatthe June reduction in continental convection cools theupper troposphere over the Atlantic but not the lowerhalf of the column. The overall effect is to enhanceCAPE and induce the intensification of precipitation inthe ITCZ, but such intensification does not match whatseen in the varying insolation experiment (Figs. 10a,c).We attribute this mismatch to the vertical structure ofthe cooling, which implies a small increase of stabilityat about 700 mb.

A better match between, on the one hand, the responseof Atlantic ITCZ rainfall to changes in continental el-evated heating and, on the other hand, the response tochanges in insolation over land (or, equivalently, tochanges in both convection and surface temperature) canbe achieved by adjusting the imposed heating over Af-rica and South America to mimic the inclusion of thecorresponding changes in radiative heating alongsidethe changes in condensation heating (not shown). By sodoing, we better match the changes in the heating ver-tical profile associated with changes in precipitation inthe varying insolation simulation. We conclude that theintensity of the ITCZ responds to remote convection,as hypothesized, and that the magnitude of the response


FIG. 10. Vertical profiles of potential temperature (u, solid), equivalent potential temperature(ue, dashed), and saturated equivalent potential temperature (ues, dotted) over the location of theMar maritime ITCZ (78N–08, 408–58W) in the (left column) varying insolation experiment, (b)the perpetual Mar with Jun Africa and South America heating run, and (d) the perpetual Marrun.

is quite sensitive to the details of the profile of theimposed remote heating.

Another interesting feature of the response to conti-nental heating is its spatial scale. Dry models (Gill 1980;Schneider 1987) would suggest a Tropics-wide responseto a deep convective heating (small, widespread sub-sidence would compensate for upward mass flux in theregion of the imposed positive isolated heating and viceversa). While our results support that the tropospherictemperature response is nearly homogeneous throughout

the Tropics (not shown), experiments in which elevatedcondensation heating is imposed over one continent ata time (Fig. 11) indicate that the response in precipi-tation, vertical motion, and low-level wind is inhomo-geneous. The reason for the inhomogeneity in the pre-cipitation response can be explained in terms of stabilitythreshold: if, for example, the profile is very stable, asmall upper-level cooling will not create a precipitationanomaly. Thus, we expect a larger response where thebasic state already has precipitation. To understand why


FIG. 11. (a) Precipitation and 850-mb wind difference between theperpetual Mar with Jun South Africa heating run and the perpetualMar run (precipitation c.i. 5 4 mm day21, starting at 62 negativevalues are dashed and shaded in dark gray, positive values are solidand shaded in lighter gray; only wind anomalies greater than 0.75m s21 are plotted). (b) Same as in (a) but for South American heating.

the circulation is also inhomogeneous, one has to invokethe most basic balance in the Tropics: diabatic heatingand adiabatic heating balance each other. The diabaticheating response depends on the basic state; therefore,the same must hold true for the vertical motion fieldand the associated divergent circulation. How the at-mosphere achieves this adjustment in our experimentshas not been established.

6. Summary and discussion

The purpose of this paper is to identify the mecha-nisms that determine the annual cycle of precipitationin an atmospheric general circulation model (AGCM)with prescribed SST. In particular, our aim is to identifyhow changes in local insolation affect precipitation overthe tropical continents, how changes in SST (which inour model experiments can be prescribed independentlyfrom insolation) affect precipitation in the ITCZ, andhow temperature and rainfall over the ocean affect pre-cipitation over the continents and vice versa.

We remark that, in the experiments presented in thispaper, the SST is prescribed, whereas in the real worldit is determined by both the local insolation and theoceanic response to atmospheric forcings through in-herently coupled processes. This is an important limi-tation and is partially addressed in a companion paper(BBS2). Nevertheless, this study offers some valuableinsight on how our AGCM models—and, by inference,nature—produce the observed annual cycle of precipi-tation in the tropical Atlantic sector. The general con-clusion that can be drawn from this study is that theobserved annual cycle of precipitation over the Atlanticsector is the product both of local insolation and ofinteractions between the landmasses and the ocean.Next, we offer a summary description of the mecha-nisms by which continental and maritime precipitationresponds to local and remote forcings.

a. Local control of continental precipitation

The response of continental precipitation in the Trop-ics to insolation can be explained in the context of theNeelin and Held (Neelin and Held 1987) simple theo-retical model, which states that a positive net energyinput in the atmospheric column (Fnet) is balanced byhorizontal energy export in a thermally direct circula-tion, and thus is associated with convection. To zerothorder, the varying insolation experiment supports theconclusions of this simple model: the seasonal changesin precipitation mimic the seasonal changes in Fnet inmost of the tropical landmasses, with the notable ex-ception of the Sahara, where Fnet never becomes positiveand convection cannot be sustained—even if the sea-sonal changes in Fnet are substantial. In the main con-tinental convection centers summer rainfall is enhancedbecause summer insolation decreases atmospheric sta-bility. In the monsoon regions, stability is unaffected byinsolation, and the enhanced moisture flow in the mon-soon is the key ingredient for rainfall.

The varying insolation experiment suggests the fol-lowing implications for modeling the annual cycle ofprecipitation over the tropical continents in a coupledGCM. Outside of the equatorial belt and the Sahel,where the influence of remote SST is important, grossshortcomings in the simulation of precipitation are mostlikely attributable to the atmosphere–land component ofthe model and not to the effect of coupling. In particular,the land albedo, determined by soil type, soil moisture,and vegetative cover, can strongly affect Fnet and thuscan have a large impact on precipitation.

b. Local control of oceanic precipitation

In the uncoupled simulations described in this study,the seasonal movement of the ITCZ is controlled byseasonal changes in SST: the ITCZ lies over the warmestwaters. The response of precipitation to the SST fieldis determined by boundary layer processes, as first de-


scribed by Lindzen and Nigam (1987). The atmosphericboundary layer hydrostatically adjusts to an increase inSST and produces a low-level thermal low. The result-ing wind converges moisture into the low, sustainingconvection. An experiment with prescribed elevatedcondensation heating in the ITCZ suggests that, al-though boundary layer processes are the fundamentalreason why the ITCZ follows the warmest SST, theITCZ is more than a passive participant, even in thisuncoupled experiment. The elevated condensation heat-ing in the ITCZ drives a substantial low-level flow andmoisture convergence.

This result is similar to the results of Chiang et al.(2001), who found that at interannual-to-decadal timescales both the SST gradient and the heating anomaliesassociated with the displacement of the ITCZ force sur-face convergence and meridional wind in the equatorialAtlantic. Differences in the details of the response toelevated heating between our study and theirs might beexplained by the difference in the forcings used. In thestudy of Chiang et al. (2001), the convective heatinganomalies used to force the model are mostly confinedto the equatorial belt, whereas in our study they extendto 158N. Another possible explanation is that the detailsof the simulated atmospheric response to elevated heat-ing are dependent on the characteristics of the model,such as the treatment of the boundary layer processesand radiation.

The strong surface wind response to elevated heatingsuggests the possibility of strong feedbacks between theITCZ and the SST in experiments in which the SST isinteractive. One implication is that we can expect theAtlantic ITCZ to have larger variability in POGA 1MXL simulations (Pacific Ocean Global Atmosphereand Mixed Layer: experiments in which the observedtime series of SST is prescribed in the tropical Pacific,and SST is calculated by a mixed layer model else-where) than in GOGA simulations (Global Ocean Glob-al Atmosphere: experiments in which the observed timeseries of SST is prescribed everywhere). A preliminaryanalysis of runs provided by P. Chang and R. Saravanansupports this conclusion.

Finally, we want to suggest some caution regardingthe interpretation of the local response to imposed heat-ing. An elevated heating source changes the stability ofthe atmosphere, so that the model precipitation (and thecirculation that sustains it) and the imposed heating arenot consistent with each other.

c. Remote control of continental precipitation

The response of continental precipitation to SSTchanges is concentrated in the coastal regions adjacentto the ITCZ and in the African Sahel. The seasonalchanges in precipitation in coastal regions appear to bean indirect response to SST changes and can be forcedby specifying the changes in elevated condensation heat-ing associated with the seasonal movement of the ITCZ.

In contrast, precipitation in the Sahel responds directlyto changes of SST.

In our experiments, we find that the influence of SSTon Sahel precipitation at the annual time scale can beexplained in terms of the influence of SST on the lo-cation of the ue maximum, determined by the SST fieldin the whole Atlantic basin (for West Africa) and thewhole tropical ocean (for East Africa). Specifically, theue profile in West Africa seems to evolve into its Sep-tember value by means of an interactive adjustment in-volving first the response of the boundary layer atmo-sphere in the Atlantic basin to SST changes directlyunderneath, then the advection of temperature anomaliesfrom the northern Atlantic over the Sahara, and finallythe advection of moisture from the Gulf of Guinea andequatorial Africa into the Sahel. The mechanisms bywhich the Indo–Pacific basin affects the annual cycleof precipitation in central Africa are less clear.

In summary, our results for the annual cycle suggestthat the covariability between Sahel rainfall and the At-lantic ITCZ is due to the presence of a common forcing,namely, SST, and not to the fact that the circulationassociated with the ITCZ forces anomalies in the Sahel.In contrast, precipitation in the Nordeste, the GuianaHighlands, and the Guinea region is affected by theelevated heating in the ITCZ, and only indirectly bySST.

d. Remote control of oceanic precipitation

In the uncoupled experiments presented here, themain response of the oceanic precipitation to remoteforcings coming from the continents is a change of in-tensity. The intensity of the maritime ITCZ responds tochanges in continental precipitation: less precipitationover the continents induces a cooling of the upper tro-posphere and a more unstable environment [or equiv-alently, more convective available potential energy,(CAPE)] over the Atlantic. It stands to reason, then, thata larger upper-level cooling will force a larger responsein the Atlantic ITCZ and, in fact, the ITCZ response isquite linear. For example, the responses to convectiveheating forcings imposed over Africa only and overSouth America only add up to the response obtained byimposing the forcings over both Africa and South Amer-ica together. What controls the limited spatial extent ofthe precipitation response as not been explored in detail.

Sensitivity experiments in which the heating was ad-justed to reproduce the vertical profile of the full con-vective forcing (i.e., the sum of the condensation andradiative heating associated with convection) suggestthat the ITCZ response to remote elevated heating isquite sensitive to the details of the heating profile, con-sistent with Schumacher et al. (2003).

That the oceanic convection responds to the intensityof precipitation in the tropical continents via its responseto changes in upper-troposphere temperature explainswhy land-induced precipitation anomalies over the


ocean are collocated with the basic-state ITCZ (BBS1).An upper-level decrease in temperature is felt as anincrease of CAPE in regions in which the atmosphereis conditionally unstable to begin with, but is not enoughto make a stable column unstable. Conversely, if theSLP changes induced by land surface temperature werethe main driver for ocean precipitation changes, wewould expect that the basic-state SST would not matteras much, because surface wind responds linearly to SLPchanges (Lindzen and Nigam 1987).

The response of the Atlantic ITCZ to continental pre-cipitation suggests the idea that the CCM3 biases inprecipitation over the tropical Atlantic sector—whereinprecipitation over land is overestimated and precipita-tion over the equatorial ocean is underestimated(BBS1)—might be related problems. This indicates thatchanging the land surface properties in a way as toreduce land precipitation might be a viable way to ad-dress both problems.

Acknowledgments. The authors wish to thank MikeWallace, Dennis Hartmann, Alessandra Giannini, andespecially John Chiang for their insightful suggestionsand editorial comments. This publication is supportedby a grant to the Joint Institute for the Study of theAtmosphere and Ocean (JISAO) under NOAA Coop-erative Agreement NA17RJ1232.

REFERENCES

Biasutti, M., 2000: Decadal variability in the tropical Atlantic assimulated by the Climate System Model and the CCM3 coupledto a slab ocean model. M.S. thesis, Dept. of Atmospheric Sci-ences, University of Washington, 43 pp.

——, D. S. Battisti, and E. S. Sarachik, 2003: The annual cycle overthe tropical Atlantic, South America, and Africa. J. Climate, 16,2491–2508.

——, ——, and ——, 2005: Terrestrial influence on the annual cycleof the Atlantic ITCZ in an AGCM coupled to a slab ocean. J.Climate, in press.

Chang, P., J. Li, and H. Li, 1997: A decadal climate variation in thetropical Atlantic Ocean from thermodynamic air–sea interac-tions. Nature, 385, 516–518.

Chiang, J. C. H., and A. H. Sobel, 2002: Tropical tropospheric tem-perature variations caused by ENSO and their influence on theremote tropical climate. J. Climate, 15, 2616–2631.

——, S. E. Zebiak, and M. A. Cane, 2001: Relative roles of elevatedheating and surface temperature gradients in driving anomalous

surface winds over tropical oceans. J. Atmos. Sci., 58, 1371–1394.

Chou, C., and J. D. Neelin, 2003: Mechanisms limiting the northwardextent of the northern summer convection zones. J. Climate, 16,406–425.

DeWitt, D. G., E. K. Schneider, and A. D. Vernekar, 1996: Factorsmaintaining the zonally asymmetric precipitation distributionand low-level flow in the Tropics of an atmospheric generalcirculation model: Diagnostic studies. J. Atmos. Sci., 53, 2247–2263.

Folland, C. K., T. N. Palmer, and D. Parker, 1986: Sahel rainfall andworldwide sea temperature. Nature, 320, 602–687.

Giannini, A., R. Saravanan, and P. Chang, 2003: Oceanic forcing ofSahel rainfall on interannual to interdecadal time scale. Science,302, 1027–1030.

Gill, A. E., 1980: Some simple solutions for heat-induced tropicalcirculation. Quart. J. Roy. Meteor. Soc., 106, 447–462.

Hack, J., J. Kiehl, and J. Hurrell, 1998: The hydrologic and ther-modynamic characteristics of the NCAR CCM3. J. Climate, 11,1207–1236.

Hastenrath, S., 1984: Interannual variability and the annual cycle:Mechanisms of circulation and climate in the tropical Atlanticsector. Mon.Wea. Rev., 112, 1097–1107.

——, and L. Greischar, 1993: Further work on the prediction ofNortheast Brazil rainfall anomalies. J. Climate, 6, 743–758.

Lindzen, R. S., and S. Nigam, 1987: On the role of sea surfacetemperature gradients in forcing low-level winds and conver-gence in the Tropics. J. Atmos. Sci., 44, 2418–2436.

McGauley, M., C. Zhang, and N. Bond, 2004: Large-scale charac-teristics of the atmospheric boundary layer in the eastern Pacificcold tongue–ITCZ region. J. Climate, 17, 3907–3920.

Mitchell, T. P., and J. M. Wallace, 1992: The annual cycle in equatorialconvection and sea surface temperature. J. Climate, 5, 1140–1156.

Neelin, J. D., and I. M. Held, 1987: Modeling tropical convergencebased on the moist static energy budget. Mon. Wea. Rev., 115,3–12.

Schneider, E. K., 1987: A simplified model of the modified Hadleycirculation. J. Atmos. Sci., 44, 3311–3328.

——, L. Bengtsson, and Z.-Z. Hu, 2003: Forcing the Northern Hemi-sphere climate trends. J. Atmos. Sci., 60, 1504–1521.

Schumacher, C., R. A. Houze Jr., and I. Kraucunas, 2003: The tropicaldynamical response to latent heating estimates derived fromTRMM precipitation radar. J. Atmos. Sci., 61, 1341–1358.

Sobel, A. H., and C. Bretherton, 2000: Modeling tropical precipitationin a single column. J. Climate, 13, 4378–4392.

——, ——, H. Gildor, and M. E. Peters, 2004: Convection, cloud-radiative feedbacks and thermodynamic ocean coupling in sim-ple models of the Walker circulation. Earth Climate: The Ocean–Atmosphere Interaction, Geophys. Monogr., No. 147, Amer.Geophys. Union, 1–19.

Wu, Z., D. S. Battisti, and E. S. Sarachik, 2000a: Rayleigh friction,Newtonian cooling, and the linear response to steady tropicalheating. J. Atmos. Sci., 57, 1937–1957.

——, E. S. Sarachik, and D. S. Battisti, 2000b: Vertical structure ofconvective heating and the three-dimensional structure of theforced circulation on an equatorial beta plane. J. Atmos. Sci.,57, 2169–2187.

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