Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Tropical and extratropical sources of hemispheric asymmetry of the Intertropical Convergence Zone in an idealized coupled general circulation model Journal: Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Manuscript ID Draft Manuscript Type: Research Article Date Submitted by the Author: n/a Complete List of Authors: Fuckar, Neven Stjepan; Barcelona Supercomputing Center, Earth Sciences Maroon, Elizabeth; University of Washington College of the Environment, Atmospheric Sciences Frierson, Dargan; University of Washington College of the Environment, Atmospheric Sciences Farneti, Riccardo; Abdus Salam International Centre for Theoretical Physics, Earth System Physics Keywords: Intertropical Convergence Zone, Meridional overturning circulation, Coupled climate general circulation model

Page 1 of 42Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Tropical and extratropical sources of hemispheric asymmetry of the Intertropical Convergence Zone in an idealized coupled general circulation model Neven S. Fučkar 1*, Elizabeth A. Maroon 2, Dargan M. W. Frierson 2 and Riccardo Farneti 3 1 Earth Sciences Department, Barcelona Supercomputing Center, Barcelona, Spain 2 Department of Atmospheric Sciences, University of Washington, Seattle, Washington, USA 3 Earth System Physics Section, International Centre for Theoretical Physics, Trieste, Italy Corresponding author address*: Neven S. Fučkar, Barcelona Supercomputing Center-Centro Nacional de Supercomputación (BSC-CNS), Earth Sciences Department, C. Jordi Girona 29, 08034 Barcelona, Spain E-mail: nevensf@gmail.com

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 2 of 42 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 ABSTRACT This study uses an idealized coupled climate general circulation model (GCM) to examine the role of ocean basin geometry on the structure of tropical precipitation. The north- south asymmetry of tropical rainfall is governed both by tropical ocean- atmosphere interactions due to the shape of tropical coastlines and by cross- equatorial energy transport that can be driven by the ocean s meridional overturning circulation (MOC). We compare these tropical and global effects in a coupled GCM with simplified atmospheric physics and an idealized land- ocean geometry that examines these two processes in competing roles. We use single closed ocean basins with two equatorially mirrored geometries, one with tropical coastlines that slant from southeast to northwest (westward) and the other with tropical coastlines that slants from southwest to northeast (eastward). When the MOC has no significant asymmetry, a slanted tropical coastline on the ocean basin s eastern side triggers cross- equatorial surface flow in the atmosphere towards the hemisphere with more land on the eastern side of the basin. This surface flow leads to higher tropical sea surface temperature (SST) and more precipitation in its destination hemisphere. However, on long time scales (decadal and longer) when the deep oceanic MOC develops significant asymmetry, the MOC s influence on energy transport and tropical circulation outweighs the tropical effect of the eastern slanted coastlines. We show that MOC evolution places the main deep- water source in the hemisphere with less tropical land on the eastern side of ocean basin; as a result, there is cross- equatorial ocean heat transport, OHT(y=0), towards the hemisphere with the dominant deep- water production. Southward (northward) OHT(y=0) in the basin with westward (eastward) slanted tropical coastlines induces an anomalous cross- equatorial Hadley circulation with a surface branch in the same direction as OHT(y=0). Surface moisture transport by this anomalous Hadley circulation places the maximum of tropical precipitation - i.e. the intertropical convergence zone (ITCZ) - in the southern (northern) hemisphere. We further test this relationship between the MOC and tropical precipitation with different oceanic initial conditions favoring deep- water production in different hemispheres. On long time scales, they demonstrate a very close dynamic association of the maximum of tropical SST and the ITCZ, 1

Page 3 of 42Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61 62 with the deep MOC asymmetry, OHT(y=0) and anomalous Hadley circulation. These results point to the need for the development of a general theory of tropical circulation and precipitation that encompasses both local and global mechanisms. 1. Introduction The intertropical convergence zone (ITCZ) is a narrow region of the most intense precipitation and most frequent deep convective clouds on Earth (Waliser and Gautier 1993, Philander et al 1996, Huffman et al. 2009). Figure 1 shows the long- term mean of precipitation, which has its zonal mean maximum located near 6 N (based on analysis of Xie and Arkin 1997). The ITCZ s precipitation comes from convergence and ascent of warm, moist air driven by the surface trade winds and, in specific locations, also guided by topography (Webster 2004, Xie 2004, Takahashi and Battisti 2007, Maroon et al. 2015). However, the location of the intense upward flow and ITCZ is also a crucial element of the large- scale coupled global circulation that is responsible for the maintenance of the planetary energy balance (Liu and Alexander 2007, Kang et al. 2008, Chiang 2009, Schneider et al. 2014). A general theory for tropical precipitation that would combine both local mechanisms with the global circulation and energetics is still a key open problem in climate dynamics. The influence of the meridional overturning circulation (MOC) and the associated ocean heat transport (OHT) on tropical precipitation is an active area of research. It was initiated to explain large and abrupt changes in the northern hemisphere (NH) climate during glacial periods in the late Quaternary (Dansgaard et al. 1993, Grootes and Stuiver 1997, Peterson et al. 2000, Koutavas and Lynch- Stieglitz 2004) that is hypothesized to stem from a weakening of the Atlantic MOC (Broecker et al. 1985, Broecker 2007). This led to investigation of water- hosing experiments with coupled general circulation models (GCM) that force a weakening of the Atlantic MOC with additional freshwater input in the northern Atlantic. The consequent reduction of Atlantic OHT leads to NH cooling and southward displacement of the ITCZ (e.g. Manabe and Stouffer 1995, Vellinga and Wood 2002, Zhang and Delworth 2005, Cheng et al. 2007, Chiang et al. 2008, Drijfhout 2010). 2

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 4 of 42 63 64 65 66 67 68 69 70 71 72 73 74 75 76 77 78 79 80 81 82 83 84 85 86 87 88 89 90 91 92 The Southern Ocean s zonally unconstrained Antarctic circumpolar current creates the most significant difference in large- scale ocean dynamics between the two hemispheres and has a profound impact on global climate. Different geometries of the Drake Passage, a narrow body of water between South America and the Antarctic Peninsula, can induce ocean states with rather different MOC and OHT (e.g. Toggweiler and Bjornsson 2000, Sijp and England 2004). Fučkar et al. (2013) show that different modeled sill depths of the Drake Passage can force different levels of MOC asymmetry. OHT(y=0) by the MOC moves heat away from the hemisphere with the circumpolar channel and towards the hemisphere with active deep water production. This ocean hemispheric asymmetry leads to cross- equatorial atmospheric heat transport AHT(y=0) in the opposite direction of OHT(y=0). At the equator, the Hadley circulation is responsible for the bulk of meridional AHT. Because of the vertical distribution of energy in the atmosphere, energy is transported in the direction of the Hadley circulation s upper branch, and moisture transport at the surface must then move in the opposite direction (e.g. Schneider et al. 2014). As a result, there is more tropical precipitation in the hemisphere with deep- water production. Frierson et al. (2013) explores the link between the ITCZ position and interhemispheric energetics and circulation in atmospheric reanalysis data and two GCMs with slab oceans. They claim that stronger heating of the NH atmosphere than the SH atmosphere is necessary to place the maximum of tropical precipitation north of the equator. This hemispheric asymmetry must be driven by the northward OHT(y=0) because the SH as a whole receives more net radiation at the top of the atmosphere than the NH. If the ocean did not transport sufficient amount of heat, having greater net TOA radiation in the SH would cause greater SH tropical precipitation. Frierson et al. (2013) also show the dependence of the ITCZ on OHT(y=0) in model configurations with and without continents. Marshall et al. (2014) also demonstrate the key role of OHT(y=0) in placing the maximum of tropical precipitation in the NH with observational analysis and numerical experiments with an idealized coupled GCM that includes a dynamical ocean. 3

Page 5 of 42Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 93 94 95 96 97 98 99 100 101 102 103 104 105 106 107 108 109 110 111 112 113 114 115 116 117 118 119 120 121 122 123 Understanding complex systems such as the ones governing tropical precipitation requires modeling and analysis of both regional and global climate domains at different levels of complexity (Held 2005). Realistic configuration of continents and complex physics sometimes obscure important processes from being distinctly identified. Using idealized land- ocean geometries with simplified boundary conditions and physics can further our knowledge at a more fundamental level. As these processes are understood, complexities in model physics and topography can be increased, building toward a fuller picture of the climate system. This study builds on the research of Fučkar et al. (2013) on coupling tropical and global climate by exploring the effect of slanted tropical coastlines on tropical rainfall. We contrast the roles of local atmosphere- ocean dynamics with the remote impact of the deep MOC. Section 2 describes our simplified coupled GCM and its two idealized ocean- land geometries. Section 3 describes the key aspects of the transient and equilibrium climate states in five coupled simulations. The final section contains conclusions and suggestions for future research. 2. Idealized coupled climate model We use a coupled intermediate complexity climate model (ICCM), which is derived from the Geophysical Fluid Dynamics Laboratory Climate Model version 2.0 (GFDL CM2.0; Delworth et al. 2006). It differs from CM2.0 through its simplifications of both atmospheric parameterizations and ocean- land geometry (Farneti and Vallis 2009; Vallis and Farneti 2009). The model solves the three- dimensional primitive equations for the atmosphere and ocean with a dynamically consistent surface exchange of momentum, heat, and freshwater fluxes. We use a coarse- resolution configuration with a sector atmosphere over flat land and a closed (no circumpolar channel) single- basin flat- bottom ocean. This model s simplifications allow us to examine key elements of coupled tropical and global dynamics in a more revealing geometrical setting, while being computationally efficient. The atmospheric component of ICCM has 3.75 x3 horizontal resolution and 7 vertical levels in a sector geometry that is 120 wide and spans meridionally from 84 S to 84 N. Our atmospheric GCM is based on a moist B- grid primitive equation dynamical core (GFDL 4

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 6 of 42 124 125 126 127 128 129 130 131 132 133 134 135 136 137 138 139 140 141 142 143 144 145 146 147 148 149 150 151 152 153 154 Global Atmospheric Model Development Team, 2004). It uses a gray radiation scheme, which calculates the radiative transfer of a single longwave band through prescribed optical depth (Frierson et al. 2006). As a result, the atmosphere s radiation does not depend on water vapor or clouds, but there is still latent heat release. A simplified Monin- Obukhov surface flux scheme and a K- profile boundary layer scheme are used. ICCM is forced with a time- independent, zonally uniform, top- of- atmosphere solar radiation that analytically fits the observed mean profile of insolation. Absorption of shortwave radiation in the atmosphere is neglected. A large- scale condensation scheme is applied with a simplified Betts Miller convection scheme (Betts 1986; Frierson 2007). By eliminating water vapor and cloud feedbacks (since radiative fluxes depend only on temperature) we focus mostly on the dynamical response of the coupled climate system to changes in the ocean. The oceanic component of ICCM is the Modular Ocean Model (MOM) version 4.0 (Griffies et al. 2004) and has 2 x2 horizontal resolution and 24 vertical levels. The ocean basin in both configurations is 60 wide, spanning from 70 S to 70 N, with a flat bottom at 3.9 km depth. In the first geometry (Exp I) the ocean basin has westward- slanted coastlines equatorward of 14 latitude, while in the second geometry (Exp II), the ocean basin has the same tropical coastline angles, but slanted eastward (Figure 2). The ocean physics parameterizations are based on the standard free surface MOM4 model incorporated into the CM2.0. We use constant vertical tracer diffusivity of 0.5 cm 2 s - 1 and the Gent McWilliams (GM) skew flux scheme combined with a downgradient neutral diffusion that parameterizes the effects of mesoscale eddies using a constant eddy tracer diffusivity of 800 m 2 s - 1 (Gent and McWilliams 1990, Griffies 1998). The dynamic thermodynamic sea ice simulator (SIS: Winton 2000) is computed on the ocean grid. The land component, LM2.0 (GFDL Global Atmospheric Model Development Team, 2004), is configured at the atmospheric horizontal resolution and is implemented as a collection of soil water reservoirs with constant water availability and heat capacity at each cell. The excess precipitated water is send back to the ocean at a prescribed nearby point. There are no lakes, mountains, glaciers or ice sheets in the system, and because there 5

Page 7 of 42Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 155 156 157 158 159 160 161 162 163 164 165 166 167 168 169 170 171 172 173 174 175 176 177 178 179 180 181 182 183 184 185 are no clouds, surface albedo is adjusted to obtain a realistic mean climate. For more details of ICCM setup, see Farneti and Vallis (2009). The latest version of ICCM is publicly available from GFDL as a part of MOM5 distribution (https://fms.gfdl.noaa.gov/modeling- systems- group- public- releases). Four Exp I simulations use the same basin configuration shown in Figure 2.a (I.zu, I.nh, I.bl and I.sh) but they differ in their oceanic no- flow initial conditions (IC). Exp I.zu is initialized from an ocean state with zonally uniform (zu) and hemispherically symmetric temperature and salinity fields. IC at the surface roughly match annual and zonal mean sea surface temperature (SST) and salinity from the World Ocean Atlas 2009 (WOA09: Locarnini, R. A. et al. 2009, Antonov, J. I. et al. 2009). They exponentially decay to the model s bottom where they roughly match the surface IC at the poleward edges of the ocean basin. Exp I.zu is integrated for 1000 years. Temperature and salinity from Exp I.zu are averaged over the last 100 years of the simulation to construct oceanic IC for the other Exp I simulations and for the Exp II simulation (details of additional simulations are in the following section). ICCM s atmosphere and land are initialized from a default uniform IC because they dynamically equilibrate with the rest of coupled system on short time scales (within a year). 3. Results Westward or eastward slanted tropical coastlines are the only boundary condition asymmetry between hemispheres in our coupled model. As such, the tropical coastlines are forcing any aspect of hemispheric climate asymmetry, directly or indirectly. We focus on the coupled ocean- atmosphere response of the tropics and the extratropics (i.e. the global system) to the different tilts of tropical coastlines. The tropical and extratropical responses have different characteristic time scales and different amplitudes of impact on the tropical atmospheric circulation and precipitation. 3.1 Experiments with westward slanted tropical coastlines 6

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 8 of 42 186 187 188 189 190 191 192 193 194 195 196 197 198 199 200 201 202 203 204 205 206 207 208 209 210 211 212 213 214 215 216 The atmosphere and ocean are strongly coupled in the tropics where atmospheric convection and precipitation are closely associated with SST and the most active part of the variability spectrum is on time scales shorter than decadal (e.g., Sarachik and Cane 2010). The average of the first 20 years of Exp I.zu has more tropical precipitation in the NH than in the Southern Hemisphere (SH, shading in Figure 3.a). This precipitation anomaly is coupled to the anomalous southerly cross- equatorial surface wind through the wind- evaporation- SST (WES) feedback (Xie and Philander 1994). The cross- equatorial C shape of the anomalous surface wind on the eastern side of the ocean basin (Figure 3.a) is a signature of the WES feedback (Xie 2004, Fučkar et al. 2013). On short time scales, the WES feedback connects the SH tropics with its stronger easterlies, stronger evaporation, and lower SST to the weaker easterlies, weaker evaporation, and higher SST in the NH tropics. This tropical asymmetry is induced by the westward slanted coastline on the eastern side of the ocean basin due to the propagation of anomalies by the trade winds. Because of ICCM s relatively coarse atmospheric resolution, annual mean insolation, and lack of water vapor and cloud feedbacks, the spatial structure of tropical precipitation does not shift very far from the equator. On long time scales, however, the sign and magnitude of anomalies in tropical surface wind, SST and precipitation in our model is not predominantly controlled by local processes. Figure 3.b shows that on multi- decadal to centennial time scales, zonal mean anomalous SST (shading), anomalous precipitation (contours), and surface wind stress (vectors) in the tropics reverse sign, which places the maximum of precipitation in the SH. The Exp I.zu simulation reaches equilibrium after approximately 300 years; hence, in Figure 3.c, we examine the average surface tropical conditions from years 381 to 400. Figure 3.c shows that the WES feedback is still present on the eastern edge of the basin, forced by the westward slanted tropical coastline. Nonetheless, there is more tropical precipitation in the SH than in the NH due to the anomalous northerly cross- equatorial surface wind over the majority of the ocean basin. Surface winds over land in the tropics do not change sufficiently to play a role in this transition. In the coupled equilibrium state, the cross- equatorial C shape of anomalous surface winds in the middle of the basin overpowers a smaller cross- equatorial C shape of anomalous northward winds on the east of the basin. 7

Page 9 of 42Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 217 218 219 220 221 222 223 224 225 226 227 228 229 230 231 232 233 234 235 236 237 238 239 240 241 242 243 244 245 246 247 This anomalous surface structure and time evolution in the tropics agrees with the results of Fučkar et al. (2013); in that study, there were no slanted coastlines, and the deep MOC asymmetry alone forced the tropical hemispheric symmetry. Figure 4.a examines the key factor in the slow transition in Exp I.zu, the evolution of the deep MOC. Because this experiment is initialized from zonally uniform oceanic IC, the MOC behaves erratically over the first 50 years as it adjusts; afterwards, the main deep- water source becomes firmly anchored in the SH (blue curve). The deep MOC in Exp I.zu develops a stable hemispheric asymmetry after roughly 400 years of integration. OHT(y=0) (Figure 4.b) is also southward when SH deep- water production becomes greater than NH deep- water production (Figure 4.a). Southward or negative OHT(y=0) accompanies the deep MOC s asymmetry and leads to greater extratropical heat release to the atmosphere in the SH than in the NH (not shown). An increased (decreased) extratropical heat release from the ocean to the atmosphere in the SH (NH) makes the SH (NH) the warmer (colder) hemisphere. The extratropical heat release decreases (increases) the meridional surface temperature gradient between the equator and extratropics, weakening (strengthening) transient eddies. The affected eddies interact with the poleward edge of the Hadley circulation, and influence its strength (Kang et al 2008). The Hadley circulation asymmetry index (blue curve) in Figure 4.b shows that MOC- forced hemispheric asymmetry strengthens (weakens) the Hadley cell in the SH (NH). In comparison to an equatorially symmetric state, an anomalous cross- equatorial Hadley cell develops with its lower (upper) branch transporting moisture (heat) to the SH (NH). At the surface, the anomalous Hadley circulation causes the mainly northerly anomalous surface winds in deep tropics over the ocean (Figure 3c). In the tropics, SST and precipitation anomalies are coupled on short time scales (Philander et al. 1996, Chang et al 2004, Xie 2004). However, Figure 4.c shows that the evolution of the interhemispheric differences of precipitation and SST on long time scales is primarily determined by the deep- MOC asymmetry and OHT(y=0) in our model. This extratropically- forced asymmetry evolves in spite of the persistent forcing of moist surface air towards the 8

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 10 of 42 248 249 250 251 252 253 254 255 256 257 258 259 260 261 262 263 264 265 266 267 268 269 270 271 272 273 274 275 276 277 278 NH at the eastern edge of the ocean basin by the WES feedback (Figure 3.c). Despite the subtlety of the precipitation pattern shifts, there is distinct hemispheric asymmetry in the amount of precipitation that develops. This relative NH versus SH precipitation asymmetry, not the absolute ITCZ position, is the focus of our work, because the mechanisms important for the hemispheric precipitation asymmetry are relevant for the observed ITCZ. After 1000 years of integration the asymmetry of the climate state of Exp I.zu did not switch hemispheres or change magnitude. Figure 5.b shows the year 801-1000 average of the MOC streamfunction. The deep MOC asymmetry controls the OHT and extratropical surface heat flux asymmetry. This surface heat flux asymmetry, in turn, drives the anomalous Hadley and Ferrel cells shown in Figure 5.a. through changes in the meridional temperature gradient at the surface. The Hadley circulation is thermally direct (Dima and Wallace 2003, Webster 2004), so coupled tropical atmosphere ocean dynamics place the ascending branch of the anomalous Hadley cell in the hemisphere warmed by the OHT, which also contains the main source of deep- water production. The maximum of tropical precipitation accompanies the ascending branch of the Hadley circulation. We can further test this relation between the deep MOC and tropical precipitation on long time scales suggested by the Exp I.zu simulation by using different oceanic IC that are conducive to deep- water production in a specific hemisphere or in neither hemisphere. We use the average ocean temperature and salinity from Exp I.zu between years 901 and 1000 and as a zero- flow IC that favors the main deep- water source in the SH from year 1 (the Exp I.sh simulation). The symmetric component with respect to the equator and the western boundary of the basin of the same Exp I.zu temperature and salinity fields are used to produce a hemispherically balanced (bl) zero- flow IC in Exp I.bl simulation. Finally, we mirror the IC of Exp I.sh across the equator to create zero- flow IC that favors placing the main deep- water source in the NH at the start of the integration in Exp I.nh simulation. We integrated these three additional IC simulations with the same ocean- land geometry as Exp I.zu for 500 years. Exp I.nh starts with vigorous deep- water production in the NH and it takes more than 100 years for the MOC to reorganize its structure and shift the main deep- 9

Page 11 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 279 280 281 282 283 284 285 286 287 288 289 290 291 292 293 294 295 296 297 298 299 300 301 302 303 304 305 306 307 308 309 water source into the SH (Figure 6.a). OHT(y=0) and the Hadley circulation asymmetry (Figure 6.c), as well as the tropical SST and precipitation asymmetry, (Figure 6.f) again closely follow the evolution of deep MOC on decadal and longer time scales. In the first 100 years, northward cross- equatorial OHT decreases (increases) the surface meridional temperature gradient in the NH (SH) leading to a weaker (stronger) Hadley cell in that hemisphere. An anomalous cross- equatorial Hadley cell has its lower (upper) branch transporting moisture (heat) toward the NH (SH); this leads to higher tropical SST and precipitation in the NH. However, by the second century of simulation, the deep MOC asymmetry reverses and forces the associated reversal of the tropical circulation and precipitation asymmetries, firmly placing the maximum of tropical precipitation in the SH. The IC of Exp I.bl are in an interhemispheric sense closest to the IC of Exp I.zu, but are dynamically quasi- balanced to avoid the random unstable MOC behavior that occurred at the beginning of Exp I.zu (Figure 4.a). As a result, the deep MOC asymmetry builds gradually in Exp I.bl (Figure 6.b). This asymmetry development takes more than 100 years for asymmetry to favor the SH over the NH. Before the significant MOC asymmetry occurs, local tropical processes place the maximum of tropical SST and precipitation in the NH (Figure 6.h). However, once the MOC anchors the main source of deep water in the SH, the asymmetry in the Hadley circulation (Fig. 6e), tropical SST, and tropical precipitation follows the MOC asymmetry. Again, the ITCZ moves to the SH once the deep MOC asymmetry fully develops (Figure 6.h). The oceanic IC in Exp I.sh favor and produce a vigorous deep- water production in the SH from the beginning of the simulation (Figure 6.c). This extratropical asymmetry forces an anomalous Hadley circulation that very quickly overpowers the local tropical processes driven by the slanted coastline (Figure 6.f). The Hadley circulation asymmetry anchors the ITCZ in the SH in the first decade of model integration (Figure 6.i). Figure 7 shows the time evolution of tropical adjustment in Exp I.nh, Exp I.bl and Exp I.sh. The long time scales involved are uncharacteristic for the tropical coupled ocean- atmosphere and follow the asymmetry in global circulation driven by the deep MOC evolution (Figure 6). The equilibrium state with southward C shaped mean wind stress in 10

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 12 of 42 310 311 312 313 314 315 316 317 318 319 320 321 322 323 324 325 326 327 328 329 330 331 332 333 334 335 336 337 338 339 340 deep tropics, SH maximum tropical SST, and SH maximum precipitation occurs at substantially different times in each these simulation. The overturning streamfunction in the atmosphere and ocean from the last 100 years of Exp I.nh, Exp I.bl and Exp I.sh (not shown) all match the equilibrium state of Exp I.zu shown in Figure 5. Overall, the time evolution of all Exp I simulations shows that the deep MOC asymmetry forces southward OHT(y=0), southward anomalous surface moisture transport by the Hadley circulation, and the maximum of tropical precipitation in the SH in equilibrium state regardless of which IC were used. 3.2 Experiment with eastward slanted tropical coastlines As compared to the Exp I. simulations, the Exp II.bl simulation uses equatorially mirrored ocean- land geometry. Exp II.bl s oceanic IC are aligned with respect to the western boundary of the ocean basin but are otherwise the same as those used in the Exp I.bl simulation. With this simulation, we additionally verify if the hemispheric asymmetry discussed in the previous section consistently switches sign with the reversed coastlines. With this additional ocean- land configuration, the mean surface conditions in the first 20 years show more tropical precipitation in the SH than in the NH (shading in Figure 8.a): the greater SH precipitation in this transient state is due to anomalous northerly equatorial surface wind on the eastern side of the ocean basin. WES feedback is initiated by the eastern tropical coastline, but in this simulation it is slanted eastward with increasing latitude, the opposite direction from all the Exp I simulations. On the eastern side of the basin in Exp II.bl, ocean- land geometry induces an anomalous interhemispheric surface pressure gradient due to anomalous interhemispheric surface temperature differences. In this case higher surface pressure over the colder ocean in the NH tropics and lower surface pressure over the warmer land in the SH tropics around 90 E forces northerly cross- equatorial surface wind there. This anomalous wind obtains easterly (westerly) component north (south) of the equator due to the Coriolis force leading to westward increase (decrease) of the trade winds. This wind speed increase (decrease) causes further intensification (reduction) of evaporative cooling, which, in turn, further decreases 11

Page 13 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 341 342 343 344 345 346 347 348 349 350 351 352 353 354 355 356 357 358 359 360 361 362 363 364 365 366 367 368 369 370 371 (increases) the surface temperature to the north (south) of the equator. This cross- equatorial pattern amplifies the initial surface temperature perturbation on the eastern side of the basin. The WES- feedback propagates this surface temperature anomaly westward, which decreases (increases) convective activity and precipitation over the ocean basin in the NH (SH) tropics. In our model, the westward propagation of the WES feedback decays about 30 west from the eastern boundary (Figures 8.a and 3.a): the WES- induced cross- equatorial dipole anomalies also propagate equatorward in the region of background easterly wind (Xie 2004) so they dissipate when they reach the equator. In Exp II.bl, the anomalies of tropical wind stress (Figure 8.b, vectors), SST (shading) and precipitation (contours) are not predominantly controlled by local processes on long time scales. After about 100 years, these anomalies reverse sign, placing the maximum of precipitation in the NH. After Exp II.bl reaches equilibrium around year 400, the average surface tropical conditions along the eastern edge of the basin still show evidence of the WES feedback. However, the maximum of tropical precipitation across the ocean basin is in the NH: anomalous southerly surface winds across the equator occur throughout most of the ocean basin, pushing precipitation northward. This anomalous surface structure and its time evolution in the tropics reflect the dominance of the anomalous cross- equatorial Hadley circulation over the ocean. This coupled circulation s asymmetry about the equator is induced by the MOC asymmetry, just as in the Exp I simulations. In Exp II.bl the main source of multidecadal changes again is driven by the evolution of the deep MOC (Figure 9.a). This ICCM experiment is initialized from dynamically quasi- balanced oceanic IC; as a result, the MOC avoids erratic behavior at the beginning of the simulation. The dominance of deep- water production in the NH over the SH is established before the end of first century. OHT(y=0) (Figure 9.b, red curve) and the Hadley circulation asymmetry index (blue curve) show less correlation in the first 50 years because the deep MOC asymmetry has not chosen a dominant hemisphere yet. As a result, tropical ocean- atmosphere processes exert the dominant control over the Hadley cells during the first 50 years (also evident in Figure 8.b). Afterwards, however, an anomalous Hadley circulation develops with its lower branch transporting moisture into the NH. Furthermore, Figure 9.c 12

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 14 of 42 372 373 374 375 376 377 378 379 380 381 382 383 384 385 386 387 388 389 390 391 392 393 394 395 396 397 398 399 400 401 402 again shows co- evolution of the interhemispheric differences of tropical SST (blue curve) and precipitation (red curve) on long time scales. Exp II.bl reaches a stable equilibrium climate after about 300 years. Figure 10.b shows the 801-1000 year average of the hemispherically asymmetric MOC streamfunction. This asymmetric streamfunction is mirrored across the equator from the Exp I simulations shown in Figure 5.b for Exp I.zu. The direction of the MOC controls the hemispheric asymmetry of the coupled global circulation and energetics on long time scales in our model. The deep MOC determines the direction and the amplitude of the anomalous Hadley and Ferrel cell shown in Figure 10.a. The ascending branch of the anomalous Hadley cell and the maximum of tropical precipitation is anchored in the NH, which has the dominant source of deep water. 4. Conclusions and future directions Five numerical experiments with the ICCM model, an idealized coupled GCM with simplified atmospheric physics and comprehensive ocean physics, are presented here to examine the competition of local effects of tropical coastlines and global coupled overturning circulation on tropical circulation and precipitation. These simulations show that slanted tropical coastlines affect tropical precipitation via the WES feedback near the eastern coast of the basin through the entire integration of all simulations. However, this local coupled ocean- atmosphere process determines the position of ITCZ only if the ocean s deep MOC does not develop a significant interhemispheric asymmetry with an opposing effect. Once the asymmetry of ocean circulation places the main source of deep- water production in one hemisphere, then the whole climate follows that asymmetry. The hemisphere with the prevailing deep- water production also contains greater tropical precipitation, irrespective of the local forcing of slanted tropical coastline on the eastern side of the ocean basin. In all numerical experiments, during both the transient evolution and in steady states, the deep MOC asymmetry (measured by our index as the maximum NH MOC streamfunction - 13

Page 15 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 403 404 405 406 407 408 409 410 minimum SH MOC streamfunction) is a useful linear predictor of the cross- equatorial OHT (Figure 11.a) and the interhemispheric asymmetry of the extratropical surface heat flux (Figure 11.b). The ocean basin geometry forces the dominant deep- water production in the hemisphere with less tropical land on the eastern side of the basin. The MOC that develops with greater deep water production in one hemisphere causes OHT(y=0) into this same hemisphere. The hemisphere heated by the ocean circulation is warmed due to greater extratropical heat release from the ocean to the atmosphere, as compared to the opposite hemisphere. The hemisphere that is relatively warmed (cooled) by the OHT develops a 411 weaker (stronger) meridional surface temperature gradient between the tropics and 412 413 414 415 416 417 418 419 extratropics. This weaker (stronger) temperature gradient, in turn, weakens (strengthens) the Hadley and Ferrel cells in that hemisphere. This change in the global atmospheric overturning circulation (Figure 11.c) is manifested in the tropics as anomalous cross- equatorial Hadley cell that transports moisture (heat) by its lower (upper) branch toward the direction of warmer (colder) hemisphere. The cross- equatorial flow of warm and moist air over the most of the ocean basin anchors the ascending branch of the Hadley circulation and the maximum of tropical precipitation in the hemisphere with the main deep- water source (Figure 11.d). 420 421 422 423 424 425 426 427 428 429 430 431 432 On Earth, the continents in all ocean basins have much more of a westward tilt than an eastward tilt. Our idealized modeling results suggest that this tropical effect on its own would cause the ITCZ to be located in the NH as in observations only if there is no significant interhemispheric asymmetry between extratropics in coupled general circulation, but such global asymmetry is clearly evident in the present climate (Liu and Alexander, 2007). We suggest that the key ingredient from these experiments is the same as that identified by Fučkar et al (2013) and Frierson et al (2013), the oceanic MOC. Specifically, in the Fučkar et al (2013) study, we showed in this same model that opening a Drake passage- like channel in the SH is sufficient to force anchoring the deep water production of the MOC in the NH, and shift the ITCZ northward. We consider this study to be strong additional evidence that whatever causes the MOC to transport heat northward across the equator in the ocean also causes the ITCZ to be in the NH. 433 14

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 16 of 42 434 435 436 437 438 439 440 441 442 443 444 445 446 447 448 449 450 451 452 453 454 455 456 457 458 459 460 461 462 463 464 Our use of an idealized model with an intermediate level of complexity, while useful for ease of interpretation, also means that there are caveats with this work that should be tested in follow- up studies with models of increasing levels of complexity. For example, would the deep MOC asymmetry still determine the main position of the ITCZ with more comprehensive model physics (e.g. clouds), model geometry (e.g., higher resolution, mountains and more realistic ocean- land configuration) and SW forcing (e.g. seasonal and diurnal solar variation). This gray atmosphere shares much code with its cousin, the comprehensive GFDL AM2 model; the tropical circulation in this gray atmosphere responds with weaker magnitude than AM2, but both respond to various forcings with the same interhemispheric asymmetry (Kang et al. 2008, 2009, Seo et al. 2014, Maroon et al. in press); we anticipate that the inclusion of more comprehensive atmospheric physics would amplify the hemispheric response that we see in this study. Furthermore, orography, especially the Andes mountain range, has a known effect on the ITCZ location in the eastern tropical Pacific with local WES and stratus cloud feedbacks involved (Takahashi and Battisti, 2007, Maroon et al. 2015). Local tropical processes induced by the Andes and other mountain ranges should be also fully tested as well in coupled climate models. While our ICCM model lacks comprehensive atmospheric physics and real- world topography and bathymetry, it does include a comprehensive ocean model; as such, ICCM, alongside more complex coupled models, is a useful tool for improving the understanding of tropical- extratropical coupled dynamics. Examining the interaction of the MOC with tropical climate at different time scales would benefit our understanding from seasonal climate to paleoclimate dynamics. Overall, the development of an encompassing global theory of tropical circulation and precipitation would benefit climate predictions and projections. Acknowledgments The authors thank Shang- Ping Xie, LuAnne Thompson, Cecilia Bitz, David Battisti, and Xiaojuan Liu for valuable discussions. D.M.W.F was supported by NSF grants AGS- 1359464, PLR- 1341497, and a UW Royalty Research Fund award. E.A.M. was supported by a NDSEG fellowship and an NSF IGERT Program on Ocean Change traineeship. 15

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Page 19 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 543 544 545 546 547 548 549 550 551 552 553 554 555 556 557 558 559 560 561 562 563 564 565 566 567 568 569 570 571 572 573 574 575 576 577 578 579 580 581 582 583 Kang, S. M., I. M. Held, D. M. W. Frierson, and M. Zhao, 2008: The response of the ITCZ to extratropical thermal forcing: Idealized slab- ocean experiments with AGCM. J. Climate, 21, 3521 3532. Kang, S.M., D. M. W. Frierson, and I. M. Held, 2009: The Tropical Response to Extratropical Thermal Forcing in an Idealized GCM: The Importance of Radiative Feedbacks and Convective Parameterization. J. Atmos. Sci., 66, 2812-2827. Kang, S. M., I. M. Held, and S.- P. Xie, 2014: Contrasting the Tropical Responses to Zonally Asymmetric Extratropical and Tropical Thermal Forcing, Clim. Dyn. 42, 2033-2043. Koutavas, A., and J. Lynch- Stieglitz, 2004: Variability of the marine ITCZ over the eastern Pacific during the past 30,000 years: Regional perspective and global context. The Hadley Circulation: Present, Past and Future, H. F. Diaz and R. S. Bradley, Eds., Springer, 347 369 Liu, Z., and M. Alexander (2007), Atmospheric bridge, oceanic tunnel, and global climatic teleconnections, Rev. Geophys., 45, RG2005, doi:10.1029/2005rg000172. Locarnini, R. A., A. V. Mishonov, J. I. Antonov, T. P. Boyer, H. E. Garcia, O. K. Baranova, M. M. Zweng, and D. R. Johnson, 2010: World Ocean Atlas 2009, Volume 1: Temperature. S. Levitus, Ed. NOAA Atlas NESDIS 68, U.S. Government Printing Office, Washington, D.C., 184 pp. Manabe, S., and R. J. Stouffer, 1995: Simulation of abrupt climate change induced by freshwater input to the North Atlantic Ocean. Nature, 378, 165 167. Maroon, E. A., D. M. W. Frierson, D. S. Battisti, 2015: The Tropical Precipitation Response to Andes Topography and Ocean Heat Fluxes in an Aquaplanet Model. J. Climate 28:1, 381-398. Maroon, E. A., D. M. W. Frierson, S. M. Kang, and J. Scheff, (in press): The precipitation response to an idealized subtropical continent. J. Climate, http://dx.doi.org/10.1175/jcli- D- 15-0616.1 Marshall, J., Donohoe, A., Ferreira, D. and McGee, D, 2013: The role of the ocean circulation in setting the mean position of the ITCZ. Clim. Dyn. http://dx.doi.org/10.1007/s00382-013- 1767- z Peterson, L. C., G. H. Haug, K. A. Hughen, and U. Rohl, 2000: Rapid changes in the hydrologic cycle of the tropical Atlantic during the last glacial. Science, 290, 1947 1951. 18

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 20 of 42 584 585 586 587 588 589 590 591 592 593 594 595 596 597 598 599 600 601 602 603 604 605 606 607 608 609 610 611 612 613 614 615 616 617 618 619 620 621 622 623 Philander, S. G. H., D. Gu, G. Lambert, T. Li, D. Halpern, N.- C. Lau, and R. C. Pacanowski, 1996: Why the ITCZ is mostly north of the equator. J. Climate, 9, 2958 2972. Sarachik, E.S., and M.A. Cane, 2010: The El Niño- Southern Oscillation Phenomenon, Cambridge University Press, 513, 45, 384 pp Schneider, T., T. Bischoff, and G.H. Haug, 2014: Migrations and dynamics of the intertropical convergence zone. Nature, doi:10.1038/nature13636 Seo, J., S. M. Kang, D. M. W. Frierson, 2014: Sensitivity of Intertropical Convergence Zone Movement to the Latitudinal Position of Thermal Forcing. J. Climate 27:8, 3035-3042. Sijp, W. P., and M. H. England, 2004: Effect of the Drake Passage throughflow on global climate. J. Phys. Oceanogr., 34, 1254 1266. Takahashi, K., and D. S. Battisti, 2007: Processes Controlling the Mean Tropical Pacific Precipitation Pattern. Part I: The Andes and the Eastern Pacific ITCZ. J. Climate, 20:14, 3434-3451 Toggweiler, J. R., and H. Bjornsson, 2000: Drake Passage and paleoclimate. J. Quat. Sci., 15, 319 328. Vallis, G.K., and R. Farneti, 2009: Meridional energy transport in the coupled atmosphere ocean system: Scaling and numerical experiments. Quart. J. Roy. Meteor. Soc., 135, 1643 1660. Vellinga, M., and R.A. Wood, 2002: Global climatic impacts of a collapse of the Atlantic thermohaline circulation. Climatic Change, 54, 3, 251-267, doi:10.1023/a:1016168827653 Waliser, D. E., and C.A. Gautier, 1993: A satellite- derived climatology of the ITCZ. J. Clim. 6, 2162 2174. Webster, P. J., 2004: The elementary Hadley circulation. The Hadley Circulation: Present, Past and Future, H. F. Diaz and R. S. Bradley, Eds., Springer, 9 60. Winton, M., 2000: A reformulated three- layer sea ice model. J. Atmos. Oceanic Technol., 17, 525 531. Xie, S.- P. and S.G.H. Philander, 1994: A coupled ocean- atmosphere model of relevance to the ITCZ in the eastern Pacific. Tellus, 46A, 340-350. 19

Page 21 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 624 625 626 627 628 629 630 631 632 633 634 635 636 637 638 Xie, S.- P., 2004: The shape of continents, air- sea interaction and the rising branch of the Hadley circulation. The Hadley Circulation: Present, Past and Future, H. F. Diaz and R. S. Bradley, Eds., Springer, 121 152. Xie, P., and P.A. Arkin, 1997: Global precipitation: A 17- year monthly analysis based on gauge observations, satellite estimates, and numerical model outputs. Bull. Amer. Meteor. Soc., 78, 2539-2558. Zhang, R., and T. L. Delworth, 2005: Simulated tropical response to a substantial weakening of the Atlantic thermohaline circulation. J. Climate, 18, 1853 1860. 639 640 641 642 643 Figure 1. The 1981-2010 averaged precipitation based on monthly NOAA CPC Merged Analysis of Precipitation (CMAP) in which rain gauge data are merged with precipitation estimates from multiple satellite- based (infrared and microwave) products. 20

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 22 of 42 644 645 646 647 648 Figure 2. Schematic outline of 60 - wide closed- ocean basins with (a) westward and (b) eastward slanted tropical coastlines (equatorward of 14 latitude), under a 120 - wide sector atmosphere. 21

Page 23 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 649 650 651 652 653 654 655 656 657 658 659 Figure 3. Results of the Exp I.zu simulation: (a) and (c) show average tropical precipitation (shading) overlaid with anomalous surface wind vectors (with respect to an equatorially symmetric zonal component and an antisymmetric meridional component) between years 1 and 20, and between years 381 and 400, respectively. Vectors outside of the deep tropics with anomalous wind speed above 2 m/s are removed for easier visualization. Blue lines mark the ocean basin boundaries. (b) Time- latitude diagram from year 1 to 400 showing annual zonal means of hemispherically anomalous SST (shading), anomalous precipitation (blue/red = positive/negative contours) and anomalous surface wind stress vectors. Anomalous precipitation was filtered with a low- pass Butterworth filter with a 40- year period cutoff. 22

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 24 of 42 660 661 662 663 Figure 4. Results of Exp I.zu simulation: (a) Maximum and absolute minimum of the NH (red curve) and the SH (blue curve) annual mean deep MOC overturning stream function 23

Page 25 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 664 665 666 667 668 669 670 671 (below 200m poleward of 20 latitude), respectively. (b) Ocean heat transport across the equator (red curve) and the Hadley circulation asymmetry index (blue curve) defined as - (minimum streamfunction value of the SH Hadley cell + maximum streamfunction value of the NH Hadley cell). (c) Tropical precipitation (red curve) and SST (blue curve) asymmetry indices (average between 20 N and the equator minus average between the equator and 20 S). Quantities in panels (b) and (c) are annual means smoothed by a 15- year boxcar filter. 672 673 674 675 676 677 Figure 5. The average of Exp I.zu simulation between year 801 and 1000 of (a) anomalous atmospheric overturning streamfunction (with respect to the hemispherically antisymmetric component), and (b) MOC streamfunction. Positive values of streamfunction indicate a clockwise circulation. 24

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 26 of 42 678 679 680 681 682 683 684 685 686 Figure 6. Annual mean results of Exp I.nh, Exp I.bl and Exp I.sh simulations in left, middle and right columns, respectively. The top row shows maximum and absolute minimum of the NH (red curve) and the SH (blue curve) deep MOC streamfunction, respectively. The middle row shows ocean heat transport across the equator (red curve) and Hadley circulation asymmetry index (blue curve). The bottom row shows tropical precipitation (red curve) and SST (blue curve) asymmetry indices. Panels in the middle and bottom rows show annual means smoothed by a 15- year boxcar filter. 25

Page 27 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 687 688 689 690 691 692 Figure 7. Results of Exp I.nh, Exp I.bl and Exp I.sh simulations in top, middle and bottom rows, respectively. Time- latitude diagrams from year 1 to 400 showing annual zonal means of anomalous SST (fill), anomalous precipitation (blue/red = positive/negative contours) and anomalous surface wind stress vectors, as in Figure 3b. 26

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 28 of 42 693 694 695 696 Figure 8. Figure 3. Results of Exp II.bl simulation: The quantities in this figure are the same as in Figure 3. 27

Page 29 of 42 Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal 697 698 699 700 Figure 9. Results of Exp II.bl simulation after 15- year box smoothing of annual means: The quantities in this figure are the same as in Figure 4. 28

Manuscripts submitted to Dynamics and Statistics of the Climate System: An Interdisciplinary Journal Page 30 of 42 701 702 703 704 705 706 Figure 10. The average of Exp II.bl simulation between year 801 and 1000 of (a) anomalous atmospheric overturning streamfunction (with respect to the antisymmetric component) and (b) MOC streamfunction. Positive values of streamfunctions show clockwise circulation. 29