For Review Only - Department of Atmospheric Sciencesdargan/papers/fmff... · 2016. 4. 20. · For...

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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

https://mc.manuscriptcentral.com/dscs

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

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Tropical and extratropical sources of hemispheric asymmetry of

the Intertropical Convergence Zone in an idealized coupled general circulation model

Neven S. Fučkar1*, Elizabeth A. Maroon2, Dargan M. W. Frierson 2

and Riccardo Farneti3

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: [email protected]

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ABSTRACT 1

2

This study uses an idealized coupled climate general circulation model (GCM) to examine 3

the role of ocean basin geometry on the structure of tropical precipitation. The north-‐south 4

asymmetry of tropical rainfall is governed both by tropical ocean-‐atmosphere interactions 5

due to the shape of tropical coastlines and by cross-‐equatorial energy transport that can be 6

driven by the ocean’s meridional overturning circulation (MOC). We compare these tropical 7

and global effects in a coupled GCM with simplified atmospheric physics and an idealized 8

land-‐ocean geometry that examines these two processes in competing roles. We use single 9

closed ocean basins with two equatorially mirrored geometries, one with tropical 10

coastlines that slant from southeast to northwest (westward) and the other with tropical 11

coastlines that slants from southwest to northeast (eastward). 12

13

When the MOC has no significant asymmetry, a slanted tropical coastline on the ocean 14

basin’s eastern side triggers cross-‐equatorial surface flow in the atmosphere towards the 15

hemisphere with more land on the eastern side of the basin. This surface flow leads to 16

higher tropical sea surface temperature (SST) and more precipitation in its destination 17

hemisphere. However, on long time scales (decadal and longer) when the deep oceanic 18

MOC develops significant asymmetry, the MOC’s influence on energy transport and tropical 19

circulation outweighs the tropical effect of the eastern slanted coastlines. We show that 20

MOC evolution places the main deep-‐water source in the hemisphere with less tropical land 21

on the eastern side of ocean basin; as a result, there is cross-‐equatorial ocean heat 22

transport, OHT(y=0), towards the hemisphere with the dominant deep-‐water production. 23

Southward (northward) OHT(y=0) in the basin with westward (eastward) slanted tropical 24

coastlines induces an anomalous cross-‐equatorial Hadley circulation with a surface branch 25

in the same direction as OHT(y=0). Surface moisture transport by this anomalous Hadley 26

circulation places the maximum of tropical precipitation -‐ i.e. the intertropical convergence 27

zone (ITCZ) -‐ in the southern (northern) hemisphere. We further test this relationship 28

between the MOC and tropical precipitation with different oceanic initial conditions 29

favoring deep-‐water production in different hemispheres. On long time scales, they 30

demonstrate a very close dynamic association of the maximum of tropical SST and the ITCZ, 31

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with the deep MOC asymmetry, OHT(y=0) and anomalous Hadley circulation. These results 32

point to the need for the development of a general theory of tropical circulation and 33

precipitation that encompasses both local and global mechanisms. 34

35

1. Introduction 36

37

The intertropical convergence zone (ITCZ) is a narrow region of the most intense 38

precipitation and most frequent deep convective clouds on Earth (Waliser and Gautier 39

1993, Philander et al 1996, Huffman et al. 2009). Figure 1 shows the long-‐term mean of 40

precipitation, which has its zonal mean maximum located near 6°N (based on analysis of 41

Xie and Arkin 1997). The ITCZ’s precipitation comes from convergence and ascent of warm, 42

moist air driven by the surface trade winds and, in specific locations, also guided by 43

topography (Webster 2004, Xie 2004, Takahashi and Battisti 2007, Maroon et al. 2015). 44

However, the location of the intense upward flow and ITCZ is also a crucial element of the 45

large-‐scale coupled global circulation that is responsible for the maintenance of the 46

planetary energy balance (Liu and Alexander 2007, Kang et al. 2008, Chiang 2009, 47

Schneider et al. 2014). A general theory for tropical precipitation that would combine both 48

local mechanisms with the global circulation and energetics is still a key open problem in 49

climate dynamics. 50

51

The influence of the meridional overturning circulation (MOC) and the associated ocean 52

heat transport (OHT) on tropical precipitation is an active area of research. It was initiated 53

to explain large and abrupt changes in the northern hemisphere (NH) climate during glacial 54

periods in the late Quaternary (Dansgaard et al. 1993, Grootes and Stuiver 1997, Peterson 55

et al. 2000, Koutavas and Lynch-‐Stieglitz 2004) that is hypothesized to stem from a 56

weakening of the Atlantic MOC (Broecker et al. 1985, Broecker 2007). This led to 57

investigation of “water-‐hosing” experiments with coupled general circulation models 58

(GCM) that force a weakening of the Atlantic MOC with additional freshwater input in the 59

northern Atlantic. The consequent reduction of Atlantic OHT leads to NH cooling and 60

southward displacement of the ITCZ (e.g. Manabe and Stouffer 1995, Vellinga and Wood 61

2002, Zhang and Delworth 2005, Cheng et al. 2007, Chiang et al. 2008, Drijfhout 2010). 62

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63

The Southern Ocean’s zonally unconstrained Antarctic circumpolar current creates the 64

most significant difference in large-‐scale ocean dynamics between the two hemispheres 65

and has a profound impact on global climate. Different geometries of the Drake Passage, a 66

narrow body of water between South America and the Antarctic Peninsula, can induce 67

ocean states with rather different MOC and OHT (e.g. Toggweiler and Bjornsson 2000, Sijp 68

and England 2004). Fučkar et al. (2013) show that different modeled sill depths of the 69

Drake Passage can force different levels of MOC asymmetry. OHT(y=0) by the MOC moves 70

heat away from the hemisphere with the circumpolar channel and towards the hemisphere 71

with active deep water production. This ocean hemispheric asymmetry leads to cross-‐72

equatorial atmospheric heat transport AHT(y=0) in the opposite direction of OHT(y=0). At 73

the equator, the Hadley circulation is responsible for the bulk of meridional AHT. Because 74

of the vertical distribution of energy in the atmosphere, energy is transported in the 75

direction of the Hadley circulation’s upper branch, and moisture transport at the surface 76

must then move in the opposite direction (e.g. Schneider et al. 2014). As a result, there is 77

more tropical precipitation in the hemisphere with deep-‐water production. 78

79

Frierson et al. (2013) explores the link between the ITCZ position and interhemispheric 80

energetics and circulation in atmospheric reanalysis data and two GCMs with slab oceans. 81

They claim that stronger heating of the NH atmosphere than the SH atmosphere is 82

necessary to place the maximum of tropical precipitation north of the equator. This 83

hemispheric asymmetry must be driven by the northward OHT(y=0) because the SH as a 84

whole receives more net radiation at the top of the atmosphere than the NH. If the ocean 85

did not transport sufficient amount of heat, having greater net TOA radiation in the SH 86

would cause greater SH tropical precipitation. Frierson et al. (2013) also show the 87

dependence of the ITCZ on OHT(y=0) in model configurations with and without continents. 88

Marshall et al. (2014) also demonstrate the key role of OHT(y=0) in placing the maximum 89

of tropical precipitation in the NH with observational analysis and numerical experiments 90

with an idealized coupled GCM that includes a dynamical ocean. 91

92

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Understanding complex systems such as the ones governing tropical precipitation requires 93

modeling and analysis of both regional and global climate domains at different levels of 94

complexity (Held 2005). Realistic configuration of continents and complex physics 95

sometimes obscure important processes from being distinctly identified. Using idealized 96

land-‐ocean geometries with simplified boundary conditions and physics can further our 97

knowledge at a more fundamental level. As these processes are understood, complexities in 98

model physics and topography can be increased, building toward a fuller picture of the 99

climate system. This study builds on the research of Fučkar et al. (2013) on coupling 100

tropical and global climate by exploring the effect of slanted tropical coastlines on tropical 101

rainfall. We contrast the roles of local atmosphere-‐ocean dynamics with the remote impact 102

of the deep MOC. Section 2 describes our simplified coupled GCM and its two idealized 103

ocean-‐land geometries. Section 3 describes the key aspects of the transient and equilibrium 104

climate states in five coupled simulations. The final section contains conclusions and 105

suggestions for future research. 106

107

2. Idealized coupled climate model 108

109

We use a coupled intermediate complexity climate model (ICCM), which is derived from the 110

Geophysical Fluid Dynamics Laboratory Climate Model version 2.0 (GFDL CM2.0; Delworth 111

et al. 2006). It differs from CM2.0 through its simplifications of both atmospheric 112

parameterizations and ocean-‐land geometry (Farneti and Vallis 2009; Vallis and Farneti 113

2009). The model solves the three-‐dimensional primitive equations for the atmosphere and 114

ocean with a dynamically consistent surface exchange of momentum, heat, and freshwater 115

fluxes. We use a coarse-‐resolution configuration with a sector atmosphere over flat land 116

and a closed (no circumpolar channel) single-‐basin flat-‐bottom ocean. This model’s 117

simplifications allow us to examine key elements of coupled tropical and global dynamics 118

in a more revealing geometrical setting, while being computationally efficient. 119

120

The atmospheric component of ICCM has 3.75°x3° horizontal resolution and 7 vertical 121

levels in a sector geometry that is 120° wide and spans meridionally from 84°S to 84°N. 122

Our atmospheric GCM is based on a moist B-‐grid primitive equation dynamical core (GFDL 123

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Global Atmospheric Model Development Team, 2004). It uses a gray radiation scheme, 124

which calculates the radiative transfer of a single longwave band through prescribed 125

optical depth (Frierson et al. 2006). As a result, the atmosphere’s radiation does not 126

depend on water vapor or clouds, but there is still latent heat release. A simplified Monin-‐127

Obukhov surface flux scheme and a K-‐profile boundary layer scheme are used. ICCM is 128

forced with a time-‐independent, zonally uniform, top-‐of-‐atmosphere solar radiation that 129

analytically fits the observed mean profile of insolation. Absorption of shortwave radiation 130

in the atmosphere is neglected. A large-‐scale condensation scheme is applied with a 131

simplified Betts–Miller convection scheme (Betts 1986; Frierson 2007). By eliminating 132

water vapor and cloud feedbacks (since radiative fluxes depend only on temperature) we 133

focus mostly on the dynamical response of the coupled climate system to changes in the 134

ocean. 135

136

The oceanic component of ICCM is the Modular Ocean Model (MOM) version 4.0 (Griffies et 137

al. 2004) and has 2°x2° horizontal resolution and 24 vertical levels. The ocean basin in both 138

configurations is 60° wide, spanning from 70°S to 70°N, with a flat bottom at 3.9 km depth. 139

In the first geometry (Exp I) the ocean basin has westward-‐slanted coastlines equatorward 140

of 14° latitude, while in the second geometry (Exp II), the ocean basin has the same tropical 141

coastline angles, but slanted eastward (Figure 2). The ocean physics parameterizations are 142

based on the standard free surface MOM4 model incorporated into the CM2.0. We use 143

constant vertical tracer diffusivity of 0.5 cm2 s-‐1 and the Gent–McWilliams (GM) skew flux 144

scheme combined with a downgradient neutral diffusion that parameterizes the effects of 145

mesoscale eddies using a constant eddy tracer diffusivity of 800 m2 s-‐1 (Gent and 146

McWilliams 1990, Griffies 1998). 147

148

The dynamic–thermodynamic sea ice simulator (SIS: Winton 2000) is computed on the 149

ocean grid. The land component, LM2.0 (GFDL Global Atmospheric Model Development 150

Team, 2004), is configured at the atmospheric horizontal resolution and is implemented as 151

a collection of soil water reservoirs with constant water availability and heat capacity at 152

each cell. The excess precipitated water is send back to the ocean at a prescribed nearby 153

point. There are no lakes, mountains, glaciers or ice sheets in the system, and because there 154

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are no clouds, surface albedo is adjusted to obtain a realistic mean climate. For more details 155

of ICCM setup, see Farneti and Vallis (2009). The latest version of ICCM is publicly available 156

from GFDL as a part of MOM5 distribution (https://fms.gfdl.noaa.gov/modeling-‐systems-‐157

group-‐public-‐releases). 158

159

Four Exp I simulations use the same basin configuration shown in Figure 2.a (I.zu, I.nh, I.bl 160

and I.sh) but they differ in their oceanic no-‐flow initial conditions (IC). Exp I.zu is initialized 161

from an ocean state with zonally uniform (zu) and hemispherically symmetric temperature 162

and salinity fields. IC at the surface roughly match annual and zonal mean sea surface 163

temperature (SST) and salinity from the World Ocean Atlas 2009 (WOA09: Locarnini, R. A. 164

et al. 2009, Antonov, J. I. et al. 2009). They exponentially decay to the model’s bottom 165

where they roughly match the surface IC at the poleward edges of the ocean basin. Exp I.zu 166

is integrated for 1000 years. Temperature and salinity from Exp I.zu are averaged over the 167

last 100 years of the simulation to construct oceanic IC for the other Exp I simulations and 168

for the Exp II simulation (details of additional simulations are in the following section). 169

ICCM’s atmosphere and land are initialized from a default uniform IC because they 170

dynamically equilibrate with the rest of coupled system on short time scales (within a 171

year). 172

173

3. Results 174

175

Westward or eastward slanted tropical coastlines are the only boundary condition 176

asymmetry between hemispheres in our coupled model. As such, the tropical coastlines are 177

forcing any aspect of hemispheric climate asymmetry, directly or indirectly. We focus on 178

the coupled ocean-‐atmosphere response of the tropics and the extratropics (i.e. the global 179

system) to the different tilts of tropical coastlines. The tropical and extratropical responses 180

have different characteristic time scales and different amplitudes of impact on the tropical 181

atmospheric circulation and precipitation. 182

183

3.1 Experiments with westward slanted tropical coastlines 184

185

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The atmosphere and ocean are strongly coupled in the tropics where atmospheric 186

convection and precipitation are closely associated with SST and the most active part of the 187

variability spectrum is on time scales shorter than decadal (e.g., Sarachik and Cane 2010). 188

The average of the first 20 years of Exp I.zu has more tropical precipitation in the NH than 189

in the Southern Hemisphere (SH, shading in Figure 3.a). This precipitation anomaly is 190

coupled to the anomalous southerly cross-‐equatorial surface wind through the wind-‐191

evaporation-‐SST (WES) feedback (Xie and Philander 1994). The cross-‐equatorial ‘‘C shape’’ 192

of the anomalous surface wind on the eastern side of the ocean basin (Figure 3.a) is a 193

signature of the WES feedback (Xie 2004, Fučkar et al. 2013). On short time scales, the WES 194

feedback connects the SH tropics with its stronger easterlies, stronger evaporation, and 195

lower SST to the weaker easterlies, weaker evaporation, and higher SST in the NH tropics. 196

This tropical asymmetry is induced by the westward slanted coastline on the eastern side 197

of the ocean basin due to the propagation of anomalies by the trade winds. Because of 198

ICCM’s relatively coarse atmospheric resolution, annual mean insolation, and lack of water 199

vapor and cloud feedbacks, the spatial structure of tropical precipitation does not shift very 200

far from the equator. 201

202

On long time scales, however, the sign and magnitude of anomalies in tropical surface wind, 203

SST and precipitation in our model is not predominantly controlled by local processes. 204

Figure 3.b shows that on multi-‐decadal to centennial time scales, zonal mean anomalous 205

SST (shading), anomalous precipitation (contours), and surface wind stress (vectors) in the 206

tropics reverse sign, which places the maximum of precipitation in the SH. The Exp I.zu 207

simulation reaches equilibrium after approximately 300 years; hence, in Figure 3.c, we 208

examine the average surface tropical conditions from years 381 to 400. Figure 3.c shows 209

that the WES feedback is still present on the eastern edge of the basin, forced by the 210

westward slanted tropical coastline. Nonetheless, there is more tropical precipitation in the 211

SH than in the NH due to the anomalous northerly cross-‐equatorial surface wind over the 212

majority of the ocean basin. Surface winds over land in the tropics do not change 213

sufficiently to play a role in this transition. In the coupled equilibrium state, the cross-‐214

equatorial “C shape” of anomalous surface winds in the middle of the basin overpowers a 215

smaller cross-‐equatorial “C shape” of anomalous northward winds on the east of the basin. 216

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This anomalous surface structure and time evolution in the tropics agrees with the results 217

of Fučkar et al. (2013); in that study, there were no slanted coastlines, and the deep MOC 218

asymmetry alone forced the tropical hemispheric symmetry. 219

220 Figure 4.a examines the key factor in the slow transition in Exp I.zu, the evolution of the 221

deep MOC. Because this experiment is initialized from zonally uniform oceanic IC, the MOC 222

behaves erratically over the first 50 years as it adjusts; afterwards, the main deep-‐water 223

source becomes firmly anchored in the SH (blue curve). The deep MOC in Exp I.zu develops 224

a stable hemispheric asymmetry after roughly 400 years of integration. 225

226

OHT(y=0) (Figure 4.b) is also southward when SH deep-‐water production becomes greater 227

than NH deep-‐water production (Figure 4.a). Southward or negative OHT(y=0) 228

accompanies the deep MOC’s asymmetry and leads to greater extratropical heat release to 229

the atmosphere in the SH than in the NH (not shown). An increased (decreased) 230

extratropical heat release from the ocean to the atmosphere in the SH (NH) makes the SH 231

(NH) the warmer (colder) hemisphere. The extratropical heat release decreases (increases) 232

the meridional surface temperature gradient between the equator and extratropics, 233

weakening (strengthening) transient eddies. The affected eddies interact with the 234

poleward edge of the Hadley circulation, and influence its strength (Kang et al 2008). The 235

Hadley circulation asymmetry index (blue curve) in Figure 4.b shows that MOC-‐forced 236

hemispheric asymmetry strengthens (weakens) the Hadley cell in the SH (NH). In 237

comparison to an equatorially symmetric state, an anomalous cross-‐equatorial Hadley cell 238

develops with its lower (upper) branch transporting moisture (heat) to the SH (NH). At the 239

surface, the anomalous Hadley circulation causes the mainly northerly anomalous surface 240

winds in deep tropics over the ocean (Figure 3c). 241

242

In the tropics, SST and precipitation anomalies are coupled on short time scales (Philander 243

et al. 1996, Chang et al 2004, Xie 2004). However, Figure 4.c shows that the evolution of the 244

interhemispheric differences of precipitation and SST on long time scales is primarily 245

determined by the deep-‐MOC asymmetry and OHT(y=0) in our model. This extratropically-‐246

forced asymmetry evolves in spite of the persistent forcing of moist surface air towards the 247

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NH at the eastern edge of the ocean basin by the WES feedback (Figure 3.c). Despite the 248

subtlety of the precipitation pattern shifts, there is distinct hemispheric asymmetry in the 249

amount of precipitation that develops. This relative NH versus SH precipitation asymmetry, 250

not the absolute ITCZ position, is the focus of our work, because the mechanisms important 251

for the hemispheric precipitation asymmetry are relevant for the observed ITCZ. 252

253 After 1000 years of integration the asymmetry of the climate state of Exp I.zu did not 254

switch hemispheres or change magnitude. Figure 5.b shows the year 801-‐1000 average of 255

the MOC streamfunction. The deep MOC asymmetry controls the OHT and extratropical 256

surface heat flux asymmetry. This surface heat flux asymmetry, in turn, drives the 257

anomalous Hadley and Ferrel cells shown in Figure 5.a. through changes in the meridional 258

temperature gradient at the surface. The Hadley circulation is thermally direct (Dima and 259

Wallace 2003, Webster 2004), so coupled tropical atmosphere–ocean dynamics place the 260

ascending branch of the anomalous Hadley cell in the hemisphere warmed by the OHT, 261

which also contains the main source of deep-‐water production. The maximum of tropical 262

precipitation accompanies the ascending branch of the Hadley circulation. 263

264

We can further test this relation between the deep MOC and tropical precipitation on long 265

time scales suggested by the Exp I.zu simulation by using different oceanic IC that are 266

conducive to deep-‐water production in a specific hemisphere or in neither hemisphere. We 267

use the average ocean temperature and salinity from Exp I.zu between years 901 and 1000 268

and as a zero-‐flow IC that favors the main deep-‐water source in the SH from year 1 (the Exp 269

I.sh simulation). The symmetric component with respect to the equator and the western 270

boundary of the basin of the same Exp I.zu temperature and salinity fields are used to 271

produce a hemispherically balanced (bl) zero-‐flow IC in Exp I.bl simulation. Finally, we 272

mirror the IC of Exp I.sh across the equator to create zero-‐flow IC that favors placing the 273

main deep-‐water source in the NH at the start of the integration in Exp I.nh simulation. 274

275

We integrated these three additional IC simulations with the same ocean-‐land geometry as 276

Exp I.zu for 500 years. Exp I.nh starts with vigorous deep-‐water production in the NH and it 277

takes more than 100 years for the MOC to reorganize its structure and shift the main deep-‐278

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water source into the SH (Figure 6.a). OHT(y=0) and the Hadley circulation asymmetry 279

(Figure 6.c), as well as the tropical SST and precipitation asymmetry, (Figure 6.f) again 280

closely follow the evolution of deep MOC on decadal and longer time scales. In the first 100 281

years, northward cross-‐equatorial OHT decreases (increases) the surface meridional 282

temperature gradient in the NH (SH) leading to a weaker (stronger) Hadley cell in that 283

hemisphere. An anomalous cross-‐equatorial Hadley cell has its lower (upper) branch 284

transporting moisture (heat) toward the NH (SH); this leads to higher tropical SST and 285

precipitation in the NH. However, by the second century of simulation, the deep MOC 286

asymmetry reverses and forces the associated reversal of the tropical circulation and 287

precipitation asymmetries, firmly placing the maximum of tropical precipitation in the SH. 288

289

The IC of Exp I.bl are in an interhemispheric sense closest to the IC of Exp I.zu, but are 290

dynamically quasi-‐balanced to avoid the random unstable MOC behavior that occurred at 291

the beginning of Exp I.zu (Figure 4.a). As a result, the deep MOC asymmetry builds 292

gradually in Exp I.bl (Figure 6.b). This asymmetry development takes more than 100 years 293

for asymmetry to favor the SH over the NH. Before the significant MOC asymmetry occurs, 294

local tropical processes place the maximum of tropical SST and precipitation in the NH 295

(Figure 6.h). However, once the MOC anchors the main source of deep water in the SH, the 296

asymmetry in the Hadley circulation (Fig. 6e), tropical SST, and tropical precipitation 297

follows the MOC asymmetry. Again, the ITCZ moves to the SH once the deep MOC 298

asymmetry fully develops (Figure 6.h). The oceanic IC in Exp I.sh favor and produce a 299

vigorous deep-‐water production in the SH from the beginning of the simulation (Figure 300

6.c). This extratropical asymmetry forces an anomalous Hadley circulation that very 301

quickly overpowers the local tropical processes driven by the slanted coastline (Figure 6.f). 302

The Hadley circulation asymmetry anchors the ITCZ in the SH in the first decade of model 303

integration (Figure 6.i). 304

305

Figure 7 shows the time evolution of tropical adjustment in Exp I.nh, Exp I.bl and Exp I.sh. 306

The long time scales involved are uncharacteristic for the tropical coupled ocean-‐307

atmosphere and follow the asymmetry in global circulation driven by the deep MOC 308

evolution (Figure 6). The equilibrium state with southward “C shaped” mean wind stress in 309

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deep tropics, SH maximum tropical SST, and SH maximum precipitation occurs at 310

substantially different times in each these simulation. The overturning streamfunction in 311

the atmosphere and ocean from the last 100 years of Exp I.nh, Exp I.bl and Exp I.sh (not 312

shown) all match the equilibrium state of Exp I.zu shown in Figure 5. Overall, the time 313

evolution of all Exp I simulations shows that the deep MOC asymmetry forces southward 314

OHT(y=0), southward anomalous surface moisture transport by the Hadley circulation, and 315

the maximum of tropical precipitation in the SH in equilibrium state regardless of which IC 316

were used. 317

318

3.2 Experiment with eastward slanted tropical coastlines 319

320

As compared to the Exp I. simulations, the Exp II.bl simulation uses equatorially mirrored 321

ocean-‐land geometry. Exp II.bl’s oceanic IC are aligned with respect to the western 322

boundary of the ocean basin but are otherwise the same as those used in the Exp I.bl 323

simulation. With this simulation, we additionally verify if the hemispheric asymmetry 324

discussed in the previous section consistently switches sign with the reversed coastlines. 325

326

With this additional ocean-‐land configuration, the mean surface conditions in the first 20 327

years show more tropical precipitation in the SH than in the NH (shading in Figure 8.a): the 328

greater SH precipitation in this transient state is due to anomalous northerly equatorial 329

surface wind on the eastern side of the ocean basin. WES feedback is initiated by the 330

eastern tropical coastline, but in this simulation it is slanted eastward with increasing 331

latitude, the opposite direction from all the Exp I simulations. On the eastern side of the 332

basin in Exp II.bl, ocean-‐land geometry induces an anomalous interhemispheric surface 333

pressure gradient due to anomalous interhemispheric surface temperature differences. In 334

this case higher surface pressure over the colder ocean in the NH tropics and lower surface 335

pressure over the warmer land in the SH tropics around 90°E forces northerly cross-‐336

equatorial surface wind there. This anomalous wind obtains easterly (westerly) component 337

north (south) of the equator due to the Coriolis force leading to westward increase 338

(decrease) of the trade winds. This wind speed increase (decrease) causes further 339

intensification (reduction) of evaporative cooling, which, in turn, further decreases 340

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(increases) the surface temperature to the north (south) of the equator. This cross-‐341

equatorial pattern amplifies the initial surface temperature perturbation on the eastern 342

side of the basin. The WES-‐feedback propagates this surface temperature anomaly 343

westward, which decreases (increases) convective activity and precipitation over the ocean 344

basin in the NH (SH) tropics. In our model, the westward propagation of the WES feedback 345

decays about 30° west from the eastern boundary (Figures 8.a and 3.a): the WES-‐induced 346

cross-‐equatorial dipole anomalies also propagate equatorward in the region of background 347

easterly wind (Xie 2004) so they dissipate when they reach the equator. 348

349

In Exp II.bl, the anomalies of tropical wind stress (Figure 8.b, vectors), SST (shading) and 350

precipitation (contours) are not predominantly controlled by local processes on long time 351

scales. After about 100 years, these anomalies reverse sign, placing the maximum of 352

precipitation in the NH. After Exp II.bl reaches equilibrium around year 400, the average 353

surface tropical conditions along the eastern edge of the basin still show evidence of the 354

WES feedback. However, the maximum of tropical precipitation across the ocean basin is in 355

the NH: anomalous southerly surface winds across the equator occur throughout most of 356

the ocean basin, pushing precipitation northward. This anomalous surface structure and its 357

time evolution in the tropics reflect the dominance of the anomalous cross-‐equatorial 358

Hadley circulation over the ocean. This coupled circulation’s asymmetry about the equator 359

is induced by the MOC asymmetry, just as in the Exp I simulations. 360

361

In Exp II.bl the main source of multidecadal changes again is driven by the evolution of the 362

deep MOC (Figure 9.a). This ICCM experiment is initialized from dynamically quasi-‐363

balanced oceanic IC; as a result, the MOC avoids erratic behavior at the beginning of the 364

simulation. The dominance of deep-‐water production in the NH over the SH is established 365

before the end of first century. OHT(y=0) (Figure 9.b, red curve) and the Hadley circulation 366

asymmetry index (blue curve) show less correlation in the first 50 years because the deep 367

MOC asymmetry has not chosen a dominant hemisphere yet. As a result, tropical ocean-‐368

atmosphere processes exert the dominant control over the Hadley cells during the first 50 369

years (also evident in Figure 8.b). Afterwards, however, an anomalous Hadley circulation 370

develops with its lower branch transporting moisture into the NH. Furthermore, Figure 9.c 371

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again shows co-‐evolution of the interhemispheric differences of tropical SST (blue curve) 372

and precipitation (red curve) on long time scales. 373

374

Exp II.bl reaches a stable equilibrium climate after about 300 years. Figure 10.b shows the 375

801-‐1000 year average of the hemispherically asymmetric MOC streamfunction. This 376

asymmetric streamfunction is mirrored across the equator from the Exp I simulations 377

shown in Figure 5.b for Exp I.zu. The direction of the MOC controls the hemispheric 378

asymmetry of the coupled global circulation and energetics on long time scales in our 379

model. The deep MOC determines the direction and the amplitude of the anomalous Hadley 380

and Ferrel cell shown in Figure 10.a. The ascending branch of the anomalous Hadley cell 381

and the maximum of tropical precipitation is anchored in the NH, which has the dominant 382

source of deep water. 383

384

4. Conclusions and future directions 385

386

Five numerical experiments with the ICCM model, an idealized coupled GCM with 387

simplified atmospheric physics and comprehensive ocean physics, are presented here to 388

examine the competition of local effects of tropical coastlines and global coupled 389

overturning circulation on tropical circulation and precipitation. These simulations show 390

that slanted tropical coastlines affect tropical precipitation via the WES feedback near the 391

eastern coast of the basin through the entire integration of all simulations. However, this 392

local coupled ocean-‐atmosphere process determines the position of ITCZ only if the ocean’s 393

deep MOC does not develop a significant interhemispheric asymmetry with an opposing 394

effect. Once the asymmetry of ocean circulation places the main source of deep-‐water 395

production in one hemisphere, then the whole climate follows that asymmetry. The 396

hemisphere with the prevailing deep-‐water production also contains greater tropical 397

precipitation, irrespective of the local forcing of slanted tropical coastline on the eastern 398

side of the ocean basin. 399

400

In all numerical experiments, during both the transient evolution and in steady states, the 401

deep MOC asymmetry (measured by our index as the maximum NH MOC streamfunction -‐ 402

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minimum SH MOC streamfunction) is a useful linear predictor of the cross-‐equatorial OHT 403

(Figure 11.a) and the interhemispheric asymmetry of the extratropical surface heat flux 404

(Figure 11.b). The ocean basin geometry forces the dominant deep-‐water production in the 405

hemisphere with less tropical land on the eastern side of the basin. The MOC that develops 406

with greater deep water production in one hemisphere causes OHT(y=0) into this same 407

hemisphere. The hemisphere heated by the ocean circulation is warmed due to greater 408

extratropical heat release from the ocean to the atmosphere, as compared to the opposite 409

hemisphere. The hemisphere that is relatively warmed (cooled) by the OHT develops a 410

weaker (stronger) meridional surface temperature gradient between the tropics and 411

extratropics. This weaker (stronger) temperature gradient, in turn, weakens (strengthens) 412

the Hadley and Ferrel cells in that hemisphere. This change in the global atmospheric 413

overturning circulation (Figure 11.c) is manifested in the tropics as anomalous cross-‐414

equatorial Hadley cell that transports moisture (heat) by its lower (upper) branch toward 415

the direction of warmer (colder) hemisphere. The cross-‐equatorial flow of warm and moist 416

air over the most of the ocean basin anchors the ascending branch of the Hadley circulation 417

and the maximum of tropical precipitation in the hemisphere with the main deep-‐water 418

source (Figure 11.d). 419

420

On Earth, the continents in all ocean basins have much more of a westward tilt than an 421

eastward tilt. Our idealized modeling results suggest that this tropical effect on its own 422

would cause the ITCZ to be located in the NH as in observations only if there is no 423

significant interhemispheric asymmetry between extratropics in coupled general 424

circulation, but such global asymmetry is clearly evident in the present climate (Liu and 425

Alexander, 2007). We suggest that the key ingredient from these experiments is the same 426

as that identified by Fučkar et al (2013) and Frierson et al (2013), the oceanic MOC. 427

Specifically, in the Fučkar et al (2013) study, we showed in this same model that opening a 428

Drake passage-‐like channel in the SH is sufficient to force anchoring the deep water 429

production of the MOC in the NH, and shift the ITCZ northward. We consider this study to 430

be strong additional evidence that whatever causes the MOC to transport heat northward 431

across the equator in the ocean also causes the ITCZ to be in the NH. 432

433

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Our use of an idealized model with an intermediate level of complexity, while useful for 434

ease of interpretation, also means that there are caveats with this work that should be 435

tested in follow-‐up studies with models of increasing levels of complexity. For example, 436

would the deep MOC asymmetry still determine the main position of the ITCZ with more 437

comprehensive model physics (e.g. clouds), model geometry (e.g., higher resolution, 438

mountains and more realistic ocean-‐land configuration) and SW forcing (e.g. seasonal and 439

diurnal solar variation). This gray atmosphere shares much code with its cousin, the 440

comprehensive GFDL AM2 model; the tropical circulation in this gray atmosphere responds 441

with weaker magnitude than AM2, but both respond to various forcings with the same 442

interhemispheric asymmetry (Kang et al. 2008, 2009, Seo et al. 2014, Maroon et al. in 443

press); we anticipate that the inclusion of more comprehensive atmospheric physics would 444

amplify the hemispheric response that we see in this study. Furthermore, orography, 445

especially the Andes mountain range, has a known effect on the ITCZ location in the eastern 446

tropical Pacific with local WES and stratus cloud feedbacks involved (Takahashi and 447

Battisti, 2007, Maroon et al. 2015). Local tropical processes induced by the Andes and 448

other mountain ranges should be also fully tested as well in coupled climate models. While 449

our ICCM model lacks comprehensive atmospheric physics and real-‐world topography and 450

bathymetry, it does include a comprehensive ocean model; as such, ICCM, alongside more 451

complex coupled models, is a useful tool for improving the understanding of tropical-‐452

extratropical coupled dynamics. Examining the interaction of the MOC with tropical climate 453

at different time scales would benefit our understanding from seasonal climate to 454

paleoclimate dynamics. Overall, the development of an encompassing global theory of 455

tropical circulation and precipitation would benefit climate predictions and projections. 456

457

Acknowledgments 458

459

The authors thank Shang-‐Ping Xie, LuAnne Thompson, Cecilia Bitz, David Battisti, and 460

Xiaojuan Liu for valuable discussions. D.M.W.F was supported by NSF grants AGS-‐1359464, 461

PLR-‐1341497, and a UW Royalty Research Fund award. E.A.M. was supported by a NDSEG 462

fellowship and an NSF IGERT Program on Ocean Change traineeship. 463

464

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References 465

466

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Huffman, G. J., Adler, R. F., Bolvin, D. T. & Gu, G, 2009: Improving the global precipitation 541 record: GPCP Version 2.1. Geophys. Res. Lett. 36, L17808 542

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Kang, S. M., I. M. Held, D. M. W. Frierson, and M. Zhao, 2008: The response of the ITCZ to 543 extratropical thermal forcing: Idealized slab-‐ocean experiments with AGCM. J. Climate, 21, 544 3521–3532. 545 546 Kang, S.M., D. M. W. Frierson, and I. M. Held, 2009: The Tropical Response to Extratropical 547 Thermal Forcing in an Idealized GCM: The Importance of Radiative Feedbacks and 548 Convective Parameterization. J. Atmos. Sci., 66, 2812-‐2827. 549 550 Kang, S. M., I. M. Held, and S.-‐P. Xie, 2014: Contrasting the Tropical Responses to Zonally 551 Asymmetric Extratropical and Tropical Thermal Forcing, Clim. Dyn. 42, 2033-‐2043. 552 553 Koutavas, A., and J. Lynch-‐Stieglitz, 2004: Variability of the marine ITCZ over the eastern 554 Pacific during the past 30,000 years: Regional perspective and global context. The Hadley 555 Circulation: Present, Past and Future, H. F. Diaz and R. S. Bradley, Eds., Springer, 347–369 556 557 Liu, Z., and M. Alexander (2007), Atmospheric bridge, oceanic tunnel, and global climatic 558 teleconnections, Rev. Geophys., 45, RG2005, doi:10.1029/2005RG000172. 559 560 Locarnini, R. A., A. V. Mishonov, J. I. Antonov, T. P. Boyer, H. E. Garcia, O. K. Baranova, M. M. 561 Zweng, and D. R. Johnson, 2010: World Ocean Atlas 2009, Volume 1: Temperature. S. 562 Levitus, Ed. NOAA Atlas NESDIS 68, U.S. Government Printing Office, Washington, D.C., 184 563 pp. 564 565 Manabe, S., and R. J. Stouffer, 1995: Simulation of abrupt climate change induced by 566 freshwater input to the North Atlantic Ocean. Nature, 378, 165–167. 567 568 Maroon, E. A., D. M. W. Frierson, D. S. Battisti, 2015: The Tropical Precipitation Response to 569 Andes Topography and Ocean Heat Fluxes in an Aquaplanet Model. J. Climate 28:1, 381-‐570 398. 571 572 Maroon, E. A., D. M. W. Frierson, S. M. Kang, and J. Scheff, (in press): The precipitation 573 response to an idealized subtropical continent. J. Climate, http://dx.doi.org/10.1175/JCLI-‐574 D-‐15-‐0616.1 575 576 Marshall, J., Donohoe, A., Ferreira, D. and McGee, D, 2013: The role of the ocean circulation 577 in setting the mean position of the ITCZ. Clim. Dyn. http://dx.doi.org/10.1007/s00382-‐578 013-‐1767-‐z 579 580 Peterson, L. C., G. H. Haug, K. A. Hughen, and U. Rohl, 2000: Rapid changes in the hydrologic 581 cycle of the tropical Atlantic during the last glacial. Science, 290, 1947–1951. 582 583

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Xie, S.-‐P., 2004: The shape of continents, air-‐sea interaction and the rising branch of the 624 Hadley circulation. The Hadley Circulation: Present, Past and Future, H. F. Diaz and R. S. 625 Bradley, Eds., Springer, 121–152. 626 627 Xie, P., and P.A. Arkin, 1997: Global precipitation: A 17-‐year monthly analysis based on 628 gauge observations, satellite estimates, and numerical model outputs. Bull. Amer. Meteor. 629 Soc., 78, 2539 -‐ 2558. 630 631 Zhang, R., and T. L. Delworth, 2005: Simulated tropical response to a substantial weakening 632 of the Atlantic thermohaline circulation. J. Climate, 18, 1853–1860. 633 634

635

636

637

638

639 Figure 1. The 1981-‐2010 averaged precipitation based on monthly NOAA CPC Merged 640

Analysis of Precipitation (CMAP) in which rain gauge data are merged with precipitation 641

estimates from multiple satellite-‐based (infrared and microwave) products. 642

643

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644 Figure 2. Schematic outline of 60°-‐wide closed-‐ocean basins with (a) westward and (b) 645

eastward slanted tropical coastlines (equatorward of 14° latitude), under a 120°-‐wide 646

sector atmosphere. 647

648

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649 Figure 3. Results of the Exp I.zu simulation: (a) and (c) show average tropical precipitation 650

(shading) overlaid with anomalous surface wind vectors (with respect to an equatorially 651

symmetric zonal component and an antisymmetric meridional component) between years 652

1 and 20, and between years 381 and 400, respectively. Vectors outside of the deep tropics 653

with anomalous wind speed above 2 m/s are removed for easier visualization. Blue lines 654

mark the ocean basin boundaries. (b) Time-‐latitude diagram from year 1 to 400 showing 655

annual zonal means of hemispherically anomalous SST (shading), anomalous precipitation 656

(blue/red = positive/negative contours) and anomalous surface wind stress vectors. 657

Anomalous precipitation was filtered with a low-‐pass Butterworth filter with a 40-‐year 658

period cutoff. 659

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660

661 Figure 4. Results of Exp I.zu simulation: (a) Maximum and absolute minimum of the NH 662

(red curve) and the SH (blue curve) annual mean deep MOC overturning stream function 663

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(below 200m poleward of 20° latitude), respectively. (b) Ocean heat transport across the 664

equator (red curve) and the Hadley circulation asymmetry index (blue curve) defined as -‐665

(minimum streamfunction value of the SH Hadley cell + maximum streamfunction value of 666

the NH Hadley cell). (c) Tropical precipitation (red curve) and SST (blue curve) asymmetry 667

indices (average between 20°N and the equator minus average between the equator and 668

20°S). Quantities in panels (b) and (c) are annual means smoothed by a 15-‐year boxcar 669

filter. 670

671

672 Figure 5. The average of Exp I.zu simulation between year 801 and 1000 of (a) anomalous 673

atmospheric overturning streamfunction (with respect to the hemispherically 674

antisymmetric component), and (b) MOC streamfunction. Positive values of streamfunction 675

indicate a clockwise circulation. 676

677

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678 Figure 6. Annual mean results of Exp I.nh, Exp I.bl and Exp I.sh simulations in left, middle 679

and right columns, respectively. The top row shows maximum and absolute minimum of 680

the NH (red curve) and the SH (blue curve) deep MOC streamfunction, respectively. The 681

middle row shows ocean heat transport across the equator (red curve) and Hadley 682

circulation asymmetry index (blue curve). The bottom row shows tropical precipitation 683

(red curve) and SST (blue curve) asymmetry indices. Panels in the middle and bottom rows 684

show annual means smoothed by a 15-‐year boxcar filter. 685

686

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687 Figure 7. Results of Exp I.nh, Exp I.bl and Exp I.sh simulations in top, middle and bottom 688

rows, respectively. Time-‐latitude diagrams from year 1 to 400 showing annual zonal means 689

of anomalous SST (fill), anomalous precipitation (blue/red = positive/negative contours) 690

and anomalous surface wind stress vectors, as in Figure 3b. 691

692

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693 Figure 8. Figure 3. Results of Exp II.bl simulation: The quantities in this figure are the same 694

as in Figure 3. 695

696

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697 Figure 9. Results of Exp II.bl simulation after 15-‐year box smoothing of annual means: The 698

quantities in this figure are the same as in Figure 4. 699

700

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701 Figure 10. The average of Exp II.bl simulation between year 801 and 1000 of (a) anomalous 702

atmospheric overturning streamfunction (with respect to the antisymmetric component) 703

and (b) MOC streamfunction. Positive values of streamfunctions show clockwise 704

circulation. 705

706

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707 Figure 11. 5-‐year mean results from full integration periods of all simulations in this study 708

as functions of the deep MOC asymmetry index (sum of the subpolar deep MOC 709

streamfunction extrema in both hemispheres). (a) Upper left panel show ocean heat 710

transport across the equator. (b) Upper right panel shows the average NH extratropical 711

surface heat flux minus average SH extratropical surface heat flux (with upward positive 712

direction). (c) Lower left panel shows the Hadley asymmetry index. (d) Lower right panel 713

shows the tropical precipitation asymmetry index. 714

715

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CPC merged analysis of precipitation (CMAP) 〈Jan 1981 - Dec 2010〉

[mm/day]

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Equator

3900m

y

x

3900m

60o60o

Exp I.zu

Exp I.nh, I.bl and I.sh

Exp II.bl

(a) (b)

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[yr]SSTa [°C]

(a)

(b)

(c)

Exp I.zu 〈1-20 〉 yr

Exp I.zu 〈381-400 〉 yr P [mm/day]

P [mm/day]

Exp I.zu

m/s

m/s

(usfc, vsfc)a

(usfc, vsfc)a

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(1010

kgs-1

)

NH deepMOC SH |deepMOC|

(a)

(b)

(c)

[yr]

[yr]

[yr]

Exp I.zu

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(a)

(b)

AAO

str

eam

func

tion

(1010

kgs

-1)

MO

C st

ream

func

tion

(Sv)

Exp I.zu

(m)

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(a)

(b)

(c)

SSTa [°C]

Exp I.sh

Exp I.bl

Exp I.nh

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(b)

SSTa [°C][yr]

(c)

(a)

Exp II.bl 〈1-20 〉 yr

Exp II.bl 〈381-400 〉 yr

P [mm/day]

P [mm/day]

m/s(usfc, vsfc)a

m/s(usfc, vsfc)a

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NH deepMOC SH |deepMOC|

(a)

(b)

(c)

[yr]

[yr]

[yr]

Exp II.bl

(1010

kgs-1

)

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(a)

(b)

AAO

str

eam

func

tion

(1010

kgs

-1)

MO

C st

ream

func

tion

(Sv)

Exp II.bl

(m)

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Exp I.zuExp I.nhExp I.blExp I.shExp II.bl

Exp I.zuExp I.nhExp I.blExp I.shExp II.bl

deep MOC asymmetry index (Sv) deep MOC asymmetry index (Sv)

deep MOC asymmetry index (Sv) deep MOC asymmetry index (Sv)

(1010

kgs

-1)

(a) (b)

(c)

(d)

trop

. Pex

trat

rop.

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