A Nonlinear Response of Sahel Rainfall to Atlantic Warming

7080 JOURNAL OF CLIMATE VOLUME 26

NARESH NEUPANE AND KERRY H. COOK Department of Geological Sciences, Jackson School of Geosciences, The University of Texas at Austin, Austin, Texas

(Manuscript received 24 July 2012, in ﬁnal form 2 February 2013)

ABSTRACT

The response over West Africa to uniform warming of the Atlantic Ocean is analyzed using idealized simulations with a regional climate model. With warming of 1 and 1.5 K, rainfall rates increase by 30%–50% over most of West Africa. With Atlantic warming of 2 K and higher, coastal precipitation increases but Sahel rainfall decreases substantially. This nonlinear response in Sahel rainfall is the focus of this analysis. Atlantic warming is accompanied by decreases in low-level geopotential heights in the Gulf of Guinea and in the large- scale meridional geopotential height gradient. This leads to easterly wind anomalies in the central Sahel. With Atlantic warming below 2 K, these easterly anomalies support moisture transport from the Gulf of Guinea and precipitation increases. With Atlantic warming over 2 K, the easterly anomalies reverse the westerly flow over the Sahel. The resulting dry air advection into the Sahel reduces precipitation. Increased low-level moisture provides moist static energy to initiate convection with Atlantic warming at 1.5 K and below, while decreased moisture and stable thermal profiles suppress convection with additional warming. In all simulations, the southerly monsoon flow onto the Guinean coast is maintained and precipitation in that region increases. The relevance of these results to the global warming problem is limited by the focus on Atlantic warming alone. However, confident prediction of climate change requires an understanding of the physical processes of change, and this paper contributes to that goal.

1. Introduction the hydrodynamics of the response to uniform warming of the Atlantic Ocean within the model domain (198S– An improved understanding of the physical processes 348N, 588W–488E; see section 3). Forcing from the At- associated with climate change in West Africa and the lantic Ocean is chosen because of the known sensitivity Sahel is needed for producing reliable predictions to aid of the monsoon system to Atlantic sea surface temper- resource management and planning as the planet warms. ature anomalies (SSTAs). Uniform warming is chosen Nonlinear responses in the climate system can lead to for simplicity and clarity, and because the modeling re- surprises and even abrupt climate change, presenting sults of Patricola and Cook (2011) indicate that much of greater challenges for adaptation. Paleoclimate proxy the global warming signal forced from the Atlantic in data show that the West African climate is prone to West Africa and the Sahel is caused by uniform warming, rapid changes between wet and dry conditions. Even the with a dependence on structure in the SSTAs present seasonal evolution of the monsoon system is character- only along the west coast. ized by abrupt transitions because the northward mi- Climate change and variability in West Africa and the gration of rainfall from the Guinean Coast into the Sahel Sahel is complicated, with demonstrated dependence on does not occur smoothly. SSTAs in the Indian, Atlantic, and Pacific Oceans, land The purpose of this paper is to contribute to an im- surface feedbacks, aerosol forcing, midlatitude–tropical provement in our physical understanding of how the interactions, and other factors—and we do not mean to West African monsoon will respond under global imply otherwise by focusing on Atlantic sea surface warming. We isolate one aspect of the problem, namely temperature (SST) forcing alone. However, confident prediction of twenty-first-century climate change in the region requires that we build a solid understanding of Corresponding author address: Kerry H. Cook, Department of Geological Sciences, Jackson School of Geosciences, The Uni- the physical processes of that change, and here we versity of Texas at Austin, Austin, TX 78712. contribute to that goal in a process study that isolates the E-mail: [email protected] response to uniform Atlantic SSTAs.

DOI: 10.1175/JCLI-D-12-00475.1

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Background is provided in section 2, including an Present-day climate records also indicate that the overview of the region’s sensitivity to Atlantic SSTs and West African climate is sensitive to Atlantic SSTs (e.g., our current understanding of the potential for change Folland et al. 1986; Fontaine and Bigot 1993; Camberlin under greenhouse gas forcing. The regional climate et al. 2001; Liu and Chiang 2012). One particular SSTA model and experimental design are described in section forcing pattern known to influence West Africa rainfall 3, and the model is validated against observations, re- is centered in the Gulf of Guinea, where the onshore analysis, and satellite measurements in section 4. Results monsoon flow originates. With positive SSTAs of 1–2 K are reported in section 5, and conclusions and implica- in the Gulf of Guinea, rainfall is enhanced along the tions are discussed in section 6. Guinean coast and suppressed in the western and central Sahel (Ward 1998; Nicholson 2000; Mo et al. 2001). An opposite pattern occurs with cool SSTAs in the Gulf of 2. Background Guinea. Vizy and Cook (2001, 2002) explain the physics Paleoclimate proxy data show that the West African of this mechanism. Evaporation is enhanced over the climate can change substantially on millennial time warm SSTAs, but a precipitation anomaly does not de- scales. For example, 14.8–5.5 thousand years (ka) before velop over the ocean because of regional subsidence as- present (BP) is known as the African Humid Period sociated with the Walker circulation. As a consequence, (AHP), when the western Sahel and Sahara received the moisture content of the northward monsoon inflow rainfall sufficient to support large lakes, small trees, and is increased along with rainfall along the coast. Anom- large fauna such as hippopotamus (e.g., Gasse and Van alous subsidence and drying in the Sahel is a secondary Campo 1994; Petit-Maire and Guo 1996). The onset of response to enhanced rainfall along the coast generated the AHP was associated with some major atmospheric by conservation of potential vorticity requirements. changes. For example, the atmospheric CO2 concentra- Although the focus of this paper is on the physics of tion increased by about 50 ppmv, the Atlantic meridional the response to uniform Atlantic Ocean warming, many overturning circulation (AMOC) weakened and the factors are known to influence precipitation over West northern ice sheets retreated during the onset, and sum- Africa. These include changes in SST patterns in the mer insolation in the Northern Hemisphere was about Atlantic (e.g., Mohino et al. 2011a), SSTAs in the Pacific 5% greater than today at 8 ka BP (White et al. 1994; and Indian Oceans (e.g., Paeth and Friederichs 2004; Monnin et al. 2001; McManus et al. 2004). Hagos and Cook 2008; Mohino et al. 2011b), changes in Of potential relevance to the development of the land surface/vegetation (e.g., Li et al. 2007; Patricola and global warming response over West Africa and the Sahel Cook 2008), aerosol forcing (e.g., Gu et al. 2012; Joseph is the ability of the region’s climate to change abruptly. and Zeng 2011; Zhao et al. 2011), and relative positions For example, the transition into and out of the AHP was of the African easterly jet and the tropical easterly jet abrupt, at least in some regions, occurring within de- (Nicholson 2009). Because the model domain does not cades to centuries despite gradual and smooth changes include the Indian and Pacific Oceans, the impacts of in the precessional forcing of insolation (e.g., deMenocal large-scale circulations such as the Madden–Julian os- et al. 2000; Russell et al. 2003; Peck et al. 2004). Also, in cillation, the El Nino–Southern~ Oscillation, and the In- a more recent model simulation, Timm et al. (2010) dian summer monsoon on the West African monsoon demonstrate that the abrupt onset of the AHP was are not included. caused by the combined effects of increases in local in- The need to improve our basic understanding of the solation and drastic melting of the Northern Hemi- physics of the African monsoon’s response to climate sphere ice sheets. forcing is evident in the uncertainty about how the re- Paleoclimate studies also provide an indication of the gion’s rainfall will change in the future (Biasutti and sensitivity of the West African climate to Atlantic SSTs. Giannini 2006; Biasutti et al. 2008; Cook 2008). Cook and The African Humid Period was interrupted by an arid Vizy (2006) examined output from 18 coupled GCMs interval corresponding to the Younger Dryas event from phase 3 of the Coupled Model Intercomparison (;12.9–11.5 ka BP), when a weakening of the AMOC Project 3 (CMIP3; Meehl et al. 2007) over West Africa. occurred in response to a sudden influx of freshwater Half of the CMIP3 simulations failed to produce the from deglaciation in North America (Alley 2000; Lowell West African monsoon at all in the sense that they did et al. 2005). Heinrich events are also reflected in African not bring the summer precipitation maximum onto the paleoclimate records; there is evidence that North At- continent. Three GCMs were selected based on their lantic cooling at 17 ka BP associated with Heinrich event ability to represent major features of the twentieth- 1 caused a widespread drought in the African monsoon century climate, including its variability. But in the fu- region (Stager et al. 2011). ture simulations, these three models diverged widely,

Unauthenticated | Downloaded 10/01/21 01:21 AM UTC 7082 JOURNAL OF CLIMATE VOLUME 26 demonstrating that a reasonable simulation of the present 3. Model description and experimental design climate is not a sufficient condition for accurate pro- jection. Inclusive analyses of the CMIP3 coupled GCMs The National Center for Atmospheric Research (Cook 2008) show that most of the CMIP3 models predict (NCAR) Advanced Research Weather Research and only small increases or decreases of annual precipitation Forecasting model, version 3.1.1 (WRF 3.1.1), is used for over West Africa by the end of the century. this study (Skamarock et al. 2008). The model is fully There is no overall improvement in simulating the compressible and nonhydrostatic, with 28 vertical levels West African monsoon system in the coupled GCMs in this case. The top of the model is set to 50 hPa. The from the CMIP5 (Vizy et al. 2013). Some GCMs, such model is used at a spatial resolution of 90 km with a time as the Community Climate System Model, version 4.0 step of 3 min. (CCSM4.0; Cook et al. 2012), produce an improved The outcome of past regional climate modeling efforts representation over West Africa compared with the over Africa (e.g., Vizy and Cook 2002, 2003; Patricola CMIP3, but in other models the simulation is worse. and Cook 2010, 2011) and extensive testing were used to Maynard et al. (2002) used a coupled GCM to simulate select parameterizations, a model domain, and lateral African climate at the end of the twenty-first century and surface boundary conditions that produce a realistic under the Intergovernmental Panel on Climate Change simulation of the present-day West African climate. (IPCC) B2 scenario and found that the summer mon- Parameterizations chosen include the Mellor–Yamada– soon precipitation over the Sahel increases. Janjic planetary boundary layer scheme (Mellor and Patricola and Cook (2010, 2011) use a regional model Yamada 1982; Janjic 1990, 1996, 2002), the Monin– at 90-km resolution centered over West Africa to pro- Obukhov–Janjic surface layer scheme (Monin and duce two 9-member ensembles for the end of the twen- Obukhov 1954; Janjic 1994, 1996, 2002), the new Kain– tieth and twenty-first centuries, with boundary conditions Fritsch cumulus scheme (Kain and Fritsch 1990, 1993), for future simulations applied as anomalies based on the Purdue–Lin microphysics scheme (Lin et al. 1983), output from nine coupled GCM. Despite the differences the Rapid Radiative Transfer Model longwave radiation in SSTAs and lateral boundary conditions in the nine parameterization (Mlawer et al. 1997), the fifth-generation GCMs, there is good agreement in the predictions over Pennsylvania State University–NCAR Mesoscale Model most of West Africa and the Sahel among the nine en- (MM5) shortwave parameterization (Dudhia 1989), and semble members. Annual precipitation changes are rel- the unified Noah land surface model (Chen and Dudhia atively small, similar to the CMIP3 multimodel ensemble 2001). mean (Cook 2008), but regional and monthly analysis Figure 1 shows the model domain selected for this reveals large changes in temperature, heat stress, pre- study. The large domain minimizes the effects of the cipitation, and extreme rainfall events that would pro- lateral boundaries on the region of interest. It includes duce pronounced impacts. In particular, rainfall in the most of northern Africa as well as tropical and eastern Sahel increases in July and August and decreases along North Atlantic. the Guinean coast in June and July. Initial and lateral/surface boundary conditions for all The reason for the agreement among the ensemble simulations are initialized at 0000 UTC 15 March using members in the regional model simulations of Patricola climatological March conditions interpolated onto the and Cook (2010) is that the first-order response over regional model grid. Lateral boundary conditions for West Africa and the Sahel at the end of the twenty- horizontal winds, temperature, relative humidity, and first century is to overall ocean warming, which is geopotential height are derived from the 1958–2001 demonstrated by an additional simulation with uniform National Centers for Environmental Prediction (NCEP)– SSTAs of 2 K. The changes in precipitation in the NCAR reanalysis (Kalnay et al. 1996) monthly climatology. uniform warming case are similar to those from the Initial surface boundary conditions for skin temperature, ensemble mean across West Africa and the Sahel, with soil temperature, soil moisture, and sea level pressure are the only exception in the far western Sahel where the derived from the 40-yr European Centre for Medium- response is sensitive to the structure of the imposed Range Weather Forecasts (ECMWF) Re-Analysis (ERA- SSTAs. 40; 1958–2001) monthly climatology (Uppala et al. 2005). Building on these results, we explore the physics of the The ERA-40 is used for the surface initialization instead of response to uniform Atlantic warming alone in a series the NCEP reanalysis climatology because the NCEP re- of idealized simulations. We are particularly interested in analysis has a cold and wet bias over the Sahara. Patricola identifying and understanding nonlinearity in the mon- and Cook (2010) show that WRF produces a more re- soon system’s response to gradual warming, because alistic simulation of the monsoon when ERA-40 surface these can be a source of abrupt change. conditions are used to initialize the simulation.

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FIG. 1. Model domain with topography (shaded; m) and average 26 Jun–26 Jul SSTs (contoured; K) from CTL.

Lateral boundary conditions are updated every 6 h structure of current Atlantic SST gradients, but create using the NCEP reanalysis climatology interpolated an overall warmer ocean basic state. Uniform positive onto the model boundaries. Monthly-mean values from SSTAs are applied in six additional simulations referred the reanalysis are taken to represent the middle of the to as 1K, 1.5K, 2K, 2.5K, 3K, and 4K. month, and linear interpolation in time is used to generate 6-hourly boundary conditions. Boundary condi- 4. Model evaluation tions produced by this method include seasonality, but diurnal and synoptic time scales are suppressed. This Figure 2 shows boreal summer monthly rainfall from ‘‘climate mode’’ approach works for large domains in the 0.258-resolution National Aeronautics and Space the tropics (Cook and Vizy 2008; Hagos and Cook 2007; Administration (NASA) Tropical Rainfall Measuring Patricola and Cook 2010) but fails for applications in Mission (TRMM 3B42V6) satellite-derived 1998–2009 middle latitudes where transients propagating into the climatology (Kummerow et al. 1998), the 0.58-resolution domain are important for determining the time-mean Climatic Research Unit (CRU; CRUTS3.0) 1979–2006 state (Patricola and Cook 2013a,b; Pu et al. 2012). Seven climatology (Mitchell and Jones 2005) over land and the 231-day simulations (15 March–31 October) are pro- 2.58-resolution Global Precipitation Climatology Pro- duced, and 3-hourly output is averaged to generate ject (GPCP; Adler et al. 2003) 1979–2006 over the daily and monthly means. This period covers the ocean, and the control simulation. Each climatology is entire evolution of the boreal summer monsoon. The presented at its original resolution to preserve as much ﬁrst 2.5 months of each simulation, from 15 March to 31 spatial information as possible. The regional model May, are devoted to model spinup and not used in the simulation is able to realistically capture the seasonal analysis. Such a long period for spinup is necessary to migration of rainfall over West Africa, including the avoid errors associated with the persisting initial soil movement of the precipitation maximum into the con- moisture properties (Simmonds and Hope 1998). tinental interior that is often missed in coupled GCMs. A control simulation (CTL) uses climatological SSTs Overall, the simulated rainfall intensity is larger than that are prescribed and updated every 6 h. These observed. This wet bias varies based on location and can 2 6-hourly SSTs are derived from the 1958–2001 ERA-40 range from 2 to more than 18 mm day 1. For example, monthly climatology using the same linear interpolation over the far western Sahel (108–128N, 68–128W) the dif- method discussed above for the lateral boundary condi- ference between the simulated and observed precipitation 2 2 tions. Six additional simulations are produced with uni- is approximately 8 mm day 1 in July and 4 mm day 1 in form Atlantic SSTAs added to the climatological SSTs, August. Also, the summer precipitation minimum over the but otherwise identical to CTL. These SSTs preserve the Guinean Coast in July and August is more pronounced

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21 FIG. 2. Monthly rainfall (mm day ) for the (a)–(d) TRMM climatology, (e)–(h) GPCP climatology over the ocean and the CRU 2 climatology over land, and (i)–(l) CTL simulation. Contours are drawn for 2 and 10 mm day 1 rainfall rates, and the shading interval 2 is 1 mm day 1. in the control simulation than the TRMM and CRU the sudden shift of the rainfall maximum from the climatologies. Guinean coast into the Sahel in early July known as the The accuracy of the simulated rainfall distribution is monsoon jump (not shown). similar to that of the most accurate GCMs in this region, The 925-hPa June geopotential heights and winds and better than most GCM simulations. This includes from the 1979–2010 ERA-Interim climatology at 1.58

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21 FIG. 3. The 925-hPa June geopotential heights (gpm) and winds (reference wind vectors provided; m s ) from the (a) ERA-Interim climatology and (b) CTL. Also, the 925-hPa geopotential heights (gpm) and winds (reference wind 2 vectors provided; m s 1) averaged over July–September from the (c) ERA-Interim climatology and (d) CTL. Contour and shading interval is every 10 m. resolution (Dee et al. 2011) and the control simulation the simulation, and formed in the reanalysis; it is stronger are shown in Figs. 3a and 3b, respectively. The summer- in the model. [Higher-resolution reanalyses available for time thermal low has developed over the central Sahel individual years capture the West African westerly jet and Sahara in the reanalysis (Fig. 3a). The model (Fig. 3b) more strongly than the ERA-Interim product.] Other produces geopotential heights of realistic magnitude, features of the low-level ﬂow are captured in the model but the low is centered farther to the north and west with good accuracy. than in the reanalysis. The position and strength of the Figures 4a and 4b show geopotential heights and North Atlantic subtropical high are realistically simu- winds at 600 hPa from the ERA-Interim climatology and lated, but the onshore ﬂow along the Guinean and west the control simulation for June. The midtropospheric coasts is stronger in the model simulation than in the Saharan anticyclone is centered near 238N, 08 in the reanalysis. The West African westerly jet (Grodsky et al. reanalysis (Fig. 4a), and only slightly to the west of that 2003; Pu and Cook 2010, 2011) near 88N is in place on the position in the simulation (Fig. 4b). The African easterly west coast in the model, earlier than its formation in the jet has formed and is realistically simulated. reanalysis. In the July–September mean, the Saharan high ex- Figures 3c and 3d show reanalysis and model 925-hPa pands northward over the Sahara in both the reanalysis geopotential heights and winds from July through Sep- and the model (Figs. 4c,d) compared with the June av- tember. The magnitude of the thermal low is captured erage, positioning the African easterly jet a little farther accurately in the simulation, but it is located about 58 to to the north, consistent with the repositioning of geo- the south of its center in the reanalysis. The West Afri- potential height, surface temperature, and moisture can westerly jet has moved a few degrees to the north in gradients (Cook 1999).

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21 FIG. 4. The 600-hPa June geopotential heights (gpm) and winds (reference wind vectors provided; m s ) for the (a) ERA-Interim climatology and (b) CTL. (c),(d) As in (a),(b), but averaged over July–September. Contour and shading intervals are 10 gpm.

2 In summary, the regional model produces a state-of- greater by up to 2 mm day 1 than in the 1K-CTL case the-art representation of the major circulation and (Fig. 5a), but the spatial structure and magnitudes of the precipitation features of the West African climate and precipitation anomalies are similar. their seasonal evolution. Precipitation increases on the western and Guinean coasts are further enhanced in 2K-CTL (Fig. 5c). The 5. Results two small regions of drying in Figs. 5a and 5b are also larger, and the drying in the central Sahel extends to the a. The dependence of West African precipitation west. The positive anomalies between 138 and 158N near anomalies on Atlantic Ocean warming the Greenwich meridian in Figs. 5a and 5b are now re- Figure 5a shows July–September precipitation dif- placed with negative anomalies. ferences between the control simulation and the SSTA Negative precipitation anomalies over the central Sahel 1K simulation in the Atlantic, denoted 1K-CTL. Rain- continuetostrengthenandextendwestwardwithaddi- 2 fall rates are enhanced by 2–6 mm day 1 along the tional Atlantic warming. Figures 5d–f display precipitation Guineanandwestcoaststothesouthof158N, and over differences for 2.5K-CTL, 3K-CTL, and 4K-CTL, respec- the western Sahel. There are negative precipitation tively. The 3K-CTL and 4K-CTL simulations produce a 2 anomalies up to 2 mm day 1 to the north of the response similar in structure to 2.5K-CTL, but with Guinean coast (near 128N, 58W) and in the central larger anomalies and more extensive Sahel drying. Sahel (148N, 108E). Differences are small to the north Rainfall rates over the western and Guinean coasts 2 of 208N. increase by more than 18 mm day 1 in 2.5K-CTL, for ex- Figure 5b shows summer rainfall anomalies for 1.5K-CTL. ample, and rainfall rates over the Sahel decrease by up to 2 Positive rainfall differences in the western Sahel are 10 mm day 1 in 3K-CTL and 4K-CTL.

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FIG. 5. Precipitation differences from CTL averaged over July–September for simulations with (a) 1-, (b) 1.5-, (c) 2-, (d) 2.5-, (e) 3-, and (f) 4-K warming of the Atlantic Ocean. Contours are drawn for precipitation differences 2 of 28, 22, 12, and 18 mm day 1.

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21 FIG. 6. Percentage differences in the average July–September precipitation (mm day ) for the (a) 1K-CTL, (b) 1.5K-CTL, (c) 2K-CTL, (d) 2.5K-CTL, (e) 3K-CTL, and (f) 4K-CTL simulations, respectively. The averaging box in each case has a width of 18 of latitude and extends from 128Wto68E. The Sahelian region from 138 to 168N is shaded.

Figure 6a shows percentage differences in summer from 38% to 140% in 2.5K-CTL and from 45% to 225% (July–September) precipitation averaged from 128Wto in 3K-CTL. The Sahel drying that was relatively modest 68E over West Africa for 1K-CTL. Each bar represents in 2K-CTL (Fig. 6c) is larger and more extensive with a18 latitude band, and Sahel latitudes (138–168N) are 2.5-K Atlantic warming, with precipitation increases to shaded. Rainfall rates increase up to 35% over West the north of 178N replaced by drying. With greater At- Africa between 88 and 108N. They do not change, or lantic warming, the same pattern of drying in the Sahel decrease slightly, between 108 and 138N. To the north persists, becoming more severe and spreading southward. of 138N, precipitation rates increase everywhere. The Figure 6 demonstrates that the precipitation response largest percentage change occurs between 168 and 198N to Atlantic SSTAs over the Sahel is nonlinear. For At- where rainfall rates increase by 90%–167%. lantic SST increases up to 1.5 K, Sahel rainfall rates in- Figure 6b displays the results for 1.5K-CTL. Between crease. When the warming reaches and exceeds 2 K, 88 and 168N, the response is similar to the 1K-CTL dif- Sahel rainfall decreases. The existence of a threshold ferences (Fig. 6a). Rainfall increases are not as large as Atlantic SSTA between 1.5 and 2 K in the model com- in 1K-CTL to the north of 178N. plicates climate prediction and suggests a possibility for With 2-K warming in the Atlantic (Fig. 6c), precip- abrupt climate change over the Sahel in association with itation increases between 88 and 128N are a little larger SST forcing from the Atlantic basin. than with 1.5-K warming but similar in structure, but Reference to Fig. 5 indicates that there is some re- there is a sign change in the precipitation anomaly over gional dependence for the threshold warming value. In the Sahel. The 1.5K-CTL has 4%–66% precipitation general, rainfall rates increase (decrease) to the west increases between 128 and 178N, but Sahel rainfall rates (east) of 68Wfrom128 to 198N in all of the warming in 2K-CTL decrease by 3%–18%. Percentage precip- simulations. In the Sahel, the rainfall rates increase from itation enhancements north of 178N are greater than in 68 to 128W in 1K and 1.5K SSTA simulations compared 1K-CTL and 1.5K-CTL. with CTL, but decrease with warming of 2 K and higher. Figures 6d–f display precipitation differences for In 1K-CTL and 1.5K-CTL, the overall increases are offset 2.5K-CTL, 3K-CTL, and 4K-CTL, respectively. Pre- byasimulateddecreasetotheeastof68W. However, cipitation differences between 88 and 118N are positive, from 178 to 198N, the decrease to the east of 68W does not

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21 FIG. 7. Average 26 Jun–26 Jul precipitation (mm day ) anomalies for the (a) 1K-CTL, (b) 1.5K-CTL, (c) 2K-CTL, and 2 (d) 2.5K-CTL simulations, respectively. Contouring indicates precipitation anomalies of 22 mm day 1. offset the increase over the western portion of the aver- Average daily precipitation over the Sahel in this period 2 aging region until SSTAs exceed 2.5 K. This regional in- in the CTL simulation is about 5 mm day 1. It increases 2 crease in rainfall rates to the west of 68W is largely to about 8 mm day 1 in 1.5K and decreases to about 2 conﬁned to the month of August. 3 mm day 1 in 2K. At the end of July, rainfall rates in Two other recent model simulations have been con- 2K are greater than in CTL. During the rest of the ducted with uniform SST increases, both with 2-K warm- summer (e.g., August and September) rainfall rates in ing. Held et al. (2005) simulated annual-mean drying in CTL, 1.5K, and 2K are similar. the Sahel, especially from 108 to 158N, in an atmospheric Figure 7 displays rainfall differences for four of the GCM. Patricola and Cook (2010, 2011) found that July– warming simulations averaged from 26 June to 26 July. September (JAS) Sahel rainfall decreased, especially As was the case for the full summer average (Figs. 5a,b), 2 during July, up to 6 mm day 1 in response to uniform the basic spatial structure of rainfall anomalies over the Atlantic warming in a regional climate model. These Sahel in 1K-CTL (Fig. 7a) and 1.5K-CTL (Fig. 7b) is results are consistent with the 2K-CTL presented here. similar. Precipitation differences in the southern Sahel To focus on the mechanisms of this nonlinear de- (near 138N) and the eastern Guinean coast region are pendence on Atlantic SSTAs, we examine the time pe- greater in 1.5K-CTL than in 1K-CTL, as is the rainfall riod from 26 June to 26 July. This time period is largely reduction in the central Sahel (along 148N east of 68W). responsible for the fact that Sahel JAS rainfall is less in Precipitation decreases in the central Sahel become 2K-CTL than in 1K-CTL and 1.5K-CTL (not shown). more pronounced with Atlantic warming of 2 and 2.5 K

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FIG. 8. (a) The 925-hPa geopotential heights (10-gpm contour interval) and winds (reference wind vector provided; 2 ms 1) from CTL. (b) Differences in 925-hPa normalized geopotential heights (3-gpm contour interval) and winds 2 (reference wind vector provided; m s 1) for 1.5K-CTL. All values are averaged over 26 Jun–26 Jul.

(Figs. 7c,d), and the dry anomalies extend to the west. eastward curvature of the flow over the Sahel (Fig. 8a) is As Atlantic SSTAs are increased above 2.5 K, the drying reduced. Warming in the Atlantic Ocean—and, in par- becomes stronger but the spatial structure is similar to ticular, in the Gulf of Guinea—decreases the large-scale that in 2K-CTL and 2.5K-CTL (not shown). Pre- meridional geopotential height gradients across West cipitation over the Gulf of Guinea increases in all the Africa. As a result, the southerly monsoon flow onto the Atlantic warming simulations. Guinean coast is not diverted to the east as strongly, and it penetrates farther north. This leads to the precipitation b. The hydrodynamic response to small Atlantic increases in the western Sahel in 1K-CTL and 1.5K-CTL, SSTAs: 1K-CTL and 1.5K-CTL and also to the small region of decreased precipitation in The hydrodynamics of the West African and Sahel the central Sahel (Figs. 5a,b). precipitation increases with Atlantic warming less than Another consequence of the decreased magnitude of 2 K is examined first, and then compared with the re- the large-scale meridional geopotential height gradients sponse to SSTAs of 2 K and greater in the following across the Sahel is a weakening of the African easterly section to understand the sign change in the response. jet. Figure 9 displays the cross section of the differ- Figure 8a shows winds and geopotential heights at ence in the zonal wind averaged from 128Wto68E for 925 hPa averaged from 26 June to 26 July from the CTL 1.5K-CTL. Easterly wind anomalies occur below simulation. The southerly monsoon flow from the Gulf 700 hPa throughout the domain, consistent with Fig. 8b. of Guinea flows into the Guinean coast region and turns The positive (westerly) anomalies centered near 138N southwesterly over the continent under the influence of and 550 hPa indicate a weakening of the African easterly 2 Coriolis accelerations. The West African westerly jet is jet by more than 2 m s 1 in 1.5K-CTL. The African in place on the west coast centered near 108N (Grodsky easterly jet is a result of strong coupling between the et al. 2003; Pu and Cook 2010). The pronounced me- atmosphere and land surface over the Sahel, and it ridional gradients that characterize the region (Fig. 3) weakens as a result of geostrophic dynamics (via the are strongest between 98 and 178N. thermal wind relation) when meridional temperature Difference fields for 1.5K-CTL are shown in Fig. 8b. gradients below the level of the jet also weaken (Cook Here, the differences in 925-hPa geopotential heights 1999). In these simulations there is no shift in the posi- are normalized by the domain-averaged difference of tion of the African easterly jet, which is known to be 13.4 gpm to highlight changes in gradients. The anom- associated with Sahel rainfall variations on interannual alous geopotential height gradient is positive, indicating time scales (Nicholson and Grist 2001). a weakening of the negative climatological gradient The increases in Guinean coast precipitation in re- (Fig. 8a). This weakening is greatest in the Sahel, cen- sponse to Atlantic Ocean warming (Figs. 5a,b) are not tered around 158N, and it is accompanied by a westward fully explained by examining differences in the low-level 2 wind anomaly of up to 2 m s 1, indicating that the flow in Figs. 8a and 8b. The southerly monsoon flow onto

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Figure 10a displays 925-hPa mixing ratios and moisture transport from the CTL simulation averaged from 26 June to 26 July. Similar to Simmonds et al. (1999), moisture flux is expressed as the sum of the mean and the transient eddy components. The transient eddy component, however, is negligible in this case. While some moisture is transported onto the continent in the southerly monsoon flow, reduced evaporation associated with the relatively cool temperatures of the At- lantic cold tongue lower the moisture content of the southerly flow. A primary source of moisture for the Sahel at this time of year is the West African westerly jet (Pu and Cook 2011). As seen in the difference fields (Fig. 10b), mixing ratio values over the Gulf of Guinea double when 1.5-K warming is imposed in the Atlantic basin. Warming of the sea surface combined with the low-level winds of the monsoon circulation (Fig. 8a) drive enhanced evaporation. The increased moisture content of the inflowing monsoon circulation fuels the Guinean coast precipitation increases. Note also the anomalous westward moisture transport from the central to the western Sahel that is associated with the wind

FIG. 9. Latitude–height cross section of the difference in zonal wind anomalies shown in Fig. 8b. speed (1.5K-CTL) for the 26 Jun–26 Jul period averaged from 128W It may seem counterintuitive that rainfall rates over 2 to 68E. The contour interval is 1 m s 1 and negative values are shaded. the Guinean coast increase when Atlantic SSTs are warmed if one thinks about this region’s rainfall as being the Guinean coast is increased in magnitude somewhat supported by a monsoon circulation that is ultimately and turned to the northwest by the warming of the related to the juxtaposition of a cooler ocean surface and Atlantic Ocean, with anomalous convergence along the a warmer land surface, because warming the ocean coast; the West African westerly jet is weaker in 1.5K than would decrease the land–sea temperature contrast in the in CTL. To fully understand the response on the coast, at- summer (all other factors remaining the same). This mospheric moisture ﬁelds must be considered. mechanism operates on large space scales and leads to

23 21 21 FIG. 10. (a) The 925-hPa horizontal moisture transport (reference vector provided; 10 gkg ms ) and mixing ratios from CTL. (b) Differences in the 925-hPa horizontal moisture transport (reference vector provided; 2 2 2 10 3 gkg 1 ms 1) and mixing ratios for 1.5K-CTL. All values are averaged over 26 Jun–26 Jul. Contour interval is 2 2 2gkg 1 in (a) and 1 g kg 1 in (b).

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rainfall rates reverse the trend and decrease when At- lantic warming reaches 2 K and higher. Figure 12 explains why the response over the Sahel is nonlinear. Increased warming in the Atlantic leads to further reductions in the large-scale meridional geopotential height gradient over West Africa as geopotential heights in the Gulf of Guinea decrease. In 2K, shown in Fig. 12a, the monsoon flow and the West Af- rican westerly jet continue to impinge on the Guinean coast supporting rainfall in that region. Farther north, the meridional geopotential height gradient is weaker, as seen in the difference fields shown in Fig. 12b, and the associated easterly wind anomalies over the Sahel become large enough to reverse the direction of the flow. In 2K-CTL, easterly flow develops in the Sahel east of 58E. This easterly flow is associated with dry air advection and precipitation reductions. FIG. 11. Average 26 Jun–26 Jul 925-hPa vertical p-velocity 2 2 With continued warming in the Atlantic—particularly (10 2 Pa s 1) from CTL. Contours are drawn for vertical p-velocity 2 2 in the Gulf of Guinea—the easterly full-field flow expands values of 3 and 6 3 10 2 Pa s 1; positive values indicating subsidence are shaded. westward and strengthens as the meridional geopotential height gradients flatten. In 3K, shown in Figs. 12c, easterly flow stretches across the western Sahel as a result of very increases in rainfall in the western Sahel in response to large easterly anomalies (Fig. 12d). these mild SSTAs, as discussed above. But along the To understand the mechanisms of the nonlinear re- Guinean coast, the precipitation increase is explained sponse, we distinguish between the nonconvective and through the moisture budget analysis. When Gulf of convective components of precipitation. Differences Guinea SSTs increase, evaporation rates and low-level in the nonconvective precipitation among all of the moisture also increase over the SSTAs, but local pre- warming simulations and the control are small (not cipitation rates are not enhanced because of the pres- shown). The differences in total precipitation, displayed ence of subsidence in this region. As seen in Fig. 11 from in Figs. 5a–d, are similar to the differences in the con- the CTL simulation, and similar to the reanalyses (not vective precipitation. To connect differences in convec- shown), subsiding air over the Gulf of Guinea sup- tive precipitation with the dynamical fields, we examine presses convection, so the result is an increase in the moist static energy (MSE) profiles. The MSE is a linear moisture flux onto the coast and precipitation increases combination of the thermal, latent, and geopotential en- to the north of the SSTAs. ergy defined as A similar mechanism is prominent in the hydrody- 5 1 1 namics of the region’s interannual variability, in which MSE cpT Lq gz, (1) warm years in the Gulf of Guinea are associated with enhanced rainfall on the Guinean coast (Vizy and Cook where cp is the specific heat of air at constant pressure, T 2002). It is also consistent with the requirements of the is air temperature, L is the latent heat of vaporization, q is region’s vorticity balance. Cook (1997) showed that the specific humidity, g is the acceleration caused by gravity, maintenance of rainfall along the Guinean coast requires and z is the geopotential height. A stable atmosphere that low planetary vorticity from the south be transported will have increasing MSE with elevation. A high value of into the region to balance the positive relative vorticity MSE at low levels destabilizes the atmosphere. tendency associated with midtropospheric condensa- Figures 13a and 13b display differences in the moist tional heating and low-level vorticity stretching. Here, we static energy profiles from the 1.5K-CTL and 2K-CTL find that the meridional advection of low relative vorticity simulations, respectively. The profiles are averaged air also contributes to this balance. from 26 June to 26 July over the region 138–168N, 108W–08 to emphasize the nonlinear signal in convective c. The hydrodynamic response to stronger Atlantic precipitation. Negative slopes of the anomalous MSE SSTAs: 2K-CTL, 2.5K-CTL, 3K-CTL, and 4K-CTL profiles (solid line), especially below 850 hPa, indicate As seen in Fig. 6, Guinean coast rainfall rates increase that the low-level atmosphere is more unstable than the with increased warming in the Atlantic Ocean, but Sahel control in both simulations. The increased instability is

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FIG. 12. (a) The 925-hPa geopotential heights (10-gpm contour interval) and winds (reference wind vector pro- 2 vided; m s 1) for 2K and (b) differences in the 925-hPa normalized geopotential heights and winds for 2K-CTL. Normalization factor is 13.4 gpm. (c),(d) As in (a),(b), but for 3K and 3K-CTL, respectively. Normalization factor is 33.5 gpm. All values are averaged over 26 Jun–26 Jul. associated with low-level moisture anomalies Lq (dotted response of Sahel rainfall to warming Atlantic SSTAs, line). Atmospheric moisture increases from the surface to a second set of seven simulations for each case was run 550 hPa in 1.5K-CTL, but it decreases at all levels in 2K- with a different initialization date in March. The same CTL. Consistent with these differences, convective precip- nonlinear response of the Sahelian rainfall to the itation increases in 1.5K-CTL and decreases in 2K-CTL. warming of the Atlantic was reproduced in the second

The thermal profiles cpT (black dashed line) show that set of simulations. there is warming of the entire atmospheric column in 1.5K-CTL and 2K-CTL (Figs. 13a,b). The profile is al- 6. Summary and conclusions most neutral in 1.5K-CTL. On the other hand, there is greater warming at 850 hPa than at the surface in 2K- To contribute to an improvement in our physical un- CTL, and this stabilizes the vertical column. Differences derstanding of how the West African monsoon will re- in the geopotential energy gz (gray-dashed line) are spond under global warming, we isolate one aspect of negligible in both. Factors associated with increases in the problem: the hydrodynamics of the response to a precipitation over the Gulf of Guinea and the Guinean uniform warming of the Atlantic Ocean. Forcing from the coasts in 2K-CTL, 3K-CTL, and 4K-CTL are similar to Atlantic Ocean is chosen because of the known sensitivity those in 1.5K-CTL, and there is no nonlinear behavior in of the monsoon system to Atlantic SSTAs. Idealized this region. To confirm the robustness of the nonlinear simulations with a regional climate model and uniform

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4 21 FIG. 13. Average 26 Jun–26 Jul MSE (310 Jkg ) proﬁles over 138–168N, 108W–08 in (a) 1.5K-CTL and (b) 2K-CTL. The solid, black-dashed, gray-dashed, and dotted lines denote MSE, cpT, gz, and Lq, respectively.

Atlantic warming ranging from 1 to 4 K are conducted. dry air advection into the Sahel and precipitation re- The model has 90-km resolution and produces a state- ductions. A moist static energy analysis shows that the of-the-art simulation of the present-day climate of West entire atmospheric column warms in all warm SSTAs Africa, including the precipitation and flow fields. simulations. The destabilization at low levels that supports With 1- and 1.5-K SSTAs prescribed in the Atlantic, convection in 1K-CTL and 1.5K-CTL is because of in- rainfall rates increase by 30%–50% over most of West creases in the low-level moisture content of the atmo- Africa, including both the Guinean coast region (88–108N) sphere. Compared with CTL the moisture content of the and the Sahel (138–168N). With Atlantic warming of 2 K atmosphere decreases, and the low-level atmosphere is and higher, precipitation still increases between 88 and stabilized, with Atlantic warming of 2 K and higher. 128N, but Sahel rainfall decreases. These decreases are In all of the simulations, the southerly monsoon flow quite substantial, with approximately 60% reductions with onto the Guinean coast is maintained, and precipitation warming of 2.5 K and 85% reductions with warming from in that region increases. These increases become more the SSTA 4K simulation. This nonlinear response in Sahel closely confined to the coast as the Atlantic warms. The rainfall is the focus of the analysis of the model simulations. hydrodynamics in this region is dominated by moisture Warming less than 2 K in the Atlantic Ocean is ac- considerations. Warmer Gulf of Guinea SSTs are asso- companied by decreases in low-level geopotential ciated with enhanced evaporation, and the moisture heights in the Gulf of Guinea and, as a consequence, content of the southerly monsoon flow increases. Precip- decreases the large-scale meridional geopotential height itation anomalies do not develop over the SST anomalies gradients across West Africa. By a geostrophic argu- because of the presence of large-scale subsidence over ment, this leads to easterly wind anomalies in the central the Gulf of Guinea. Warming even up to 4 K does not Sahel and the southerly monsoon flow onto the Guinean break through this subsidence. coast is not diverted as strongly to the east. The moisture Climate change and variability in West Africa and the from the Gulf of Guinea is able to penetrate farther to Sahel is complicated and regional. In this analysis, we the north over the western Sahel, and this leads to pre- focus only on Atlantic SST forcing to develop a deeper cipitation increases in that region. With greater Atlantic understanding of the region’s hydrodynamic response. warming, the weakening of the continental-scale meridi- The application of these results to the global warming onal geopotential height gradient continues, and the problem is limited by this focus. For example, we do not easterly anomalies strengthen. With 2-K warming—the take into account Indian Ocean warming, which is known threshold value in these simulations—the easterly anom- to be a prominent forcing factor for decadal variability alies reverse the westerly flow over the Sahel. The result is over the Sahel, and we do not increase CO2 levels in these

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