Abrupt rainfall transitions over the Greater Horn of Africa: Observations and regional model simulations



[1] We examine the yearly occurrence of a monsoon jump of approximately 20° latitude during the boreal spring and summer rainy seasons over the Greater Horn of Africa (GHA). This jump is in contrast with a simple model of a smoothly varying ITCZ over the region. The rainfall jump is observed annually during April and May in three precipitation data sets and in regional climate model simulations using the PSU/NCAR Mesoscale Model (MM5). The MM5 simulations show the rainfall jump to be roughly coincident with abrupt circulation changes that occur as the Somali jet develops during April, May, and June. In particular, the cross-equatorial (meridional) branch of the Somali jet forms along the East African coast in April, bringing moisture (and rainfall) northward to the southern slopes of the Ethiopian plateau. This meridional branch forms well before the northern zonal branch of the jet, which diverts moisture eastward from southern Ethiopia and feeds the Indian monsoon. These results establish a framework for understanding the precipitation cycle over the GHA, and provide a foundation for improving subseasonal forecasts over drought-prone regions of eastern Africa.

1. Introduction

[2] The seasonal cycle of rainfall in the tropics is ultimately driven by seasonal changes in solar radiation. While the solar cycle varies smoothly with latitude, rainfall cycles can often respond with sudden jumps, particularly over the continents where varying surface properties (e.g., topography, soil moisture, land/sea contrasts) can create important circulations and nonlinear feedbacks in the climate system. Abrupt jumps in the latitude of rainfall occur annually during the onset of the West African and East Asian monsoons [Sultan and Janicot, 2000; Hagos and Cook, 2007; Xue et al., 2004]. Here we show that similar sudden transitions occur during the evolution of the East African monsoon.

[3] Understanding abrupt transitions in rainfall is crucial to generating useful seasonal and subseasonal forecasts. While most prediction studies focus on the total seasonal rainfall amount [e.g., Thiaw et al., 1999], the timing of that rainfall is equally important to agriculture and water resource management. For example, pastoralists in northern Kenya and southern Ethiopia have identified the spring onset date as the single most desirable piece of forecast information [Luseno et al., 2003]. The timing of the seasonal transitions can also be an important control on drought, since a late onset or early termination can initiate or exacerbate drought conditions [e.g., Segele and Lamb, 2005]. As an example, anomalously short spring rainy seasons in recent years (2000–2004) have led to recent food shortages in southern Ethiopia [Verdin et al., 2005].

[4] The purpose of this study is to understand abrupt seasonal and subseasonal transitions that occur during the spring and summer over the Greater Horn of Africa (GHA), focusing particularly on the onset of the rainy seasons over Ethiopia, Eritrea and Djibouti. Observations and regional climate model simulations are used to investigate the timing of the jumps and examine the associated atmospheric circulation patterns.

2. Background

[5] Rainfall over the Greater Horn of Africa (GHA) demonstrates strong seasonality. In the southern GHA (e.g., Kenya, Uganda, Northern Tanzania, and Somalia) the rainy seasons come during the boreal spring (MAM) and fall (ON). These seasons are referred to as the “long” and “short” rains, respectively. Further north, in Ethiopia, Djibouti and Eritrea, the primary rainy seasons are in the spring (belg) and summer (kiremt) with a dry period (bega) in the winter [e.g., Beltrando and Camberlin, 1993].

[6] Atmospheric circulation patterns also demonstrate a strong seasonal cycle over the GHA, guided by changes in insolation and SSTs. A prominent feature is the development of the low-level Somali jet [e.g., Findlater, 1969] in the boreal summer. The East African topography plays a crucial role in confining this jet to the coast of Somalia [Slingo et al., 2005; Rodwell and Hoskins, 1995]. This flow pattern is reversed during the winter monsoon season.

[7] Much of literature concerning rainfall over the GHA has focused on establishing predictive relationships between seasonal rainfall totals and remote indicators. In the northern GHA (Ethiopia, Djibouti and Eritrea), the majority of these studies examine the predictability of summer (JJAS) rainfall. Rainfall in this region is negatively correlated with El Niño [e.g., Beltrando and Camberlin, 1993; Camberlin, 1995; Gissila et al., 2004; Segele and Lamb, 2005; Korecha and Barnston, 2007] which is also generally associated with basin-wide heating of the Indian Ocean. Studies disagree on the importance of independent modes of SST variability in the Indian Ocean, such as the dipole mode [Gissila et al., 2004; Korecha and Barnston, 2007]. Beltrando and Camberlin [1993] find that influence of the Pacific and the Indian Oceans is strongest during June and September, while, in July and August, rainfall intensity is additionally influenced by local convection and the West African monsoon [Flohn, 1987].

[8] Other studies find very strong correlations between Ethiopian summer rainfall and the strength of the Indian monsoon, independent of a mutual dependence on ENSO [e.g., Bhatt, 1989; Camberlin, 1997]. Vizy and Cook [2003] discuss the mechanisms of this connection using regional climate model simulations.

[9] In contrast to the Ethiopian summer rains, the short rains during October and November over the southern GHA (Kenya, Uganda, Tanzania) are positively correlated with El Niño events [e.g., Ogallo, 1988; Hastenrath et al., 1993; Indeje et al., 2000; Mutai and Ward, 2000], and also with the Indian Ocean dipole mode [e.g., Behera et al., 2005; Clark et al., 2003; Bowden and Semazzi, 2007], showing strong dependence on Indian Ocean surface temperatures [e.g., Goddard and Graham, 1999; Black, 2005]. Severe flooding off the coast of East Africa in October and November is sometimes associated with strong dipole events [Black et al., 2003]. Recent modeling studies have explored some of these teleconnections [Sun et al., 1999; Anyah and Semazzi, 2006, 2007; Conway et al., 2007].

[10] The spring (MAM) rainy season, which occurs in both the northern and southern GHA, is the least studied and the most complex. In general MAM rainfall does not exhibit very strong relationships with external modes of variability such as El Nino or the Indian Ocean dipole mode [e.g., Ogallo, 1988; Hastenrath et al., 1993]. This is partly because rainfall anomalies in boreal spring are less coherent spatially and temporally than during the summer or fall. Several authors have demonstrated that early and late stages of the MAM rainy season exhibit different rainfall variability patterns [e.g., Camberlin and Philippon, 2002; Zorita and Tilya, 2002], suggesting that correlations may be more meaningful at smaller temporal or spatial scales.

[11] Several studies have nonetheless made progress toward establishing some of the features controlling the evolution of the spring rainy season. Spring rainfall events in northern Ethiopia and Eritrea may be connected to excursions of the midlatitude ridge-trough systems [e.g., Camberlin and Philippon, 2002]. Rainfall in eastern Ethiopia is negatively correlated with tropical cyclone activity in the southwestern Indian Ocean [Shanko and Camberlin, 1998]. Other studies have noted that strong MAM rainfall over the southern GHA is associated with a westerly anomaly in the lower to midtropospheric winds across equatorial Africa during March [e.g., Camberlin and Wairoto, 1997; Okoola, 1999; Camberlin and Philippon, 2002; Zorita and Tilya, 2002; Camberlin and Okoola, 2003].

[12] Camberlin and Okoola [2003] examined the timing of the onset and cessation of the long rains over the southern GHA (Kenya, Uganda and Northern Tanzania). They found the onset date to vary by almost 2 months, and be an important determinant of the cumulative rainfall for the season. In a more recent study, Pohl and Camberlin [2006a, 2006b], suggest that the MJO is an important source of intraseasonal variability modulating the MAM rainfall over the southern GHA.

3. Methods

[13] A combination of observed, analyzed and model data is used to characterize the seasonal precipitation cycle over the GHA and examine the associated atmospheric dynamics. Precipitation data sets serve to establish a climatology of the seasonal precipitation cycle over the GHA, and to validate the model-generated precipitation rates. Reanalyzed data are used for validation of the three-dimensional model fields. Regional climate model simulations provide a high resolution, self-consistent representation of the regional momentum, energy, and moisture budgets. The regional climate model, run at 30 km resolution, is better able to capture the complex topography and land surface features over East Africa than existing three-dimensional data sets (e.g., the NCEP reanalysis). The primary results of this study are duplicated in the observations, the reanalysis, and model and thus do not depend solely on the accuracy of any one approach.

3.1. Observations and Reanalysis

[14] Three observational precipitation data sets are used, based on varying combinations of satellite-derived precipitation and rain gauge measurements. The Climate Research Unit (CRU CL 2.0) precipitation climatology [New et al., 2002] is 30 years long (1961–1990) and is derived exclusively from an extensive network of rain gauge measurements. The data are gridded at 10-min (0.17° × 0.17°) resolution, though the actual resolution of the underlying data over the GHA is considerably lower. These data provide a long term precipitation record, but do not extend over the oceans and are available only on a monthly timescale.

[15] In contrast, the Tropical Rainfall Measure Mission (TRMM) 3B42 V6 product is derived using a combination of satellites (TRMM and other), calibrated using monthly averaged rain gauge data. It is available every 3 h from 1998–2006 at a resolution of a 0.25° × 0.25° [Huffman et al., 2007]. Another satellite-based product, the NCEP Climate Prediction Center's African Rainfall Estimate (CPC RFE2.0), is used by the Famine Early Warming System (FEWS) over Africa [Herman et al., 1997]. This data set (FEWS from here forward) is available daily from 2001–2006 at 0.1° × 0.1° resolution.

[16] Both the TRMM and FEWS data sets offer high spatial and temporal resolution. However, both are relatively short, being only 9 and 5 years long, respectively, so one cannot construct a true climatology. Nonetheless, we use short-term climatologies created from these data sets for comparison with the model simulation. This can be justified by looking at longer term data sets (e.g., CRU TS 2.1 [Mitchell and Jones, 2005]) which show that the most prominent synoptic-scale patterns in the seasonal cycle over the GHA have not changed significantly in recent decades.

[17] The NCEP/DOE AMIP-II Reanalysis [Kanamitsu et al., 2002], based on the NCAR/NCEP reanalysis, is available from 1979 to present and is used to validate the model geopotential height and wind fields. It is also used to supply initial and boundary conditions to the model. Compared with the regional model simulations, however, the reanalysis has a much coarser resolution both in the horizontal (2.5° × 2.5°) and the vertical dimensions, and so it cannot resolve the finer-scale circulations associated with, for example, the East African topography.

3.2. Regional Climate Model

[18] The regional climate model simulations enhance existing three-dimensional data sets (e.g., NCEP-2 reanalysis) by linking the atmospheric circulation with a realistic precipitation cycle. Previous studies [e.g., Anyah and Semazzi, 2006] have suggested that high-resolution simulations that resolve the complex topography of East Africa are necessary in this region to accurately capture the precipitation cycle.

[19] The regional climate model used for this study is the PSU/NCAR Mesoscale Model (MM5 [Grell et al., 1994]). MM5 is a limited area, nonhydrostatic model with 23 terrain-following sigma levels with a range of physical parameterization options. All simulations in this study use the Kain-Fritsch convective scheme [Kain, 2004], the Rapid Radiative Transfer Model (RRTM) longwave radiation scheme [Mlawer et al., 1997], the Blackadar planetary boundary layer scheme [e.g., Blackadar, 1979], and the Dudhia simple ice explicit moisture scheme [Dudhia, 1989]. These particular physics options have been found in previous studies to successfully represent the hydrological cycle over Africa [e.g., Vizy and Cook, 2002, 2003; Patricola and Cook, 2007].

[20] The domain for the simulations is chosen to include most of eastern Africa and the Arabian Peninsula (Figure 1). The model resolution is set to 30 kilometers to resolve the primary topographic features of the region (shaded in Figure 1). The domain is centered on the Ethiopian Plateau. To the south is the Turkana Channel, which separates the Ethiopian Plateau from the high topography of Kenya and Tanzania, including Mt. Kilimanjaro. This break in the topography is thought to be an important conduit for moisture reaching the continental interior via moisture transport by the Turkana jet [Kinuthia and Asnani, 1982; Indeje et al., 2001].

Figure 1.

Regional map showing the model domain (box) and topography (shading) on a 30-km grid.

[21] Land surface conditions (e.g., soil moisture and surface albedo) for MM5 are specified using 24 categories of land surface types from the United States Geological Survey (USGS) data set (Figure 2). These surface conditions are fixed throughout the integrations. Sea surface temperatures (SSTs) are based on Shea et al. [1992], and land surface temperatures (including lake surfaces) are initialized based on the NCEP-2 reanalysis. Initial and lateral boundary conditions are also taken from the NCEP-2 reanalysis.

Figure 2.

USGS land surface types used to determine surface conditions throughout the model simulations.

[22] Three MM5 integrations are discussed. The first two are 7-month (March–September) simulations of the evolution of the rainy seasons in 2002 and 2003. The first fifteen days of the simulations are discarded as the model is spinning up. SSTs and lateral boundary conditions are updated every 6 h to match the 2002 and 2003 analyzed conditions. Model output is recorded and stored every 3 h.

[23] The third integration is a 9-month (March–November) climate simulation. In this simulation, the boundary and surface conditions are taken from the monthly climatological mean values of the NCEP-2 reanalysis averaged over 1979 to 2004, interpolated linearly to 12-h intervals. This isolates seasonal forcing of the climate within the domain by removing the effects of interannual variability and transient perturbations entering the domain. While this climate-mode simulation is theoretically different from an observational climatology, previous studies [e.g., Vizy and Cook, 2002] have shown that in the tropics this method can accurately reproduce observed climatological fields.

4. Observations and Simulations of the GHA Rainy Seasons

[24] To be valuable for this study, the MM5 simulations must capture the primary rainfall and circulation patterns associated with the spring and summer rainy seasons of the GHA. This section will identify particular strengths and weaknesses of the model in capturing these features.

4.1. Rainy Season Precipitation

[25] Figures 3a, 3b, and 3c display observed precipitation for the spring rainy season (MAM) from the TRMM, FEWS and CRU “climatologies” respectively (described in section 3.1). All three data sets have been interpolated to the same 1° × 1° grid for the comparison. Spring rainfall rates are largest over the southern half of the domain, with local maxima over the eastern Congo, Lake Victoria, coastal Tanzania, and southwestern Ethiopia. The three data sets agree fairly well, despite being derived from different measurements (satellite versus ground-based) and averaged over different time periods.

Figure 3.

MM5 and observed precipitation (mm/d) during the spring (MAM) rainy season based on (a) TRMM, (b) FEWS, (c) CRU data sets, and (d) MM5 simulated precipitation. CRU data is not available over the ocean. Selected contour lines are shown at 1 mm/d (black solid), 5 mm/d (black dashed) and 9 mm/d (white).

[26] Springtime precipitation from the MM5 climate-mode simulation is shown in Figure 3d. As in the observations (Figures 3a3c), nearly all of the precipitation is in the lower half of the domain. While modeled precipitation rates are greater than the observed, the model captures the observed precipitation maxima over the Congo (perhaps with interference from the lateral boundaries), Lake Victoria, and southern Ethiopia following the southern topographic boundary. Excessive rainfall over Lake Victoria is often seen in regional climate simulations [e.g., Anyah and Semazzi, 2004, 2007] due to lake-induced local convection.

[27] One concern in this simulation is that model precipitation over the Indian Ocean and along the coast of Tanzania is too strong. The excess precipitation over this region is caused by a series of spurious “storms” which form near the eastern domain boundary south of the equator and propagate into the domain before dissipating when hitting the coastline. These depressions extend up to 500 hPa and are most prominent in March and April during which time they deposit a large amount of rainfall over the Western Indian Ocean. We suggest that these disturbances are being created near the eastern edge of the domain due to inconsistencies at the lateral boundary. While these storms are worrisome, they do not appear to interfere with the results that will be presented in section 5.

[28] Observed precipitation for the summer rainy season (JJAS) is shown in Figures 4a, 4b, and 4c. Again, the three observational data sets agree well, although precipitation rates in the CRU data are significantly greater than those in the satellite-based observations. The observed precipitation has shifted north of the equator, and a primary precipitation maximum is located over the western Ethiopian highlands. Additional rainfall extends to the west over southern Sudan, and the Congo Basin maximum has moved about 7° farther north.

Figure 4.

MM5 and observed precipitation (mm/d) during the summer (JJAS) rainy season based on (a) TRMM, (b) FEWS, (c) CRU data sets, and (d) MM5 simulated precipitation. CRU data is not available over the ocean. Selected contour lines are shown at 1 mm/d (black solid), 5 mm/d (black dashed) and 9 mm/d (white).

[29] The MM5 climate-mode simulation (Figure 4d) captures the important features of the summer season precipitation over the GHA. The large-scale shift of the rainfall to the north is well represented in the model, and the primary rainfall maximum is located over the western Ethiopian highlands. The zonal structure of the precipitation across Ethiopia and Sudan is also reasonable. However, there is too much rainfall north of Lake Victoria and in the eastern Congo Basin. The anomalous rainfall over the western Indian Ocean is still present, though weaker than in the spring.

[30] In both spring and summer the model precipitation is consistent with the observations, especially compared with the GCM simulations [e.g., Anyah and Semazzi, 2006] which are on coarser resolution and fail to resolve the topography accurately. Overall, the validation is stronger over Ethiopia, Eritrea, and Djibouti than in the extended domain. However, at subdegree spatial scales (not shown), the model precipitation field shows unrealistic spatial fluctuations over distances of 30–90 km, which are associated with local topographic and surface variations. This suggests that at these smaller scales, the model precipitation is too sensitive to local surface features. This sensitivity carries over to the larger scales as well, producing too much rainfall over lakes (e.g., Lake Victoria) and topographic features. Despite these deficiencies, the model is able to sufficiently capture the seasonal rainfall cycle as will be further demonstrated along with the results in section 5.

4.2. Rainy Season Circulation Patterns

[31] The mean low-level circulation patterns associated with the spring rainy season (MAM) are shown at 850 hPa for the NCEP-2 reanalysis (Figure 5a) and at 825 hPa for the MM5 climate mode simulation (Figure 5b). Similarly, circulation patterns for the summer rainy season (JJAS) are shown in Figures 5c and 5d. (The model and reanalysis are shown at slightly different pressure levels, however these levels are near enough to exhibit very similar features.) The MM5 simulation results are interpolated to the NCEP 2.5° horizontal grid for these comparisons. Topography (plotted at 30 km resolution) is shaded black.

Figure 5.

Geopotential height and wind from (a) the NCEP-2 reanalysis spring climatology (MAM), (b) the springtime MM5 climate mode simulation (MAM), (c) the NCEP-2 reanalysis summer climatology (JJAS), and (d) the summer MM5 climate mode simulation (JJAS). The NCEP-2 geopotential heights and winds (a and c) are shown at 850 mb, while MM5 heights and winds (b and d) are shown at 825 mb. The MM5 output is interpolated to the NCEP 2.5° grid. Black shading indicates topography extending above the plotted pressure surfaces.

[32] For the spring season (Figures 5a and 5b), MM5 and NCEP reanalysis show similar circulation features. A region of high geopotential heights extends into the domain from the south. In the south, the flow is predominantly easterly or southeasterly, and more geostrophic in the east than in the west. Across the central part of the domain the wind tends to be easterly, including the flow impinging on the Ethiopian highlands and in the Turkana channel (i.e., the Turkana jet). In the north, the flow is essentially geostrophic and dominated by anticyclonic circulation around highs over the Sahara and the Arabian Peninsula. These regions are more distinct from each other in the reanalysis, inducing stronger zonal gradients than in the MM5 simulation. A closed depression over the western Indian Ocean is present only in the model output in association with the spurious storms moving over this region (described in section 3).

[33] The summertime height field is dominated by a deep Indian monsoon trough extending across the Arabian Peninsula, and, weakly, into the Sudan (Figures 5c and 5d). In addition, the high to the south is significantly stronger than in the spring. The most prominent and stable feature of the summertime circulation is the low-level Somali jet, which carries moisture across the equator and eventually feeds the Indian monsoon. Over the Ethiopian highlands, the flow is primarily westerly, in contrast to the easterly flow of the spring season (Figures 5a and 5b). Despite this change in wind direction, the Turkana jet remains southeasterly. Note the inflow of low level winds from the south (Indian Ocean), southwest (Congo basin), and north (Mediterranean) over Ethiopia.

[34] The NCEP and MM5 circulation and geopotential height fields show remarkably good agreement in the summertime, with all of the major features reproduced by the model. Particularly in the summertime, this agreement may be partially due to the dominance of the lateral forcing derived from the NCEP reanalysis data.

[35] These comparisons suggest that the model is able to capture many important features associated with the spring and summer rainy seasons. The model validates particularly well over Ethiopia, a primary focus area of this study. Further validation will be provided along with results in the following sections.

5. Results

5.1. Seasonal Rainfall Cycle

[36] Agricultural calendars in the central GHA are closely tied to the timing of local rainfall, therefore it is useful to classify regions with respect to the similarity of their seasonal cycles. One way to do this is to classify regions by the month in which they receive their peak rainfall. This is done in Figure 6, in which the spring or summer month of peak precipitation is plotted for the observations and model over the focus region (Ethiopia, Eritrea, and Djibouti). In the observations (Figures 6a6c), note the sharp divide between southeastern Ethiopia, which mostly receives maximum rainfall in April, and the northern and western parts of Ethiopia and Eritrea, which receive maximum rainfall in July and August. The MM5 simulation captures this divide as well, though the spatial transition region is somewhat less distinct. These results are contradictory to a simplified model of the ITCZ following the solar cycle; such a model would show a series of increasingly narrow horizontal latitude bands for April, May, June, and July as the ITCZ migrated northward.

Figure 6.

Month of peak spring rainfall during March through August over 1° × 1° pixels in Ethiopia and Eritrea as recorded by (a) TRMM, (b) FEWS, (c) CRU, and (d) the MM5 climate-mode simulation.

[37] To investigate this further, we examine the seasonal cycle at each 1° × 1° grid point, starting with the CRU data set. On the basis of the similarity of the CRU seasonal cycles, four regions are identified as shown in Figure 7. The resulting regional classification scheme is then applied to the remaining two observational data sets (TRMM and FEWS) and the model output. Figure 8 shows each of these seasonal cycles sorted by region.

Figure 7.

Regions with similar annual precipitation cycles according to the CRU precipitation climatology.

Figure 8.

Annual rainfall cycles at 1° × 1° grid points over the (a) northwest, (b) northeast, (c) southwest, and (d) southeast study regions (Figure 7), as recorded by TRMM, FEWS, CRU, and the MM5 climate-mode simulation. Regional averages are indicated with solid black lines.

[38] Each region has a unique cycle of rainfall. The northwestern highlands (Figure 8a), which receive the heaviest overall rainfall, experience only one large peak during the months of July and August. The northeast (Figure 8b) experiences a peak in April, with a second peak in August, after a very dry month in June. In a small region in the southwest (Figure 8c), elevated precipitation levels are observed throughout the entire rainy season (March–October). Finally, the low-lying southern and eastern parts of Ethiopia (Figure 8d) experience two peaks in precipitation, with the larger peak in April, a second, smaller peak in October and a dry period in between. These regions are similar to rainfall “clusters” identified over Ethiopia by Gissila et al. [2004].

[39] The three observational data sets agree well not only on the mean seasonal cycle in each of these regions, but also on the degree of variability of the cycle within the regions. For the MM5 climate simulation, however, the variability is too large and the classification scheme, which is based on the CRU data set, is not entirely appropriate. This is particularly true for the small, transitional southwestern region and the northwestern highlands where the topography is highly variable. However, the model does capture the mean seasonal cycle in each of the four regions, and reproduces the seasonal peaks correctly in the mean. This suggests that, while the model cannot capture local variability accurately, it can capture the primary features of the annual precipitation cycle at a regional scale (∼1000 km).

[40] Together, Figures 68 demonstrate that the cycle of precipitation in this region is complicated, with large variations from east to west as well as north to south. For example, from late May to June rainfall declines over northeastern Ethiopia, while at the same latitude the summer rainy season intensifies over the northwestern highlands. Thus there is a westward shift in rainfall in addition to a northward shift. Again, these patterns are not consistent with a smoothly varying, zonally symmetric model of precipitation in this region.

5.2. Monsoon Jump Over the GHA

[41] To further resolve intraseasonal shifts in GHA rainfall, daily time resolution is needed. Figure 9 shows the annual cycle of TRMM rainfall averaged between 33° and 43° east longitude for 1998 through 2006, plus the 9-year climatology. This figure ignores zonal asymmetries, but is instructive in understanding abrupt seasonal transitions.

Figure 9.

Daily TRMM precipitation (mm/d) averaged between 33° and 43°E longitude for 9 years (1998–2006) and a climate-mode simulation.

[42] All of the 9 years demonstrate some common features. In January through mid-March, the primary rainfall band is located near 10°S. In many years and the climatology, a weaker rainfall band appears south of 10°N during late February or early March. This weak rainfall band is associated with an early phase of the Ethiopian spring rainy season (belg), primarily over the southwestern focus region (Figures 7 and 8).

[43] An abrupt transition occurs every year between late March and the end of May. During the transition, the rainfall maximum jumps approximately 20° of latitude, from 10°S to 10°N, a monsoon jump. This jump is larger than over West Africa, where a monsoon jump of approximately 5° latitude is observed [Sultan and Janicot, 2000, 2003; Sultan et al., 2003; Hagos and Cook, 2007]. This East African monsoon jump is a persistent climatological feature. It is clearly seen in the climatology though, as expected, the abruptness of the transition is somewhat smoothed by averaging. Similar features can be observed in the FEWS data set (not shown).

[44] Looking more closely at the timing of the jump (April and May), one can see that the transition occurs in several stages which vary from year to year. In 2003, for example, a “first jump” in early April brings the precipitation maximum from south of the equator to the equatorial region (0°N–7°N). In late May, another jump occurs, bringing rainfall to 10°N. A similar two-stage transition is also clearly observed in 2002 and 2005. In other years (e.g., 2000), one jump from approximately 10°S to 7°N occurs in April, after which the rainfall north of the equator migrates slowly northward and intensifies; in these years the second jump is not pronounced. The timing of the first jump can vary between 20 March and 20 April. The subsequent northward shift from 7°N to 10°N occurs between early May and early June.

[45] During the transition period (April and May), two rainfall bands are often present: one band persisting south of the equator and second developing north of the equator. In this sense, the transition is not precisely a “jump” in rainfall, but rather the rapid development of a second branch of rainfall north of the equator. Later in the season, the southern branch dissipates or, in another interpretation, migrates rapidly northward to merge with the northern branch (Semazzi, personal communication). The two rainfall bands may represent a double ITCZ, similar to that observed by Anyah and Semazzi [2006] during October and November. For simplicity we will continue to refer the abrupt latitude transitions over the GHA as “jumps”, though it is understood that rainfall south of the equator may persist for a short time after the onset of rainfall in the north.

[46] During some years (e.g., 1999, 2003, and 2004), a reversed jump is also present during boreal fall, while in other years (e.g., 2001, 2005, and 1998) the southern recession of rainfall is smoother. When the reverse jump is present, the latitude of maximum rainfall shifts abruptly from 10°N to the equator in mid-October, and then from the equator to about 10°S in early December.

[47] The MM5 simulations capture the dominant features of these jumps in 2002 and 2003 (Figure 10). The timing of the transitions for 2002 and 2003 agrees very well with the observations. In the climate simulation, the day-to-day variability is suppressed because of the smoothly varying lateral and surface boundary conditions. Nonetheless, relatively abrupt shifts in the latitude of rainfall occur, suggesting a nonlinear response to the smooth external forcing. In contrast to the individual years, the initial jump in the climate simulation is somewhat later than in the observations. In all three model simulations, persisting summer rainfall south of the equator is a result of the anomalous precipitation near Lake Victoria and over the western Indian Ocean (Figure 4).

Figure 10.

Daily MM5 simulated precipitation (mm/d) averaged between 33° and 43°E longitude for two years (2002–2003) and a climate-mode simulation.

[48] Based on the timing of these jumps, it is useful to divide the evolution of the boreal summer rainy season into three stages. The first stage occurs in March before the first jump. The second, transitional stage occurs in between the jumps during the last half of April and the first half of May. The final stage, associated with the summer rainy season persists from June through early September. In this three-stage description of the monsoon onset, the Ethiopian belg rainy season (MAM) is divided between the first and second stages described above, while the summer rainy season (kiremt) is associated with the final stage.

[49] Figures 11a11f show precipitation maps over Ethiopia for each stage at quarter degree resolution. During Stage 1 (March), precipitation is weak and primarily confined to the southwestern part of the country. In the Stage 2 (April/May), precipitation covers southern and eastern Ethiopia. This is the primary rainy season in the south. In Stage 3 (June), the rainfall is strongest in the northwestern highlands. The summer precipitation pattern shown here for June continues through September.

Figure 11.

(Top) Maps of precipitation (mm/d) based on TRMM precipitation data for (a) Stage 1, (b) Stage 2, and (c) Stage 3 of the rainy season onset. (Bottom) MM5 climate-mode precipitation for (d) Stage 1, (e) Stage 2, and (f) Stage 3.

[50] The hydrodynamical features that accompany these jumps are examined below to better understand why they occur, and how they are related to larger-scale circulation patterns and moisture sources.

5.3. Monsoon Onset Hydrodynamics

[51] Figures 12a, 12b, and 12c show low-level (850 hPa) geopotential heights and moisture flux vectors from the MM5 climate mode simulation for Stages 1, 2, and 3, respectively. Before the first jump, in Stage 1 (March), the low-level moisture flux is primarily easterly across the Horn of Africa, including Ethiopia. Between the equator and about 10°N the zonal geopotential height gradient is positive. The resulting flow down the height gradient transports modest amounts of moisture through the Turkana channel from the Arabian Sea. Recall that precipitation rates are quite low north of the equator during this period (Figure 11).

Figure 12.

Maps of 825 hPa geopotential height (m) and moisture flux (g/m2/s) based on the MM5 climate-mode simulation for (a) Stage 1, (b) Stage 2, and (c) Stage 3 of the rainy season onset. Grid points below the topography are shaded black.

[52] In Stage 2 (April/May), low geopotential heights over the western part of the domain persist, but are located further north. Southerly flow organizes along the equatorial coast forming the meridional branch of the Somali jet. The jet brings moisture from the Southern Hemisphere into southern Ethiopia, producing rainfall on the southern slopes of the Ethiopian plateau.

[53] In June (Stage 3), the geopotential height field is dominated by two features: (1) the Indian monsoon trough which extends over the Arabian Peninsula and (2) a ridge located along the East African topography in the south of the domain. These features generate a very strong geopotential height gradient from the southwest to the northeast corner of the domain. A trough in the western part of the domain persists as well, located still further north over the Sudan.

[54] The fully formed Somali jet flows across the southwest-northeast pressure gradient. Both the meridional and zonal branches of the jet are now present and the meridional cross-equatorial branch of the jet is wider than in April/May. The zonal (northern) branch, not present in April/May, carries moisture across the Arabian Sea toward the Indian monsoon, and thus diverts the primary moisture source that fueled the Ethiopian spring rains.

[55] During the summer, moisture is drawn into northern Ethiopia by a combination of factors. First, southeasterly flow associated with the trough over the Sudan brings moisture northward from the Congo basin. Second, the deep monsoon trough to the east introduces a negative geopotential height gradient across northern Ethiopia and the Red Sea, drawing the moisture eastward toward northwestern Ethiopia and Eritrea. Over the northwestern Ethiopian highlands, this inflow provides moisture for the summer rainy season. These results are consistent with those of Camberlin [1997] who shows summer rainfall over northern Ethiopia to be highly correlated with the strength of Indian monsoon indices. A deepening of the Indian monsoon trough leads both to increased monsoon rainfall over India, and to stronger westerlies across northern Ethiopia, drawing increased moisture from central Africa [Vizy and Cook, 2003].

[56] Figure 13 provides validation of these results using the NCEP reanalysis. As before, Figures 13a, 13b, and 13c show low-level (925 hPa) geopotential heights and moisture flux vectors from the NCEP reanalysis for Stages 1, 2, and 3, respectively. All of the primary features discussed above are present in the reanalysis. However, many detailed features, such as the horizontal structure of the coastal Somali jet and the flow through the Turkana channel, are not resolved in the 2.5° reanalysis fields.

Figure 13.

Maps of 850 hPa geopotential height (m) and moisture flux (g/m2/s) based on the NCEP reanalysis for (a) Stage 1, (b) Stage 2, and (c) Stage 3 of the rainy season onset. Grid points below the topography are shaded black.

[57] In both the model and the reanalysis (Figures 12 and13), the meridional branch of the Somali jet forms along the East African topography almost a month before the zonal branch. This is in contrast to the original description of the springtime development of the jet by Findlater [1971] who suggested that the full jet forms offshore in February and slowly moves onshore throughout the spring.

[58] The results presented in this section are robust in both the model and the observations/reanalysis. However, the model simulation is important in that it solidifies the physical link between circulation and precipitation processes that would otherwise only be observed in independent data sets.

5.4. Timing of the Monsoon Jumps

[59] Figure 14 demonstrates the association between the timing of first monsoon jump, which initiates the spring rains in southern Ethiopia, and the formation of the meridional branch of the Somali jet. The top row of Figure 14 shows zonally averaged precipitation (33°E to 43°E) for the 2002, 2003, and climate-mode simulations. The bottom row displays the 910 hPa wind direction averaged over a box between 38°E–45°E and 2.5°S–2.5°N. The flip in these equatorial winds from easterly to southerly is used as a proxy for the formation of the meridional branch of the Somali jet. Grey vertical lines mark the approximate date of the first jump in precipitation for each simulation. It is clear that the timing of the most-prominent precipitation jump corresponds with a shift in the direction of the equatorial winds.

Figure 14.

(Top) Zonally averaged precipitation (33°–43°E) and (bottom) 910 hPa equatorial wind direction averaged over 38°E–45°E longitude and 2.5°S–2.5°N latitude based on 2002, 2003, and climate-mode simulations. Gray scale is the same as in Figure 9. Vertical grey lines indicate the timing of the first precipitation jump.

[60] In 2002 and 2003, the equatorial wind direction changes abruptly in both the model simulations and the NCEP reanalysis (not shown). In 2002, equatorial winds oscillate for several weeks between easterly and southerly before stabilizing in mid-April. Rainfall north of the equator is likewise intermittent in early April. The final northward jump in rainfall (grey vertical line) occurs roughly 1 week after winds have stabilized. In 2003, the flip in wind direction is remarkably abrupt, occurring over a 2-d period. The onset of rainfall north of the equator is similarly abrupt and coincident with the change in wind direction.

[61] In the climate-mode simulation the wind direction changes more smoothly, over the course of about a month. This suggests that winds within the domain are influenced by the wind field at the lateral boundaries, which is smoothly interpolated from monthly mean reanalysis. The jump in rainfall is nonetheless present in the climate simulation, but does not occur until the end of April, after the meridional jet has completely formed.

[62] The sample size is too small to infer any statistical correlation between the two processes (the flip in equatorial wind direction and jump in precipitation). However, the similarity in timescales, along with the simple physical mechanism linking cross-equatorial moisture transport to southern Ethiopian precipitation, strongly suggests that the two events are connected. This analysis does not, however, establish a cause-and-effect relationship between the two events.

[63] Figure 15 is similar to Figure 14 but focused on the second precipitation jump, which brings rainfall to the northwestern Ethiopian highlands, and the formation of the zonal branch of the Somali jet. To capture zonal jet formation, the wind direction is averaged between 50°E–55°E and 10°N–15°N, off of the tip of Somalia. In both 2002 and 2003, the wind direction in this region is chaotic during March and April, before the stable southwesterly jet is established in mid-May. In both years, the jump in rainfall (grey vertical lines) occurs several weeks after the jet is first established. However, initially, the jet does not extend across the Arabian Sea, but feeds a local circulation off the coast of Africa (not shown). The jump in rainfall occurs only after the zonal branch of the jet strengthens, and the winds across northern Ethiopia switch from easterly to westerly.

Figure 15.

(top) Zonally averaged precipitation (33°–43°E) and (bottom) 910 hPa equatorial wind direction averaged over 50°–55°E longitude and 5°N–10°N latitude based on 2002, 2003, and climate-mode simulations. Gray scale is the same as in Figure 9. Vertical grey lines indicate the timing of the second precipitation jump.

[64] In the climate-mode simulation, the wind direction changes smoothly over a month and a half. The onset of rainfall occurs only after the wind direction has stabilized and the full jet has formed.

6. Conclusions

[65] Improved forecasts of rainy season onset and termination over the GHA would greatly benefit people in the region, for example, by allowing farmers to plan when to plant their crops and communities to better manage their water resources. However, forecasting the timing of intraseasonal changes in rainfall is difficult, and a better understanding of the processes governing these changes is necessary before useful forecasts can be achieved.

[66] In this study, we have examined the onset of the boreal spring and summer rainy seasons over the GHA, using observed and analyzed data sets and regional climate model simulations. Precipitation climatologies were generated from three precipitation data sets, two primarily satellite-based (TRMM and FEWS), and one surface-based (CRU). While each of these data sets has disadvantages (the satellite-based records are short, and the surface stations are sparse and available only over land), combined with the regional climate model simulations they provide a consistent and reliable picture of the rainy season evolution.

[67] The MM5 simulations using 2002, 2003, and climatological lateral and surface boundary conditions were carried out at 30 km resolution. These simulations augment existing three-dimensional data sets (e.g., NCEP-2 reanalysis) by providing high enough spatial resolution to capture the complex topography of East Africa (e.g., the Turkana Channel), and linking the atmospheric circulation with a realistic precipitation cycle.

[68] The rainfall climatologies and simulations both show that rainfall over the GHA is not zonally uniform as would be suggested by a simple migration of a zonally symmetric ITCZ. Because of this, the rainfall cycles and the relative importance of different rainy seasons vary significantly from one region to another. Over Ethiopia, for example, four separate regions can be identified with unique seasonal cycles. These complexities must be considered when generating regional forecasts.

[69] A primary result of this paper is the documentation of a 20° latitude jump in rainfall that occurs in the boreal spring over the GHA. The jump occurs during the months of April and May at which time the rainfall moves abruptly from approximately 10° south of the equator to 10° north of the equator. The precipitation jump is larger than the 5° latitude jump documented during the onset of the West African monsoon [Sultan and Janicot, 2000].

[70] The full jump is often composed of two smaller jumps. A large first jump occurs between late March and early April when precipitation moves from 10°S into the Northern Hemisphere, bringing rain to the southern slopes of the Ethiopian plateau. The second smaller jump occurs in May, bringing the rainfall band further north into the western Ethiopian highlands. Thus the onset of the monsoon over the GHA is divided into three stages: (1) early spring (March) before the initial jump, (2) a transitional period (April/May) between the two jumps, and (3) the summer rainy season (June) after the second jump.

[71] The associated circulation patterns were investigated with the MM5 simulations, using the NCEP-2 reanalysis for a course-scale validation. When examined separately, each of the onset stages represents a different phase in development of the Somali jet. During Stage 1, the equatorial winds are easterly and the Somali jet has not yet begun to form. During Stage 2, only a meridional branch of the Somali jet is present; equatorial winds are southerly, but not yet diverted eastward toward the Arabian Sea. This provides a moisture source for the southern slopes of the Ethiopian plateau. The Somali jet is fully formed in Stage 3.

[72] The circulation changes associated with the formation of the jet are often abrupt and roughly coincident with the jumps in precipitation. This is particularly true in 2003 when the equatorial winds flip direction over a period of 2 d at the same time as the precipitation jumps north of the equator. This process is very different than over West Africa [Sultan and Janicot, 2000], where the monsoon jump is thought to be caused by a local inertial instability along the Gulf of Guinea [Hagos and Cook, 2007].

[73] To our knowledge, the two-phase development of the Somali jet and its association with GHA precipitation has not been previously documented in the literature. From early observations [Findlater, 1971], the jet was thought to form offshore in the early spring, and slowly move onshore as the season develops.

[74] This study identifies a monsoon jump over the GHA, and shows that the precipitation jump coincides with the formation of the Somali jet. However, to support prediction, cause-and-effect relationships must be established between the background climate regime, the Somali jet development, and the associated changes in precipitation. The model simulations presented here lay the groundwork for exploring these subjects in greater depth.


[75] This research was supported by the NASA Earth Systems Science Fellowship program and NSF grant ATM-0415481. We extend our gratitude to Edward Vizy for his valuable assistance with the model simulations and analysis, and also to our three anonymous reviewers for their very careful and insightful comments which have greatly improved this manuscript.