Guinean coastal rainfall of the West African Monsoon

Quarterly Journal of the Royal Meteorological Society Q. J. R. Meteorol. Soc. 137: 1828–1840, October 2011 A

Guinean coastal rainfall of the West African Monsoon

Hanh Nguyen,a* Chris D. Thorncrofta and Chidong Zhangb

aDepartment of Atmospheric and Environmental Sciences, University at Albany, SUNY, New York, USAbRosenstiel School of Marine and Atmospheric Science, University of Miami, Miami, Florida, USA

*Correspondence to: H. Nguyen, Centre for Australian Weather and Climate Research, Melbourne, Victoria, Australia.E-mail: [email protected]

The nature and variability of the springtime rainfall onset at the Guinean coastis explored. The coastal onset is defined as the time when the oceanic extent ofthe rain band crosses the Equator. The oceanic extent is defined as the locationwhere the rain rate falls below 2 mm day−1. The mean coastal onset averaged over1979–2009 from Global Precipitation Climatology Project (GPCP) rainfall data ison 11 May with a standard deviation of 14.5 days. This result is robust and seenin different rainfall products. The coastal rainfall demise is determined as the timewhen simultaneously the peak rainfall at the coast weakens and the rainfall over theSahel intensifies. The mean coastal demise is 26 June with a standard deviation of9.5 days. This implies that the mean length of the coastal phase is 47 days with astandard deviation of 13 days. The coastal onset and its demise are primarily drivenby changes in sea-surface temperature (SST) between the Guinean coast and theEquator. There exists a 301 K threshold below which the equatorial cold tonguedevelops rapidly and leads to the coastal onset 10 days later. The same thresholdapplies to SSTs near the coast, where the water cooling precedes the end of thecoastal rain phase by 16 days. Sea-surface temperature anomalies in the southeastAtlantic in winter are potentially a predictor for the coastal rainfall variability inspring. Copyright c© 2011 Royal Meteorological Society

Key Words: West African coastal rainfall; monsoon onset; cold tongue

Received 24 September 2010; Revised 16 May 2011; Accepted 20 May 2011; Published online in Wiley OnlineLibrary 25 July 2011

Citation: Nguyen H, Thorncroft CD, Zhang C. 2011. Guinean coastal rainfall of the West African Monsoon.Q. J. R. Meteorol. Soc. 137: 1828–1840. DOI:10.1002/qj.867

1. Introduction

The West African continent is characterized by two distinctrainy seasons: one is boreal spring when rainfall occurs nearthe Guinea coast (∼5◦N) (e.g. Gu and Adler, 2004) andthe other is summer as the main rain band moves inlandover the Sahelian region (∼10◦N) (e.g. Sultan and Janicot,2000, 2003; Le Barbe et al., 2002; Lebel et al., 2003). Despitethe fact that the rainy season near the coast is the wettestperiod of the whole West African monsoon (WAM) (Gu andAdler, 2004; Thorncroft et al., 2011), most research effort,especially in recent years, has been devoted to the natureand variability of the Sahelian rainfall (e.g. Hagos and Cook,2007; Fontaine et al., 2008). There is a continued lack ofdetailed knowledge and understanding of the nature and

variability of the coastal rainy season. The main aim of thispaper is to shed more light on this understudied part of theWAM, especially its onset.

According to Le Barbe et al. (2002), the first rainy seasonstarts on the coast in February and the rain moves regularlynorthward until reaching 13◦N in May. On the otherhand the Sahelian season starts with a surge in mid-June.Lebel et al. (2003) showed that the oceanic rainy seasonis characterized by a progressive increase of the moist airflow from the ocean into the continent, associated with theseasonal migration of the intertropical convergence zone(ITCZ) from its southern position in boreal winter to itsnorthern position in boreal summer. Interestingly, Sultanand Janicot (2003) determined the pre-onset date of thesummer monsoon in the Sahel to be 14 May when the ITCZ

Copyright c© 2011 Royal Meteorological Society

Guinean Coastal Rainfall of West African Monsoon 1829

is centred at 5◦N. As will be shown, this pre-onset is actuallyconcomitant with the coastal onset that we are exploring inthis paper.

Other studies have shown evidence for the relationshipbetween the development of the cold tongue and the WAMrainfall. Gu and Adler (2003) showed that the presenceof a rainy season near the Guinean coast (5◦N) duringApril–June roughly matches the seasonality of sea-surfacetemperatures (SSTs) in the tropical eastern Atlantic. Theyargued that acceleration of the low-level winds in the Gulf ofGuinea during March–June is a direct response to the rainfallnear the coast. Equatorial upwelling is then immediatelyenhanced due to the shallow thermocline. Consequently SSTdecreases, which eventually forms the cold tongue complex.With continued SST decrease, rainfall in the equatorialregion is suppressed in summer. In a modelling study,Okumura and Xie (2004) showed that the equatorial coolingintensifies the southerly monsoon flow in the Gulf of Guineaand pushes the rain band inland from the Guinea Coast. Arecent study of Caniaux et al. (2011) described in more detailthe formation of the Atlantic cold tongue and emphasized itspossible role in influencing the Sahelian onset. They showedthat the cold tongue exhibits large interannual variability interms of its commencement, horizontal extension, durationand intensity. They suggested that the strengthening of thelow-level circulation between the Equator and the WestAfrican land can impact the Sahelian onset. Thorncroftet al. (2011) further emphasized the role of the cold tonguein regulating the timing and intensity of the coastal rainfallin spring. Thorncroft et al.(2011) described the migration ofthe rain band in the West African region between 10◦W and10◦E in terms of four key phases: (i) the oceanic phase fromNovember to April; (ii) the coastal phase from mid-May toJune; (iii) the transitional phase during the first half of July;and (iv) the Sahelian phase from mid-July to September.They suggested that variability in the coastal rainfall islikely to be associated with variability in the Atlantic coldtongue development (timing, intensity) and the shallowmeridional circulation (SMC) related to the Saharan heatlow (intensity, location). Based on these studies we proposethat it is important to examine more closely how the coldtongue formation relates to the nature of the coastal rainfall,which is missing in most previous studies.

The main aim of this paper is to improve ourunderstanding of the nature and interannual variabilityof the coastal rainfall onset in the WAM. We hypothesizethat a late coastal onset would be associated with a latecold tongue formation and a longer oceanic phase, and/or astrong heat low, and vice versa. Previous studies suggestingthe heat low influence on the Sahelian onset (e.g., Sultanand Janicot, 2003; Ramel et al., 2006; Lavaysse et al., 2009)did not pay much attention to the coastal onset. We willexplore the possible joint influences of the cold tongue andheat low on the coastal onset.

The paper is organized as follows. Section 2 brieflydescribes the data used in this study. Section 3 introducesthe method of detecting the coastal rainfall onset. In section4 we revisit the mean annual cycle of the rain band and theassociated cold tongue to provide a context for the variabilityanalysis. Section 5 shows the interannual variability of thecoastal rainfall including its onset, length and intensity.These characteristics are compared with those of the coldtongue to explore their relationship. Section 6 discussesthe relationship between the coastal onset and the Sahelian

onset. The role of the Saharan heat low in both coastaland Sahelian phases is discussed in section 7. Summary andconcluding remarks are given in the final section.

2. Data

A combination of observations and reanalysis data are usedin this study. Three rainfall datasets are used:

• The Tropical Rainfall Measuring Mission (TRMM)pentad data (Huffman et al., 2007) based on acombination of multiple satellites and gauge analysesat resolutions of 0.25◦ and 3 h. The data cover theperiod from 1998 onward.

• The Climate Prediction Center Merged Analysis ofPrecipitation (CMAP) data (Xie and Arkin, 1997) arealso derived from a merge of multiple satellites andgauge data but include precipitation from the NationalCenter for Environmental Protection/National Centerfor Athmospheric Research reanalysis primarily to fillin gaps at high latitudes. The resolution is 2.5◦ andthe data record covers from 1979 onward.

• The Global Precipitation Climatology Project (GPCP)pentad data (Xie et al., 2003) are estimated frommonthly analysis (Adler et al., 2003) that mergesstation rain-gauge observations, satellite geostationaryand low-orbit infrared and passive microwaveretrievals, and sounding observations. Two versionsof GPCP rainfall datasets are used. One with highresolution (1◦) is available from 1997 onwards(Huffman et al., 2001) and the other with lowerresolution (2.5◦) is available from 1979 onwards (Xieet al., 2003).

These datasets are used to study the annual cycle of therain band as well as to determine the rainfall phases and onsetdates. The locations of the peak rainfall during the oceanicand coastal phases are very close, as shown in Thorncroftet al.(2011). The high resolution data are needed to detail thenature of these two phases. The longer time series are usedto examine the statistical robustness of the gross features inthe coastal rainfall evolution.

Agreement and discrepancies among the different rainfalldata provide some confidence limits on our analysis andavoid biases due to a particular dataset. The period usedfor this comparison is from 1998 to 2007. For consistencythe TRMM data is interpolated to the same spatio-temporalresolution as the other two datasets.

The NOAA Optimum Interpolation 1/4 degree daily SeaSurface Temperature (SST) data Version 2 (OIV2 SST)(Reynolds et al., 2007) are used. The data are availablefrom September 1981 onwards. The mean sea-level pressure(MSLP) used here comes from the third generation of theEuropean Centre for Medium-Range Weather Forecastsreanalysis daily dataset (ERA-interim) at 1.5◦ spatialresolution for the period 1989–2009 (Simmons et al., 2007;Uppala et al., 2008). All the data are smoothed by a 15-day boxcar filter to eliminate high-frequency fluctuations(synoptic disturbances such as equatorial waves). All zonalaverages are undertaken between 10◦W and 10◦E.

3. Definition of the WAM rainfall onsets

The climatological rainfall averaged between 10◦E and 10◦Wdisplays two pronounced peaks associated with a coastal

Copyright c© 2011 Royal Meteorological Society Q. J. R. Meteorol. Soc. 137: 1828–1840 (2011)

1830 H. Nguyen et al.

phase and a Sahelian phase of the WAM (Figure 1(a)).Figures 1b and c are latitudinal distributions of annualrainfall peak frequency and associated amounts based ontwo datasets with different resolutions. The peak just offthe coast (3–4◦N), more frequent than that over the Sahelregion (11–12◦N), clearly highlights the importance of thecoastal phase. The closeness of the location of the peak inrainfall during the oceanic (1–2◦N) and coastal phases andthe relative small amplitude of the oceanic peak (Figure (1c))explains why we see only one preferred peak location nearthe coast in Figure (1b) and may also explain why previouslythey have been considered as one single ‘oceanic regime’(Lebel et al., 2003). In contrast, the coastal and continentalregimes are separated by 8◦ latitude, making the two regimesclearly distinct, even with low resolution rainfall data (Sultanand Janicot, 2000, 2003).

The present analysis indicates that the coastal onset cannotbe determined using the same method as in Sultan andJanicot (2003) for the Sahelian onset (called ‘monsoon onset’there). The Sahelian onset is defined by them as simultaneousrainfall decrease at the coast (5◦N) and rainfall increase inSahel (10◦N), as the rain band maxima rapidly shifts from5◦ to 10◦N, and also with the increase of the positive slope ofthe rainfall at 15◦N. This method is sensitive to the data used(see Table I). There is a 10–15 day uncertainty, which isthe time necessary for the rainfall environment to shift fromthe coast to the Sahel (Sultan and Janicot 2003; Fontaineet al., 2008). This 10–15 day period actually coincides withthe transitional phase during the first half of July defined byThorncroft et al. (2011). In contrast, the distance betweenthe location of the oceanic and the coastal rainfall peaksis only 2◦ of latitude (Figure (1c)), with the oceanic peakrainfall being much weaker than the coastal one. Indeedthere is some inconsistency between the different rainfalldatasets with respect to the rainfall over the ocean. Thisresults from a lack of in situ observations over the ocean,which are needed for validation of the datasets (Adler, Xieand Huffman, 2009, personal communication).

This inconsistency can be highlighted by comparing thethree rainfall products at 2.5◦ and 5 day resolutions over theperiod of 1998–2007 (this comparison period is limited bythe TRMM data availability). Figure 2 shows climatologicaltime–latitude diagrams of mean rainfall (averaged between10◦W and 10◦E) together with latitudes of peak rainfall basedon the three precipitation datasets. Each of them highlightsthe presence of a wide rain band over the ocean fromFebruary to April. Meridional excursions in the location ofpeak rainfall during the oceanic phase are not well separatedfrom rainfall in the coastal region, which complicates anydefinition of the coastal phase in terms of local rainfall. Allthree feature comparable patterns but clearly have differentintensities. For example, in June TRMM has the strongestcoastal peak rainfall with 16 mm day−1, about 3 mm day−1

higher than the other two. The CMAP precipitation is ingeneral stronger over the ocean and weaker over the landthan GPCP. These differences are consistent with Gruberet al. (2000), Yin et al. (2004) and Fontaine and Louvet(2006).

The locations of peak rainfall also differ between thethree datasets. The shift of peak rainfall from the oceanicphase at around 1.25◦N to the coastal phase at 3.75◦Nvaries between the end of March in GPCP and the endof April in CMAP (Figure 2). This disagreement is mainlyrelated to the differences in rainfall intensity over the ocean.

For example GPCP oceanic rainfall is weak compared withCMAP. Therefore the coastal onset, if defined by the rapidshift of the peak rainfall, occurs on 1 April, which is morethan a month earlier than in the CMAP dataset (5 May).The coastal onset thus defined is also earlier in the TRMMdataset (17 April) compared with CMAP, but in this caseit is due to the much wetter coastal region in the TRMMdataset.

Because of this inconsistency we propose that a morepractical working definition of the coastal onset associatedwith the demise of the oceanic phase should be based onthe meridional extent of the rain band, which is a robustfeature in the time series and is clearly an intimate part ofthe coastal onset (Figure 2). We are especially interestedin the rapid shift of the southern boundary of the rainband during the oceanic–coastal transition when it movesfrom 7.5◦ –5◦S to the Equator within a month that alwaysaccompanies the increase in rainfall at the coast. This rapidtransition is clearly illustrated by the rainfall time series at1.25◦S shown in Figure 3(a). Despite the slight discrepanciesin intensity, the time series from the three datasets all showthe same sharp weakening at this latitude, with a decreasefrom about 5 mm day−1 at the beginning of May to lessthan 1 mm day−1 at the end of May. This clearly highlightsthe demise of the oceanic phases but does not necessarydetermine the time when peak rainfall shifts from the oceanto the coast, which is less consistent among the differentrainfall datasets. With this definition, the coastal onset isnot necessarily characterized by when it starts to rain at thecoast but rather by when the oceanic phase ends.

We thus monitor the time evolution of this sharpweakening with the oceanic extent of the rain band �y.Following the annual mean rain band, we chose this to bedefined as when the rainband values are greater than orequal to 2 mm day−1 over the ocean. The coastal onset isthus defined as when �y crosses the Equator. Annual cycleof the oceanic extent �y depicted in Figure 3(b) shows a verygood agreement between the three datasets in stark contrastto the timing of the shift in peak rainfall discussed earlier.Using this criterion, the coastal onset is defined for eachyear. The interannual variability of the coastal onset datedefined this way is also consistent between the three datasets(Figure 4). The difference from one dataset to another is 0–3pentads. The mean coastal onset is close to the 27th pentad(i.e., close to 11 May) with a standard deviation of around 3pentads (see Table II for further details). For a given rainfalldataset, the interannual variability of the onset is important,ranging from the 20th pentad (6 April) to the 33rd pentad(10 June). This variability will be discussed in more detail insection 5.

In the rest of this paper we will focus our analysis on thelow resolution GPCP data because it does not suffer fromthe adjustment of the oceanic precipitation estimates overthe global oceans as in CMAP.∗ Table II shows the meancoastal onset dates based on GPCP. The mean coastal onsetdate, tC, for the period 1979–2009 is the 27th pentad, whichcorresponds to 11 May with a standard deviation of 2.9

∗In CMAP the adjustment is based on comparisons of the satelliteestimates over the tropical Pacific atoll region where rain gauge dataexist. In contrast GPCP does not attempt to reduce the bias, except that itadjusts the IR-based rainfall estimates to the passive microwave-derivedrainfall estimates with the implication that the latter are bias-free (E. J.Zipser, personal communication).



(a)

(b)

(c)

Figure 1. (a) Latitude–time diagram of GPCP pentad 2.5◦ precipitation (mm day−1). (b) Latitudinal distribution of peak rainfall frequency (bars, rightaxis, number of pentads per year) and associated amounts (line, left axis, mm yr−1) based on the GPCP pentad 2.5◦ precipitation averaged from 10◦Wto 10◦E for February to September 1979–2007. (c) Same as (b) except based on daily 1◦ GPCP precipitation for 1997–2009.

Table I. Mean Sahelian onset dates tS determined from GPCP low-resolution rainfall data and from various sources∗.

Source* Period tS Standard deviation

GPCP 1979–2009 6 July 1.8 pentadsSJ03 (IRDr) 1968–1990 24 June 8 daysFL06 (CMAP) 1979–2004 28–29 June 8.5 daysFLR08 (OLR) 1970–2001 30 June 15.6 daysDGB08 (VIMT) 1979–2004 25–29 June 2 pentads

* SJ03, Sultan and Janicot (2003); FL06, Fontaine and Louvet (2006); FLR08, Fontaine et al. (2008); DGB08, Dalu et al. (2009).

pentads. It is interesting to note that this date is close tothe pre-onset date of 14 May found in Sultan and Janicot(2003) (Table II) using a dynamical variable (meridionalwind at 925 hPa). This coincidence is actually consistentwith the link between low-level southerly acceleration andthe coastal phase (Gu and Adler, 2004; Okumura and Xie,2004; Thorncroft et al.(2011)). The coastal onset date foreach year will be used as a reference for the compositeanalysis in the following sections.

The end of the coastal phase is defined as the time whenrainfall time series at the coast (i.e., at 3.75◦N) and overthe Sahel (i.e., at 11.25◦N) intersect. Its mean date tEC is 26June with a standard deviation of 1.9 pentads, coincidentwith the Sahelian onset date in Sultan and Janicot (2003).This will be discussed further in section 7. Note that thereis a delay of 5–7 days in Thorncroft et al. (2011), becausethe dates in Thorncroft et al.(2011) are defined from annualmean analyses.



TRMM

(a)

(b)

(c)

CMAP

GPCP

Figure 2. Time (pentad)–latitude (2.5◦) diagram of (10◦W–10◦E) zonal mean precipitation (mm day−1) averaged over 10 yr (1998–2007) from(a) TRMM, (b) CMAP and (c) GPCP products. Blue curves indicate the latitude of peak rainfall. This figure is available in colour online atwileyonlinelibrary.com/journal/qj

Table II. Mean coastal onset dates tC for three different datasets: TRMM, CMAP and GPCP. The resolution is 2.5◦ and 5daily (pentad).

Dataset Period tC Standard deviation

TRMM 1998–2007 9 May (26.5 pentads) 3.3 pentadsCMAP 1998–2007 10 May (26.9 pentads) 2.6 pentadsGPCP 1998–2007 10 May (26.9 pentads) 3.4 pentadsGPCP 1979–2009 11 May (27 pentads) 2.9 pentadsSJ03* 1968–1990 14 May 9.5 days

*SJ03, Sultan and Janicot (2003) pre-onset date defined from low-level southerly winds.



(a)

(b)

Figure 3. Time series of zonal mean (a) precipitation at 1.25◦S (mm day−1) and (b) oceanic rain band extent (meridional location of the 2 mm day−1

southern rain band contour �y) averaged over 10 yr (1998–2007) from TRMM (solid), CMAP (dash) and GPCP (dot).

Figure 4. Interannual variability of coastal rainfall onset dates tC (number of pentads) for TRMM (solid), CMAP (dash) and GPCP (dot).

For completeness, and for later comparison with thecoastal onset date, we have calculated the Sahelian onsetdate, tS. The Sahelian onset is defined here to be when thepeak rainfall shifts to a location north of 7.5◦N. Table I liststhe Sahelian onset dates based on the different data andfor different periods in the literature. The mean Sahelianonset based on our analysis occurs on 6 July with a standarddeviation of 1.8 pentads. This Sahelian onset is 1 to 2 weekslater than in the literature. Different Sahelian onsets couldresult from different methods, data and periods used inprevious studies. However we argue that this 1–2 week lagis likely to correspond to the transitional phase in the firsthalf of July that precedes the Sahelian onset (Thorncroftet al., 2011) and was considered as part of the Sahelianphase in previous studies. For instance in Sultan and Janicot(2003), a composite based on the Sahelian onset shows thatrainfall at 5◦N starts to decline at the time of the Sahelianonset and simultaneously rainfall at 10◦N rapidly increases(see their figures 4(c and d)). However the Sahelian rainyseason is established only 15 to 20 days after that, after thetransitional phase as described in Thorncroft et al.(2011).The transitional phase is also omitted (or included in theSahelian phase) in Fontaine and Louvet (2006) and Fontaineet al. (2008).

4. Annual cycle of the rainfall and associated sea surfaceconditions

To examine the time evolution and spatial patterns of rainfalland improve our understanding of the role of the coldtongue in the coastal onset, a time-lag composite analysis is

performed for 28 years (1982–2009). The coastal onset timetC is the reference point of day 0. The composite enables usto better characterize the different phases of the rain bandthan using a simple annual mean, especially when largevariability in the onset dates is observed (see Figure 4).

Figure 5 is a time–latitude diagram of composite rainfalland SST. The four rainfall phases described by Thorncroftet al. (2011) are clearly shown: (i) the oceanic phase(t < tC = 0) in which rainfall is present over a wide rangeof latitudes over the tropical ocean and part of the coastalregion; (ii) the coastal phase (tC < t < tC + 46 = tEC), inwhich the rain band is much narrower with a peak locatednear the coast at about 4◦N (although there is actually a meanpeak at the coast before this date that is consistent with thebias in GPCP discussed in section 3); (iii) the transitionalphase (tC + 46 < t < tC + 60) characterized by weakenedrainfall and a steady poleward movement; and (iv) theSahelian phase (t > tC + 60). It is interesting to observethat the Sahelian onset is 60 days after the coastal onset,suggesting a possible link between the two. This suggests aprediction potential of 2 months in advance for the Sahelianonset here, if the relationship holds at interannual time-scales. We will discuss this relationship in more detail insection 7.

Figure 5 shows that before the coastal onset tC, thewarm waters are collocated with the oceanic rainfall in theequatorial and coastal regions, while after the onset the coldtongue development with its peak just south of theEquator is associated with suppressed rainfall. Note thatthe equatorward boundary of the rain band (1 mm day−1)during the period between tC − 60 and tC − 10 is roughly



Time (days)

tC tEC tS

Figure 5. Composite diagrams for the period 1982–2009 of NOAA OIV2 SST (shading, K) and GPCP precipitation (contours with values �7 mm day−1,bold). The white contour represents the SST 301 K isotherm. Vertical lines at day 0, day +46 and day +60 represent the coastal onset tC, end of coastalphase tEC and Sahelian onset tS respectively. Data are zonally averaged between 10◦W and 10◦E.

collocated with the 301 K isotherm. Based on this andthe analysis presented in section 6 we determine the onsetof the cold tongue using this 301K threshold applied toSSTs averaged between 10◦W–10◦E and 0–5◦S. With thisdefinition the cold tongue initiates about 10 to 12 days priorto the coastal onset. This is characterized by the rapid shiftof the 301 K contour from 5◦S to the Equator. The cooledwaters (below 301 K) lead to a suppression of rainfall there.

After the cold tongue is established there remains aprominent region of warm water between the West Africancoastline (∼5◦N) and a few degrees of latitude south of it.This warm water is collocated with the wettest part of thecoastal phase. While persistent into the coastal phase thewater steadily cools, reaching 301 K about 35 days after thecoastal onset tC. As the coastal water cools the rainfall off thecoast weakens. Interestingly the peak rainfall moves inlandaround the time when the coastal water reaches the 301 Kthreshold seen to be important in the equatorial region (Xieand Carton, 2004). This strongly suggests a thermodynamiccontrol for the demise of the coastal phase. One simpleexplanation for the thermodynamic control is that thecore of the rain band during the coastal phase is locatedimmediately off the coast. Note that the formation of the coldtongue itself is established in response to the accelerationof surface cross-equatorial southerly winds (Mitchel andWallace, 1992).

5. Interannual variability of the coastal phase

Figure 6(a) shows time series of coastal onset tC togetherwith the length of the coastal phase TC for the period1979–2009. The coastal onset exhibits large interannualvariability. Its occurrence ranges from 6 April in 2005 to10 June in 1991, consistent with the large variability of thecold tongue onset (Figure 7; and Caniaux et al., 2011). Thenature and origin of such variability is explored below. Alsoincluded in Figure 6(a) is the time series of the end of coastalphase tEC. This date varies between 5 June in 2005 and 15July in 1984 and 1999. Correlations between the variousindices (defined in Table III) that are significant at the 99%level are given in Table IV and are discussed below. Theend of the coastal phase tends to follow the evolution ofthe coastal onset and is positively correlated with the coastalonset with a correlation coefficient of 0.49 (significant at the

99% level). However, this is more likely to be influenced bythe Sahelian onset.

Consistent with the composite pattern (Figure 5) themean TC is 9.2 pentads with a standard deviation of 2.6pentads. It ranges between 3 pentads in 2008 and 14 pentadsin 1983. The TC tends to vary out of phase with the coastalonset with a correlation of −0.77 significant at the 99% level.In contrast, the contribution of the end of coastal phase isless marked, consistent with weaker interannual variabilitycompared with the coastal onset (Table IV).

For completeness we briefly highlight the variability inrainfall intensity during the coastal phase. The coastal rainfallintensity IC averaged over the coastal phase TC taken at thecentre of the rain band (4◦N) is shown in Figure 6(b)together with the rainfall amounts for a fixed period basedon the climatologically wettest period (mid-May to mid-June). The IC exhibits large variability, ranging from just5.1 mm day−1 in 1987 to 11.1 mm day−1 in 1988 with amean of 7.9 mm day−1. Somewhat consistent with this arethe rainfall totals for the climatological wet period varyingsubstantially from 6.2 mm day−1 to 10.5 mm day−1 for thoseyears respectively.

6. Influence of the SSTs on the coastal phase

The composite analysis demonstrates the tendency foroceanic rainfall to decrease rapidly once the SST beneathdrops below 301 K at the end of the oceanic phase and thecoastal phase (Figure 5). Given this, we hypothesize thatinterannual variability in SST in this region will impact boththe onset and termination of the coastal phase.

6.1. Role of the Atlantic cold tongue onset on the coastal onset

The cold tongue onset tCT described in section 4 occurs on1 May, and varies between 22 March in 1997 and 22 Mayin 2009 with a standard deviation of 16.3 days (Figure 7).Interestingly tCT leads the coastal onset tC by 10 days, as seenin the composite pattern (Figure 5). Moreover the time seriesof tC and tCT are positively correlated with a correlation of0.61 significant at the 99% level (Figure 8(a)).

Caniaux et al. (2011) explored the evolution of the coldtongue formation for the period 1982–2006. In contrast toour study, they only considered the relationship between



Table III. Estimates of all indices analysed in the study.

Variable Indice Mean Minimum Maximum Standard deviation

Cold tongue onset tCT 1 May 22 March 1997 22 May 2009 16.3 daysCoastal onset tC 11 May 6 April 2005 10 June 1991 2.9 pentadsWarm water cooling tWC 10 June 26 May 1997 23 June 2009 8 daysEnd of coastal phase tEC 26 June 5 June 2005 15 July 1984/1999 1.9 pentadsLength of coastal phase (pentads) TC 9.2 3 14 2.6Saharan heat low onset tSHL 24 June 29 May 2003 11 July 1997 10.2 daysSahel onset tS 6 July 20 June 2005 25 July 1979 1.8 pentads

(a)

(b)

Figure 6. (a) Time series of coastal onset (tC, black), coastal demise or transitional phase (tEC, red) and Sahelian onset (tS, green) dates (pentads) andlength of the coastal phase (TC, blue). (b) Coastal rainfall (mm day−1) averaged over the TC (bar charts) and between mid-May and mid-June (blue).

Table IV. Correlation coefficients between different indicesdefined in Table III that are significant at the 99% level.

Indices Correlation r

tC – tEC 0.49tC –TC 0.77tC – tCT 0.61tEC – tWC 0.61tC – tS 0.5tEC – tS 0.7tS – tWC 0.63

cold tongue variability and the Sahelian onset. Their coldtongue index was defined as the surface area occupied bySSTs <298 K in the domain of 30◦W–12◦E and 5◦S–5◦N.The cold tongue onset was defined by a threshold area of0.4 × 106 km2. They show a correlation of 0.59 betweenthe cold tongue formation and the Sahelian onset. Thecold tongue formation defined by Caniaux et al. (2011) isequivalent to a surface area 3 K colder than the 301 Kthreshold used here. The mean date for the onset of the coldtongue according to Caniaux et al. (2011) is 10 June, roughlya month after the coastal onset. They also mentioned that theresults are independent of the threshold chosen. Indeed thecorrelation between our tCT and the cold tongue formationdefined by Caniaux et al.(2011) is 0.74, suggesting that SST

anomalies in the cold tongue region in May often persistinto June. As expected, the correlation between the coldtongue formation in Caniaux et al. (2011) and our coastalonset tC is 0.71, significant at the 99% level for the sameperiod (Figure 8b). This also suggests that the SST coolingin May and in June is somewhat linear.

6.2. Role of the coastal water on the demise of the coastalphase

We have hypothesized above that cooling of the coastalwater is coupled to the suppression of the coastal rainfall(Figure 5). We explore this by correlating the end ofthe coastal phase tEC and the date when SSTs near thecoast (averaged between 10◦W–10◦E and 2◦ –5◦N) reach aparticular threshold (Figure 9). The correlation is significantat the 99% level for SST values between 300.6 K and 299.2 K,with the highest correlation of 0.61 obtained for a thresholdof 300 K. This result suggests that the lead time betweenwhen the SST reaches 301 and 300 K will likely benefitthe prediction of the end of the coastal phase. Indeed, onaverage, the date when the warm tongue typically reachesthe 301 K threshold is 10 June, 21 days earlier than forthe 300 K threshold is reached and 16 days earlier thanthe typical end of the coastal phase. This analysis furthersuggests a thermodynamic influence on the demise of thecoastal rainy season and, by inference, the beginning of the



Figure 7. Time series of the cold tongue onset (tCT, solid) and the warm water cooling (tWC, dash) (see text for detail).

(a) (b)

Figure 8. Scatter plot of coastal rainfall onset tC and (a) cold tongue onset tCT over 28 years (1982–2009), and (b) tC and cold tongue formation fromCaniaux et al. (2011) over 25 years (1982–2007). The black line represents the linear regression and its correlation coefficient is indicated on top of eachfigure.

Figure 9. Correlation between the cooling of the coastal water tWC and theend of coastal phase (tEC, solid) and between tWC and the Sahelian onset (tS,dash) as a function of the cooling SST threshold over 28 years (1982–2009).The horizontal black lines represent the 95% and 99% confidence levels.

transitional season and potentially even the beginning of theSahelian rainy season.

6.3. Relationship between the coastal rainy season and globalSSTs

Given that a contemporaneous relationship between rainfalland SSTs exists, it is of particular interest to explore theextent to which these SST anomalies were present in monthsbefore this. If present this would motivate using SSTs aspredictors in seasonal forecasting efforts for the region

(e.g. Ward, 1998). Correlations between the coastal rainfallamounts in mid-May–mid-June and global monthly SSTsfor different time-lags are displayed in Figure 10. Thesecoastal rainfall amounts are correlated with SSTs in thePacific, consistent with previous studies on the causes ofWest African rainfall variability – with wet years generallyoccurring during La Nina conditions in the Pacific and whenthe equatorial Atlantic is anomalously warm (Ward, 1998).Figure 10 highlights the fact that significant SST anomaliesin the southeast Atlantic were present roughly 4 monthsahead of the coastal rainy season. We recommend thatpredictability suggested by these results be taken advantageof in prediction efforts in the region.

7. Relationship between the coastal phase and Sahelianonset

The composite of rainfall (Figure 5) shows that the Sahelianonset, tS, occurs about 60 days after the coastal onset. Itsinterannual variability is much less than that of the coastalonset tC (Figure 6(a)). Its occurrence ranges from between20 June in 2005 and 25 July in 2002. The correlation betweenthe coastal and Sahelian onsets is 0.5, significant at the 99%confidence level (Figure 11). This significant correlation,together with the composite, highlights the potential for tC

to be a useful predictor of the Sahelian onset. The simplestexplanation for this relationship is that both onsets arestrongly influenced by the cold tongue and that, as discussedabove, variability in the cold tongue in Spring is correlated



Figure 10. Lag-correlation between coastal rainfall in mid-May–mid-June and monthly SSTs for the period 1979–2009. Correlation values significantat the 95% level are shaded.



Figure 11. Scatter plot between coastal rainfall onset tC and Sahelianrainfall onset tS (pentads) over 31 yr (1979–2009).

with that in Summer. On the other hand, as suggested inFigure 6(a) the end of the coastal phase tEC and the Sahelianonset tS vary in phase with a correlation of 0.7. The latteroccurs about 2 weeks later. This suggests that the Sahelianonset does not start until the coastal phase ends, whichwould be driven by the coastal SST.

Consistent with the coastal onset, the Sahelian onset tS

is correlated with the cold tongue onset tCT. However thecorrelation coefficient is 0.42, significant only at the 95%level. This is consistent with the hypothesis that the coldtongue may influence the Sahelian onset but is not the mainfactor triggering it. On the other hand a better correlation isobtained between the Sahelian onset tS and the coastal watercooling tWC (see Figure 9). The highest correlation of 0.63is obtained for the 300 K threshold. Consistent with the endof the coastal phase tEC, the correlation drops dramaticallyfor warmer SST thresholds and relatively rapidly with colderones. This further highlights the condition that the coastalphase ends before the Sahelian onset. With the mean coastalwater cooling occurring roughly 1 month before the Sahelianonset, this result suggests the potential for using coastal SSTanomalies in spring to predict the Sahel onset.

8. Role of the Saharan heat low on the rain band

We have pointed out that much of the variability in thecharacteristics of the coastal phase of the WAM is associatedwith the variability in tropical Atlantic SSTs directly to thesouth. Meanwhile, we consider the potential role played bythe Saharan heat low (SHL). This is motivated by previousstudies that have suggested an important role for the heatlow on the Sahelian onset (e.g., Sultan and Janicot, 2003;Ramel et al., 2006; Sijikumar et al., 2006; Lavaysse et al.,2009).

Figure 12 shows the composite evolution of the heat lowwith respect to the coastal onset. The mean heat low isstrongest in April, prior to the mean coastal onset on 11May. This is consistent with Zhang et al. (2008) who showedthat the heat-low-associated shallow meridional circulation

(SMC) is most intense in April. The location of the heatlow moves steadily poleward until the end of the coastalphase. At the end of this phase, the heat low reaches itsnorthernmost location around 22◦N, where it remains untilafter the Sahelian phase. Based on similar analyses, Ramelet al. (2006) and Lavaysse et al. (2009) have suggested thatthis shift (or Saharan heat low onset) may be important forencouraging the Sahelian onset.

8.1. Impact of the heat low on the coastal phase

The relationship between the heat low and the coastalrainfall is investigated through correlation analysis. A prioriwe would expect the location, intensity and shift of the heatlow potentially to affect the coastal phase in terms of bothits intensity and demise. However all correlations are foundto be weak (not shown). The highest correlation found was−0.25 for the SHL intensity and the coastal onset, significantonly at the 70% level. Motivated by the lag seen in thecomposite pattern (Figure 6), we computed lag-correlationbetween coastal rainfall in mid-May–mid-June and the heatlow (not shown). The highest correlation obtained is −0.44for June and July significant at the 95% level. In fact in2005 where the coastal onset is early (6 April), the heatlow maximum occurs during the coastal phase (not shown).This combined with the heat low maximum occurring inApril, suggests that the intensity of the heat low does notinfluence the coastal rainfall.

Our analysis therefore suggests a very marginal role, ifany, for the heat low in terms of its impact on interannualvariability of the coastal phase. It is perhaps surprising that,given the intensity of the heat low SMC in spring (Zhanget al., 2008), it has little impact on any aspect of the coastalphase. Our conclusion from this is that while the heat-low-associated SMC does interact with the coastal regionduring spring, its impact is overwhelmed by the impact ofthe adjacent ocean. In contrast, the correlation hints to apossible impact of the coastal rainfall on the weakening ofthe heat low in June–July.

8.2. Impact of the heat low on the Sahelian onset

Although not central to the aim of this paper, we brieflyexplore the potential role of the heat low on the Sahelianonset. Previous studies, and composite analysis (Figure 12)suggested that the onset of the SHL observed at the endof the coastal phase may be related to the Sahelian onset.Indeed, this coincidence has been observed in a modellingstudy based on one single year by Ramel et al. (2006) andin climatological data analysis by Lavaysse et al. (2009).The date of the SHL onset tSHL is obtained here as thedate when its location crosses the 20◦N latitude for thatyear. With this definition, the SHL onset occurs on 24June (std. dev. = 10.2 days), which is 12 days prior to themean Sahelian onset. Note that in Lavaysse et al. (2009)the heat low shift is on 20 June ±9 days. The differentdate found here might be related to the different data (andthus different resolution) and period used compared tothat of Lavaysse et al. (2009) who used ERA-40 data forthe period 1984–2001. Figure 13 highlights the interannualevolution of the Sahelian onset together with the SHL onset.Despite the close average dates, the two time series clearlyshow no coherence when considering individual years (withr = 0.127). However, note that lag-correlation of Sahelain



tC tEC tS

Figure 12. As in Figure 5 except that the grey shading is mean sea-level pressure (MSLP, hPa) from ERA-interim for the period 1989–2009. The whitecontours represent the SST 301 K isotherm and the location of the Saharan heat low based on the lowest MSLP, respectively.

onset shows highest correlation with the heat low intensityin May (0.36 significant at the 90% level) and with the heatlow location in August (0.58 significant at the 99% level).Therefore we conclude that the SHL onset does not appearto be important for the Sahelian onset.

In addition to the timing of the Sahelian onset it has alsobeen suggested that the heat low may have an impact onthe rainfall intensity during the Sahelian rainy season. Therelationship between the rainfall amounts in the Sahel andthe location of the heat low has been established in Lele andLamb (2010) and in Biasutti et al. (2009). Consistent withthis, our analysis indicates a strong lag-correlation betweenthe location of the heat low and the poleward extent of therain band (not shown). This combined with previous studiesconfirms the relationship between the poleward extent ofthe rain band and the location of the heat low as suggestedin Thorncroft et al. (2011). Moreover the lag-correlationshows that the poleward extent of the rainfall slightly leadsthe location of the heat low (by 1 pentad), suggesting thatthe rainfall can influence the location of the heat low.

9. Summary and concluding remarks

This paper highlights the nature and variability of the coastalphase of the West African monsoon, which has been littlestudied. A comparison of three different rainfall datasetsraised some issues regarding their disagreement, especiallyover the ocean due to the lack of in situ observations. Becauseof this, and also because of intermittent rainfall occurringat the coast even in winter when the climatological peak isstill over the ocean, we define the coastal onset in terms ofrapid reduction of the meridional width of the rain bandover the ocean due to the northward retreat of its southernboundary. The mean coastal onset thus defined occurs on11 May with a standard deviation of 2.9 pentads.

The coastal phase exhibits marked interannual variability;the onset date varies between 6 April and 10 June and, asexpected, is intimately linked to variability of the equatorialcold tongue onset. The coastal onset tends to occur about10 days after the cold tongue onset. It would be crucial forprediction purposes to be able to quantify the impact of SSTvariability on the timing of the coastal onset.

The length of the coastal phase also varies markedly,ranging between 3 and 14 pentads. We showed that the

length of the coastal phase is more related to the date itstarts rather than to the date it ends. We also showed thatthe date for the end of the coastal phase is influenced bythe cooling of the coastal water north of the equatorial coldtongue. A relevant threshold for the coastal rainfall demiseand potentially for the Sahelian onset, appears to be 301 K.Again the predictability of the timing of different rainfallphases depends on SST variability.

The variability in coastal rainfall rates and rainfall totalsduring the wettest period of the year (mid-May–mid-June)was also discussed. Consistent with previous work on theWest African region (e.g., Ward 1998), variability in totalrainfall is positively correlated with SSTs in the Atlanticand negatively correlated with SSTs in the West Pacific.More importantly significant correlations between rainfallamounts and southeast Atlantic SSTs suggest predictabilitywith several months lead time.

While SSTs are clearly important for determining thenature and variability of the coastal phase, in contrast, therole of the heat low remains unclear. There is no evidence inour analysis for the influence of the heat low on the coastalphase, suggesting that the cold tongue and coastal water aredominant factors for influencing rainfall in spring.

In this study the Sahelian onset is also analysed in relationto the coastal phase. The mean Sahelian onset occurs about2 months after the coastal onset and 2 weeks after the end ofthe coastal phase. The correlation of 0.53 between the coastalonset and the Sahelian onset suggests that both onsets arestrongly controlled by the cold tongue development. Thecorrelation of 0.63 between the coastal SST cooling in Juneand the Sahelian onset also suggests prediction potential.

To fully understand the mechanism of the coastal rainfall,further study is needed to explore the causes of the coldtongue variability. Different origins of the cold tongueformation can be found in the literature, such as southerlycross-equatorial acceleration due to surface temperaturegradients (Gu and Adler, 2004); the westward propagationof easterly wind enhancement (Okumua and Xie 2004); theinitiation of the cold tongue results from the intensificationof the southeastern trades associated with the Saint Helenaanticyclone (Caniaux et al., 2011); and the SMC (Thorncroftet al., 2011). There is a clear need for more in situobservations of the cold tongue evolution between winterand summer and, given the disagreements between available



Figure 13. Time series of the Sahelian onset tS (solid) and SHL onset tSHL (dash).

rainfall estimates over the ocean, this should ideally includecontemporaneous measurements of rainfall.

Finally, the strong relationship between the SSTs in theequatorial region and the coastal onset and between coastalSSTs and the end of the coastal phase suggests that these SSTsshould be closely monitored in real time to supplement theprediction of the WAM onsets. More than this, our analysishas also indicated that equatorial SST anomalies persist4 months ahead of the wettest season, strongly suggestingthat the coastal rainfall intensity is predictable with a notablelead time.

Acknowledgements

We are thankful to Robert Adler, John Janowiak, EdwardZipser, Pingping Xie, George Huffman and Chuntao Liu fora very constructive discussion on the rainfall estimate issue.We thank the reviewers whose comments greatly helpedto improve the manuscript. This study was supportedby NOAA Grant NA05OAR4311130 and by NASA grantNNX09AI69G.

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Guinean coastal rainfall of the West African Monsoon

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