Six oceanographic cruises (Figure 1) named “Etude de la circulation océanique et de savariabilitédans le Golfe de GuinEE” (EGEE) project [12] was conducted (2 cruises/year, see Table 1) from June 2005 to September 2007 in the eastern equatorial Atlantic. This project was the oceanic component of the African Monsoon Multidisciplinary Analyses (AMMA) program. The yearly cruises have consisted of two legs starting and ending in Cotonou (Benin, West Africa) during June and September [13] .

A total of 195 samples surface seawater was collected for Dissolved Inorganic Carbon (DIC) and TA Total Alkalinity (TA) analyses. Samples were poisoned with a saturated HgCl₂ solution to stop biological activities. DIC and TA were measured using potentiometric titration that derived from the method developed by Edmond [14] with a closed cell. The equivalent points were estimated using a non?linear regression method [15] while Certified Reference Materials (CRMs) provided by Prof. A. Dickson (Scripps Institution of Oceanography, San Diego, USA) was used for calibration [16] . The accuracy of DIC and TA in this work reaches ±2 µmol・kg⁻¹.

The fCO₂ measured during EGEE 3 was used to determine the best set of dissociation constants for the calculation of fCO₂ at stations where DIC and TA were recorded. It was also used to validate fCO₂ derived from our extrapolated DIC and TA along the EGEE 3 cruise track. Then, EGEE data were supplemented by data from FOCAL 4, 6 and 8, CITHER 1, and EQUALANT 99 cruises provided by several authors [1] - [3] [17] . These oceanographic cruises were carried out from the 1980s to the 2000s in the eastern equatorial Atlantic (Table 2).

The precipitation dataset was extracted from the Global Precipitation Climatology Project (GPCP) [18] on a 1˚ × 1˚ regular grid over the Tropical Atlantic Ocean for the period 2005-2007. The GPCP dataset was composed by a combination of in situ observations with satellite microwave and infrared measurements.

The air-sea CO₂ fluxes (F) is calculated using Equation (1):

Whereas

(in cm・h⁻¹) is the CO₂ gas transfer velocity, is the difference between seawater fCO₂ and atmospheric fCO₂, and is the solubility of CO₂ in seawater (in mol・L⁻¹・atm⁻¹) with regard to Sea Surface Salinity (SSS) and Sea Surface Temperature (SST) [19] ; the weekly mean CO₂ exchange coefficient in the 6˚N - 10˚S and 10˚E - 10˚W is derived from QuikSCAT wind speeds [20] using the quadratic relationship of [21] at a spatial resolution of 1˚ × 1˚. represents the wind speed at 10 m above sea surface and Sc is the Schmidt number [21] [22] . The atmospheric fCO₂ was calculated from the monthly CO₂ molar fraction recorded at the Ascension Island at 7.92˚S - 14.42˚W from 2005 to 2007 [23] with a mean value of about 373.2 ± 0.75 µatm.

In addition, the fluorescence was measured using the CTD sensor during the EGEE cruises while chlorophyll a was sampled only during EGEE 3 cruise and analyzed according to the HPLC standard technique [24] . However,

fluorescence data will be used in this study instead of chlorophyll a since it is available in all cruises.

The following paragraphs outline the chemistry of carbon dioxide in the ocean. When it is dissolved in the seawater, the carbonate system can be described by the Equations (2), (3), (4) and (5).

where, K₁ and K₂ represent stoichiometric equilibrium constants for the description of the carbonate system in the seawater. The different sums of the dissolved forms (i.e. CO₂, , and) give the DIC and the carbonate alkalinity (CA) following Equations (4) and (5):

In Equation (5), the carbonate ion is counted twice because it has a double negative charges. The CA is the part of the TA, which includes boron compounds and more:

Unfortunately, the concentrations of the individual species of the carbon dioxide system in solution cannot be measured directly. The Equations (2), (3), (4) and (5) have six unknown variables (i.e. CO₂, , , [H⁺], [DIC] and [CA]) and only four variables (i.e. [CO₂], [H⁺], [DIC] and [TA]) can be measured directly. The system is determined when two variables at least are known and all the parameters can then be calculated with ancillary information. That allows to get a complete description of the carbon dioxide system in seawater [15] [25] .

Moreover, oceanic fCO₂ was estimated from TA and DIC using the different dissociation constants. In order to choose the best dissociation constants, measured fCO₂ [26] are compared with the values calculated from TA and DIC during the EGEE 3 cruise. The dissociation constants of [27] refitted by [28] , present small values of the root mean square error (~7 µatm) and mean bias error (~ +4 µatm). Thereafter, they will be used to calculate the oceanic fCO₂.

For each year, the relationships established by [13] were used to estimate fCO₂ of FOCAL 6, CITHER 1 and EQUALANT 99 cruises (Table 3). These authors found that fCO₂ ranged between 330 µatm and 420 µatm in the GG even if it could exist extreme values associated to the equatorial upwelling [29] . The fugacity of CO₂ was estimated during three cruises carried out in boreal winter (January-February 1984 and January-March 1993) and in boreal summer (July-August 1999) using these relationships. Although measured fCO₂ data were few, the values of the fugacity of CO₂ were well reproduced [13] and the mean estimated fCO₂ was very close to the mean measured fCO₂ during these three cruises.

The annual and monthly fluxes of CO₂ in GG (6˚N - 10˚S; 10˚E - 10˚W) derived from the [13] relationships were calculated at a 1˚ × 1˚ spatial resolution using the SST data of the TMI (Tropical Rainfall Measuring Mission Microwave Imager). The climatological and others mean air-sea CO₂ fluxes were calculated using the CO₂ gas transfer velocity relationship of [21] . The gridded mean air-sea CO₂ fluxes (mmol・m⁻²・d⁻¹) for each day was computed by multiplying daily mean values of and. The daily mean value of K was obtained

by dividing its weekly (resp. monthly) mean value by seven (resp. by 30 days). [11] noted a low seasonal variability of the mean values of K in the tropical Atlantic belt (14˚N - 14˚S). The mean zonal flux was computed by multiplying the gridded air-sea CO₂ fluxes by the surface representing each grid. Furthermore, the spatial mean air-sea CO₂ fluxes (in Pg-C y⁻¹) was the sum of the air-sea CO₂ fluxes of each grid.

Figure 5 shows the seasonal cycle of the air-sea CO₂ fluxes in both regions (6˚N - 0˚ and 0˚ - 10˚S) in 2006, since the patterns were quite similar in 2005 and 2007 (not shown). Low values (resp. high values) were noted in January - May (resp. July - October). In the 6˚N - 0˚ area (resp. 0˚ - 10˚S area), the mean air-sea CO₂ fluxes ranges between −0.16 ± 0.66 mmol・m⁻²・d⁻¹ (resp. 0.78 ± 0.89 mmol・m⁻²・d⁻¹) in March to 3.76 ± 1.12 mmol・m⁻²・d⁻¹ (resp. 4.76 ± 1.90 mmol・m⁻²・d⁻¹) in August. When looking at the two areas, a north-south gradient was observed in the distribution of the air-sea CO₂ fluxes in the Gulf of Guinea. Not surprisingly the ocean CO₂ source was higher in south of the Equator than in the north. This gradient was due to the GC, which allowed transporting eastward−unsalted waters due to precipitations, and contributed to the dilution and the decrease of CO₂ in the northern region. Furthermore, at the south of the Equator, upwelled waters that made the air-sea CO₂ fluxes mainly driven by the oceanic fCO₂ supply the CO₂.

The Table 4 shows the air-sea CO₂ fluxes during the EGEE cruises. Each year has been divided into two semesters to take into account the non-upwelling (January - June) and the upwelling (included in July - December) periods in the GG. The air-sea CO₂ fluxes was higher (~2.5 times) in July - December in January - June for all the three years (2005, 2006 and 2007). The upwelled water was responsible for the high CO₂ out gassing from July to September, which increased the air-sea CO₂ fluxes. Low year-to-year variability of CO₂ fluxes was also observed for each semester. However, a slight drop appeared in 2006 compared to 2005 and 2007. This could be explained by the relative high SST in 2006, which was due to the time shift in the development of the cold tongue and the intensity of the upwelling [39] . Moreover, [26] found that the low concentration of CO₂ in 2006 could be caused by the transport of water, which was in contact with the atmosphere long time enough to come close to equilibrium.

To remove the seasonal and interannual variability, the anomalies of the air-sea CO₂ fluxes were computed

and standardized for 2005, 2006 and 2007. During these periods, no relationship was found between ENSO index and the air-sea CO₂ fluxes (not shown). The Hovmöller diagram (Figure 6) shows a slight variation of the CO₂ fluxes in the south of the Equator. At 5˚S, the ocean behavior was not homogenous because of the local drop of the SST (~26˚C), which could create a sink of the CO₂ fluxes. At the northern part of the Equator, a significant year-to-year variability of the costal upwelling appeared during the monsoon period. The coldest months were observed in 2005 and consequently imply a high value (>2 mmo・m⁻²・d⁻¹) of the air-sea CO₂ fluxes.

This sub-section highlights the differences between the CO₂ fluxes estimated during EGEE 3 in 2006, the climatology of [40] , and the FOCAL 4 (F4, see Table 2) and FOCAL 8 (F8, see Table 2) cruises. EGEE 3 is chosen because it is the cruise during which fCO₂ were measured underway (see Figure 1(c)). The climatology of [40] was referenced to the year 2000 and has been built by averaging the CO₂ fluxes in our study area previously defined. The mean climatological fluxes in the northern (resp. in the southern) region reached 0.56 ± 0.24 mmol・m⁻²・d⁻¹ (resp. 0.74 ± 0.33 mmol・m⁻²・d⁻¹). The weak difference between the northern and the southern fluxes (~0.18 mmol・m⁻²・d⁻¹) implied that the north-south gradient was not well reproduced by the climatology of [40] . This could be due to the coarse resolution (4˚ latitude × 5˚ longitude) used in the climatology, which

tended to smooth the difference between northern and southern waters. In the case of F4 and F8, the comparison is made by averaging the CO₂ fluxes between 5˚N - 5˚S and along 4˚W where data were available.

From June to December, the climatology of [40] (Figure 7), F4 and F8 mean fluxes reached 0.95 ± 0.5 mmol・m⁻²・d⁻¹, 0.56 ± 1.33 mmol・m⁻²・d⁻¹ and 0.82 ± 1.96 mmol・m⁻²・d⁻¹ respectively. Our estimated fluxes by using EGEE 3 (see Figure 7) were 2.30 ± 1.55 mmol・m⁻²・d⁻¹. It was two times higher than that of the climatology of Takahashi et al. [40] and four times higher than that of the F4 cruise and finally, three times higher than that of F8. This difference of fluxes values could be explained by the local phenomena that were not taken into account by the coarse horizontal resolution of the climatology of [40] . During January - May, the mean fluxes from the climatology of [40] were almost constant and remained on average around 0.5 mmol・m⁻²・d⁻¹. From June to December, the climatological mean fluxes had the same value (~1 mmol・m⁻²・d⁻¹), then falls to 0.5 mmol・m⁻²・d⁻¹. Over the same period, our estimated fluxes were higher than 1 mmol・m⁻²・d⁻¹.

During F4 and F8 cruises, the fluxes reached −0.25 ± 0.79 mmol・m⁻²・d⁻¹ and −0.16 ± 1.27 mmol・m⁻²・d⁻¹ (resp. 1.28 ± 1.32 mmol・m⁻²・d⁻¹ and 1.71 ± 2.09 mmol・m⁻²・d⁻¹) in the north (resp. in the south). The weak values in the north implied an equilibrium state with the atmosphere, while the high values in the south indicated that the region was a source. The climatology of [40] and the EGEE 3 estimated fluxes during the same months as F4 and F8, showed positive values indicating that both regions were sources. The evolution of the northern region from an equilibrium state (FOCAL, in 1983-1984) to a source area (EGEE 3 cruises, in 2006) is in agreement with the works of [41] . This author showed the increase of oceanic fCO₂ faster than the atmospheric CO₂ in the southern ocean. This fact implied the decrease of the oceanic CO₂ sink.