Using TRMM precipitation data, we calculate the rainy season onsets and ends (RSOs and RSEs, respectively) in the equatorial Congo Basin similar to the approach used in Chakraborty et al. (2021), originally developed by Li and Fu (2004), for RSOs and RSEs in the Amazon. This definition is designed to capture persistently high rainfall during the rainy season and persistent low rainfall in the dry season using the annual mean rainfall as the threshold. We first calculate the climatological pentad (5 d mean) rainfall for each year over our domain, then determine the RSOs and RSEs in the following steps.

For the spring rainy season, the onset is defined as the first pentad of each year that meets the following criterion: four out of the previous six pentads are less than the climatological annual mean, and four out of the following six pentads are more than the climatological annual mean. Likewise, the end of the spring rainy season is defined as the first pentad of each year that meets the following criterion: four out of the previous six pentads are more than the climatological annual mean, and four out of the following six pentads are less than the climatological annual mean.

For the fall rainy season, the onset is defined as the first pentad of each year that meets the following criterion: five out of the previous eight pentads are less than the climatological annual mean, and five out of the following eight pentads are more than the climatological annual mean. The fall RSE is defined as the first pentad of each year that meets the following criterion: six out of the previous eight pentads are more than the climatological annual mean, and six out of the following eight pentads are less than the climatological annual mean. We note that this criterion for the fall rainy seasons differs from that of the spring rainy season because the latter is shorter and weaker than that of the fall rainy season (Fig. 2). Therefore, a shorter period and less persistent threshold was needed to best capture the starts and ends of persistently high and low rainfall for the spring RSO and RSE.

Figure 2 shows the RSOs and RSEs for the spring and fall rainy seasons between the years 2000–2020. We excluded the year 2000 when calculating the spring RSO because we needed to consider the ends of the prior years in our RSO calculations. Similarly, we excluded the year 2020 when calculating the fall RSE because we needed to consider the start of the next years in our RSE calculations. We additionally excluded from our following analyses any years during which a RSO or RSE was unable to be identified, or if the length of a rainy season was less than 5 pentads. The spring RSO ranges from the 6th to 14th pentad (with Days 1–5 defined as Pentad “0”), corresponding to early-February to mid-March. On average, the spring RSO occurs on 20–24 February. The spring RSE ranges from the 25th–32nd pentads, corresponding to mid-May to mid-June. On average, the spring RSE occurs on 21–25 May. The fall RSO ranges from the 40th–50th pentads, corresponding to mid-July to early-September. On average, the fall RSO occurs on 18–23 August. The fall RSE ranges from the 64th–71st pentads, corresponding to mid-November to late-December. On average, the fall RSE occurs on 2–6 December. We show 15 pentads prior to the spring RSO and 19 pentads prior to the fall RSO as these periods capture the key changes in processes that drive the dry to rainy season transition as in Li and Fu (2004) and Wright et al. (2017).

We calculate the area-mean equivalent potential temperature (θe) from AIRS gridded data at daily time resolution to examine atmospheric instability during and after the transition to the rainy season. We calculate θe using the following equation from Bolton (1980): 1θe=Tk1000p0.28541-0.28×10-3⋅r×exp⁡3.376TL-0.00254×r1+0.81×10-3r Where Tk is the absolute temperature (K) at pressure level p, p is pressure (hPa), r is the mixing ratio (g kg⁻¹), TL is the absolute temperature at the lifting condensation level (K) calculated by the following: 2TL=28403.5ln⁡Tk0-ln⁡e-4.805+55 Where e=p⋅r622+r is the water vapor pressure, and here Tk0 is the absolute temperature at the surface, defined here as 925 hPa as the surface of the equatorial region is above sea level. We use the AIRS mass mixing ratio, temperature, and pressure in these calculations. We similarly calculate saturated equivalent potential temperature θes by letting r=rs, where rs is the saturated mixing ratio. We calculate rs by the following:

We calculate es as in Bolton (1980): 3es=6.112exp⁡17.67TT+243.5 Where T is the temperature in Celsius (°C). We calculate saturated vapor pressure by: 4rs=εesp-es Where ε=0.622 is ratio of molar masses of vapor and dry air, and es is the saturation vapor pressure, calculated as in Bolton (1980): 5esT=6.112exp⁡17.67TT+243.5 We also calculate moist static energy (MSE) where MSE is represented by: 6MSE=cpT+Lq+gz. Where cp is the specific heat capacity at constant pressure, L is the latent heat of evaporation and g is gravity.

Finally, we examine changes in moisture in the equatorial Congo with the following. We use vertically integrated moisture flux MF=∫pspTOAq⋅Vdp/g, where q is the specific humidity, and V=(u,v), is the horizontal wind, with u the zonal, and v the meridional wind components, respectively. Ps denotes the surface pressure and PTOA denotes the pressure of the top of the atmosphere. In particular, we calculate the net zonal and meridional MF as follows: We take the zonal vertically integrated MFzonal=∫pspTOAq⋅udp/g across the western and eastern boundaries of the equatorial Congo basin, and the meridional vertically integrated MFmeridional=∫pspTOAq⋅vdp/g across its northern and southern boundaries. For simplicity, we choose the northern boundary along 2° N and between 14–30° E, and the southern boundary along 2° S and between 14–30° E. The western boundary is along 14° E and between 2° S–2° N and the eastern boundary is along 30° E and between 2° S–2° N. To be able to compare to vertically integrated MF convergence, we calculate the length between the western/eastern boundaries and the northern/southern boundaries, respectively, and integrate average MFzonal and MFmeridional across these boundaries as in Satyamurty et al. (2013): 7QE=∫MFzonal,EdyQW=∫MFzonal,WdyQN=∫MFmeridional,NdxQS=∫MFmeridional,Sdx Where MFzonal,E is MFzonal integrated along the eastern boundary, MFzonal,W is MFzonal integrated along the western boundary, MFmeridional,N is MFmeridional integrated along the northern boundary, and MFmeridional,S is MFmeridional integrated along the southern boundary. y is the length of the northern/southern boundaries, and x is the length of the western/eastern boundaries. To calculate the net zonal vertically integrated MF, we calculate Qzonal=QW-QE. To calculate the net meridional vertically integrated MF, we calculate Qmeridional=QS-QN. This ensures that positive values of Qzonal and Qmeridional means net convergence, and negative means net divergence.

We calculate the first component of the moisture flux convergence term (named the moisture convergence term, q⋅∇V) and moisture transport vectors (q⋅V) averaged between different pressure levels to represent the lower and middle tropospheric moisture transport. We do not examine the evolution of moisture advection term (V⋅∇q) as it is small compared to that of the moisture convergence term (Cook and Vizy, 2022; Figs. S1–S4 in the Supplement).

Due to a lack of reliable ET measurements, we use the deuterium content of water (HDO) derived from satellite measurements and an isotope mixing model to estimate the fractional contribution of ET to atmospheric moisture. HDO is expressed as the relative ratio of the number of HDO molecules to the total number of H2O molecules in parts per thousand (‰) relative to the isotopic composition of ocean water as shown below: 8δD=1000×R-RstdRstd Where R is of the ratio of HDO molecules to the total number of H2O molecules and Rstd is the corresponding ratio in a reference standard, taken here to be the Vienna Standard Mean Ocean Water: Rstd=3.11×10-4 (e.g. Wright et al., 2017, and references therein). The isotopic composition of water vapor in the free troposphere is due to a mixture of air parcels originating from different sources (Galewsky, 2018; Galewsky and Hurley, 2010). We can use the isotopic composition of an air mass to trace its source to either vegetation or ocean because δD values from ocean evaporation are distinctively different from those by rainforest ET. Atmospheric moisture coming from transpiration will generally be more isotopically enriched (or heavier) than evaporation because lighter isotopes preferentially evaporate (Risi et al., 2020; Tremoy et al., 2014; Worden et al., 2007). However, observed values of free-tropospheric deuterium content also depend on atmospheric processes, such as the type of convection present, described further in Galewsky et al. (2016). This is because they vary by the first order with changes in specific humidity (Bailey et al., 2017).

Therefore, we compare the observed isotopic composition of water vapor in the free troposphere to two models, a mixing model and a Rayleigh model, to identify which air parcels are likely influenced by land (i.e., transpiration) versus ocean (i.e., evaporated water; Noone, 2012). A mixing model describes what happens to a mixture of two air masses with different water vapor isotopic compositions. Meanwhile, under the Rayleigh distillation model, as an air mass moves upward (or towards cooler conditions), condensate enriched with heavy isotope is completely removed immediately after it forms under the assumption of pseudo adiabatic process (Galewsky et al., 2016; Wright et al., 2017). A further description of these models can be found in the Supplement and in Worden et al. (2021b).

We examine the seasonal evolution of key dynamic features that transport moisture and energy within and around the Congo Basin. They are discussed further in several review papers: Nicholson (2022), Pokam et al. (2026).

Low Level Westerlies (LLWs). These winds transport moisture from the Atlantic Ocean into Central Africa, with varying seasonal intensity (Pokam et al., 2012, 2014; Vondou et al., 2010)

The first part of the spring transition, i.e., the early-transition period, starts with a turning point from strengthening to weakening vertically integrated MF divergence (Fig. 3a) 10 pentads (50 d) before the spring RSO. To explain this weakening, we evaluate Qzonal and Qmeridional (Fig. 3c). Negative values indicate net moisture leaving the basin for both Qzonal and Qmeridional. Variations in Qzonal during the early transition period are driven by increases in low-level moisture transport from the Atlantic Ocean, by the lower branch of the Congo Basin Cell, which goes through the equatorial Congo, towards the West African Heat Low (Lavaysse et al., 2009) in the northern Congo (Figs. 4a, b; B1a, e). This is driven by increasing near-surface temperature gradients between the Atlantic Ocean and West African Heat Low (Figs. S5a, b; S6). These increases in low-level moisture transport, combined with increases in the second half of the transition period of mid-level moisture transport into the basin across its eastern boundary, acts against zonal moisture leaving the basin at mid-levels, across the western boundary (Fig. B1a, b), but is not enough to lead to Qzonal convergence (Fig. 3c). Meridionally, mid-level moisture entering the basin across the northern boundary does not change significantly, although low-level moisture leaving the basin across the northern boundary increases in response to increased low-level moisture transport towards the West African Heat Low (Figs. 4a, b; B1c, f). Across the southern boundary, moisture transport entering the basin at low levels increases, while moisture transport leaving the basin at mid-levels does not significantly change (Fig. B1d, f). Ultimately, it is changes in low-level meridional moisture transport that drive changes in Qmeridional divergence (Fig. 3c).

Meanwhile, the contribution of ET to atmospheric moisture increases somewhat compared to the pre-transition (fmixpre=0.611±0.093; fmixearly=0.772±0.12), and is clearly the dominant contributor to moisture (Fig. 5a, b). This is consistent with Worden et al. (2021b) for a similar domain. Therefore, ET is important for maintaining high background moisture while weakening in vertically integrated MF divergence initiates the transition to the spring RSO.

Furthermore, the atmosphere appears thermodynamically primed for deep convection, with large convective available potential energy (CAPE) and conditional instability as shown by dθedp<0 (Fig. 6c, d). However, weak ascent or even subsidence still exists above 650 hPa (Fig. 6e). Additionally, the level of free convection (LFC; Fig. 7a), and therefore convective inhibition energy (CIN, Fig. 6c) do not change significantly throughout the early transition period. This indicates that the atmospheric conditions, especially the dynamic conditions, are still not ready for large-scale increases of deep convection and subsequently, rainfall.

We explain this as following: with weakening atmospheric vertically integrated MF divergence, boundary layer specific humidity begins to increase (Fig. 7d), due to increases in moisture transport from the equatorial Atlantic Ocean into the equatorial Congo (Fig. 4a, b). The increase of boundary layer moisture, rather than increase of near-surface temperature, drives boundary layer increases in MSE (Fig. 7b–d). However, at the mid-levels, MSE initially decreases and then increases in mid-level specific humidity (Fig. 7b), which is driven by mid-level moisture divergence. Therefore, while the boundary layer MSE is favourable for deep convection, additional changes are needed to support deep convection in the mid-layers of the atmosphere.

The late-transition period starts with a rapid increase in atmospheric column water vapor (CWV; Fig. 3b), a key condition for deep convection (Bretherton et al., 2004; Schiro et al., 2018). This is accompanied by rapid increases in precipitation and vertically integrated MF convergence (Fig. 3a). ET does not change significantly (fmixlate=0.778±0.139; Fig. 5c), revealing the importance of large-scale circulation changes in altering atmospheric moisture over the equatorial Congo prior to the spring rainy season.

Qzonal and Qmeridional rapidly switch from divergence towards convergence (Fig. 3c). Zonally, further increases in near surface temperatures over the West African Heat Low drive stronger LLWs from the Atlantic and across the equatorial region along the way (Figs. B1a, S5c, S6, 4c). This, together with increases in mid-level (above 800 hPa) zonal moisture transport across the eastern boundary (Fig. B1b) compensate the increases in mid-level moisture transport leaving the basin across the western boundary (Fig. B1a). The latter is likely due to the formation of the AEJ-N (Fig. 6f). Therefore, Qzonal becomes convergent, although weakly so (Fig. 3c). Meridionally, moisture transport entering the basin at low levels through the southern boundary increases (Fig. B1c), bringing with it moist air from the southern Congo (Fig. S7c), which is experiencing its rainy season during this time (Worden and Fu, 2025). However, moisture transport leaving the basin across its northern boundary towards the West African Heat Low increases as well (Fig. B1d), leading to net low level meridional moisture transport out of the basin (Fig. 4c). In contrast, moisture convergence at mid-levels increases (Fig. 4f) as mid-level moisture leaving the basin across the southern boundary decreases (Fig. B1d, f). Altogether, this leads to Qmeridional convergence (Fig. 3c).

These changes in atmospheric moisture transport increase boundary layer moisture, as indicated by increasing relative humidity (Fig. 8a, b), but subsidence from the surface to ∼850 hPa (Figs. 6e, 8a, b) between 10–26° E (west of the East African Rift) likely prevents this moisture from being uplifted within large parts of the interior equatorial region. Instead increases in westerly wind within the interior of the region pushes this moisture towards the East African Rift (Fig. 8), where orographic lifting forces this moist air into the free troposphere above 800 hPa. Additionally, moisture is uplifted by the ascending branch of the Congo Basin Cell over the equatorial coastline (highlands of Gabon and Cameroon) at the western edge of the equatorial Congo (Fig. 8). Therefore, this results in increasing mid-level moisture and mid-level MSE while decreasing LFC and CIN (Figs. 6c, 7a, b, d). Upward vertical velocity increases in strength starting in the mid-levels and reaches 300 hPa for the first time (Fig. 6e). Meanwhile, shear increases with the increasing strength of the AEJ-N (Fig. 6f), thus creating conditions favourable for deep convection and the start of the spring rainy season. This is indicated by increases in high clouds, increases in cloud top height, and decreases in cloud top temperature (Fig. 6a, b).

As in the transition to the spring RSO, the early-transition period starts 14 pentads before the fall RSO with weakening in vertically integrated MF divergence over the equatorial Congo (Fig. 9a), indicated reduced moisture export out of the region. This is mainly due to decreases in Qzonal divergence (Fig. 9c), which becomes convergent about halfway through the early transition period. This acts against the change of Qmeridional, which starts convergent and switches to divergence during the transition period (Fig. 9c).

Changes in initial Qzonal divergence are primarily due to increases in low-level moisture transport into the basin, across the western boundary (Figs. 10a, b; B2a), which combined with steady moisture transport between 800–600 hPa into the basin across the eastern boundary (Fig. B2b), acts against weakly decreasing moisture transport out of the western boundary above 800 hPa (Fig. B2a) via the return branch of the Congo Basin Cell (Longandjo and Rouault, 2020). Increases in low-level atmospheric moisture across the western boundary are driven by the near-surface temperature gradient between the Atlantic Ocean and the equatorial and southern Congo (Figs. S8a, b; S9), which strengthens the lower branch of the Congo Basin Cell (Longandjo and Rouault, 2020).

Qmeridional convergence initially weakens towards zero (Fig. 9c). Across the southern boundary, moisture transport leaving the basin at low levels increases (Fig. B2d, f), moving towards the Congo Air Boundary (aqua contours in Fig. 10b denote sharp surface humidity gradients that show the rough location of the Congo Air Boundary; Howard and Washington, 2019) and associated low-level convergence band (Fig. 10b). Additionally, moisture transport into the equatorial Congo across the southern boundary, above 850 hPa, decreases in strength (Fig. 10d, e). This acts against increases in low-level moisture entering the basin across the northern boundary (Fig. B2d). Together, this causes the Qmeridional divergence by the end of the transition period.

Meanwhile, the contribution of ET to atmospheric moisture does not change significantly: fmix=0.612±0.082 during the “before” period, while fmix=0.665±0.168 during the early transition (Fig. 11a, b). Therefore, it is the increases of low-level moisture from the Atlantic Ocean that initiate the early transition period prior to the fall RSO.

Despite an overall unstable lapse rate (dθedp<0), the thermodynamic condition of tropospheric column does not yet support deep convection. In the first half of the early transition, decreases in low-level atmospheric moisture (Fig. 13d) weakens negative dθedp, coinciding with rising CIN and LFC, as well as decreasing CAPE and low and mid-level MSE (Figs. 12c, d; 13b, c). Changes in MSE at both levels are mainly due to changes in Lvq rather than CpT (Fig. 13b, c, d). In the second half of the early transition, low-level moisture begins to increase, corresponding to increases in θe in the low-levels and strengthening negative (destabilizing) dθedp (Figs. 12c, 13d). This corresponds to the LFC stabilizing (no longer consistently increasing) between -7 to -2 pentads prior to the RSO (Fig. 13a) and decreases in CIN (Fig. 12c). Additionally, the vertical velocity is near zero above 600 hPa (Fig. 12e). Therefore, atmospheric dynamic conditions are not yet ready to facilitate increase of deep convection, consistent with relatively low and constant mid-high and high cloud covers (Fig. 12a), cloud top temperatures, and cloud top heights (Fig. 12b).

The late transition period to the fall rainy season starts with sharp increases in precipitation, vertically integrated MF convergence, and CWV (Fig. 9a, b). Meanwhile, the contribution of ET to atmospheric moisture (f=0.613±0.194) does not change significantly (Fig. 11c). Qzonal convergence sharply increases, while Qmeridional divergence weakens to near zero (Fig. 9c).

Zonally, the speed of low-level westerly wind that carry moisture from the Atlantic Ocean further increases (Fig. 10b, c). This is due to further increases in the zonal temperature gradient as the Atlantic cold tongue develops (Figs. S8b, c; S9; Caniaux et al., 2011), which increases the strength of the lower branch of the Congo Basin Cell (Longandjo and Rouault, 2020). Meanwhile, the strength of the mid-level zonal moisture transport does not change as much compared to its low-level counterpart, indicating that the sharp increase in Qzonal convergence is driven by increases low-level moisture transport (Figs. 9c; B2a, b, e).

Meridionally, in the low-levels, moisture transport out of the southern boundary into the southern Congo continues to increase as the low-level winds move towards the Congo Air Boundary (Fig. 10b, c). However, this is compensated by moisture transport into the equatorial Congo across the southern boundary above 850 hPa (Figs. 10e, f, h, g; B2d). Across the northern boundary, low level moisture transported into the basin increases, acting against mid level moisture leaving the basin (Fig. B2c). Therefore, Qmeridional divergence decreases (Fig. 9c).

At the low-levels, atmospheric moisture increases rapidly due to increases in westerly winds (Fig. 14) and further destabilizes (strengthens negative dθedp; Fig. 12d). The increased atmospheric moisture lifts up into the mid-troposphere above 800 hPa via the ascending branch of the Congo Basin Cell (located between 10–20° E) as well as orographically along the slope of the East African Rift (Fig. 14a, b), Therefore, mid-tropospheric relative humidity increases (Fig. 14b) despite subsidence in the low troposphere in the interior of the basin (Figs. 12e; 14b). As the mid-troposphere moistens, the thermodynamic atmospheric condition becomes favourable for deep convection rapidly, as shown by decreases in CIN (Fig. 12c) and LFC (Fig. 13a), and increases in CAPE (Fig. 12c).

Meanwhile, despite the absence of the AEJ-N or AEJ-S over the equatorial region (Fig. 12f), the return branch of the Congo Basin Cell provides the vertical wind shear (Fig. 12f, on average within the equatorial region 8.30 m s⁻¹ during the late transition period prior to the fall RSO compared to 8.55 m s⁻¹ during the late transition period prior to the spring RSO) needed for favourable dynamic conditions for mesoscale convective systems (MCSs). These conditions enable rapid increase of P (Fig. 9a), and lead to the fall RSO as indicated by increases in the strength of vertical velocity from the lower atmosphere to 300 hPa (Fig. 12e), sharp increases (decreases) in high/mid-high (low) clouds, sharp increases in cloud top height, and sharp decreases in cloud top temperatures (Fig. 12a, b).

In summary, the transition from the boreal summer dry season to the fall RSO is mainly initiated by increasing westerly moisture transport from Atlantic Ocean into the equatorial Congo basin in the boundary layer via the lower branch of the Congo Basin Cell (925–875 hPa). These changes start 14 pentads (around mid-June) before the fall RSO (around late August). This weakens the vertically integrated MF divergence, which, in turn, slows down the drying trend during the dry season. Low-level moisture increases due to increases in moisture transport from the Atlantic Ocean during the early transition (14–3 pentads before the fall RSO). Mid-level atmospheric moisture increases as a result of increased uplift via the ascending branch of the Congo Basin Cell and orographic uplift against the East African Rift during the late transition period. This increases CWV, which subsequently reduces CIN and lows the LFC, while increasing CAPE, MSE and vertical wind shear. Subsequently, the fall rainy season begins.