Methane paradox in tropical lakes? Sedimentary fluxes rather than water column production in oxic waters sustain methanotrophy and emissions to the atmosphere

Despite growing evidence that methane (CH4) formation could also occur in well-oxygenated surface freshwaters, its significance at the ecosystem scale is uncertain. Empirical models based on data gathered at high latitude predict that the contribution of oxic CH4 increases with lake size and should represent the majority of CH4 emissions in large lakes. However, such predictive models could not directly apply to tropical lakes which differ from their temperate counterparts in some fundamental characteristics, such as year-round elevated water temperature. We conducted stable isotope tracer experiments 15 which revealed that oxic CH4 production is closely related to phytoplankton metabolism, and is a common feature in five contrasting African lakes. Nevertheless, methanotrophic activity in surface waters and CH4 emissions to the atmosphere were predominantly fuelled by CH4 generated in sediments and physically transported to the surface. Indeed, measured CH4 bubble dissolution flux and diffusive benthic CH4 flux were several orders of magnitude higher than CH4 production in surface waters. Microbial CH4 consumption dramatically decreased with increasing sunlight intensity, suggesting that the freshwater “CH4 20 paradox” might be also partly explained by photo-inhibition of CH4 oxidizers in the illuminated zone. Sunlight appeared as an overlooked but important factor determining the CH4 dynamics in surface waters, directly affecting its production by photoautotrophs and consumption by methanotrophs.

in a liquid N2 trap, CO2 and H2O were removed with a soda lime and a magnesium perchlorate traps, and the CH4 was converted to CO2 in an online combustion column similar to that in an elemental analyzer (EA). The resulting CO2 was subsequently preconcentrated in a custom-built cryo-focussing device by immersion of a stainless-steel loop in liquid N2, passed through a micro-packed GC column (HayeSep Q 2 m, 0.75mm ID; Restek), and finally measured on a Thermo Scientific Delta V Advantage isotope ratio mass spectrometer (IRMS). CO2 produced from certified reference standards for δ 13 C analysis (IAEA-85 CO1 and LSVEC) were used to calibrate δ 13 C-CH4 data. Reproducibility of measurement estimated based on duplicate injection of a selection of samples was typically better than 0.5 ‰, or better than 0.2‰ when estimated based on multiple injection of standard gas.

Diffusive CH4 flux calculation
Surface CH4 concentrations were used to compute the diffusive air-water CH4 fluxes (FCH4) according to eq. (1): 90 Where k is the gas transfer velocity of CH4 computed from wind speed (Cole & Caraco 1998) and the Schmidt number of CH4 in freshwater (Wanninkhof 1992), and ΔCH4 is the air-water gradient. Wind speed data were acquired with a Davis 95 Instruments meteorological station located in Mweya peninsula (0.11°S 29.53°E).

CH4 ebullition flux
CH4 ebullition flux was investigated in In L. Edward, George, and Nyamusingere only. Bubble traps made with an inverted funnel (24 cm diameter) connected to a 60 ml syringe were deployed for a period between 24 h and 48 h at 0.5 m below the water surface (4 replicates). Measurements were performed at sites with water depth of 20 m, 2.5 m and 3 m for L. 100 Edward, George and Nyamusingere, respectively. After measuring the gas volume collected within the trap during the sampling period, the gas bubbles were transferred in a tightly closed 12 ml Exetainer vial (Labco) for subsequent analysis of their CH4 concentration. Variability of the gas volume in the 4 replicates was less than 10%. We used the SiBu-GUI software (McGinnis et al. 2006, Greinert et al. 2009) to correct for gas exchange within the water column during the rise of bubbles and thus obtained the CH4 ebullition and CH4 bubble dissolution fluxes. Calculations were made following several scenarios: two 105 extreme bubble-size scenarios considering a release of many small (3 mm diameter) bubbles or fewer large (10 mm) bubbles, and an intermediate scenario of release of 6 mm diameter bubbles.

CH4 flux across the sediment-water interface
CH4 flux across the sediment-water interface was determined from short-term intact core incubations in L. Edward, L. George and L. Nyamusingere only. CH4 flux was quantified from the change of CH4 concentration in overlying waters at 5 110 different time steps, every 2 hours. Briefly, in every lake, 2 sediment cores (6 cm wide; ~ 30cm sediment and 30cm of water) were collected taking care to avoid disturbance at the sediment-water interface. Cores were kept in the dark until back in the laboratory, typically 6h later. Overlying water was carefully removed and replaced by bottom lake water filtered through 0.2µm polycarbonate filters (GSWP, Millipore) in order to remove water column methanotrophs. It was then degassed with helium during 20 minutes in order to remove background O2 and CH4, and gently returned in the core tubes, on top of the 115 sediments. Core tubes were tightly closed with a thick rubber stopper equipped with two sampling valves. A magnetic stirrer placed ~ 10 cm above the sediments was allowed to rotate gently in order to homogenize the overlying water layer during the incubation. At each time step, 60 ml of overlying water was sampled by connecting a syringe to the first sampling valve while an equivalent volume of degassed water was allowed to flow through the second valve in order to avoid any pressure disequilibrium. Subsamples of overlying water were transferred into a two 20 ml serum bottles filled without headspace and 120 https://doi.org/10.5194/bg-2020-142 Preprint. Discussion started: 5 May 2020 c Author(s) 2020. CC BY 4.0 License. poisoned with HgCl2. Determination of the dissolved CH4 concentration was performed with a GC-FID following the same procedure as described above.

Primary production and N2 fixation
Primary production and N2 fixation rates were determined from dual stable isotope photosynthesis-irradiance experiments using NaH 13 CO3 (Eurisotop) and dissolved 15 N2 (Eurisotop) as tracers for incorporation of dissolved inorganic 125 carbon (DIC) and N2 into the biomass. The 15 N2 tracer was added dissolved in water (Mohr et al. 2010). Incident light intensity was measured by a LI-190SB quantum sensor during day time during the entire duration of the sampling campaign. At each station a sample of surface waters (500 ml) was spiked with the tracers (final 15 N atom excess ~5%). Three subsamples were preserved with HgCl2 in 12-mL Exetainers vials (Labco) for the determination of the exact initial 13 C-DIC and 15 N-N2 enrichment. The rest of the sample was divided into nine 50-ml polycarbonate flasks, filled without headspace. Eight flasks 130 were placed into a floating incubation device providing a range of light intensity (from 0 to 80% of natural light) using neutral density filter screen (Lee Filters). The last one was immediately amended with neutral formaldehyde (0.5% final concentration) and served as killed control sample. Samples were incubated in situ during 2 hours around mid-day just below the surface at lake surface temperature. After incubation, biological activity was stopped by adding neutral formaldehyde into the flasks, and the nine samples were filtered on pre-combusted GF/F filters when back in the lab. Glass fiber filters were decarbonated with 135 HCl fumes overnight, dried, and their δ 13 C-POC and δ 15 N-PN values were determined with an EA-IRMS (Thermo FlashHTdelta V Advantage). For the measurement of the initial 15 N2 enrichment, a 2-ml helium headspace was created, and after 12h equilibration, a fraction of the headspace was injected into the above-mentioned EA-IRMS equipped with a Cu column warmed at 640°C and a CO2 trap. Initial enrichment of 13 C-DIC was also measured.
Photosynthetic (Pi) (Hama et al. 1983) and N2 fixation (N2fixi) (Montoya et al. 1996) rates in individual bottles were 140 calculated, and corrected for any abiotic tracer incorporation by subtraction of the killed control value. For each experiment, the maximum photosynthetic and N2 fixation rates (Pmax, N2fixmax) and the irradiance at the onset of light saturation (Ik_PP, Ik_N2fix) were determined by fitting Pi and N2fixi to the light intensity gradient provided by the incubator (Ii) using the equation (eq. 2) for photosynthesis activity (Vollenweider 1965) and (eq. 3) for N2 fixation (Mugidde et al. 2003).
2.7. Determination of CH4 oxidation rates. 150 CH4 oxidation rates in surface waters (1m depth) were determined from the decrease of CH4 concentrations measured during short (typically < 24h) time course experiments. Samples for CH4 oxidation rate measurement were collected in 60 mL glass serum bottles filled directly from the Niskin bottle with tubing, left to overflow, and immediately closed with butyl stoppers previously boiled in milli-Q water, and sealed with aluminum caps. The first bottle was then poisoned with a saturated solution of HgCl2 (100 µl) injected through the butyl stopper with a polypropylene syringe and a steel needle and corresponded 155 to the initial CH4 concentration at the beginning of the incubation (T0).
The remaining bottles were incubated in the dark, at in situ (~26°C) temperature during ~12h or ~24h except in L.
George and Nyamusingere where the incubation was shorter (~6h). At 4 different times step one bottle was poisoned with 100 µL of HgCl2 and stored in the dark until measurement of the CH4 concentrations with the above-mentioned GC-FID. CH4 oxidation rates were calculated as a linear regression of CH4 concentrations over time (r² generally better than 0.80) during the 160 course of the incubation. https://doi.org/10.5194/bg-2020-142 Preprint. Discussion started: 5 May 2020 c Author(s) 2020. CC BY 4.0 License.

Sunlight inhibitory effect on CH4 oxidation
The influence of light intensity on methanotrophy was investigated in Lake Edward and Lake George by means of a stable isotope ( 13 CH4) labelling experiment. For each experiment, 12 serum bottles (60 mL) were filled with lake surface waters (1m) as described above. All bottles were spiked with 100 µL of a solution of dissolved 13 CH4 (50 µmol L -1 final concentration, 165 99% enrichment) added in excess. Half of the bottles were amended with 3-(3,4-dichlorophenyl)-1,1-dimethylurea (DCMU, 0.5 mg L -1 ) in order to inhibit photosynthesis (Bishop 1958) and investigate the hypothetical inhibitory effect of dissolved O2 production by phytoplankton. Two bottles were poisoned immediately with pH-neutral formaldehyde (0.5% final concentration) and served as killed controls. The ten others were incubated during 24h at 26°C in a floating device providing 5 different light intensities (from 0 to 80% of natural light using neutral density filter screens (Lee Filters). For every bottle at 170 the end of the incubation, one 12-mL vial (Labco Exetainer) was filled with the water sample and preserved with 50 µL HgCl2.
The rest of the sample (~50 mL) was filtered on a precombusted GF/F filter for subsequent δ 13 C-POC measurement.
δ 13 C-DIC and δ 13 C-POC were determined with an EA-IRMS as described above. The methanotrophic bacterial production, defined at the CH4-derived 13 C incorporation rates into the POC pool was calculated as in eq. (4) (Morana et al. 2015): 175 Where POCt is the concentration of POC after incubation, % 13 C-POCt and % 13 C-POCi are the final and initial percentage of 13 C in the POC, t is the incubation time and % 13 C-CH4 is the percentage of 13 C in CH4 after the inoculation of 180 the bottles with the tracer. Similarly, the methanotrophic bacterial respiration rates, defined as the CH4-derived 13 C incorporation rates into the DIC pool, were calculated as in eq. (5): Where DICt is the concentration of DIC after the incubation, % 13 C-DICt and % 13 C-DICi are the final and initial percentage of 13 C in DIC and % 13 C-CH4 is the percentage of 13 C in CH4 after the inoculation of the bottles with the tracer.
Potential CH4 oxidation rates (MOX) were calculated as the sum of MBP and MBR rates. The fraction (%) of MOX inhibited by light was calculated at every light intensity as (eq.6): 190 Where MOXi is the potential CH4 oxidation for a given light treatment and MOXdark is the potential CH4 oxidation in the dark.
Time course 13 C tracer experiments were carried out in well oxygenated surface waters at every sampling site. 195 Measurement of the isotopic enrichment of the CH4 during this experiment allowed to estimate production rates of CH4 issued from 3 different precursors: 13 C-DIC (NaH 13 CO3), 13 C(1,2)-acetate and 13 Cmethyl-methionine. Serum bottles (60 ml) were spiked with 1 ml of 13 C tracer solution, or with an equivalent volume of distilled water for the control treatment. NaH 13 CO3 was added in the bottles at a tracer level (less than 5% of ambient HCO3concentration) while 13 C(1,2)-acetate and 13 Cmethyl-methionine were added largely in excess (>99% of ambient concentration). Therefore, we assume the CH4 production rates measured from 200 13 C-DIC could be representative of in-situ rates, but the production rates measured from 13 C-acetate and 13 C-methionine should https://doi.org/10.5194/bg-2020-142 Preprint. Discussion started: 5 May 2020 c Author(s) 2020. CC BY 4.0 License. instead be viewed as potential rates. The exact amount of 13 C-DIC added in the bottles was determined filling a borosilicate 12 ml exetainer vials preserved and analysed for δ 13 C-DIC as described above.
The control bottles and the bottles amended with the different 13 C tracer were incubated under constant temperature conditions (26°C) following three different treatments : (1) one third were incubated under constant light (PAR of ~ 200 µmol 205 photon m -2 s -1 ), (2) another third were incubated under the same light intensities conditions but were first amended with DCMU (0.5 mg L -1 ; final concentration), an inhibitor of photosynthesis, (3) and the last third were incubated in the dark.
At each time step (typically every 6-12h, 5-time steps), the biological activity was stopped by adding 100 µL of a saturation solution of HgCl2. Bottles were kept in the dark until CH4 concentration measurement and δ 13 C-CH4 determination as described above. 210 The term CH4_prod (nmol L -1 h -1 ) defined as the amount of CH4 produced from a specific tracer during a time interval t (h), was calculated following this equation (eq. 7) derived from Hama et al. (1983): 215 Where CH4_t and % 13 CCH4_t, represent the CH4 concentration (nmol L -1 ) and the % 13 C atom of the CH4 pool at a given time step, respectively. % 13 CCH4_i represent the % 13 C atom of the pool of CH4 at the beginning of the experiment.
% 13 Ctracer represent the % 13 C atom of the isotopically enriched pool of the precursor molecule tested (NaHCO3, methionine or acetate, depending of the treatment). % 13 C-tracer was assumed constant during the full course of the incubation given the high concentration of ambient DIC in the sampled lakes (~ 2 mmol L -1 in L. George, > 6 mmol L -1 in the other lakes) and that 220 acetate and methionine were spiked in large excess (>99%).

DNA extraction
Surface water sample for DNA analysis (between 1 L and 0.15 L, depending on the biomass) were first filtered through 5.0 µm pore size polycarbonate filters (Millipore). The eluent was then subsequently filtered through 0.2 µm pore size polycarbonate filters (Millipore) to retain free living prokaryotes. Filters were stored frozen (-20°C) immerged in a lysis buffer 225 until processing in the laboratory. Total DNA was extracted from the 0.2 µm and 5.0 µm 47 mm filters using DNeasy PowerWater kit (Qiagen) following the manufacturer's instructions. Quality and quantity of the extracted DNA were estimated using the NanoDrop ND-1000 spectrophotometer (ThermoFisher) and the Qubit 3.0 fluorometer (Life technology). Extracted DNA was stored at -20 °C until further use. identity. Generated OTU table was used to calculate relative abundances of each OTU per sample.
Surface waters were super-saturated in CH4 in all lakes, with surface concentrations (at 1 m) ranging between 78 and 652 nmol L -1 (atmospheric equilibrium ~ 2 nmol L -1 ). Vertical patterns of CH4 and stable carbon isotope composition of CH4 (δ 13 C-CH4) were variable among the different lakes. In L. Kyamwinga and Katinda, higher CH4 concentrations and lower δ 13 C-CH4 values were observed in the well-oxygenated epilimnion compared to the metalimnion showing a source of relatively 13 C-275 depleted CH4 to the epilimnetic CH4 pool (Fig. 1). The CH4 concentrations and δ 13 C-CH4 were homogeneous in the water column of L. Edward that is much larger than the other studied lakes (2300 km², Table S1) and characterized by a higher wind exposure and a substantially weaker thermal stratification (Fig. 1). However, a clear horizontal gradient in CH4 concentration and δ 13 C-CH4 occurred between the littoral and pelagic zones (Fig. S5). Vertical gradients were also observed at much smaller scale in the near sub-surface (top 0.3 m) in the shallow and entirely well oxygenated L. George and L. Nyamusingere (Fig. 2). 280 In both lakes CH4 concentrations were relatively modest in the hypolimnion (< 50 nmol L -1 ) but increased abruptly in the thermal gradient (0.3 m interval) to reach a surface maximum > 240 nmol L -1 (Fig. 2). δ 13 C-CH4 mirrored this pattern with https://doi.org/10.5194/bg-2020-142 Preprint. Discussion started: 5 May 2020 c Author(s) 2020. CC BY 4.0 License.
significantly lower values in surface than at the bottom of the water column indicating that a source of relatively 13 C-depleted CH4 contributed to the higher epilimnitic CH4.

Occurrence of microbial CH4 production in surface waters 285
Despite the prevalence of oxic conditions, 13 C-labelling experiments revealed that CH4 was produced in surface waters of each lake with the exception of L. Kyamwinga (Fig. 3). The kinetic of incorporation of NaH 13 CO3 into the CH4 pool revealed that a substantially higher amount of CH4 was produced from dissolved inorganic carbon (DIC) in illuminated waters, and this mechanism of CH4 formation appears to be related to photosynthesis, as none or only modest quantities of CH4 were produced from 13 C-labelled DIC under darkness or when photosynthesis was inhibited by DCMU (Figs. 3a and S6). 290 Furthermore, CH4 production from DIC appeared strongly correlated (r² = 0.91) to the photosynthetic activity (Fig. 4a) and N2 fixation rates (Fig 4b), supporting the view that CH4 formation in oxic waters was directly linked to phytoplankton metabolism (Bizic et al. 2020).
Aside from DIC, an appreciable amount of CH4 was generated in all lakes from the sulfur bonded methyl group of methionine when bottles were incubated under light, irrespective of the addition of DCMU ( Fig. 3b and S6), that were 295 approximately 4 times higher than in the dark. In addition, a positive relationship between CH4 production from methionine in the light and the photosynthetic activity was found (Fig. 4c). 13 C-labelled acetate, the substrate of acetoclastic methanogenesis, supported the production of CH4 in all lakes with the exception of L. Kyamwinga, but at much lower rates compared to light-dependent CH4 production from DIC (50 times lower, n=7) or methionine (10 times lower, n=4) ( Fig. 3c and S6). δ 13 C analysis of the DIC in the bottles spiked with 13 C-300 labelled acetate showed that the acetate was mineralized at rates of 5-6 orders of magnitude higher than acetoclastic methanogenesis so that added acetate appeared to be used almost exclusively by heterotrophic micro-organisms other than methanogens. Pattern of acetate-derived production of CH4 were similar in light and dark treatments (Figs. 3c and S6) and this mode of CH4 production appeared unrelated to phytoplankton activity (Fig. 4d).

Mechanisms of epilimnitic CH4 production 305
Only a minimal fraction of the CH4 produced under aerobic conditions originated from acetate in contrast with several earlier studies (Bogard et al. 2014, Donis et al. 2017) which proposed, based on the apparent fractionation factor of δ 13 C-CH4, that acetoclastic methanogenesis linked to phytoplankton production of organic matter would be the dominant biochemical pathway of pelagic CH4 production in oxic freshwaters. Instead, our results suggest that epilimnetic CH4 production in welloxygenated conditions was related to DIC fixation by photosynthesis (Fig. 3), and correlated to primary production (Fig. 4a) 310 and N2 fixation (Fig 4b). When normalized to POC concentrations, the average DIC-derived CH4 production rates (0.08 ± 0.05 nmol mmolPOC -1 h -1 n = 7) was remarkably similar to the CH4 production rates recently reported in Cyanobacteria cultures (0.04 ± 0.02 nmol mmolPOC -1 h -1 ) grown at 30°C, among which the freshwater Microcystis aeruginosa (Bizic et al. 2020), the dominant Cyanobacterium species in the tropical lakes investigated in our study (see Fig S2). These CH4 production rates are 2 orders of magnitude higher than rates reported in an axenic culture of the eukaryote Emiliania huxleyi (0.19 ± 0.07 pmol 315 mmolPOC -1 d -1 ) (Lenhart et al. 2016), but they are 4 orders of magnitude lower than typical anoxic CH4 production rates by methanogenic Archaea (Mountford & Asher 1979). Although it seems improbable that 13 C-DIC acted as a direct precursor molecule for the CH4 released by phytoplankton (Lenhart et al. 2016, Klintzsch et al. 2019) 13 C-DIC could have been taken up by phytoplankton cells and then used as a C source for the synthesis of many different organic molecules that may serve as the actual CH4 precursors. Indeed, healthy phytoplankton cells actively release a variety of low molecular weight molecules which 320 are generally highly labile and rapidly consumed (Baines & Pace 1991, Morana et al. 2014. Phytoplankton metabolism could have fuelled CH4 production pathways, at least partially, excreting substrates involved in CH4 production via biochemical https://doi.org/10.5194/bg-2020-142 Preprint. Discussion started: 5 May 2020 c Author(s) 2020. CC BY 4.0 License. While the source of methylphosphonate in freshwaters is obscure and its actual natural abundance remains to be 325 determined, dissolved free amino acids would represent up to 4% of the DOC produced by phytoplankton and are rapidly consumed by heterotrophic bacteria (Sarmento et al. 2013). Our incubations indeed demonstrated that the methyl group of methionine was a potential precursor of CH4 in all lakes investigated, in line with recent findings showing that Emiliania huxleyi could act as a direct source of CH4 in oxic conditions using methionine as precursor, without involvement of any other micro-organisms (Lenhart et al. 2016). We found that CH4 production from methionine was clearly stimulated under light, 330 even when photosynthetic activity was inhibited by DCMU, while little CH4 from methionine was produced in darkness (Fig.   3b). DCMU notably prevents reduction of plastoquinone at photosystem II and generates singlet oxygen (Petrillo et al. 2014).
The mechanism of CH4 production from methionine is still unclear, but its residue in proteins is particularly sensitive to oxidation to methionine sulfoxide by radical oxygen species (ROS) (Levine et al. 1996) so that methionine would act as an effective ROS scavenger and play important protective roles under photooxidative stress conditions, as shown in vascular 335 plants (Bruhn et al. 2012). The side chain of methionine sulfoxide is identical to dimethyl sulfoxide which is known to react with hydroxyl radicals (OH) to form CH4 (Repine et al. 1979). Besides its photoprotective role for phytoplankton, methionine could also be catabolized by a wide variety of microorganisms to methanethiol, which could in turn be transformed to CH4 as shown in Arctic Ocean surface waters (Damm et al. 2010). Nevertheless, occurrence of this latter mechanism in the tropical lakes investigated seems unlikely as this mode of CH4 production would be expected to be insensitive to light irradiance and 340 no CH4 was produced from methionine in the dark during the incubations.
3.4. Relevance of epilimnitic CH4 production compared to CH4 loss terms at ecosystem scale Net CH4 oxidation was detected in all 5 investigated lakes ranging from 11 to 5212 nmol L -1 d -1 (Fig. 5), and was by far the largest loss term of dissolved CH4 at ecosystem scale (8 to 46 times higher than the diffusive emission to the atmosphere). Surface water CH4 turnover times were particularly short in the shallow and eutrophic L. George (2h) and L. 345 Nyamusingere (3h) and slightly longer in the deeper and less productive L. Katinda (11h), L. Kyamwinga (77h) and L. Edward (100h). In all studied lakes, pelagic CH4 production rates measured during the stable isotope tracer experiments represented between 0.1% and 8.5% of net CH4 oxidation rates, regardless of the CH4 precursors tested.
All of the major sources and sinks of CH4 at ecosystem scale were experimentally determined offshore in three lakes (L. Edward at 20 m depth, George and Nyamusingere) (Fig. 6). In these three lakes, surface CH4 production rates from the 350 diverse precursors molecules investigated were modest relative to the diffusive CH4 efflux to the atmosphere (0.4 -20.0 %) and microbial CH4 oxidation (0.1 -13.2 %). In opposition, the combined CH4 bubble dissolution flux and diffusive benthic CH4 flux were several orders of magnitude higher than CH4 production in surface waters, and met the CH4 loss terms (emission and oxidation) (Fig. 6). These results gathered in tropical lakes of various size (from 0.44 to 2300 km²) and depth are in sharp contrast with the estimation of an empirical model (Gunthel et al. 2019) which proposed that mechanisms of oxic CH4 355 production represents the majority of CH4 emissions in lakes larger than 1 km². This discrepancy highlights the need to consider the unique limnological characteristics of a vast region of the world that harbours 16% of the total surface of lakes (Lehner & Doll 2004). One of the most distinctive features of tropical aquatic environment is the persistent elevated water temperature in the hypolimnion and at the water-sediment interface which favours methanogenic activity in sediment and decreases CH4 solubility, enhancing bubbles formation. 360

Origin of 13 C depleted CH4 in surface waters
Epilimnetic CH4 production was a marginal flux at ecosystem scale and could not explain alone the accumulation of 13 C-depleted CH4 in the epilimnion of most of the lakes of our dataset (Figs. 1, 2), for which we propose two other alternative https://doi.org/10.5194/bg-2020-142 Preprint. Discussion started: 5 May 2020 c Author(s) 2020. CC BY 4.0 License. mechanisms: dissolution of arising CH4 bubbles in the epilimnion combined with inhibition by light of CH4 oxidation. The partial dissolution of the CH4 bubbles as they rise in the epilimnion should allow a rapid transport of 13 C-depleted CH4 from 365 the sediment, bypassing the hotspot of CH4 oxidation at the sediment-water interface and representing an alternative source of 13 C-depleted CH4 in water column. The shallower L. George and L. Nyamusingere were notably characterized by sharp thermal density gradients (Fig. 2) and extreme phytoplankton biomass largely dominated by Microcystis aeruginosa (Chlorophyll a up to 190 µg L -1 ). Microcystis aeruginosa cells form large aggregates (>1 mm) embedded in a matrix of extracellular polymeric substance that might act as a barrier to trap small CH4 bubbles arising from the sediment (Fig S7). Dissolution of CH4 bubbles 370 could be enhanced at the very near surface due to the entrapments of bubbles at the air-water interface by abundant surface organic films that delay the bubble "burst". The presence of a sharp sub-surface temperature gradient would further enhance CH4 accumulation during day-time near the air-water interface (by blocking vertical redistribution of CH4 by mixing). We hypothesize that this process could be widespread in shallow tropical lakes which are characterized by high productivity and are susceptible to be simultaneous large benthic CH4 sources. 375 The influence of light on methanotrophy was investigated in the deep L. Edward, and shallow L. George and L.