Biogeosciences, 17, 5209–5221, 2020https://doi.org/10.5194/bg-17-5209-2020© Author(s) 2020. This work is distributed underthe Creative Commons Attribution 4.0 License.

Methane paradox in tropical lakes? Sedimentary fluxes rather thanpelagic production in oxic conditions sustain methanotrophy andemissions to the atmosphereCédric Morana1,2, Steven Bouillon1, Vimac Nolla-Ardèvol1, Fleur A. E. Roland2, William Okello3, Jean-Pierre Descy2,Angela Nankabirwa3, Erina Nabafu3, Dirk Springael1, and Alberto V. Borges2

1KU Leuven, Leuven, Belgium2Université de Liège, Liège, Belgium3National Fisheries Resources Research Institute, Jinja, Uganda

Correspondence: Cédric Morana (cedric.morana@kuleuven.be)

Received: 21 April 2020 – Discussion started: 5 May 2020Revised: 20 August 2020 – Accepted: 8 September 2020 – Published: 29 October 2020

Abstract. Despite growing evidence that methane (CH4) for-mation could also occur in well-oxygenated surface fresh wa-ters, its significance at the ecosystem scale is uncertain. Em-pirical models based on data gathered at high latitude pre-dict that the contribution of oxic CH4 increases with lakesize and should represent the majority of CH4 emissions inlarge lakes. However, such predictive models could not di-rectly apply to tropical lakes, which differ from their temper-ate counterparts in some fundamental characteristics, such asyear-round elevated water temperature. We conducted stable-isotope tracer experiments, which revealed that oxic CH4production is closely related to phytoplankton metabolismand is a common feature in five contrasting African lakes.Nevertheless, methanotrophic activity in surface waters andCH4 emissions to the atmosphere were predominantly fu-elled by CH4 generated in sediments and physically trans-ported to the surface. Indeed, CH4 bubble dissolution fluxand diffusive benthic CH4 flux were several orders of mag-nitude higher than CH4 production in surface waters. Micro-bial CH4 consumption dramatically decreased with increas-ing sunlight intensity, suggesting that the freshwater “CH4paradox” might be also partly explained by photo-inhibitionof CH4 oxidizers in the illuminated zone. Sunlight appearedas an overlooked but important factor determining the CH4dynamics in surface waters, directly affecting its productionby photoautotrophs and consumption by methanotrophs.

1 Introduction

Emissions from inland waters are an important component ofthe global methane (CH4) budget (Bastviken et al., 2011), inparticular from tropical latitudes (Borges et al., 2015). Whileprogress has been made in evaluating the CH4 emissionrates, much less attention has been given to the underlyingmicrobial production (methanogenesis) and loss (methaneoxidation) processes. It is generally assumed that CH4 inlakes originates from the degradation of organic matter inanoxic sediments. Because most methanogens are consid-ered to be strict anaerobes, and net vertical diffusion of CH4from anoxic bottom waters is often negligible (Bastviken etal., 2003), physical processes of CH4 transport from shal-low sediments are usually invoked to explain patterns of localCH4 concentration maxima in surface waters (Fernández etal., 2016; Peeters et al., 2019; Martinez-Cruz et al., 2020). In-deed, CH4-rich pore water is regularly released from littoralsediment into the water column during resuspension eventsassociated with surface waves (Hofmann et al., 2010).

The view that CH4 is formed under strictly anaerobicconditions has been challenged by several recent studies,which have proposed that acetoclastic methanogens directlyattached to phytoplankton cells are involved in epilimneticCH4 production (Grossart et al., 2011; Bogard et al., 2014)and are responsible for distinct near-surface peaks of CH4concentration in certain thermally stratified, well-oxygenatedwaterbodies (Tang et al., 2016). It has also been shown thatCyanobacteria (Bizic et al., 2020) and widespread marine

Published by Copernicus Publications on behalf of the European Geosciences Union.

5210 C. Morana et al.: Sedimentary fluxes rather than pelagic production

phytoplankton (Klintzsch et al., 2019) are able to release asubstantial amount of CH4 during a culture study, and thisCH4 production mechanism might be linked to photosyn-thesis. From a model-based approach, epilimnetic CH4 pro-duction was shown to sustain most of the CH4 oxidation in14 Canadian lakes (DelSontro et al., 2018) and would evenrepresent up to 90 % of the CH4 emitted from a temperatelake (Donis et al., 2017). Further, empirical models basedon data gathered in boreal and temperate lakes predict thatthe contribution of oxic CH4 increases with lake size (Gün-thel et al., 2019) and should represent the majority of CH4emissions in lakes larger than 1 km2. Still, aerobic CH4 pro-duction has so far only been documented in temperate andboreal lakes so that such predictive models could not di-rectly apply to tropical lakes, which differ from their temper-ate counterparts in some fundamental characteristics, such asyear-round elevated water temperature. Primary production,methanogenic and methanotrophic activities, and cyanobac-terial dominance are potentially much higher in tropical lakesdue to favourable temperature (Lewis, 1987; Kosten et al.,2012). It has also been shown that CH4 emissions are pos-itively related to temperature at the ecosystem scale (Yvon-Durocher et al., 2014)

Here, we tested the hypothesis that phytoplanktonmetabolism could fuel CH4 production in well-oxygenatedwaters in five contrasting tropical lakes in East Africa cover-ing a wide range of sizes, depths, and productivity (Lake Ed-ward, Lake George, Lake Katinda, Lake Nyamusingere andLake Kyambura). Phytoplankton activity could provide di-verse substrates required for CH4 production mediated bymethanogenic Archaea, or alternatively CH4 could be di-rectly released by phytoplankton cells. Additionally, the sig-nificance of epilimnetic CH4 production at the scale of theaquatic ecosystem was assessed by quantifying CH4 releasefrom sediments, CH4 production and oxidation rates in thewater column, and CH4 diffusive and ebullitive emissions tothe atmosphere.

2 Material and methods

2.1 Site description

The sampled lakes cover a wide range of size (<1 to2300 km2), maximum depth (3–117 m), mixing regimes,phytoplankton biomass and primary productivity (Ta-ble S1 in the Supplement). Oligotrophic Lake Kyamwinga(−0.18054◦ N, 30.14625◦ E) and eutrophic Lake Katinda(−0.21803◦ N, 30.10702◦ E) are stratified, small but deeptropical lakes located in western Uganda. Neighbour-ing Lake Nyamusingere (−0.284364◦ N, 30.037635◦ E)is a small but shallow and polymictic eutrophic lake.Lake George is a larger (250 km2), hypereutrophic, shallowlake located at the Equator (−0.02273◦ N, 30.19724◦ E). Asingle outlet (Kazinga Channel) flows from Lake George to

the neighbouring Lake Edward (−0.28971◦ N, 29.73327◦ E),a holomictic, mesotrophic large lake (2325 km2).

Water samples from pelagic stations of Lake Katinda(0.3 km from shore, 14 m deep), Lake George (2 km fromshore, 2.5 m deep) and Lake Edward (15 km from shore, 20 mdeep) were collected in April 2017 (rainy season) and Jan-uary 2018 (dry season). Pelagic sites of Lake Kyamwinga(0.5 km from shore, 40 m deep) and Lake Nyamusingere(0.2 km from shore, 3 m deep) were sampled only once, inApril 2017 and January 2018, respectively.

2.2 Environmental setting of the study sites

Conductivity, temperature and dissolved-oxygen concentra-tion measurements were performed with a Yellow Spring In-strument EXO II multi-parametric probe. Samples for par-ticulate organic carbon (POC) concentration were collectedon glass fibre filters (0.7 µm nominal pore size) and anal-ysed with an elemental analyser coupled to an isotope ratiomass spectrometer (EA–IRMS; Morana et al., 2015a). Pig-ment concentrations were determined by high-performanceliquid chromatography (Descy et al., 2017) after filtrationof water samples through glass fibre filters (0.7 µm nominalpore size).

Water samples for determination of dissolved CH4 con-centration were transferred with tubing from the Niskin bot-tle to 60 mL borosilicate serum bottles that were poisonedwith 200 µL of a saturated solution of HgCl2, closed witha butyl stopper and sealed with an aluminium cap. Theconcentrations of dissolved CH4 were measured with theheadspace equilibration technique (20 mL headspace) usinga gas chromatograph with flame ionization detection (GC–FID, SRI8610C).

Samples for δ13C–CH4 determination were collected in60 mL serum bottles following the same procedure as sam-ples for CH4 concentration determination. In the labora-tory, δ13C–CH4 was measured as described in Morana etal. (2015b). Briefly, a 20 mL helium headspace was createdin the serum bottles, then samples were vigorously shakenand left to equilibrate overnight. The sample gas was flushedout through a double-hole needle and purified of non-CH4volatile organic compounds in a liquid N2 trap, CO2 and H2Owere removed with a soda lime and magnesium perchlo-rate traps, and the CH4 was converted to CO2 in an onlinecombustion column similar to that in an elemental analyser(EA). The resulting CO2 was subsequently pre-concentratedin a custom-built cryo-focussing device by immersion of astainless-steel loop in liquid N2, passed through a micro-packed GC column (HayeSep Q 2 m, 0.75 mm ID; Restek),and finally measured on a Thermo Scientific Delta V Advan-tage isotope ratio mass spectrometer (IRMS). CO2 producedfrom certified reference standards for δ13C analysis (IAEA-CO1 and LSVEC) were used to calibrate δ13C–CH4 data. Re-producibility of measurements estimated based on duplicateinjection of a selection of samples was typically better than

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C. Morana et al.: Sedimentary fluxes rather than pelagic production 5211

0.5 ‰ or better than 0.2‰ when estimated based on multipleinjection of standard gas.

2.3 Diffusive CH4 flux calculation

Surface CH4 concentrations were used to compute the diffu-sive air–water CH4 fluxes (FCH4) according to Eq. (1):

FCH4 = k × 1CH4, (1)

where k is the gas transfer velocity of CH4 computed fromwind speed (Cole and Caraco, 1998) and the Schmidt numberof CH4 in fresh water (Wanninkhof, 1992), and 1CH4 is theair–water gradient. Wind speed data were acquired with aDavis Instruments meteorological station located on Mweyapeninsula (0.11◦ S, 29.53◦ E).

2.4 CH4 ebullition flux

CH4 ebullition flux was investigated in lakes Edward (at20 m), George (at 2.5 m), and Nyamusingere (at 3 m) only,at the same sampling sites where biogeochemical processeswere measured. Bubble traps made with an inverted funnel(24 cm diameter) connected to a 60 mL syringe were de-ployed for a period between 24 and 48 h at 0.5 m belowthe water surface (four replicates). Measurements were per-formed at sites with water depth of 20, 2.5 and 3 m forlakes Edward, George and Nyamusingere, respectively. Af-ter measuring the gas volume collected within the trap dur-ing the sampling period, the gas bubbles were transferred ina tightly closed 12 mL Exetainer vial (Labco) for subsequentanalysis of their CH4 concentration. Variability of the gasvolume in the four replicates was less than 10 %. We usedthe SiBu-GUI software (McGinnis et al., 2006; Greinert andMcGinnis, 2009) to correct for gas exchange within the watercolumn during the rise of bubbles and thus obtained the CH4ebullition and CH4 bubble dissolution fluxes. Calculationswere made following several scenarios: two extreme bubble-size scenarios considering a release of many small (3 mm di-ameter) bubbles or fewer large (10 mm) bubbles and an inter-mediate scenario of release of 6 mm diameter bubbles. Del-wiche and Hemond (2017) experimentally determined thata large majority (∼ 90 %) of the bubbles released from lakesediments fall in this size interval.

2.5 Potential CH4 flux across the sediment–waterinterface

The potential CH4 flux across the sediment–water interfacewas determined from short-term intact core incubations inLake Edward, Lake George and Lake Nyamusingere only,at the same sampling sites where biogeochemical processeswere measured. CH4 flux was quantified from the change inCH4 concentration in overlying waters at five different timesteps, every 2 h. Briefly, in every lake, two sediment cores(6 cm wide; ∼ 30 cm of sediment and 30 cm of water) were

collected, taking care to avoid disturbance at the sediment–water interface. Cores were kept in the dark until back in thelaboratory, typically 6 h later. Overlying water was carefullyremoved with a syringe, taking care to avoid any disturbanceof the sediments. It was replaced by bottom lake water, whichhad been previously filtered through 0.2 µm polycarbonatefilters (GSWP, Millipore) in order to remove water columnmethanotrophs and degassed with helium for 20 min in or-der to remove background O2 and CH4, and the filtered anddegassed water was gently returned to the core tubes on topof the sediments. Core tubes were tightly closed with a thickrubber stopper equipped with two sampling valves. A mag-netic stirrer placed∼ 10 cm above the sediments was allowedto rotate gently in order to homogenize the overlying waterlayer during the incubation. At each time step, 60 mL of over-lying water was sampled by connecting a syringe to the firstsampling valve while an equivalent volume of degassed wa-ter was allowed to flow through the second valve in orderto avoid any pressure disequilibrium. Subsamples of over-lying water were transferred into two 20 mL serum bottlesfilled without headspace and poisoned with HgCl2. Deter-mination of the dissolved CH4 concentration was performedwith a GC–FID following the same procedure as describedabove. The removal of oxygen might have inhibited the bio-geochemical activity of aerobic methanotrophs present at thesediment–water interface; hence we considered that the re-sults gathered from this experiment are representative of apotential (maximal) CH4 flux.

2.6 Primary production and N2 fixation

Primary production and N2 fixation rates were determinedfrom dual stable-isotope photosynthesis–irradiance experi-ments using NaH13CO3 (Eurisotop) and dissolved 15N2 (Eu-risotop) as tracers for incorporation of dissolved inorganiccarbon (DIC) and N2 into the biomass. The 15N2 tracer wasdissolved in water (Mohr et al., 2010). Incident light in-tensity was measured by a LI-190SB quantum sensor dur-ing daytime during the entire duration of the sampling cam-paign. At each station a sample of surface waters (500 mL)was spiked with the tracers (final 15N atom excess ∼ 5 %).Three subsamples were preserved with HgCl2 in 12 mL Ex-etainer vials (Labco) for the determination of the exact ini-tial 13C–DIC and 15N–N2 enrichment. The rest of the sam-ple was divided into nine 50 mL polycarbonate flasks, filledwithout headspace. Eight flasks were placed into a floatingincubation device, providing a range of light intensity (from0 % to 80 % natural light) using neutral density filter screen(LEE Filters). The last one was immediately amended withneutral formaldehyde (0.5 % final concentration) and servedas a killed control sample. Samples were incubated in situfor 2 h around midday just below the surface at lake sur-face temperature. After incubation, biological activity wasstopped by adding neutral formaldehyde into the flasks, andthe nine samples were filtered on pre-combusted glass fi-

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5212 C. Morana et al.: Sedimentary fluxes rather than pelagic production

bre (GF/F grade) filters when back in the lab. Glass fibrefilters were decarbonated with HCl fumes overnight, dried,and their δ13C–POC and δ15N–PN (particulate nitrogen) val-ues were determined with an EA–IRMS (Thermo Flash HT –Delta V Advantage). For the measurement of the initial 15N2enrichment, a 2 mL helium headspace was created, and after12 h equilibration, a fraction of the headspace was injectedinto the above-mentioned EA–IRMS equipped with a Cu col-umn warmed at 640 ◦C and a CO2 trap. Initial enrichment of13C–DIC was also measured.

Photosynthetic (Pi; Hama et al., 1983) and N2 fixation(N2fixi; Montoya et al., 1996) rates in individual bottles werecalculated and corrected for any abiotic tracer incorporationby subtraction of the killed control value. For each exper-iment, the maximum photosynthetic and N2 fixation rates(Pmax, N2fixmax) and the irradiance at the onset of light sat-uration (Ik_PP, Ik_N2fix) were determined by fitting Pi andN2fixi to the light intensity gradient provided by the incu-bator (Ii) using Eq. (2) for photosynthesis activity (Vollen-weider, 1965) and (Eq. 3) for N2 fixation (Mugidde et al.,2003).

Pi = 2Pmax

[Ii/2Ik_PP

1+ (Ii/2Ik_PP)2

](2)

N2fixi = 2N2fixmax

[Ii/2Ik_N2fix

1+(Ii/2Ik_N2fix

)2]. (3)

2.7 Determination of CH4 oxidation rates

CH4 oxidation rates in surface waters (1 m depth) were de-termined from the decrease in CH4 concentrations measuredduring short (typically<24 h) time course experiments. Sam-ples for CH4 oxidation rate measurement were collected in60 mL glass serum bottles filled directly from the Niskin bot-tle with tubing, left to overflow, and immediately closed withbutyl stoppers previously boiled in Milli-Q water and sealedwith aluminium caps. The first bottle was then poisoned witha saturated solution of HgCl2 (100 µL) injected through thebutyl stopper with a polypropylene syringe and a steel nee-dle and corresponded to the initial CH4 concentration at thebeginning of the incubation (T0).

The remaining bottles were incubated in the dark at insitu (∼ 26 ◦C) temperature for ∼ 12 h or ∼ 24 h except inlakes George and Nyamusingere, where the incubation wasshorter (∼ 6 h). At four different time steps, one bottle waspoisoned with 100 µL of HgCl2 and stored in the dark un-til measurement of the CH4 concentrations with the above-mentioned GC–FID. CH4 oxidation rates were calculated asa linear regression of CH4 concentrations over time (r2 gen-erally better than 0.80) during the course of the incubation.O2 consumption was followed in a parallel incubation ofwater samples in 60 mL biological oxygen demand bottles(Wheaton) to ensure the samples were incubated under oxicconditions during the full course of this short-term experi-ment.

2.8 Sunlight inhibitory effect on CH4 oxidation

The influence of light intensity on methanotrophy was in-vestigated in Lake Edward (April 2017 and January 2018),Lake Nyamusingere and Lake George (January 2018) bymeans of a stable-isotope (13CH4) labelling experiment.For each experiment, 12 serum bottles (60 mL) were filledwith lake surface waters (1 m) as described above. Allbottles were spiked with 100 µL of a solution of dis-solved 13CH4 (50 µmol L−1 final concentration, 99 % enrich-ment) added in excess. In Lake Edward in January, half ofthe bottles were amended with 3-(3,4-dichlorophenyl)-1,1-dimethylurea (DCMU; 0.5 mg L−1) in order to inhibit pho-tosynthesis (Bishop, 1958) and investigate the hypotheticalinhibitory effect of dissolved O2 production by phytoplank-ton. Two bottles were poisoned immediately with pH-neutralformaldehyde (0.5 % final concentration) and served as killedcontrols. The 10 others were incubated for 24 h at 26 ◦Cin a floating device providing five different light intensities(from 0 % to 80 % of natural light using neutral density filterscreens; LEE Filters). For every bottle at the end of the incu-bation, one 12 mL vial (Labco Exetainer) was filled with thewater sample and preserved with 50 µL HgCl2. The rest ofthe sample (∼ 50 mL) was filtered on a pre-combusted glassfibre (GF/F grade) filter for subsequent δ13C–POC measure-ment.δ13C–DIC and δ13C–POC were determined with an EA–

IRMS following the method described in Gillikin and Bouil-lon (2007) and Morana et al. (2015), respectively. Themethanotrophic bacterial production (MBP), defined as theCH4-derived 13C incorporation rates into the POC pool, wascalculated as in Eq. (4) (Morana et al., 2015):

MBP=POCt ×

(%13CPOCt/%13CPOCi

)t ×

(%13CCH4/%13CPOCi

) , (4)

where POCt is the concentration of POC after incubation,%13C–POCt and %13C–POCi are the final and initial per-centage of 13C in the POC, t is the incubation time, and%13C–CH4 is the percentage of 13C in CH4 after the inoc-ulation of the bottles with the tracer. Similarly, the methan-otrophic bacterial respiration (MBR) rates, defined as theCH4-derived 13C incorporation rates into the DIC pool, werecalculated as in Eq. (5):

MBR=DICt ×

(%13CDICt/%13CDICi

)t ×

(%13CCH4/%13CDICi

) , (5)

where DICt is the concentration of DIC after the incubation,%13C–DICt and %13C–DICi are the final and initial percent-age of 13C in DIC, and %13C–CH4 is the percentage of 13Cin CH4 after the inoculation of the bottles with the tracer.

Potential CH4 oxidation rate (MOX) was calculated as thesum of MBP and MBR rates. The fraction (%) of MOX in-hibited by light was calculated at every light intensity as

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C. Morana et al.: Sedimentary fluxes rather than pelagic production 5213

(Eq. 6):

MOXinihibition(%)= (1− MOXi/MOXdark) × 100, (6)

where MOXi is the potential CH4 oxidation for a given lighttreatment, and MOXdark is the potential CH4 oxidation in thedark.

2.9 Determination of pelagic CH4 production rates

Time course 13C tracer experiments were carried out in well-oxygenated surface waters at every sampling site. Measure-ment of the isotopic enrichment of the CH4 during this ex-periment allowed us to estimate production rates of CH4 is-sued from three different precursors: 13C–DIC (NaH13CO3),13C(1,2)–acetate and 13Cmethyl–methionine. Serum bottles(60 mL) were spiked with 1 mL of 13C tracer solution orwith an equivalent volume of distilled water for the con-trol treatment. NaH13CO3 was added in the bottles at atracer level (less than 5 % of ambient HCO−3 concentra-tion), while 13C(1,2)–acetate and 13Cmethyl–methionine wereadded largely in excess (>99 % of ambient concentration).Therefore, we assume the CH4 production rates measuredfrom 13C–DIC could be representative of in situ rates, butthe production rates measured from 13C–acetate and 13C–methionine should instead be viewed as potential rates. Theexact amount of 13C–DIC added in the bottles was deter-mined filling borosilicate 12 mL Exetainer vials preservedand analysed for δ13C–DIC as described above.

The control bottles and the bottles amended with the dif-ferent 13C tracer were incubated under constant tempera-ture conditions (26 ◦C) following three different treatments:(1) one-third were incubated under constant light (PAR of∼ 200 µmol photon m−2 s−1); (2) another third were incu-bated under the same light intensity conditions but were firstamended with DCMU (0.5 mg L−1 ; final concentration), aninhibitor of photosynthesis; (3) and the last third were incu-bated in the dark.

At each time step (typically every 6–12 h, five time steps),the biological activity was stopped by adding 100 µL of asaturation solution of HgCl2. Bottles were kept in the darkuntil CH4 concentration measurement and δ13C–CH4 deter-mination as described above. O2 consumption was followedin a parallel incubation of water samples in 60 mL biologi-cal oxygen demand bottles (Wheaton) to ensure the sampleswere incubated under oxic conditions during the full courseof the experiment.

The term CH4_prod (nmol L−1 h−1), defined as the amountof CH4 produced from a specific tracer during a time intervalt (h), was calculated following this equation (Eq. 7) derivedfrom Hama et al. (1983):

CH4_prod =CH4_t ×

(%13CCH4_t/%13CCH4_i

)t ×

(%13Ctracer/%13CCH4_i

) , (7)

where CH4_t and %13CCH4_t represents the CH4 concen-tration (nmol L−1) and the %13C atom of the CH4 pool

at a given time step, respectively. %13CCH4_i representsthe %13C atom of the pool of CH4 at the beginning ofthe experiment. %13Ctracer represent the %13C atom ofthe isotopically enriched pool of the precursor moleculetested (NaHCO3, methionine or acetate, depending on thetreatment). %13Ctracer was assumed constant during thefull course of the incubation given the high concentrationof ambient DIC in the sampled lakes (∼ 2 mmol L−1 inLake George, >6 mmol L−1 in the other lakes) and that ac-etate and methionine were spiked in large excess (>99 %).

2.10 DNA extraction

Surface water samples for DNA analysis (between 1 and0.15 L, depending on the biomass) were first filtered through5.0 µm pore size polycarbonate filters (Millipore). The eluentwas then subsequently filtered through 0.2 µm pore size poly-carbonate filters (Millipore) to retain free living prokaryotes.Filters were stored frozen (−20 ◦C) and immersed in a lysisbuffer until processing in the laboratory. Total DNA was ex-tracted from the 0.2 and 5.0 µm 47 mm filters using DNeasyPowerWater kit (Qiagen) following the manufacturer’s in-structions. Quality and quantity of the extracted DNA wereestimated using the NanoDrop ND-1000 spectrophotometer(Thermo Fisher) and the Qubit 3.0 fluorometer (Life Tech-nologies). Extracted DNA was stored at −20 ◦C until furtheruse.

2.11 Quantification of mcrA via qPCR

Quantification of mcrA gene copies was performed by quan-titative polymerase chain reaction (qPCR) on the total ex-tracted DNA. The used primer pair consisted of forwardprimer qmcrA-F 5′-TTCGGTGGATCDCARAGRGC-3′ andqmcrA-R 5′-GBARGTCGWAWCCGTAGAATCC-3′ (Den-man et al., 2007). The reaction mixture contained 3 µL oftotal community DNA extract, 7.5 µL of ABsolute qPCRSYBR Green Mix (Thermo Fisher, Cat. AB1158B), 0.3 µLof 10 µM forward primer mcrF, 0.3 µL of 10 µM reverseprimer mcrR, 1.5 µL of a 1 % w/v bovine serum albuminsolution (Amersham Bioscience), and 2.4 µL of nuclease-and DNA-free water. The qPCR was performed in a Ro-torgene 3000 (Corbett Research) using the following condi-tions: 95 ◦C (15 min) followed by 40 cycles of 20 s at 95 ◦C,20 s at 58 ◦C, 20 s at 72 ◦C and a final extension step of 5 s at80 ◦C. Standard curves were prepared from serial dilutions ofa pre-quantified mcrA PCR fragment amplified using primersmcrF and mcrR from a plasmid extract carrying the com-plete mcrA gene using concentrations ranging from 1× 102

to 1×108 copies µL−1. Samples were analysed in triplicates.

2.12 16S rRNA gene amplicon sequencing

Sequencing of the 16S rRNA gene was done on thetotal extracted DNA to assess community composition.The 16S rRNA gene sequencing was done with the Il-

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5214 C. Morana et al.: Sedimentary fluxes rather than pelagic production

lumina MiSeq v3 chemistry following the “16S Metage-nomic Sequencing Library Preparation” protocol with thefollowing universal 16S rRNA gene primers targetingthe V4 region, forward UniF/A519F-(S-D-Arch-0519-a-S-15) 5′-CAGCMGCCGCGGTAA-3′ and reverse UniR/802R-(S-D-Bact-0785-b-A-18) 5′-TACNVGGGTATCTAATCC-3′

(Klindworth et al., 2013). Sequenced read quality waschecked using FastQC v0.11 (https://www.bioinformatics.babraham.ac.uk/projects/fastqc/, last access: 4 June 2018).Short reads were trimmed to 250 bp with FastX Toolkitv0.0.13 (http://hannonlab.cshl.edu/fastx_toolkit/, last access:4 June 2018) in order to remove trailing Ns and low-qualitybases. Operational taxonomical units (OTUs) for each anal-ysed sample were obtained from the quality-trimmed readsusing mothur v1.39.5 (Kozich et al., 2013) and followingthe online MiSeq SOP (https://mothur.org/wiki/MiSeq_SOP,last access: 4 June 2018) using the Silva v128 16S rRNAdatabase with the following parameters: maxambig= 0 bp,maxlength= 300 bp, maxhomop= 8 and classify OTUs to97 % identity. The generated OTU table was used to calcu-late relative abundances of each OTU per sample.

3 Results

3.1 Environmental settings

The sampled lakes cover a wide range of sizes (<1 to2300 km2), maximum depths (3–117 m), mixing regimes,phytoplankton biomass and primary productivity (Table S1,Fig. S1). Phytoplankton biomass (chlorophyll a from 3.6to 190.2 µg L−1) was dominated by Cyanobacteria (>95 %)in the most productive lakes, while diatoms (<20 %) andchrysophytes (<40 %) also contributed in the less produc-tive ones (Fig. S2). Maximum potential photosynthetic ac-tivity (Pmax) varied from 1.5 µmol C L−1 h−1 in Lake Ed-ward to 199.0 µmol C L−1 h−1 in Lake George and was lin-early related to chlorophyll a concentration. Light-dependentN2 fixation was detected in every lake with the exceptionof Lake Kyamwinga. No significant N2 fixation rates weremeasured in the dark. Maximum potential N2 fixation rates(N2fixmax) ranged between 1 and 128 nmol L−1 h−1 and werepositively related to Pmax (Fig. S1).

We detected and quantified the abundance of the archaealalpha subunit of methyl-coenzyme M reductase gene (mcrA),a proxy for methanogens, in the surface waters of eachlake. mcrA gene copy abundance (mcrA copy ng DNA−1)ranged between 319± 41 (Lake Edward) and 7537± 476(Lake Katinda) in the fraction of seston <5 µm and between541±19 (Lake Edward) and 7968±167 (Lake Katinda) in thefraction of seston >5 µm (Fig. S3). Illumina 16S rRNA geneamplicon sequencing indicated that methanogens accountedfor a small fraction of the prokaryotic community in thesurface waters of Lake Edward (0.01 %), Lake Kyamwinga(0.03 %) and Lake Nyamunsingere (0.08 %). They repre-

sented a substantially higher fraction of the community inLake Katinda (0.38 %) and Lake George (0.57 %; Fig. S4).In all lakes, hydrogenotrophic (Methanomicrobiales andMethanobacteriales) were always more abundant than aceto-clastic (Methanosarcinales) microorganisms, representing atleast 65 % of the methanogens (up to 95 % in Lake Katinda;Fig. S4).

3.2 Water column CH4 concentration and δ13C–CH4patterns

Surface waters were supersaturated in CH4 in all lakes,with surface concentrations (at 1 m) ranging between 78and 652 nmol L−1 (atmospheric equilibrium ∼ 2 nmol L−1).Diffusive CH4 emissions varied between 0.05 (Lake Ed-ward) and 0.40 mmol m−2 d−1 (Lake Katinda). The ben-thic CH4 flux across the sediment–water interface was el-evated in comparison with the diffusive CH4 emissionsin the three lakes where it was measured: Lake Ed-ward (0.96 mmol m−2 d−1), Lake George (9 mmol m−2 d−1)and Lake Nyamusingere (5 mmol m−2 d−1). CH4 ebulli-tion was the dominant pathway of CH4 evasion to the at-mosphere in the three lakes (Lake Edward, Lake George,Lake Nyamusingere), where ebullitive fluxes were in-vestigated. Ebullitive flux in Lake Edward ranged be-tween 0.16 and 0.24 mmol m−2 d−1, depending on thebubble-size scenario considered (see material and meth-ods), being at least 4 times higher than the diffusive CH4flux. This discrepancy was even larger in the shallowerLake George (13.26–13.9. mmol m−2 d−1) and Lake Nya-musingere (19.03–19.09 mmol m−2 d−1) where CH4 ebulli-tion appeared ∼ 100 and ∼ 50 times higher than diffusiveCH4 emissions. During the ascent of CH4 bubbles to the sur-face, gas exchange occurs, and we estimated that, depend-ing on the CH4 bubble size considered, between 0.04 and0.21 mmol CH4 m−2 d−1 dissolved in the water column ofLake Edward during bubble ascent. This bubble dissolutionflux ranged between 0.70 and 3.30 and between 1.21 and5.55 mmol m−2 d−1 in Lake George and Lake Nyamusin-gere, respectively.

Vertical patterns of CH4 and stable carbon isotope com-position of CH4 (δ13C–CH4) were variable among the dif-ferent lakes. In lakes Kyamwinga and Katinda, higher CH4concentrations and lower δ13C–CH4 values were observedin the well-oxygenated epilimnion compared to the metal-imnion showing a source of relatively 13C-depleted CH4to the epilimnetic CH4 pool (Fig. 1). The CH4 concentra-tions and δ13C–CH4 were homogeneous in the water col-umn of Lake Edward, which is much larger than the otherstudied lakes (2300 km2; Table S1) and characterized by ahigher wind exposure and a substantially weaker thermalstratification (Fig. 1). However, a clear horizontal gradi-ent in CH4 concentration and δ13C–CH4 occurred betweenthe littoral and pelagic zones (Fig. S5). Vertical gradientswere also observed at a much smaller scale near the sub-

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Figure 1. Depth profile. Depth profile of the temperature (◦C; dashed line), CH4 concentration (nmol L−1 ; red symbols), dissolved-oxygensaturation (%; solid line) and stable-isotope carbon composition of CH4 (δ13C–CH4; ‰; blue symbols) in Lake Kyamwinga (a), LakeKatinda (b) and Lake Edward (c).

Figure 2. Depth profile, focus on the surface. Depth profile of the temperature (◦C; dashed line), CH4 concentration (nmol L−1 ; redsymbols), dissolved-oxygen saturation (%; solid line) and stable-isotope carbon composition of CH4 (δ13C–CH4; ‰; blue symbols) in LakeGeorge (a), Lake Nyamusingere (b) and the surface waters (0–2 m) of Lake Edward (c).

surface (top 0.3 m) and in the shallow and entirely well-oxygenated Lake George and Lake Nyamusingere (Fig. 2).In both lakes CH4 concentrations were relatively modest inthe hypolimnion (<50 nmol L−1) but increased abruptly inthe thermal gradient (0.3 m interval) to reach a surface max-imum >240 nmol L−1 (Fig. 2). δ13C–CH4 mirrored this pat-tern with significantly lower values at the surface than at thebottom of the water column, indicating that a source of rela-tively 13C-depleted CH4 contributed to the higher epilimneticCH4.

3.3 Occurrence of microbial CH4 production in surfacewaters under oxic conditions

Despite the prevalence of oxic conditions, 13C-labelling ex-periments revealed that CH4 was produced in surface watersof each lake with the exception of Lake Kyamwinga (Fig. 3).The kinetic of incorporation of NaH13CO3 into the CH4 pool

revealed that a substantially higher amount of CH4 was pro-duced from dissolved inorganic carbon (DIC) in illuminatedwaters, and this mechanism of CH4 formation appears to berelated to photosynthesis as no or only modest quantitiesof CH4 were produced from 13C-labelled DIC in darknessor when photosynthesis was inhibited by DCMU (Figs. 3aand S6). Furthermore, CH4 production from DIC appearedstrongly correlated (r2

= 0.91) to the photosynthetic activity(Fig. 4a) and N2 fixation rates (Fig 4b), supporting the viewthat CH4 formation in oxic waters was directly linked to phy-toplankton metabolism (Bizic et al., 2020).

Aside from DIC, an appreciable amount of CH4 was gen-erated in all lakes from the sulfur-bonded methyl group ofmethionine when bottles were incubated under light, irre-spective of the addition of DCMU (Figs. 3b and S6). CH4production rates were approximately 4 times higher in thebottles incubated in the light than in the dark. In addition, a

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5216 C. Morana et al.: Sedimentary fluxes rather than pelagic production

Figure 3. Tracer experiments show CH4 production in well-oxygenated surface waters. CH4 production rates (nmol L−1 d−1) from dissolvedinorganic carbon (a), the methyl group of methionine (b) and acetate (c) measured in the surface waters (0.3 m) of a variety of African tropicallakes. Green, grey and dark bars, respectively, represent rates measured under light, light in the presence of a photosynthesis inhibitor(DCMU) and darkness. Values showed for Lake Edward, Lake George and Lake Katinda are the average of 2017 and 2018 samplingcampaign measurement; n.s.= not significant, <0.1= below 0.1 nmol L−1 d−1.

Figure 4. Direct link between CH4 production and phytoplanktonmetabolism. Relationship between the maximum photosyntheticactivity (Pmax; µmol C L−1 h−1) or maximum nitrogen fixationrates (N2max; nmol L−1 h−1) and surface CH4 production rates(nmol C L−1 h−1) from dissolved inorganic carbon (a, b), methylgroup of methionine (c) and acetate (d).

positive relationship between CH4 production from methio-nine in the light and the photosynthetic activity was found(Fig. 4c).

13C-labelled acetate, the substrate of acetoclasticmethanogenesis, supported the production of CH4 in alllakes with the exception of Lake Kyamwinga but at muchlower rates compared to light-dependent CH4 production

from DIC (50 times lower, n= 7) or methionine (10 timeslower, n= 4; Figs. 3c and S6). δ13C analysis of the DICin the bottles spiked with 13C-labelled acetate showedthat the acetate was mineralized at rates of 5–6 orders ofmagnitude higher than acetoclastic methanogenesis so thatadded acetate appeared to be used almost exclusively byheterotrophic microorganisms other than methanogens. Pat-terns of acetate-derived production of CH4 were similar inlight and dark treatments (Figs. 3c and S6), and this mode ofCH4 production appeared to be unrelated to phytoplanktonactivity (Fig. 4d).

3.4 Microbial methane oxidation

Net CH4 oxidation was detected in all five investigatedlakes ranging from 11 to 5212 nmol L−1 d−1 (Fig. 5), andwas by far the largest loss term of dissolved CH4 at theecosystem scale (8 to 46 times higher than the diffusiveemission to the atmosphere). Surface water CH4 turnovertimes were particularly short in the shallow and eutrophicLake George (2 h) and Lake Nyamusingere (3 h) and slightlylonger in the deeper and less productive Lake Katinda (11 h),Lake Kyamwinga (77 h) and Lake Edward (100 h). In allstudied lakes, the volumetric CH4 oxidation rates were al-ways much higher than the volumetric CH4 production wemeasured during the stable-isotope tracer experiments, re-gardless of the CH4 precursors tested. Pelagic CH4 produc-tion rates represented 8.5 %, 2.6 %, 0.2 % and 0.1 % of CH4oxidation rates in lakes Edward, George, Katinda and Nya-musingere, respectively.

The influence of light on methanotrophy was investi-gated in the deep Lake Edward and shallow Lake Georgeand Lake Nyamusingere, revealing that CH4 oxidationrates decreased dramatically with increasing light intensity(Fig. 5). For instance, when exposed to full sunlight inten-sity, methanotrophs consumed only 42 % (Lake Edward),

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Figure 5. Light inhibition patterns of CH4 oxidation in surface waters. (a) CH4 oxidation rates (nmol L−1 d−1) measured in the dark in thesurface waters (0.3 m) of a variety of African tropical lakes. (b) Relationship between illumination (fraction of incident sunlight irradiance;%) and CH4 oxidation inhibition (fraction of CH4 oxidation in the dark, inhibited at a given irradiance; %) in Lake Edward, Lake Georgeand Lake Nymusingere. Symbols represent the mean, and error bars represent the maximum and minimum of duplicate experiments.

54 % (Lake Nyamusingere) or 74 % (Lake George) of theCH4 they were able to oxidize in the dark. The magnitudeof this sunlight-induced inhibition decreased substantiallywith decreasing sunlight intensities (Fig. 5). In Lake Edward(April 2017) sunlight inhibition of methanotrophy followedthe same pattern of lower rates at high sunlight intensitiesin the bottles where O2 production (via photosynthesis) wasstopped by DCMU addition.

4 Discussion

4.1 Mechanisms of CH4 production under aerobicconditions

The results from the stable-isotope labelling experimenthighlight that only a minimal fraction of the CH4 pro-duced under aerobic conditions originated from acetate incontrast with several earlier studies (Bogard et al., 2014;Donis et al., 2017), which proposed, based on the ap-parent fractionation factor of δ13C–CH4, that acetoclasticmethanogenesis linked to phytoplankton production of or-ganic matter would be the dominant biochemical pathwayof pelagic CH4 production in oxic fresh waters. Instead, ourresults support the study of Bizic et al. (2020) and sug-gest that epilimnetic CH4 production in well-oxygenatedconditions was related to DIC fixation by photosynthesis(Fig. 3) and correlated to primary production (Fig. 4a) andN2 fixation (Fig. 4b). When normalized to POC concentra-tions, the average DIC-derived CH4 production rates (0.08±0.05 nmol mmol−1

POC h−1 n= 7) were remarkably similar tothe CH4 production rates recently reported in Cyanobacte-ria cultures (0.04±0.02 nmol mmol−1

POC h−1) grown at 30 ◦C,among which the freshwater Microcystis aeruginosa (Bizicet al., 2020), the dominant Cyanobacterium species in thetropical lakes investigated in our study (see Fig. S2). TheseCH4 production rates are 2 orders of magnitude higher thanrates reported in an axenic culture of the eukaryote Emilia-

nia huxleyi (0.19± 0.07 pmol mmol−1POC d−1; Lenhart et al.,

2016), but they are 4 orders of magnitude lower than typi-cal anoxic CH4 production rates by methanogenic Archaea(Mountfort and Asher, 1979). Although it seems improba-ble that 13C–DIC acted as a direct precursor molecule forthe CH4 released by phytoplankton (Lenhart et al., 2016;Klintzsch et al., 2019), 13C–DIC could have been taken upby phytoplankton cells and then used as a C source for thesynthesis of many different organic molecules that may serveas the actual CH4 precursors. Indeed, healthy phytoplank-ton cells actively release a variety of low-molecular-weightmolecules, which are generally highly labile and rapidly con-sumed (Baines and Pace, 1991; Morana et al., 2014). Phy-toplankton metabolism could have fuelled CH4 productionpathways, at least partially, excreting substrates involved inCH4 production via biochemical processes such as demethy-lation of a variety of organic molecules like methionine,one of the S-bonded methylated amino acids (Lenhart et al.,2016), trimethylamine (Bizic et al., 2018) or methylphospho-nate (Yao et al., 2016).

While the source of methylphosphonate in fresh waters isobscure, and its actual natural abundance remains to be de-termined, dissolved free amino acids would represent up to4 % of the DOC produced by phytoplankton and are rapidlyconsumed by heterotrophic bacteria (Sarmento et al., 2013).Our incubations indeed demonstrated that the methyl groupof methionine was a potential precursor of CH4 in all lakesinvestigated, in line with recent findings showing that Emil-iania huxleyi could act as a direct source of CH4 in oxic con-ditions using methionine as a precursor without the involve-ment of any other microorganisms (Lenhart et al., 2016).We found that CH4 production from methionine was clearlystimulated under light, even when photosynthetic activitywas inhibited by DCMU, while little CH4 from methio-nine was produced in darkness (Fig. 3b). DCMU notablyprevents reduction in plastoquinone at photosystem II andgenerates singlet oxygen (Petrillo et al., 2014). The mecha-

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5218 C. Morana et al.: Sedimentary fluxes rather than pelagic production

nism of CH4 production from methionine is still unclear, butits residue in proteins is particularly sensitive to oxidationto methionine sulfoxide by radical oxygen species (ROSs;Levine et al., 1996) so that methionine would act as an ef-fective ROS scavenger and play important protective rolesunder photo-oxidative stress conditions, as shown in vascu-lar plants (Bruhn et al., 2012). The side chain of methioninesulfoxide is identical to dimethyl sulfoxide, which is knownto react with hydroxyl radicals (OH) to form CH4 (Repineet al., 1979). Besides its photoprotective role for phytoplank-ton, methionine could also be catabolized by a wide varietyof microorganisms to methanethiol, which could in turn betransformed to CH4 as shown in Arctic Ocean surface waters(Damm et al., 2010). Nevertheless, occurrence of this lattermechanism in the tropical lakes investigated seems unlikelyas this mode of CH4 production would be expected to be in-sensitive to light irradiance, and no CH4 was produced frommethionine in the dark during the incubations.

4.2 Relevance of pelagic CH4 production compared tomethanotrophy and CH4 emissions at theecosystem scale

The stable-isotope labelling experiments revealed that micro-bial CH4 oxidation largely exceed the pelagic CH4 produc-tion. All of the major sources and sinks of CH4 at the ecosys-tem scale were experimentally determined offshore in threelakes (lakes Edward at 20 m depth, George at 2.5 m and Nya-musingere at 3 m; Fig. 6). In contrast with high-latitude lakeswhere no ebullition could be detected in locations with a wa-ter column depth higher than 2 m (DelSontro et al., 2018), wefound that the CH4 ebullition flux dominates over the CH4diffusion flux in every sampling site. However, due to the dis-solution of arising bubbles, the contribution of ebullition tothe total CH4 emissions might be lower in deeper (>20 m)locations of Lake Edward (maximum depth of 113 m), asshown elsewhere (DelSontro et al., 2015). Comparison of theCH4 production, consumption and emission fluxes (Fig. 6)shows that the depth-integrated CH4 production rates deter-mined from diverse precursor molecules were modest rela-tive to the diffusive CH4 efflux to the atmosphere and thedepth-integrated microbial CH4 oxidation. In opposition, thecombined CH4 bubble dissolution flux and diffusive benthicCH4 flux were several orders of magnitude higher than CH4production in surface waters and were sufficient to supportthe microbial CH4 oxidation and the emissions to the atmo-sphere(Fig. 6). These results gathered in tropical lakes of var-ious size (from 0.44 to 2300 km2) and depth are in sharpcontrast with the estimation of an empirical model (Gun-thel et al., 2019) which proposed that mechanisms of oxicCH4 production represents the majority of CH4 emissions inlakes larger than 1 km2. This discrepancy highlights the needto consider the unique limnological characteristics of a vastregion of the world that harbours 16 % of the total surfaceof lakes (Lehner and Döll, 2004). One of the most distinc-

Figure 6. Epilimnetic CH4 production is a marginal source of CH4compared to sedimentary sources and CH4 sinks in several con-trasting African lakes. Summary of the different CH4 flux experi-mentally measured in Lake Edward, Lake George and Lake Nya-musingere. Values of CH4 oxidation in brackets are values not con-sidering CH4 photoinhibition. Pelagic CH4 production are valuesdetermined from NaH13CO3 (∼ 5 % final enrichment) and 13C–acetate (99 % final enrichment), as described in the methods sec-tion. 13C-labelling experiment carried out under constant light ir-radiance. CH4 flux at the water–air and sediment–water interfacewas determined experimentally as described in the methods section.CH4 bubble dissolution and CH4 ebullition flux were determinedusing the SiBu-GUI software (Greinert and McGinnis, 2009); mini-mum and maximum represent the values obtained from two extremebubble-size scenarios considering a release of many small (3 mm di-ameter) bubbles or fewer large (10 mm) bubbles.

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C. Morana et al.: Sedimentary fluxes rather than pelagic production 5219

tive features of the tropical aquatic environment is the per-sistent elevated water temperature in the hypolimnion and atthe water–sediment interface, which favours methanogenicactivity in sediment and decreases CH4 solubility, enhancingbubble formation.

Epilimnetic CH4 production appeared as a marginal fluxat the ecosystem scale and could not explain alone the accu-mulation of 13C-depleted CH4 in the epilimnion of most ofthe lakes of our dataset (Figs. 1, 2), for which we proposea combination of two other alternative mechanisms: dissolu-tion of arising CH4 bubbles in the epilimnion combined withinhibition by light of CH4 oxidation. The partial dissolutionof the CH4 bubbles as they rise in the epilimnion would al-low a rapid transport of 13C-depleted CH4 from the sediment,bypassing the hotspot of CH4 oxidation at the sediment–water interface and representing an alternative source of 13C-depleted CH4 in water column. This mechanism of bubble-mediated transport to the epilimnion would be especially im-portant in shallow lakes as CH4 ebullition is widely vari-able as a function of water column depth (DelSontro et al.,2015). The shallower Lake George and Lake Nyamusingerewere notably characterized by sharp thermal-density gra-dients (Fig. 2) and extreme phytoplankton biomass largelydominated by Microcystis aeruginosa (chlorophyll a up to190 µg L−1). Microcystis aeruginosa cells form large aggre-gates (>1 mm) embedded in a matrix of extracellular poly-meric substance that might act as a barrier to trap small CH4bubbles arising from the sediment (Fig. S7 in the Supple-ment). Dissolution of CH4 bubbles could be enhanced at thevery near surface due to the entrapments of bubbles at the air–water interface by abundant surface organic films that delaythe bubble “burst”. The presence of a sharp subsurface tem-perature gradient would further enhance CH4 accumulationduring daytime near the air–water interface (by blocking ver-tical redistribution of CH4 by mixing). We hypothesize thatthis process could be widespread in shallow tropical lakes,which are characterized by high productivity and are simul-taneously susceptible to being large benthic CH4 sources.

The stable-isotope labelling experiment carried out to in-vestigate the inhibitory effect of light on methanotrophic ac-tivity demonstrated that CH4 consumption dramatically de-creased with increasing light intensities (Fig. 5), as alreadyreported in a tropical reservoir (Dumestre et al., 1999) andLake Biwa (Murase and Sugimoto, 2005). The inhibitory ef-fect of light on freshwater methanotrophs remains surpris-ingly understudied since it was first reported 20 years ago(Dumestre et al., 1999), so the physiological mechanismof photoinhibition is still not understood. The physiologicalmechanism of photoinhibition of CH4 oxidation could be re-lated to the fact that the copper-containing methane mono-oxygenase enzyme and be structurally close to the ammo-nia mono-oxygenase enzyme, and it might be inactivated byROS produced during photo-oxidative stress, as shown forammonium oxidizers (French et al., 2012; Tolar et al., 2016).Altogether, our results emphasize the role of sunlight irradi-

ance as an important but frequently overlooked environmen-tal factor driving the CH4 dynamics in lake surface watersand possibly contributing to the occurrence of 13C-depletedCH4 in surface waters.

Data availability. All data included in this study are available uponrequest by contacting the corresponding author.

Supplement. The supplement related to this article is available on-line at: https://doi.org/10.5194/bg-17-5209-2020-supplement.

Author contributions. This study was designed by CM, AVB andSB. All authors participated in sample collection, data acquisitionand analysis. The manuscript was drafted by CM and substantiallyimproved with the input of all authors. All authors approved thefinal version of the paper.

Competing interests. The authors declare that they have no conflictof interest.

Acknowledgements. We are grateful to Marc-Vincent Commarieuand Dries Grauwels for their help in the lab and during the fieldsampling. We also thank the Uganda Wildlife Agency for researchpermission in the Queen Elizabeth National Park (Uganda), the staffof the Tembo Canteen for the use of their incubation room andthe crew of the Katwe Marine research vessel for their help dur-ing the Lake Edward sampling. We also thank the associate editor(Tyler Cyronak), D. F. McGinnis and one anonymous reviewer, whohelped to improve the initial version of this paper. This work wasfunded by the Belgian Federal Science Policy Office (BELSPO;HIPE project, BR/154/A1/HIPE) and by the Fonds Wetenschap-pelijk Onderzoek (FWO-Vlaanderen, Belgium), with travel grantsawarded to Cédric Morana and Steven Bouillon. Alberto V. Borgesis a research director at the Fond National de la Recherche Scien-tifique (FNRS; Belgium).

Financial support. This research has been supported by the Bel-gian Federal Science Policy Office (grant no. BR/154/A1/HIPE).

Review statement. This paper was edited by Tyler Cyronak and re-viewed by D. F. McGinnis and one anonymous referee.

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