Introduction

Substantial increases in tropospheric CH4 and CO2 over the past decades and the attendant climatic implications have prompted renewed studies on the productivity and emission of greenhouse gases from marine and terrestrial ecosystems.

Estuaries, where marine and fresh waters mix, are one of the largest sources of atmospheric CO2. The high production of CO2 is a diagenetic consequence of the high delivery and remineralization of organic and inorganic C from upstream rivers and from intense bacterial activity within estuaries (Bauer et al. 2013). Estuaries release from 0.27 to 0.60 Pg CO2–C y−1 (Borges and Abril 2011), which is more than the emission from rivers and lakes (Borges et al. 2006; Cole et al. 2007) as well as wetlands (Bauer et al. 2013). In contrast, the contribution of estuaries to the global CH4 budget is uncertain but assumed to be minor (Kirschke et al. 2013). Estimates of total estuarine CH4 emissions vary but are in the range of 0.8–6.6 Tg CH4 y−1 (Borges and Abril 2011 and references therein). For comparison, wetlands and freshwater lakes and rivers emit 217 and 40 Tg CH4 y−1, respectively (Kirschke et al. 2013). The low contribution of estuaries to global CH4 production and emissions is caused by the effect of salinity on the activity on methanogenic archaebacteria (MA). It is well established for saline environments, replete in dissolved sulphate, that the MA are outcompeted in the sulphates reduction zone (SRZ) by sulphate-reducing bacteria (SRB) for many methanogenic substrates (Whiticar 1999). In addition, anaerobic CH4 oxidation by SRB in the SRZ effectively consumes CH4 that fluxes into the SRZ from deeper layers or is formed locally from non-competitive substrates. Consequently, in saline waters, the zone of methanogenesis and CH4 accumulation is displaced to deeper sediment layers beneath the SRZ. In these environments CH4 is typically produced from more refractory substrates, including hydrogenotrophic methanogenesis (CO2 reduction).

Due to their small global areal footprint, estuaries are thought to represent only a small fraction of the global atmospheric CH4 emissions. However, individual estuaries can be highly methanogenic. For example, Bange (2006) and Harley et al. (2015) reported high CH4 saturations in estuarine waters from 62 to 25,700,000% along the southwestern Baltic coast and from 100 to 13,100% along the north-western North Sea coast. The biogeochemical literature contains extensive discussion of the potential impact that estuaries will have in response to climate change and the sea level rise (Kjerfve et al. 1994; Scavia et al. 2002; Borges et al. 2006; Day et al. 2008). However, due to uncertainties with regard to the present day emissions of CH4 and CO2 from estuaries the prediction of future emissions of these gases is very complex (Borges and Abril 2011 and references therein). The uncertainties in the CH4 and CO2 fluxes mainly derive from divergent evaluations of the area covered by estuaries throughout the world, as well as from the complexity of the estuarine systems (Borges and Abril 2011).

Spatial and temporal variability in salinity and hydrodynamic conditions in estuaries make it challenging to select representative sites for CH4 and CO2 measurement. There is a need to make measurements at diverse field sites to recognize the controls and distribution of microbial gas production within estuaries. On the Polish, southern coast of the Baltic Sea, estuaries are typically coastal lakes. Cieśliński and Drwal (2005) pointed out that these lakes act as quasi-estuaries because there are no tides on this coast and the lakes are only episodically inundated by brackish waters inflowing from the sea during specific anemometric conditions. Woszczyk et al. (2011) suggested that these lakes are potentially a source for CH4. However neither saturation levels nor CH4 fluxes out of the lakes have been determined so far.

This paper provides the first comprehensive study on the sediment gases produced in coastal lakes. The aim of this study is to extend the knowledge of the genesis and distribution of microbial gases in coastal lakes, with special attention to (a) effects of salinity on gas compositions and spatial distribution; (b) climate-related seasonality in the compositions of sediment gases; (c) pathways of methanogenesis in the saline-freshwater mixing zone; and (d) CH4 oxidation in these shallow water and dynamically mixed environments.

We investigated the bubbles of gas released from the lake sediments via ebullition, the most efficient pathway of gas emission from shallow (<50 m) and turbulent water bodies (Engle and Melack 2000; Kankaala et al. 2004; McGinnis et al. 2006; Del Sontro et al. 2010; Bastviken et al. 2011). Our approach was to combine the analysis of the molecular and stable isotope compositions of the gas collected in seven lakes along the southern Baltic area over a study period of several months. We provided evidence that the molecular and isotopic compositions of sediment gas show reasonable spatio-temporal variability and that they are prone to substantial transformation in the water column.

Methods

Study area

The lakes investigated are located along the 150 km-long section of the Poland Baltic Sea coast (Fig. 1). These coastal lakes were formed by isolation from the brackish Baltic Sea in the early-middle Atlantic period (Miotk-Szpiganowicz et al. 2008) and thus preserved some features of lagoons. The lakes are relatively large and shallow (mean water depth is less than 2 m) (Table 1) and are oriented parallel to the predominant wind directions (W–SW). These features result in high sensitivity of coastal lakes to wind-driven turbulence (Woszczyk et al. 2014). Consequently, the surface and bottom waters of the lakes were highly oxygenated throughout the sampling period. The dissolved O2 varied from 0.14 to 0.47 μmol L−1, and O2 saturation was between 51 and 165% (Table 1). The highest oxygenation occured during winter, when the highest wind speeds (up to 24 m s−1 in Ustka) were observed, and the minimum of dissolved O2 was obtained in summer. Due to the connection to the Baltic Sea, most of the lakes have brackish conditions, however, the salinity has appreciable spatial and temporal variability (Table 1). The salinity (expressed as total dissolved solids (TDS)) ranged from 132 to 5542 ppm, but the majority of observations had TDS < 1000 ppm. The coastal lakes have high biological productivity, indicated by high chlorophyll-a and low Secchi disc visibility (Table 1) throughout the study period.

Fig. 1
figure 1

Location of the seven lakes studied and distribution of sample sites in each lake

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Table 1 Location and physical–chemical characteristics of the studied lakes
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Sampling

The lake sediment gas was sampled every 2–3 months between May 2013 and October 2014. In each lake, there were 2–6 sample stations (Fig. 1) located in different parts of the lakes displaying differences in salinity, water depth, macrophyte cover, and wind fetch (and thus wave action). The lowest salinity (TDS between 132 and 250 ppm) occurred throughout Lake Wicko and Lake Jamno, while Lake Resko and Lake Łebsko had the highest salinity (TDS > 1000 ppm). The water depth varied between c.a. 50 cm in Lake Kopań (site KOP1 and KOP4), Lake Wicko (WIC1), and Lake Jamno (JAM1) to >300 cm in Lake Łebsko (LEB5) and Lake Wicko (WIC8). Dense macrophyte cover was developed in western parts of Lake Wicko (WIC1), Lake Kopań (KOP1, KOP4), and Lake Bukowo (BUK6). WIC1, KOP1, KOP4 and BUK6 were located in wind-protected embayments and had low wind wave action. The highest wind wave action occurred in the sites located in the eastern (upwind) parts of the lakes studied (JAM8, BUK8, WIC8, GAR2, LEB5, and LEB4). The sediment gas was sampled from an anchored boat. The gas was released by gentle agitation of the sediments with a hammer suspended from a string. Bubbles escaping from the sediment were trapped at the lake surface with a funnel and by displacement into 0.5 L dark glass bottles filled prior to sampling with lake water. After sampling, some water (~20–25 mL) was left in the neck of the sealed bottle to minimize gas escape. The surface 5 cm layer of lake sediments was sampled using a gravity sampler (Limnos). The physical–chemical properties of lake waters (temperature, TDS, pH, Red-ox potential, dissolved O2) were determined in situ with HI 9828 (Hannah Instruments) and HQD40 (Hach–Lange) probes. Measurements were made at the lake surfaces (5–10 cm below the water surface), near-bottom (5–10 cm above the lake bottom), and in sediment pore waters (uppermost 5–10 cm of sediments). The Redox potential measured in situ was corrected with the reference potential of the sensor to obtain the potential of a standard hydrogen electrode (Eh). Water depth was measured with a Humminbird 150 sonar. Wind action was determined only approximately in terms of wind fetch (F) and the depth of water mixed layer (DWML). The higher F and DWML, the deeper wave motion in lake water column. For calculating F and DWML we used formulae proposed by Douglas and Rippey (2000).

Analytical procedures

The bulk composition of lake sediments was analyzed using lyophilized and powdered samples, as were the stable C isotope ratio measurements of TOC (total organic carbon, δ 13CTOC). The total carbon (TC), total nitrogen (TN), and total sulfur (TS) contents were determined by an elemental analyser (EA) (Elementar VarioMax CNS). TOC was also determined by an EA, except that prior to the measurement carbonates were removed by 1 M HCl. The total inorganic carbon (TIC) was calculated as TIC [wt%] = TC−TOC. The TOC/N was calculated on a molar (at./at.) basis. The samples were analysed in duplicate. Quality control was performed using certified reference materials for peaty (BN225497), chalky (BN230229), and sandy soils (BN230227) (Elementar).

For δ 13CTOC, the carbonate fraction was removed beforehand with 10% HCl and rinsed with water until pH 7. Carbon isotope ratio analyses of TOC were performed with a Flash EA 2000 connected online to a ThermoFinnigan Delta V Plus mass spectrometer. All carbon isotope ratios are reported in the usual δ-notation (δ 13CTOC) relative to Vienna Pee Dee Belemnite (VPDB). The accuracy and reproducibility of the analyses were checked by replicate analyses of laboratory standards calibrated to international standards USGS 40 and 41. The δ 13CTOC reproducibility was ±0.2‰.

The molecular composition of hydrocarbons (HC) in the sediment gas bubbles was analysed with a Fissons Instrument GC 8160 gas chromatograph with a flame ionization detector (FID-GC). The volumetric concentrations (vol%) of CH4, CO2, O2, N2, and higher hydrocarbons (HHC ≈ ∑C2+) were measured using a Carlo Erba Instruments GC 6300 gas chromatograph. Details of the procedure are given by Sechman et al. (2009). The precision of concentration measurements was equal: HC ± 0.1 ppm; CO2: ±0.01 vol%; O2: ±0.02 vol% for concentration <5 vol% and ±0.1 vol% for concentration >5 vol%; and N2: ±0.05 vol% for concentration <10 vol% and ±0.1 vol% for concentration >10 vol%. The CHC (Bernard parameter) and CO2–CH4 index (CDMI), used to decipher the origin of sediment gases (Bernard et al. 1978; Kotarba and Nagao 2015), were calculated as follows:

$$C_{HC} = \frac{{CH_{4} }}{{\left( {C_{2} H_{6} + C_{3} H_{8} } \right)}}$$
([1])
$$CDMI (\% ) = \frac{{CO_{2} }}{{\left( {CO_{2} + CH_{4} } \right)}} \times 100$$
(2)

where CH4, C2H6, C3H8, and CO2 denominate volumetric percentages of the gases (Kotarba and Nagao 2015).

CH4 and CO2 were isolated chromatographically for carbon and hydrogen isotope analyses. CH4 was combusted over hot CuO (850 °C) and converted to CO2. Stable isotope analyses were performed using Finnigan Delta Plus and Micromass VG Optima isotope ratio mass spectrometers (IRMS). The stable C isotope data are expressed in the δ-notation (δ 13C, ‰) relative to VPDB, taking the reference NBS-22 (oil) to be −30.03‰. The stable H isotope data are reported in δ-notation (δ 2H, ‰) relative to the international standard, VSMOW (Vienna-standard mean ocean water = 0.0‰), and linearized to standard light arctic precipitation (SLAP) as recommended by Coplen (2011). Molecular N2 was separated chromatographically for stable N isotope analyses and was measured by on-line IRMS. The results of stable N isotope analyses were presented in δ-notation (δ 15N, ‰) relative to the air N2 standard. Analytical precision is estimated to be ±0.2‰ for C, ±3‰ for H2, and ±0.4‰ for N2.

Statistical calculations were performed with Statistica 12 (StatSoft, Inc. 2014).

Results

Sediments

The sediments containing CH4 and CO2 bubbles had TOC contents from 7.5 to 21.1 wt%, and the TIC was between 0.4 and 4.9 wt%. The contents of TS were 1.0–2.6 wt%. The molar TOC/TN was between 7.9 and 10.0 (Table 2). The stable C isotope signatures in bulk organic carbon (δ 13CTOC) varied from −29.1 to −26.6‰, except for Lake Kopań which displayed less negative values of −24.7 and −24.4‰ (Table 2). There were appreciable differences in the sediment composition between the lakes and within each lake. The highest TOC values were obtained in Lake Wicko. Lake Gardno was distinguished by the highest TIC and the lowest TOC contents. Throughout the study period, the surface sediments were anoxic (dissolved O2 = 0 μmol L−1) and displayed appreciable seasonal changes in Eh. During cold periods (Nov. 2013–Feb. 2014) the Eh was positive (+8 to +181 mV) at all but one study site, while during the warm season (Aug. 2014) Eh dropped to negative values (−32 to −255 mV).

Table 2 Bulk chemical composition of surface sediments of the lakes studied
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Molecular composition of sediment gas

Up to 16 gas species were identified in the sediments, with CH4, N2, and CO2 being by far the most abundant (see Table 3 in Appendix). The remaining fraction was comprised of O2 and HHC from saturated (C2H6), unsaturated (C2H4), and aromatic HC along with single-chain and branched chain compounds (see Table 4 Appendix). On average, the relative abundances were CH4 ~72.0 vol%, N2 ~23.8 vol%, CO2 ~2.6 vol%, and O2 ~1.2 vol%. The relative abundances of CH4 were strongly negatively correlated to N2 (r = −0.98; n = 114; p < 0.001) and showed an overall positive relationship with CO2 (r = 0.54; n = 114; p < 0.001).

Table 3 Molecular and stable isotope compositions of sediment gas from Polish coastal lakes
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Table 4 Molecular composition of higher hydrocarbons in the sediment gas from Polish coastal lakes
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Although the mean CH4 relative abundances for the lakes were similar (see Table 3 in Appendix), within each lake the CH4 displayed variations with water depth. In most lakes, except for Lake Łebsko and Lake Gardno, the concentrations of CH4 increased with decreasing water depth (Fig. 2a). The CO2 and CDMI showed differences between lakes. By far the highest values of both parameters occurred in Lake Kopań and the lowest occurred in Lake Gardno (see Table 3 in Appendix). Within lakes the changes in CO2 were negatively related to water depth (Fig. 2b), while CDMI was different in vegetated and non-vegetated sites (Fig. 3a). In the former the CDMI was higher (i.e. there was more CO2 in relation to CH4—4.72 on average) than in the latter (2.04 on average). The CDMI displayed a weak seasonal pattern, with a slight tendency to increase with increasing water temperatures (Fig. 3b).

Fig. 2
figure 2

CH4 (A) and CO2 (B) concentrations versus lake water depth for the seven lakes studied. In most lakes, except for Lake Łebsko and Lake Gardno, there is an overall negative relationship between these parameters

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Fig. 3
figure 3

The CO2–CH4 index (CDMI) differences between vegetated and unvegetated sites (A) and the relationship between the CDMI values and lake water temperature (B). The horizontal lines in A indicate average CDMI for vegetated and unvegetated sites (mean veg and mean unveg, respectively). Unvegetated sites in Lake Resko are represented by one measurement

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The composition of HC (in terms of CHC ratio) displayed some spatial and temporal variation. The highest values of CHC occurred in Lake Wicko, and Lake Gardno, while the lowest values were obtained in Lake Resko and Jamno. The mean CHC showed overall (statistically insignificant) negative relationship to salinity (Fig. 4a) and temperature (Fig. 4b) i.e. the high values of CHC (>20,000) occurred at low salinity (<2000 ppm) and low temperature (<5 °C).

Fig. 4
figure 4

Bernard index (CHC) versus lake water salinity (expressed as total dissolved solids, TDS) (A) and temperature (B). The CHC was low throughout major parts of the TDS and temperature ranges except for TDS <2000 ppm and temperatures <5 °C. At low salinity and temperature the CHC increased sharply

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Stable isotope composition of sediment gas

N2

The δ 15N2 signatures showed minor changes from −2.1 to +1.7‰ (see Table 3 in Appendix). The δ 15N2 showed differences between lakes. The highest values were measured in Lake Jamno, and Lake Gardno and the lowest in Lake Bukowo, and Lake Łebsko.

CO2

The δ 13CCO2 varied over a very broad range from −51.5 to +6.0‰ (see Table 3 in Appendix), with a slight tendency towards higher values during colder seasons (Fig. 5). There was also a clear increasing trend in δ 13CCO2 values with decreasing water depth (except for Lake Łebsko) (Fig. 5) and increasing CO2 concentrations (Fig. 6). The δ 13CCO2 values obtained in vegetated sites were higher than in non-vegetated sites (Fig. 7).

Fig. 5
figure 5

Stable C isotope composition of CH4 and CO2 in the lakes studied during different sampling campaigns and at different depth intervals. Lake water depth is indicated by grey dashed lines. The number attached to the lines denominate the lake water depth in cm. In Lake Łebsko there was no relationship between the stable isotope composition of gases and water depth, and therefore the depth lines were not shown. The grey numbers next to data points show the collection depth of outliers. Note that there is an overall tendency for δ 13CCO2 and δ 13CCH4 to increase with decreasing water depth. αc denominate apparent C isotope fractionation factors (according to Whiticar 1999). Stable C isotope composition indicate that the CH4 is predominantly produced via acetoclastic methanogenesis (AM) with only minor contribution of hydrogenotrophic methanogenesis (HM). In a few sites the CH4 is prone to oxidation (OX). The colors indicate seasons: white—winter, green—spring, red—summer, yellow—autumn

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Fig. 6
figure 6

Stable C isotope signatures of CO2 (δ 13CCO2) versus concentration of CO2 in the sediment gas. There is a clear negative relationship between δ 13CCO2 and the CO2 concentration, explained by CO2 removal from rising bubbles owing to dissolution and gas exchange. The more strongly 13C-depleted δ 13CCO2 values can only be explained by methanotrophy. The 13C-enriched δ 13CCO2 values extending to +6‰ were obtained in shallowest parts of the lakes and resulted from methanogenesis. The δ 13CCO2 values centered around −25‰ can be derived from a variety of sources and processes, including the respiration of CO2 from organic matter

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Fig. 7
figure 7

Stable C isotope signatures of CO2 (δ 13CCO2) and CH4 (δ 13CCH4) in vegetated and non-vegetated sites. The former display higher δ 13CCO2 and δ 13CCH4 signatures, which provides evidence for stable C isotope fractionation accompanying bioventilation. Mean δ 13CCO2 and δ 13CCH4 for vegetated (δ 13CCO2veg; δ 13CCH4veg) and unvegetated (δ 13CCO2unveg; δ 13CCH4unveg) sites are shown by horizontal and vertical dashed lines, respectively

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CH4

The δ 13CCH4 values ranged from −77.9 to −55.1‰ (see Table 3 in Appendix) and showed differences between lakes (Fig. 5). The highest mean δ 13CCH4 values were obtained in Lake Kopań, while the lowest were measured in Lake Bukowo, and Lake Resko. The mean δ 13CCH4 revealed overall seasonality. A decrease in δ13C in CH4 occurred in warmer seasons (δ 13CCH4 for summer months was between −77.9 and −55.4‰), while during autumn and winter the CH4 become isotopically heavier (δ 13CCH4 between −71.1 and −57.4‰).

The δ 13CCH4 showed changes within each lake. Except for Lake Łebsko and Lake Resko, spatial and temporal changes in δ 13CCH4 were strongly related to the concentrations of CH4. The more abundant the CH4, the higher the δ 13CCH4 signatures (the r values for this relationship are significant at p < 0.01 and range from 0.71 to 0.88). In addition, in most of the lakes studied (except for Lake Gardno and Łebsko) the δ 13CCH4 revealed an overall negative relationship to water depth (Fig. 5), and the correlation coefficients for this relationship were between −0.52 and −0.85 (at p < 0.04). We obtained differences between the δ 13CCH4 signatures in vegetated and non-vegetated sites (Fig. 7): −63.6‰ for vegetated sites and −67.2‰ for unvegetated sites.

The δ 2HCH4 values varied between −378 and −261‰ (Fig. 8). There were no clear differences between the lakes.

Fig. 8
figure 8

Combination of δ 13CCH4 and δ 2HCH4 to delineate CH4 sources. The δ 2HCH4 and δ 13CCH4 indicative of acetoclastic methanogenesis (AM). Genetic fields after Whiticar (1999). AM acetoclastic methanogenesis, HM hydrogenotrophic methanogenesis

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Discussion

Factors affecting the concentration of CH4, CO2, N2, and O2 in sediment gas

The collected gas was enriched in N2 by up to 51%, and contained appreciable amounts of O2 (up to 5.5%). Furthermore, the δ 15N2 signatures were too high for N2 produced via denitrification, a highly fractionating process resulting in δ 15N2 between −40 and −5‰ (Talbot 2001). Instead, the δ 15N2 values obtained were close to the values in atmospheric N2 (in which δ 15N2 is ~0‰; Talbot 2001). N2 exchange between the atmosphere and water is accompanied by a very low fractionation (Talbot 2001), and consequently dissolved N2 preserves the δ 15N2 values of atmospheric N2. Hence, the δ 15N2 values obtained argued for an atmospheric origin of the N2 occurring in the gas from the sediments of the lakes studied. Anoxia of pore waters indicated that the O2 present in the sediment gas was unlikely to be derived from processes occurring within the sediments (e.g. photosynthesis). It seems therefore that both N2 and O2 invaded the sediment gas via three possible routes.

First, the macrophytes act as conduits for gas exchange between the sediment and the atmosphere (Makhov and Bazhin 1999; Chanton et al. 2005; Laanbroek 2010). Plant ventilation was invoked to explain the sediment gas composition in the White Oak River estuary (Chanton et al. 1989). Chanton et al. (1989) demonstrated that N2 and CH4 were strongly negatively correlated and that sediment gas from densely vegetated sites was clearly N2-enriched compared with “purely” microbial gas. Despite the fact, that in our study the sampling stations were located in both vegetated and non-vegetated parts of the coastal lakes, there were no differences in N2 concentrations between these sites. In addition, the δ 15N2 signatures from vegetated and non-vegetated sites were similar. In the former, δ 15N2 ranged between −1.3 and 0.7‰ (−0.3‰ on average), while in the latter it ranged from −2.1 to 1.7‰ (−0.2‰ on average). Moreover, vegetated and non-vegetated sites had the same O2 distribution. These features indicated that the N2 and O2 concentrations were likely little affected by plant ventilation.

Second, N2 and O2 can be supplied to the sediments in the form of large bubbles entrapped by solid particles settling in aerated water after a storm (Boudreau 2012) and/or via diffusion from overlying aerated water (Chanton et al. 1989; Walter et al. 2008). These bubbles act as nucleation sites (heterogenous nucleation) for larger bubbles containing microbial gas. The CH4 and CO2 from methanogenesis are delivered to these initial bubbles via diffusion in the pore waters; and when the critical size of a bubble is achieved, ebullition commences. There are some data indicating that the sediments of coastal lakes are deeply mixed due to wind-induced turbulence. Woszczyk et al. (2014) established that in Lake Sarbsko (located 3 km east of Lake Łebsko) the mixing was up to 50 cm depth. Unpublished data (210Pb, 137Cs and 7Be distributions) from Lake Gardno, Lake Bukowo and Lake Resko showed mixing depth between 13 and 22 cm. These values supported the hypothesis on the potential effect of storm redeposition on air bubble entrapment in the sediments.

Third, the N2 and O2 can invade ascending gas bubbles in a lake water column due to exchange between CH4/CO2 and N2/O2 dissolved in the column (Leifer and Patro 2002; McGinnis et al. 2006; Ostrovsky et al. 2008; Yamamoto et al. 2009; Del Sontro et al. 2010; McGinnis et al. 2011). N2 and O2 build up in lake waters due to dissolution of air bubbles forced by breaking wind waves, Langmuir currents, and rain drop impacts (Chiba and Baschek 2010; Liang et al. 2011), and the concentration (supersaturation) of these gases is positively related to wind speed (Nakayama et al. 2002). At high wind speeds (>15 m s−1) the air–water gas exchange is predominantly via bubbles (Chiba and Baschek 2010). Coastal lakes on the Polish Baltic coast reveal very high vulnerability to wind mixing. The Baltic coast is the windiest part of Poland (Lorenc 1996), and in 2013/2014 there were between 8 (meteorological station in Koszalin; Fig. 1) and 105–111 days (meteorological station in Łeba and Ustka, respectively; Fig. 1) with a strong wind (with speed vw > 10 m s−1). In addition, due to the parallel orientation of coastal lakes to the predominant and the strongest W–SW winds, the wind fetch (F) and the depth of water mixed layer (DWML) for the lakes were both high. The F varied between 0.9 km for RP5 and 5.9 km for LEB4, and the DWML ranged from 0.5–2.8 m (at vw = 1 m s−1) to 2–6 m (at vw = 10 m s−1) which was higher than the water depth in most of the sampling stations. The obtained DWML values indicated that the whole water column in the lakes was mixed by wind.

The air entrapped in water can be transported to appreciable depths in the lake water column, sometimes to more than 20 m below the water surface (Farmer et al. 1993). A rough estimation of the depth of bubble entrainment by wind waves in the distal (eastern) parts of the lakes studied, according to the formula by Chiba and Baschek (2010), showed that during strong winds the bubbles could be drawn to c.a. 40 cm below the water surface. This value could be even higher when the Langmuir circulation was active (Langmuir circulation was observed in the lakes during field campaigns) because it is known to entrain air bubbles even deeper than wind waves (Liang et al. 2011). It is thus likely that in some parts of the coastal lakes there is a very deep air bubble penetration in the water.

The aeration of water during vigorous mixing leads to supersaturation with respect to N2 and O2. According to Liang et al. (2011), the degree of supersaturation is inversely related to the solubility of the gases. Therefore N2 tends to be more supersaturated than O2 and the N2/O2 ratio for equilibrium concentrations of both gases in water is c.a. 1.8–1.9 (at temperatures between 0 and 20 °C and 0‰ salinity) (Weiss 1970). However, the values obtained for N2/O2 of 3–7000 are much higher than the equilibrium. Taking into account that some O2 is added to lake waters by the photosynthesis of phytoplankton (Ford et al. 2002; Walter et al. 2008), it appears that O2 is significantly underrepresented, particularly during summer. Indeed, the mean seasonal values of N2/O2 displayed a positive correlation to lake water temperature (r = 0.70, n = 7, p < 0.08), and during the summer there were the lowest concentrations of dissolved O2 in the lake waters. A plausible reason for O2 depletion might be the enhanced oxidation of organic matter (Schubert et al. 2010; Woszczyk et al. 2011).

It seems that gas exchange between bubbles and surrounding water strongly affected the molecular composition of the sediment gas from the coastal lakes. This influence explains the strong negative correlation between CH4 and N2 (r = −0.98, n = 114, p < 0.001), as well as the overall negative relationship between the CH4 or CO2 and water depth observed in most of the lakes studied (Fig. 2). The greater the water depth, the longer the exposure of bubbles to highly aerated water and the higher the loss of the original gas composition.

The relationship between lake water depth and CO2 concentrations is stronger than the relationship between depth and CH4 concentrations (Fig. 2), which can be explained using the model by McGinnis et al. (2011). These authors claimed that CO2 “lifetime” in rising bubbles is very short compared with CH4 due to rapid dissolution of the former. Solubility of CO2 in water is c.a. 24 times higher than CH4 (Yamamoto et al. 1976), and consequently the concentration of CO2 in the bubbles decreases more rapidly than CH4 over the same distance in the water column.

However, the amount of CH4 exchanged seems surprisingly high in such shallow water bodies. Assuming an initial 0% contents of N2 and O2 and a predominance of CH4 and CO2 in the sediment gas, the fraction of lost CH4/CO2 varies between 10 and 67% (26% on average). Despite 30% bubble CH4 dissolution being reported from Lake Wohlen (Switzerland), a shallow (10 m on average) reservoir lake (DelSontro et al. 2010), up to 67% CH4 loss in coastal lakes seems high in relation to the depth of the sampling stations (maximum 3 m). On the basis of the model proposed by McGinnis et al. (2006), losses of CH4 are strongly dependent on the initial bubble size and the release depth. In our study we were unable to measure bubble diameters. However, for typical bubbles, 6–10 mm in diameter (McGinnis et al. 2006; Ostrovsky et al. 2008; Yamamoto et al. 2009), a trajectory 10–30 m long is required to cause such a depletion in CH4. On the other hand, for bubbles 20 mm in diameter, reported as being representative of gassy sediments (Haeckel et al. 2007), a dissolution of 10–67% CH4 would occur at water depths between ca. 20 and 150 m.

There are different possible explanations for the high loss of CH4 from the bubbles collected in shallow coastal lakes. It seems that the dissolution of bubbles is enhanced by the low concentrations of dissolved CH4 (CH4diss) in lake water (Yamamoto et al. 2009). In the lakes studied, the concentrations of CH4diss vary from 0.024 to 3.770 μmol L−1 (0.495 μmol L−1 on average) (unpubl. data) which is rather low compared with anoxic bottom waters of productive lakes where the CH4diss can be as high as several hundred μmol L−1 (Liu et al. 1996; Juutinen et al. 2008; Schubert et al. 2010). Moreover, the high CH4 loss may be triggered by the supersaturation with respect to N2 and O2 in intensely wind-agitated surface waters (Nakayama et al. 2002). In the seven coastal lakes studied, the saturation of O2 can be as high as 165% (Table 1). The concentrations of CH4 in sediment gas can be modified by the rate of ebullition, which is water depth-dependent. Engle and Melack (2000), Liikanen et al. (2003), Kankaala et al. (2004), and Boudreau (2012) suggested enhancement of CH4 ebullition under low hydrostatic pressure (i.e. in shallow waters). In turn, the intensity of bubble release is known to control the CH4/N2 ratio. Chanton et al. (1989) and Walter et al. (2008) showed that intense ebullition removes N2 from pore waters and results in higher CH4/N2 ratios in sediment gas. In some of the coastal lakes studied, there is an overall negative relationship between CH4/N2 and water depth (LW r = −0.62, n = 13, p < 0.03; LK r = −0.62, n = 14, p < 0.02; LJ r = −0.45, n = 17, p < 0.07 and LB r = −0.59, n = 15, p < 0.03), arguing for a more intense ebullition in the shallowest parts of these lakes. Chanton et al. (1989) and Walter et al. (2008) used the CH4/N2 ratio to document spatial and temporal changes in ebullition rate. Ebullition usually intensifies during summer (Makhov and Bazin 1999; Liikanen et al. 2003; Sapulveda-Jauregui et al. 2015). However, in the seven coastal lakes studied, the CH4/N2 displayed irregular temporal changes. We believe that the seasonality was obscured by the forcing of gas ebullition by wind action which is known to enhance bubble releasement (Leventhal and Guntenspergen 2004; Bussmann 2005). On the other hand, the distribution of CH4 seems to be related to the distribution of macrophytes. This is corroborated by differences in CDMI between vegetated and non-vegetated sites (Fig. 3a). The highest CDMI (the more CO2 in relation to CH4) in Lake Kopań are explained by the fact that a large part of the lake bed (sites KOP1 and KOP4) is covered by macrophyte meadows, with Potamogeton as the predominant species. Aquatic (mainly emergent) plants are known to channel considerable amounts of CH4 from the sediments (Chanton 2005; Borges and Abril 2011). CO2 is removed in much lower rates (Olsson et al. 2015). Consequently, bioventilation leads to the increase in the CDMI in the sediment gas. Emergent plants (i.e. Phragmites, Typha etc.) transfer CH4 directly from the rhizosphere to the atmosphere, while submersed plants (i.e. Potamogeton) transmit CH4 to the water column, where it is oxidized (Laanbroek 2010). It was shown that in pore waters, in densely vegetated sites, the concentrations of dissolved CH4 were c.a. 50% lower than those in non-vegetated areas. The lower CH4 concentrations reduce the CH4 bubble formation and gas ebullition (Chanton 2005).

The transformation of the molecular composition of sediment gas in the water column of the coastal lakes studied conflicts with the findings of Del Sontro et al. (2010) and Katsman et al. (2013). These authors claimed that ascending gas bubbles rise quickly through the sediment and water columns, essentially by-passing the zones of anaerobic and aerobic microbial oxidation and, as a consequence, only a small fraction of gas fluxing into the boundary layer is lost.

Higher hydrocarbons (HHC) in sediment gas

CHC values in the gas samples from Polish coastal lakes are typical for microbial gases produced via methanogenesis (Bernard et al. 1978). Trace amounts of HHC were previously reported from microbial systems in lakes (Oremland and Des Marais 1983; Woszczyk et al. 2011), estuaries (Oremland 1981), and the ocean (Hinrichs et al. 2006). However, to the best of our knowledge, there are few data showing the production of HHC in lake sediments, and the mechanisms of gas formation remain unclear. Hinrichs et al. (2006) showed that CH4 and HHC can be co-produced during bacterial degradation of OM and established that the production of HHC is favored under low salinity (low sulfate) conditions. Therefore, the increase in salinity would result in an increase in CHC, which conflicts with our observations (Fig. 4a). On the other hand, Whiticar and Faber (1986) found that changes in CHC values might be caused by selective removal of CH4 during anaerobic oxidation of hydrocarbons. CH4 shows lower resistance to this process than HHC (Whiticar and Faber 1986). In contrast, during aerobic oxidation HHC are preferentially decomposed compared with CH4 (Kinnaman et al. 2007). Thus the overall negative relation between the average CHC values and the lake water salinity in our lakes (Fig. 4a) may indicate that saltwater ingressions enhance anaerobic oxidation of CH4 in the sediments. This process is triggered by archaea in concert with SRB (Schouten et al. 2001). The latter were found to be one of the most predominant microbial communities in coastal lakes (Skórczewski and Mudryk 2005). The oxidation of CH4 becomes possible when the concentrations of dissolved sulfate are >1 mmol SO4 2− L−1 (Segers 1998). During the study period, the dissolved SO4 2− in the lakes varied between 0.17 and 3.9 mmol L−1 (unpubl. data), which made the CH4 oxidation thermodynamically possible.

In accord with previous reports (Utsumi et al. 1998; Whiticar 1999; Karim et al. 2011; Harley et al. 2015), the oxidation of CH4 is enhanced by higher water temperatures. This conclusion is supported by the decrease in CHC in warmer seasons (Fig. 4b) and a concomitant increase in CDMI (Fig. 3b). However, the increase in CDMI is not accompanied by a decline in δ 13CCO2, which would be expected during CH4 oxidation (Jędrysek 1999). It appears, therefore, that anaerobic oxidation plays only a minor role in the CH4 cycling in coastal lakes albeit its effect on the CHC may be strong. The concentrations of CH4 are 4–5 orders of magnitude higher that the HHCs and even a small decline in CH4 due to oxidation drives the CHC towards higher values.

Mechanisms of methanogenesis in coastal lakes

The two primary methanogenic pathways suspected in the lakes are acetoclastic methanogenesis (AM) and hydrogenotrophic methanogenesis (HM). Whiticar and Faber (1986) and others argued that in marine sediments HM predominates. This is due to the utilization of potential methanogenic precursor organic substrates by SRB in the sulfate reduction zone, before they reached the diagenetic depth for methanogenesis. To complicate the issue, aerobic and anaerobic CH4 oxidation consume some of the CH4, while the remainder can flux out of the lakes into the troposphere.

The CH4 production and consumption pathway can be recognized on the basis of the apparent C isotope fractionation factor between CO2 and CH4 (Whiticar 1999):

$$\alpha_{\text{C}} = \, \left( {\delta^{13} {\text{C}}_{\text{CO2}} \; + \;10^{3} } \right) \, / \, \left( {\delta^{13} {\text{C}}_{\text{CH4}} \; + \;10^{3} } \right)$$
(3)

as well as by using the δ 2HCH4 signatures. The δ 13CCO2 versus δ 13CCH4 relationships assume so-called co-existing pairs, i.e., that CO2 and CH4 are mechanistically linked. When this condition is fulfilled, αC > 1.065 is characteristic for HR, while αC < 1.055 indicates the predominance of AM (Conrad 2005; Hershey et al. 2014). On the other hand, in the C/H isotope systematics of CH4, the demarcation lines between AM and HM are made at δ 13CCH4 ~ −60‰ and δ 2HCH4 ~ −250‰ (Whiticar 1999). The lower values denote AM, while HM has δ 13CCH4 > −60‰ and δ 2HCH4 > −250‰. For CH4 oxidation, in turn, the C isotope fractionation factor is usually 1.005 < αC < 1.031 (Whiticar and Faber 1986).

In the previous section we showed that the molecular composition of gas in rising bubbles can change significantly in a lake water column. It seems, however, that the gases preserve their original stable C isotope signatures because the isotope effects associated with dissolving CH4 and CO2 in water are small (0.3–0.6 and 1‰, respectively) (Bergamaschi 1997; Sansone et al. 1999; McGinnis et al. 2011).

Most of the gas samples from coastal lakes reveal the stable C and H isotopic signatures indicative of AM (Figs. 5, 8). However, several samples show αC > 1.055, arguing for minor contribution from HM; and a few samples display αC < 1.031, characteristic for CH4 oxidation. In the case of a few samples from Lake Bukowo, Lake Wicko and Lake Kopań, both δ 13CCO2–δ 13CCH4 and δ 13CCH4–δ 2HCH4 signatures consistently indicate some HM contribution. The dominance of AM in the lakes is consistent with low sulphate, freshwater conditions. The shift towards HM for some samples may indicate the exhaustion of the labile AM substrates and a shift in methanogenic pathway. Hornibrook et al. (2000) claimed that recalcitrance of sedimentary organic matter (SOM) acts as one of key factors controlling the mechanism of CH4 production. The more labile the organic matter delivered to the methanogenic zone, the higher the preference for AM.

The δ 13CCO2 and δ 13CCH4 show weak seasonality, with a tendency to enhanced HM during colder periods and enhanced AM during warmer periods (Fig. 5). This agrees with previous reports (Hornibrook et al. 2000; Avery et al. 2003; Chanton et al. 2005) and may be related to formation and abundance of precursor substrates for AM. However, the temporal patterns of δ 13CCO2 and δ 13CCH4 are different in each lake; consequently, there is no compelling evidence of a significant climatic (temperature) control on methanogenesis.

Based solely on δ 13CCO2 vs δ 13CCH4, the methanogenic pathway in the lakes would appear to be influenced by water depth (Fig. 5). The δ 13CCO2, and to a lesser extent the δ 13CCH4, become 12C-enriched with increasing water depth. There is also a shift in αC to lower values with depth. This would suggest enhanced AM over HM at depth. This conclusion contrasts with observations by Jędrysek (1999) and Alstad and Whiticar (2011) who argued that HM should dominate in deeper waters. The transition from AM to HM with depth is hypothesized to be due to a depletion by utilization of labile AM precursor substrates in the deeper waters, and by higher availability of dissolved CO2 in deeper waters. We believe that such reasoning is not valid for the seven lakes studied, because in our data there is neither clear evidence for substantial differences in dissolved CO2 within lakes nor differences in the decomposition of SOM throughout the lakes (the TOC/N and δ 13CTOC show only minor variability within each lake) (Table 2). In addition, because the δ 2HCH4 values are not related to water depth, the overall isotope evidence does not support a shift from HM to AM with depth. Therefore, we suggest that the stable isotope composition of microbial gases is altered by secondary processes.

Methane oxidation

The observed shifts in αC in the lakes are determined more by changes in δ 13CCO2 than in δ 13CCH4, and with increasing water depth the δ 13CCO2 values decline to <−35‰ (the minimum δ 13CCO2 is −51‰) (Fig. 5). Such low δ 13CCO2 values can only be explained by methanotrophy, i.e., the oxidation of 13C-depleted microbial CH4. Other remineralization reactions, such as organic matter respiration, could not result in the observed highly negative δ 13CCO2 because these processes produce CO2 in which δ 13C signatures are close to δ 13CTOC (i.e. between −24 and −29‰; Table 2). The oxidation of CH4 occurred during summer and autumn (Fig. 5). We hypothesize that enhanced oxidation in deeper water is related to the intense wind-induced lake water mixing which is known to drive methanogenesis to greater depths within the sediments (Blair 1998). The oxidation of CH4 can be performed within the pore waters and in the water column. In the former case the oxidation is under reducing conditions and enhanced SO4 concentration. Earlier in this paper we postulated that the oxidation of CH4 at the expense of SO4 could be responsible for the lowered CHC values in the gas collected during saltwater inflows to coastal lakes. However, the lack of correlation between δ 13CCH4 and lake water salinity shows that this process has a negligible isotope effect.

Methanotrophic bacteria existing in lake water column consume methane during its diffusion and/or bubble transport via the oxidized water (Schubert et al. 2010). However, in the case of ascending bubbles the oxidation is minor due to the very short contact of bubble CH4 with O2 in the water column (Liikanen et al. 2003; Chanton 2005). Interestingly 12C-enriched δ 13CCO2 values correspond to low concentrations of CO2 (Fig. 6) which seems inconsistent with methanotrophy because the enhanced oxidation of CH4 and OM is expected to deliver more CO2 than CH4 (Hornibrook et al. 2000) and thus leads to higher CDMI. Interpreting the low CO2 and 12C-enriched δ 13CCO2 values in terms of methanotrophy would imply that all the excess CO2 was removed from the gas phase. In light of the model proposed by McGinnis et al. (2011), the plausible explanation for the CO2 removal is rapid dissolution of CO2 from ascending bubbles in the water column.

The effect of CH4 oxidation fades out in gas samples from the shallow water sections of the lakes where less negative or even highly positive δ 13CCO2 values of 0 to +6‰ were obtained. These signatures are attributed to methanogenesis. Methanogenic origin of shallow water CO2 is corroborated by the mass balance calculation relating stable C isotope composition of CO2 and CH4 to the δ 13COM of the precursor organic material. From the equation proposed by Corbett et al. (2013):

$$\delta^{13} {\text{C}}_{\text{OM}} = \, 0.5\;\cdot\;\delta^{13} {\text{C}}_{\text{CO2}} \; + \;0.5\;\cdot\;\delta^{13} {\text{C}}_{\text{CH4}} \;$$
(4)

we derived δ 13COM values ranging from −26.2 to −29.9‰, which are close to the measured δ 13CTOC of −24.4 to −29.1‰ (Table 2) thus indicating that δ 13CCO2 fits well the expected stable C isotope composition of CO2 produced from the organic matter in the lakes’ sediments.

Shallow water sites are less affected by oxidation, presumably due to location in wind-protected embayments and/or dissipation of water energy by macrophytes.

The effect of ebullition on δ13CCH4 and δ13CCO2

The stable isotope composition of the gas can be influenced by the rate of ebullition. Ebullition itself causes rather low C fractionation, with a tendency for the dissolved forms of CH4 and CO2 (CH4diss and CO2diss, respectively) to be slightly enriched in 12C (Hornibrook et al. 2000). However, when methanogenically derived CH4 and CO2 are constantly removed from the sediments by intense ebullition δ 13C in remaining gases becomes progressively heavier in accordance with the Rayleigh relationship (Whiticar 1999). Moreover, since the solubility of CO2 and αC for this reaction are higher than for CH4, the increase in δ 13CCO2 is expected to be higher than in δ 13CCH4. We believe that this process contributes to the variability in the stable C isotope composition of gases throughout the coastal lakes, because there are differences in the rate of ebullition as shown by the CH4/N2 ratio. The shallower parts of the lakes display more intense ebullition as well as higher δ 13CCH4 and δ 13CCO2 compared with the deeper parts. On the other hand, the lower rate of ebullition in deeper parts of the lakes may be responsible for the prolonged contact of CH4 with oxygenated water and thus lead to more intense oxidation of CH4.

The effect of bioventilation on δ13CCH4 and δ13CCO2

It was shown earlier that the composition of gas may be modified due to ventilation by macrophytes (Fig. 3a). It has been recognized that the removal of CH4 by macrophytes is an isotope fractionating process discriminating against 13C and leaving residual CH4 with higher δ 13CCH4 signatures (Chanton 2005). Whether or not there is an analogous effect for methanogenically derived CO2 in the rhizosphere is uncertain, as is the magnitude of the isotope fractionation. On the other hand, the macrophytes are known to shift δ 13CCH4 towards higher values by the enhanced deposition of metabolizable OM in the proximity of vegetation (Blair 1998). The influence of macrophytes on the stable C isotope composition of CO2 and CH4 can be seen in Fig.7, where there is a noticeable difference between vegetated and non-vegetated sites. The former have 3.6‰ higher δ 13CCH4 and 10.5‰ higher δ 13CCO2 values than the latter.

Conclusions

This paper provides a comprehensive study of microbial gas production in the ecosystem of Polish coastal lakes located on the non-tidal Baltic coast. The molecular and stable isotope compositions of the gases show differences between the seven lakes studied and within each lake, due to variations in the biogeochemical processes. These are modulated by the overall complexity of coastal lakes, including the high biological activity, spatial and temporal salinity changes, the very high water column turbulence and vertical mixing of sediments. The increasing lake depth results in a depletion in CH4 and CO2 in the bubbles in favor of N2 and O2. A large part of microbially produced CH4 and CO2 in the sediments is removed in the lake water column by the dissolution of rising bubbles and gas exchange with the ambient water. The aquatic macrophytes also affect the compositions of microbial gas. From our data it emerged that in vegetated parts of the lakes, the CH4 is depleted in favor of CO2 and that the residual CH4 and CO2 are enriched in 13C. The removal of CH4 from the bubbles via dissolution and bioventilation acts as a primary source of the dissolved CH4 pool in the water column of coastal lakes. Despite the fact that the coastal lakes display highly oxidizing conditions in the water column and that the bottom sediments are prone to intense mixing by wind waves, the CH4 reveals rather low oxidation, except for the CH4 produced in the deepest parts of the lakes. The oxidation of CH4 in deeper waters seems to be related to lower rates of ebullition and thus longer exposure to oxygenated waters, as well as to the penetration of O2 to the pore waters during vertical mixing of sediments.

In the seven lakes studied there is no clear relationship between salinity and the stable C and H isotope composition of sediment gas. It seems, however, that salinity may influence the molecular composition of hydrocarbons via preferential oxidation of CH4.

Given the remarkable dissolution of gas bubbles in the water column, the effect of salinity on CH4 oxidation and the C isotope effects accompanying bioventilation both require further investigations to decipher the mechanisms driving these processes.

On the other hand, the data obtained in the current study will assist in the assessment of the emissions of CH4 and CO2 from coastal lakes. Our observations have provided the basis for establishing the representative monitoring sites in which reliable estimations of ebulitive and diffusive emission of gases from the lakes can be performed.