Seismic data were collected using a Geometrics GeoEel Streamer of 150 and 183.5 m length and 96 and 112 channels, respectively. Two generator-injector guns in harmonic mode (105/105 cubic inch) served as the seismic source. Processing included navigation processing (1.5625 m crooked line binning), 20, 45, 250, and 400 Hz frequency filtering, and post-stack Stolt migration with water velocity yielding an approximately 2 m horizontal and 5 m vertical resolution close to the seafloor.

At seepage and vent sites, the video-guided MUC was used to discover recent fluid release, which was indicated by typical chemosynthetic biological communities at the seafloor (microbial mats, bivalves, etc.; Sahling et al., 2002). However, small-scale, patchy distributions of active seepage spots and the visibility of authigenic carbonate concretions made it difficult to select the best locations for coring. Hence, the comparison of results from different seeps might be biased in this regard, as not all seepage areas could be sampled at their most active places. GC deployments were typically performed at sites initially investigated with the MUC video system or at the center of suspected seeps (based on bathymetry and seismic data).

In total, we present pore fluid and gas data collected at the seepage sites North (GC01, MUC11), Central (GC03, GC13, GC15, MUC04), and Ring (MUC05), one reference site (see Reference Site below; GC04, MUC02), and the hydrothermal vent field (Smoker Site; GC09, GC10, MUC15, MUC16). The Reference Site that did not show active seepage or faults indicated by seismic data was chosen to obtain geochemical background values. In addition, the slope towards the Mexican mainland was sampled as well (Slope Site; GC07) (Fig. 1, Table 1). Immediately after core retrieval, GCs were cut, split, and subsampled. Samples were transferred into a cold lab at 4 ∘C and processed within 1 or 2 h. Pore fluids were obtained by pressure filtration (e.g., Jahnke et al., 1982). After MUC retrieval, bottom water was sampled and immediately filtered for further analyses. The sediment was transferred into a cold lab and sampling was executed in an argon-flushed glove bag. Pore fluids were retrieved by centrifugation and subsequent filtration using 0.2 µm cellulose acetate membrane filters (e.g., Jahnke et al., 1982). Sediment samples (2 cm3) for hydrocarbon analyses were taken using cut-off 3 mL syringes. All hydrocarbon samples were taken immediately after sediment surfaces were exposed after core cutting or sectioning, ensuring minimal disturbance to sediment surfaces prior to sampling and transferred to vials containing concentrated NaCl solution (as seen in Sommer et al., 2009). MUCs were extruded and sampled from the top. GCs were sampled at the bottom ends of 1-m core sections, either at the core catcher or at freshly cut section ends. In some cases additional samples were taken from within GC core sections by cutting the core liner with an oscillating saw and inserting cut-off syringes into the sides of core sections.

Temperature gradients and thermal conductivity were measured at the North Seep, Central Seep, Reference Site, and Smoker Site as well as along a transect across the newly discovered hydrothermal vent field and the rift valley (Graben Site). Miniaturized temperature loggers (MTL) were attached to GCs or to a 5 m long sediment lance at a sampling rate of 1 measurement per second. The absolute accuracy of these temperature measurements was about 0.1 K and the temperature resolution was 0.001 K (Pfender and Villinger, 2002).

Thermal conductivity was measured on recovered core material in close vicinity to the MTLs using the KD2 Pro Needle Probe instrument. For temperature measurements obtained by a lance, a constant thermal conductivity of 0.7 W m-1 K was assumed. Data processing was done according to Hartmann and Villinger (2002).

Water samples were taken by using a Niskin water sampler (Rosette System), equipped with a video camera designed for near-seafloor sampling (Schmidt et al., 2015) in order to study water column chemistry (i.e., dissolved CH4) and oceanographic parameters (i.e., temperature, salinity, turbidity). Eight water sampling locations were chosen in the vicinity of MUC and GC stations and were termed the North Seep (VCTD03), Central Seep (VCTD02), Ring Seep (VCTD01), Graben Site (CTD01; no video-guided sampling), Smoker Site (VCTD06 and 10), and Slope Site (VCTD07). The (V)CTDs were either used in a towed mode (VCTD03, 06, 09, 10) or in station (CTD01; VCTD01, 02, 07) keeping hydrocast mode. The water depth was controlled based on pressure readings, altitude sensors (< 50 m distance to bottom), and online video observation (1–2 m above the seafloor).

At the Central Seep a block (approx. 1×0.5×0.3 m) was recovered using a video-guided hydraulic grab (VgHG, GEOMAR) in 1843 m water depth from the surface of a typical cold seep environment (close to high abundance of tube worms) . The block consisted mainly of a solidified carbonate matrix covered by a whitish carbonate rim and was characterized by coarse open pore space in the mm to cm scale (see Supplement Fig. S1).

The water column above the newly discovered hydrothermal vent field exhibits elevated CH4 concentrations (up to 400 µM) and pCO2 data (> 6000 µatm) (Berndt et al., 2016). The range of the measured stable isotope signature of methane (δ13CCH4 between -39 ‰ and -14.9 ‰) and the helium isotope anomaly (3He / 4He ratio of 10.8×10-6) clearly indicate the existence of gas exhalations from thermogenic organic matter degradation with contributions from a mantle source (see Berndt et al., 2016). These northern trough hydrothermal fluids are comparable in their gas geochemistry to the southern trough (Lupton, 1979; Von Damm et al., 1985; Berndt et al., 2016). However, the highest heat flow values of up to 10 835 mW m-2 measured in this study are found close to the Smoker Site and are much higher than those observed in earlier studies (maximal 2000 mW m-2, Fisher and Becker, 1991). The high heat flow at the Smoker Site even exceeds the more hydrothermally active southern trough, where heat flow values of 2000 to 9000 mW m-2 were measured (Lonsdale and Becker, 1985; Fisher and Becker, 1991). This might indicate that hydrothermal activity at the northern trough is younger and possibly a more recent process compared to the southern trough.

Hydrothermal fluids are typically depleted in Mg and are highly enriched in fluid-mobile elements like Li caused by high-temperature reactions with mafic rocks (here sills) and/ or sediments through which they percolate (e.g., Einsele et al., 1980; Gieskes et al., 1982; Kastner, 1982; Von Damm et al., 1985; Lizarralde et al., 2010; Teske et al., 2016). Such compositions were reported from DSDP site 477 (Gieskes et al., 1982) and fluids obtained by Alvin dives (Von Damm et al., 1985) (see Fig. 1 for location of Site DSDP 477). Although strongly diluted, CTD samples from the hydrothermal plume in the northern trough show this trend (Berndt et al., 2016).

An indication for the presence of hydrothermal fluids in pore waters in the vicinity of the hydrothermal vent field is found at about 4 m depth in core GC09. Here, positive Li and negative Mg concentrations (Fig. 4a) are probably caused by the weak admixing of hydrothermal solutions (Gieskes et al., 1982; Hensen et al., 2007). Likewise, 87Sr / 86Sr isotope ratios decrease to a value of 0.708949 (Fig. 5) and thus tend towards the 87Sr / 86Sr ratio of the local hydrothermal endmember (87Sr / 86Sr =0.7052; Von Damm, 1990). Hydrothermal endmember Li concentrations in the Guaymas Basin have been reported in a range between 630 and 1076 µM (Von Damm et al., 1985) and are 20 to 30× higher than those measured at the Smoker Site (∼34 µM; Fig. 4a, Table S1). Here, hydrothermal fluids account for about 3 % of the mix with seawater (Fig. 7). The sediments in this core section also change from diatomaceous clay to unconsolidated, coarse-grained hydrothermal deposits (Fe-rich sulfides; see also Sect. 3.3), which may facilitate the circulation of hydrothermal fluids.

Despite the proximity of the remaining GCs and MUCs to the hydrothermal vent field (∼500 m distance; temperatures immediately after retrieval are up to 60 ∘C) typical pore fluid indicators such as Mg, Li, and 87Sr / 86Sr do not show major excursions from seawater values (Fig. 4). Similarly, NH4, an indicator for a diagenetic or catagenetic breakdown of organic matter, is only poorly enriched in sediments surrounding the hydrothermal vents (NH4≤0.3 mM). NH4 remains well below the value reported from the southern trough (20 mM; Von Damm et al., 1985) and the Slope Site (GC07) where 10 mM were already reached at subsurface depths of only a few meters (Fig. 4). The pore fluid geochemistry around the hydrothermal vent field therefore confirms that early diagenetic processes are not intense (Fig. 7) and that the shallow sediments are not significantly affected by hydrothermal fluids. We hypothesize that hydrothermal venting causes a shallow convection cell drawing seawater through the sediments towards the hydrothermal vent field, while the sediments become heated by lateral heat conduction (cf. Gamo et al., 1991; Henry et al., 1996; Kinoshita and Yamano, 1997).

The diatomaceous clay might act as a seal to fluids migrating upwards, which are channeled to the catchment area of the rising hydrothermal fluids of the hydrothermal vent field (see also Fig. 4 in Berndt et al., 2016). The geochemical composition of these fluids is likely influenced by a high temperature-chemical alteration of the sediment caused by the intruded sills (Fig. 2b). However, shallower pore fluids of surface sediments at the Smoker Site (i.e., 0–4 m) are not affected much by contributions from these fluids and show predominantly ambient diagenetic fluid signatures.

The selection of sampling sites at presumed seep locations was based on existing published data (Lizarralde et al., 2010) and information from seismic records (Fig. 2). Seismic amplitude blanking along vertical zones below the seafloor indicates active fluid and/ or gas conduits at the North and Central seeps. Given that sill intrusions and related high-temperature alteration of sediments are driving the seepage, the expectation was to find deeply sourced (average sill depth ∼400 m) fluids with a typical geochemical signature analogous to findings at hydrothermal vents in the Guaymas Basin (Von Damm et al., 1985; Von Damm, 1990; Berndt et al., 2016). Such characteristics are, e.g., a high concentration of thermogenic hydrocarbon gases formed by organic-matter degradation, enrichments in NH4, a depletion in Mg, and a strong enrichment in fluid-mobile tracers like Li and B (e.g., Aloisi et al., 2004; Scholz et al., 2009). The hydrocarbon formation caused by abiogenic processes plays only a minor role in the hydrothermal vent field (McDermott et al., 2015; discussion in Berndt et al., 2016).

Samples obtained using a video-guided MUC revealed the highest methane concentrations at the North, Central, and Ring seeps (Fig. 4b). In conjunction with visual evidence (abundant chemosynthetic biological communities), this confirms that we have hit active seepage areas during our sampling campaign. At the two most active sites, North and Central, high methane levels are accompanied by a significant drop in sulfate and an increase in TA and TH2S, providing evidence for AOM, according to the following net reaction; (R2)CH4+SO42-→HCO3-+HS-+H2O (e.g., Nauhaus et al., 2005; see Wegener et al., 2016 for a recent update).

These pore-water trends are even more pronounced in GC01 (North) where the AOM zone was completely penetrated and gas hydrate was found at about 2.5 m b.s.f. Unfortunately, GCs from similarly active sites could not be obtained from the Central and Ring seeps, mainly because of patchiness of seepage spots and the widespread authigenic mineralizations at the seafloor preventing sufficient penetration. Nevertheless, active methane seepage at all three investigated sites is evident. The methane flux is, however, not accompanied by any significant excursion of pore-water constituents typical for deeply sourced, high-temperature sediment–water interactions (e.g., Mg, Cl, Li). Also, Sr concentrations show seawater values at all seep sites (Fig. 5), except for the North Seep, where values drop together with Ca due to co-precipitation during carbonate formation. The 87Sr / 86Sr ratios show predominantly seawater signatures as well (Fig. 5, Table S2). Similarly, low NH4 concentrations of < 1 mM indicate a low intensity of organic matter decomposition (as discussed in Sect. 4.1.1). Taken together, our data show that, with the exception of methane and sulfate, the pore water corresponds to ambient diagenetic conditions that are typically met in this shallow subsurface depth. An explanation for the decoupling of methane levels and pore-water composition is that only methane is rising to the seafloor as a free gas. This assumption requires a closer look at the composition of dissolved hydrocarbons in general, which is given below.

The origin of hydrocarbon gases can be deciphered by plotting CH4 / (C2H6+C3H8) ratios versus δ13CCH4 data in a modified Bernard diagram (Schmidt et al., 2005 and literature therein) (Fig. 8a) and δ13CCH4 versus δDCH4 after Whiticar (1999) and Welhan (1988) (Fig. 8b). Most of the measured stable isotope data of pore-water methane indicate a microbial origin or a mixed microbial and thermogenic origin (Fig. 8). By contrast, hydrocarbons venting at the hydrothermal vent field reflect a mixture of thermogenic methane and abiogenic methane derived from water–rock interactions (Berndt et al., 2016).

It has to be considered though that, except of three samples from the North Seep, all δ13CCH4 measurements were performed on samples located above the AOM zone (see Fig. 4). This implies that the upward-rising methane has likely undergone fractionation due to methane oxidation by sulfate in the AOM zone underneath. AOM enriches DIC in 12C and results in progressively increasing δ13CCH4 values in the residual methane (Whiticar, 1999). Considering the δ13CCH4 at the Slope Site as a microbial endmember composition for the Guaymas Basin (Fig. 8a), most of the data fall on calculated fractionation lines for AOM, following a Rayleigh trend (Whiticar, 1999). Methane sampled close to the Smoker Site (MUC15) is obviously also affected by AOM (Fig. 8a). This is in line with recent studies on hydrothermal sediments of the southern trough of the Guaymas Basin, where bacterial and archaeal communities catalyze the oxidation of methane and higher hydrocarbons, shifting δ13CCH4 values to heavier signatures (Dowell et al., 2016).

The origin of methane and oxidation effects can further be identified in the δ13CCH4 versus δDCH4 plot after Whiticar (1999) and Welhan (1988) (Fig. 8b). Slope Site samples plot in the field of microbial CO2 reduction while hydrothermal plume samples plot in the thermogenic field. One sample even points to a mantle signature and thus shows potential endmember isotope signatures (Berndt et al., 2016). North Seep samples (pore fluids and gas hydrates) plot in the mixing region while samples from the Central Seep clearly shift away from the microbial field and are considered to be affected by bacterial oxidation (Whiticar, 1999).

Considering the methane below the AOM as being unaltered, three North Seep samples and the majority of the Slope Site samples show a clear microbial source of methane (Fig. 8a). All other samples appear to be affected by major oxidation following a Rayleigh fractionation process and show that only a fraction between 2 % (MUC04, Central Seep) and 0.05 % (GC15, Central Seep) remains as unoxidized methane (Fig. 8a).

The δ13CCH4 versus δDCH4 plot of the unaltered North Seep samples suggests a mixing of microbial and thermogenic methane (Fig. 8b). Similar signals have also been observed at Hydrate Ridge (Milkov et al., 2005) and seem to be a common phenomenon in hydrothermal and cold-seep-affected sediments. In a few samples from the North and Central seeps ethane concentrations have been high enough to measure stable carbon isotopes, and the δ13CC2H6 values point to a thermogenic origin (Table S3).

Based on our data set, no deep-sourced fluid is currently migrating upwards at the cold seeps investigated (compare deep-sourced seepage sites from the Gulf of Cadiz in Fig. 7). Hence, in terms of the original hypothesis that fluid emanation is directly linked to recent sill intrusions, these cold seep sites cannot be considered active as claimed by Lizarralde et al. (2010). These authors argued that thermogenic carbon is currently released up to 50 km away from the rift axis, causing a maximum carbon flux of 240 kt C yr-1. Further, Lizarralde et al. (2010) showed temperature anomalies, high methane concentrations, and helium isotopic anomalies in the water column potentially indicative of a magmatic source. These anomalies were detected in close vicinity to bacterial mats, tubeworms, and authigenic carbonates, situated above areas of sill intrusions. Comparable structures have been identified in this study by video-guided MUCs and seismic data (Fig. 2). Our detailed results on pore fluid, water column, and gas geochemistry now show that most methane was of microbial origin (Fig. 8) and only traces of thermogenic methane were found up to ∼20 km off axis (North Seep). Even pore fluids taken close to the hydrothermal vent field are dominated by shallow microbial degradation processes, indicating that the hydrothermal fluid flow in the Guaymas Basin is rather localized and bound to focused fluid pathways. The temperature and chemical anomalies detected by Lizarralde et al. (2010) might also arise from the deep water layer in the Guaymas Basin itself, which is influenced by hydrothermal fluids (Campbell and Gieskes, 1984). Hydrothermal activity in the Guaymas Basin produces hydrothermal plumes, rise 100–300 m above seafloor and spread out along density gradients throughout the basin (Campbell and Gieskes, 1984). Our results nevertheless show that the Guaymas Basin has a well-mixed bottom seawater layer, with temperatures ranging between 2.8 and 3.9 ∘C in > 1000 m depth and off-axis methane concentrations that vary quite considerably (e.g., 6 to 28 nM at Ring Seep, Fig. 9). These bottom seawater variabilities are bigger than the reported anomalies by Lizarralde et al. (2010) and might indicate that thermogenic methane release might not be as widespread as previously suspected.

Pore fluids taken in a transect up to ∼30 km away from the rift axis show no evidence for high-T reactions (Figs. 4, 7). We can still not exclude the possibility that thermogenic methane is released in other areas of the basin, but the lack of evidence for high temperature geochemical processes at our sites is evident and clearly contradicts the conclusions drawn by Lizarralde et al. (2010). Our findings suggest that a projection of the thermogenic methane release based on the number of detected sills (Lizarralde et al., 2010) represents a maximum estimate, as it neither considers the time of the emplacement of a sill nor the lifetime of such magmatic systems. Today, shallow microbial degradation processes determine pore fluid signatures (Figs. 4, 8). Whereas high temperature thermogenic reactions have certainly been active during sill emplacement and once released large amounts of carbon, these processes have apparently ceased. However, pipe structures may still act as high-permeability pathways and facilitate the advection of gas. As a result, small amounts of thermogenic carbon might be released as reflected by the signatures of δ13CCH4 and the thermogenic δ13CC2H6 isotope data at the North and Central seeps. However, present methane advection rates are slow (probably < 1 cm yr-1), as observed by low methane gradients in the pore fluid profiles (Fig. 4). These conditions favor an effective turnover of CH4 to bicarbonate and authigenic carbonates by AOM (Wallmann et al., 2006; Karaca et al., 2010).

The porous authigenic carbonate block recovered from the seafloor at the Central Seep can preserve long-term information about seepage in this area. The predominant biomarkers found in the seep carbonate from the Central Site (56-VgHG-4) are consistent with microbial consortia performing AOM. In particular, high abundances of crocetane and sn2-hydroxyarchaeol indicate major contributions from the methanotrophic archaea of the ANME-2 cluster, whereas DAGE originate from syntrophic sulfate-reducing bacteria, probably of the Desulfosarcina–Desulfococcus group (Blumenberg et al., 2004; Niemann and Elvert, 2008). These consortia gain energy from AOM, with sulfate as the final electron acceptor (see Eq. 2).

At the Central Seep, the increase in TA due to the AOM reaction plausibly explains the precipitation of isotopically depleted authigenic carbonates. Particularly, ANME-2 biomarkers have been reported in association with abundant fibrous, often botryoidal aragonite cements (Leefmann et al., 2008), which is in line with the observations made at the Central Seep (see Sect. 3.3). Moreover, the high abundance of ANME-2 indicates that seep carbonate formation took place under high sulfate concentrations and strong advective methane flow, but there were no elevated water temperatures (c.f. Nauhaus et al., 2005; Peckmann et al., 2009; Timmers et al., 2015). Minor amounts of typical marine sterols also show that the seep carbonates also captured detritus from the surrounding sediment and water column during their ongoing cementation.

The bulk carbonate carbon isotope signature (δ13CV-PDB=-46.6 ‰) overlaps with the shallow, heavy δ13CCH4 values (-27.5 ‰ and -48.6 ‰) in the pore fluids at the Central Seep and confirms a dominant AOM signature with a minor planktonic and potentially δ13C-diluting background signal. The oxygen isotopes point to a low formation temperature of about 3 ∘C, consistent with a precipitation in bottom waters (2.8 to 3.0 ∘C (Figs. 6, 9; Table S4). The 87Sr / 86Sr analyses support this assumption by values within uncertainty identical to modern seawater. U-Th carbonate dating provide ages younger than 240 yrs BP. In summary, authigenic carbonates originate from shallow methane and were subrecently formed in ambient seawater.

The seismic data taken across the seep locations indicate that the disrupted sediment layers do not reach the sediment surface (Fig. 2a, c). This implies that fluid mobilization ceased at some time before the uppermost sediment layers were deposited. The doming above the Central Seep provides some clues on the timing of fluid migration (Fig. 2c). Assuming that the doming is the result of buoyancy-related uplift (Koch et al., 2015), it represents the time when intrusion-related gas reached the seafloor. Assuming further a sedimentation rate of 1.7 m per 1000 years (Central Seep; Table 1) and maxima and minima deposition depths of 48 and 12 m below seafloor, respectively (see Fig. 2c), this would imply that most of the gas reached the seafloor between 28 and 7 kyr ago. Even at maxima and minima sedimentation rates of 3.5 m (North Seep) and 0.5 m (Ring Seep) per 1000 years, the gas flow would have ceased between 14 and 3 kyr ago at the earliest and between 96 and 24 kyr ago at the latest. Accordingly, this finding further supports the results of the pore fluid and gas geochemistry, which show no sign of active fluid flow from a depth at the cold seep sites in the northern Guaymas Basin.

We agree with Lizarralde et al. (2010) that hydrothermal activity in the Guaymas Basin is an important driver for CH4 (and CO2) emissions into bottom waters. However, our data set shows that there is no deep fluid advection at the investigated sites. Our interpretation is that hydrothermal activity at these off-axis locations has ceased and previously formed pathways seem to mediate the advection of biogenic gas at present. It is not unlikely that seep-induced, hydrothermal activity is still ongoing in other places than those investigated in this study, but in order to provide more accurate predictions for (thermogenic) carbon fluxes and the potential impact on the climate, sill emplacement mechanisms need to be better constrained. Apart from their spatial distribution, the most important and currently unknown factors are the determination of the time of their emplacement and the longevity of the sill-systems that require further investigation.