In natural coastal wetlands, high supplies of marine
sulfate suppress methanogenesis. Coastal wetlands are, however, often
subject to disturbance by diking and drainage for agricultural use and can
turn to potent methane sources when rewetted for remediation. This suggests
that preceding land use measures can suspend the sulfate-related methane
suppressing mechanisms. Here, we unravel the hydrological relocation and
biogeochemical S and C transformation processes that induced high methane
emissions in a disturbed and rewetted peatland despite former brackish
impact. The underlying processes were investigated along a transect of
increasing distance to the coastline using a combination of concentration
patterns, stable isotope partitioning, and analysis of the microbial
community structure. We found that diking and freshwater rewetting caused a
distinct freshening and an efficient depletion of the brackish sulfate
reservoir by dissimilatory sulfate reduction (DSR). Despite some legacy
effects of brackish impact expressed as high amounts of sedimentary S and
elevated electrical conductivities, contemporary metabolic processes
operated mainly under sulfate-limited conditions. This opened up favorable
conditions for the establishment of a prospering methanogenic community in
the top 30–40 cm of peat, the structure and physiology of which resemble
those of terrestrial organic-rich environments. Locally, high amounts of
sulfate persisted in deeper peat layers through the inhibition of DSR,
probably by competitive electron acceptors of terrestrial origin, for
example Fe(III). However, as sulfate occurred only in peat layers below
30–40 cm, it did not interfere with high methane emissions on an ecosystem
scale. Our results indicate that the climate effect of disturbed and
remediated coastal wetlands cannot simply be derived by analogy with their
natural counterparts. From a greenhouse gas perspective, the re-exposure of
diked wetlands to natural coastal dynamics would literally open up the
floodgates for a replenishment of the marine sulfate pool and therefore
constitute an efficient measure to reduce methane emissions.
Introduction
Coastal wetlands play an important role in climate change mitigation and
adaption as they can efficiently accrete organic sediments, adjust coastal
elevations to sea level rise and protect low-lying areas in the hinterland.
Further, while freshwater wetlands constitute the largest natural source of
the greenhouse gas methane (CH4, Zhang et al., 2017), the efficient
accumulation of autochthonous C in coastal wetlands comes without the
expense of high CH4 emissions (Holm et al., 2016). Methane is a
potent greenhouse gas that is formed as a terminal product of organic matter
breakdown under strictly anaerobic conditions typically in the absence of
electron acceptors other than carbon dioxide (CO2) (Segers and
Kengen, 1998). In coastal environments, methane production and emission are
effectively suppressed by sulfate-rich seawaters: methanogens are
outcompeted by sulfate-reducing bacteria (SRB) for acetate-type precursors
and hydrogen (Schönheit et al., 1982; Lovley and Klug, 1983). This
shifts the prevailing anaerobic C metabolic pathways from methanogenesis
towards dissimilatory sulfate reduction (DSR) (Martens and Berner, 1974; King and Wiebe, 1980). In addition, sulfate (SO42-) operates as an electron
acceptor for anaerobic methane oxidation by a syntrophic consortium of
anaerobic methanotrophs (ANMEs) and SRB (Iversen
and Jorgensen, 1985; Boetius et al., 2000). Anaerobic methane oxidation has been specifically
described for brackish wetland sediments but is not exclusively confined to
the utilization of sulfate as an electron acceptor (Segarra et al., 2013, 2015).
Human activities such as diking and drainage place intensive pressure on
coastal landscapes with sometimes irreversible impairments of their
biogeochemical cycles and ecosystem functions (Karstens et al., 2016;
Zhao et al., 2016). Dikes separate coastal wetlands from resupply of
seawater, and drainage for agricultural use induces the aerobic
decomposition of organic-rich sediments, resulting in substantial CO2
losses and land subsidence (Deverel and Rojstaczer,
1996; Miller, 2011; Deverel et al., 2016; Erkens et al., 2016). As sea levels are expected to
rise, the controlled retreat from flood-prone areas becomes an essential
strategy of integral coastal risk management to complement conventional
technical solutions such as diking (Sánchez-Arcilla et al., 2016).
Rewetting may re-establish the ability of abandoned coastal wetlands to
efficiently accrete organic matter under anaerobic conditions and represents
a promising management technique to reverse land surface subsidence caused
by drainage-induced peat oxidation (Deverel et al., 2016; Erkens et
al., 2016). Moreover, while freshwater wetlands may become methane sources
upon rewetting (Wilson et al., 2009; Vanselow-Algan et
al., 2015; Franz et al., 2016; Hemes et al., 2018), sulfate-rich seawater could potentially
reduce post-rewetting methane release in coastal wetlands. However, recent
work on a degraded brackish peatland has revealed high post-rewetting
CH4 emissions (Hahn et al., 2015; Koebsch et al., 2015) and
methanogen abundance (Wen et al., 2018), thereby challenging the common
notion of coastal wetlands as negligible methane emitters. In fact, diking
and the drainage-rewetting cycle may induce hydrological shifts and
biogeochemical transformation processes that are so far not well understood.
In particular, the transformation and/or relocation of the marine sulfate
reservoir in the sediments of diked wetlands are of vital importance to
understand the implications of anthropogenic intervention on coastal wetland
biogeochemistry and to better constrain the climate effect of coastal
wetland remediation.
Here, we investigate the mechanisms that allow for high methane production
in disturbed and remediated coastal wetlands. We therefore address the fate
of brackish compounds and the emerging S and C transformation processes in a
rewetted, freshwater-fed peatland that was naturally exposed to episodic
intrusions from the Baltic Sea. In the past, the peatland had been subject
to intense human intervention including diking and drainage for agricultural
use. After rewetting by freshwater-flooding, the site turned into a strong
methane source. The underlying hydrological and biogeochemical processes
were investigated along a brackish–terrestrial transect that spans between
300 and 1500 m in distance from the coastline using hydrogeochemical element
patterns, stable isotope biogeochemistry, and microbiological analyses.
The specific goals were to
retrace the marine legacy effect remaining after diking and freshwater
rewetting in the peat pore space using salinity, the isotope composition of
water, and a suite of inert dissolved constituents that may be indicative for
the intermingling of brackish and terrestrial waters.
track the fate of Baltic Sea-derived sulfate and uncover potential S
transformation pathways using concentration patterns, stable isotope
measurements of pore water SO42- (δ34S and δ18O), and solid S compounds as well as the bacterial community
structure.
describe evolving methane cycling processes using concentration and stable
isotope measurements of CH4 (δ13C, δ2H) and
dissolved inorganic C (DIC, δ13C) as well as the abundance and
community structure of methane-cycling microbes.
We hypothesized the marine legacy effect to be represented by a lateral gradient in
electrical conductivity (EC) and pore water sulfate along the
brackish–terrestrial transect. We further expected increasing terrestrial
impact to promote the deprivation of the brackish sulfate pool and to induce
complementary patterns of methane production.
Material and methodsStudy site and sampling design
The study site is part of the nature reserve “Heiligensee und
Hütelmoor”, a 490 ha coastal peatland complex located in NE Germany
directly at the SW Baltic coast with an elevation between -0.3 and +0.7 m
above sea level (Dahms, 1991) (latitude 54∘12′, longitude
12∘10′, Fig. 1). Climate is transitional maritime with
continental influence from the east. The area receives a mean annual
precipitation of 645 mm with a mean annual temperature of 9.2 ∘C
(reference period 1982–2011, data from the German Weather Service, DWD).
Peat formation was initiated by the Littorina Sea transgression and the
postglacial sea level rise around 5400 BC. Presently, the Hütelmoor is
fed by a 15 km2 forested catchment dominated by gley over fine sands.
Originally, the fen exhibited 0.2–2.3 m deep layers of sulfidic reed-sedge
peat underlain by Late Weichselian sands over impermeable till (Voigtländer et al., 1996; Bohne
and Bohne, 2008). A total of 40 years of drainage for
grassland use caused severe degradation of the peat, which was recently
identified as sapric Histosol (Koebsch et al., 2013). Since the
rewetting by flooding in 2010 through the construction of a weir at the
outflow of the catchment, more than 80 % of the area has been permanently
inundated with freshwater from the surrounding forest catchment
(Miegel et al., 2016). Current vegetation of the Hütelmoor is
dominated by patches of competitive emergent macrophytes such as reed and
sedges (Phragmites australis (Cav.) Trin. ex Steud and Carex acutiformis Ehrh.) that increasingly supersede species indicative for brackish
conditions (Bolboschoenus maritimus (L.)
Palla, Schoenoplectus tabernaemontani (C. C. Gmel.) Palla) (Koch et al., 2017).
(a) The study site Hütelmoor is located directly at the
southwestern Baltic coast at an altitude between -0.2 and +0.2 m a.s.l. In its pristine state, the site was exposed to episodic brackish
water intrusion by storm surges. (b) Profiles of sediments and pore waters
were taken along a transect with 300–1500 m distance to the coastline.
Deviations of the transect from the straight normal to the Baltic coastline
arose due to the restricted accessibility of the site. (c) A former study
located close to spot 2 in the center of the current sampling transect
revealed high pore water sulfate concentrations in 30–60 cm below the surface
with annual means up to 24±3 mM (red circles indicate annual means
while dashed circle lines represent the standard deviation over the year).
Map data are copyrighted under OpenStreetMap contributors and available from
https://www.openstreetmap.org (last access: 23 November 2016).
Under natural coastal dynamics, the Hütelmoor is episodically flooded by
storm surges. Low outflow and high evapotranspiration rates promote brackish
conditions. Major brackish water intrusions were reported for 1904, 1913,
1949, 1954, and 1995 (Bohne and Bohne, 2008) though flooding frequency
is reduced since the site was diked in 1903. Additional brackish input
occurs through underground flow and atmospheric deposition as well as
through high water situations at the Baltic Sea when backwater of the
interconnected Warnow River delta enters the fen. However, potential
brackish water entry paths other than storm surges have revealed a negligible
effect on peat salinity (Selle et al., 2016). The last flooding event in
1995 raised EC in the drainage ditches up to 8 mS cm-1, but the EC
decreased to the pre-flooding level of 2 mS cm-1 within the following
5 years (Bohne and Bohne, 2008).
Samples were collected at four spots along a transect with increasing
distance to the Baltic Sea (300–1500 m, Fig. 1b) within 2 weeks in
October/November 2014. The transect included the area of a former study
which revealed high concentrations of brackish SO42- with annual
means up to 23.7±3.2 mM (Koch, unpublished, Fig. 1c). At the time of
sampling, water depth above the peat surface spanned from 9 to 19 cm, which
presented the lowest range within the seasonal water level fluctuation.
Sampling depth ranged from 45 to 65 cm, which was in most cases sufficient to
cover the full peat depth including the underlying mineral soil.
Pore water analysis
Pore waters were collected from distinct depth below the surface (cm b.s.f.)
with a stainless steel push-point sampler attached to a syringe to draw the
sample from a distinct penetration depth. Temperature, pH, EC, and salinity
were measured directly after sampling (Sentix 41 pH probe and a TetraCon 325
conductivity-measuring cell attached to a WTW multi 340i handheld; WTW,
Weilheim). Samples were filtered (0.45 µm membrane syringe filters)
in situ and transferred without headspace into vials (except for dissolved
CH4). Vials had been previously preconditioned with 1 MHCl and
subsequent 1 MNaOH and were filled with a compound-specific preservative
(see below).
Dissolved CH4 concentration was measured with the headspace approach.
Therefore, 5 mL of pore water was transferred into 12 mL septum-capped
glass vials under atmospheric pressure. Before taking them to the field, the
sampling vials were flushed with Ar and filled with 500 µL saturated
HgCl solution to prevent further biological activity. After sampling, the
punctuated septum was covered with lab foil and the vials were stored upside
down to minimize CH4 loss. Headspace gas concentrations after
equilibration were measured in duplicates with an Agilent 7890A gas
chromatograph equipped with a flame ionization detector and with a carbon
plot capillary column or HP-Plot Q (Porapak Q) column. Helium was used as the
tracer gas. Gas sample analyses were performed after calibration of the gas
chromatograph with gas standards (accuracy > 98.5 %). The measured headspace CH4 concentration was then converted
into dissolved CH4 concentration using the temperature-corrected
solubility coefficient (Wilhelm et al., 1977).
Samples for anion concentrations (SO42-, Cl-, Br-) were
filled in 20 mL glass vials preserved with 1 mL 5 % ZnAc solution to
prevent sulfide oxidation. Anion concentrations were analyzed by ion chromatography (Thermo
Scientific Dionex) in a continuous flow of 9 mM NaCO3 eluent in an IonPac AS9-HC 4 column, partly after dilution of the sample. The device was
calibrated with NIST standard reference material solutions freshly prepared before each run
to span the concentration ranges of the (diluted) samples. Reproducibility
between sample replicates was smaller than ±5 %.
For H2S analysis, pore water was filled into 5 mL polypropylene vials
and preserved with 0.25 mL 5 % ZnAc solution. H2S concentration was
measured photometrically (Specord 40, Analytic Jena) using the methylene
blue method (Cline, 1969).
The metal and total dissolved S (TSdiss) concentrations were analyzed
by ICP-OES (inductively coupled plasma optical emission spectrometry, iCAP 6300 DUO Thermo Fisher Scientific) after appropriate
dilution. Since high amounts of dissolved organic carbon (DOC) may cause severe interferences in the
ICP-OES element measurements, samples were boiled in Teflon beakers with
65 % HNO3 and subsequently 19 % HCl prior to analysis. The accuracy
and precision were routinely checked with the certified CASS standards as
described previously (Kowalski et al., 2012). The residual,
non-specified S fraction (ResS resulting from the difference between
TSdiss, H2S, and SO42- is suggested to consist primarily of
dissolved organic S, polysulfides, and S intermediates.
δ13C and δD values of methane were analyzed using the gas
chromatography–combustion technique (GC-C) and the gas
chromatography–high-temperature-conversion technique (GC-HTC). The gas was
directly injected in a gas chromatograph Agilent 7890 (Agilent Technologies,
Germany), the peaks were separated using a CP-PoraBOND Q GC column
(50 m × 0.32 mm × 5 µm, isotherm 60 ∘C, Varian). Methane was
quantitatively converted to the analysis gases CO2 and H2 in the
GC–Isolink interface (Thermo Finnigan, Germany) and directly transferred via
open split interface (ConFlo IV, Thermo Finnigan, Germany). The δ13C and δD values of both gases were then measured with the
isotope ratio mass spectrometer MAT 253 (Thermo Finnigan, Germany). Results
for δ13C ratios of methane are given in the usual δ notation versus the Vienna PeeDee Belemnite (VPDB) standard. δD–CH4 ratios were referenced to the Vienna Standard Mean Ocean Water (V-SMOW).
The carbon isotope values (δ13C) of DIC were measured from a
HgCl-preserved solution using a Thermo Finnigan MAT 253 gas mass
spectrometer coupled to a Thermo Electron Gas Bench II via a Thermo Electron
ConFlo IV split interface. NBS19 and LSVEC were used to scale the isotope
measurements to the VPDB standard. Based on replicate measurements of
standards, reproducibility was better than ±0.1 ‰ (Winde et al., 2014).
For the determination of sulfate isotope signatures, dissolved sulfate was
precipitated with 5 % barium chloride as barium sulfate (Böttcher
et al., 2007). After precipitation the solid was filtered, washed and dried,
and further combusted in a Thermo Flash 2000 EA elemental analyzer that was
connected to a Thermo Finnigan MAT 253 gas mass spectrometer via a Thermo
Electron ConFlo IV split interface with a precision of better than ±0.2 ‰. Isotope ratios are converted to the Vienna Canyon Diablo Toilite (VCDT) scale
(Mann et al., 2009). For oxygen isotope analyses, BaSO4 was
decomposed by means of pyrolysis in silver cups using a high-temperature
conversion elemental analyzer (HTO-, Hekatech, Germany) connected to an
isotope gas mass spectrometer (Thermo Finnigan MAT 253) (Kornexl et al.,
1999). The calibration took place via the reference materials IAEA-SO-5 and
IAEA-SO-6 and 18O/16O values were referenced to the V-SMOW standard.
Replicate measurements agreed within ±0.5 ‰.
Stable oxygen (O) isotope measurements of pore waters were conducted using a
CRDS system (Picarro L2140-i) versus the V-SMOW standard. International
V-SMOW, SLAP, and GISP in addition to in-house standards were used to scale the
isotope measurements.
Sediment analysis
Intact peat cores were collected with a perspex liner (ID: 59.5 mm) and
subsequently punched out layer by layer. The peat section protruding from
the end of the liner was divided into three subsamples for the analysis of (i) total reduced inorganic S (TRIS), (ii) total solid S (TSsolid) and
reactive iron, and (iii) the microbial community structure. In order to
minimize oxygen contamination, the outer layer of the peat core was omitted
and subsamples were immediately packed. The aliquot for TRIS analysis was
preserved with 1:1 (v/v) 20 % ZnAc. Subsamples for microbial analysis were
immediately stored in RNAlater stabilization solution to preserve DNA. A second core was taken for
the analysis of water content and dry bulk density. TSsolid and TRIS
samples were frozen within 8 h after collection. Aliquots for TSsolid
elemental analysis were further freeze-dried and milled in a planet-ball
mill.
TSsolid contents were analyzed by means of dry combustion using an Eltra CS-2000
after combustion at 1250 ∘C. The device was previously calibrated
with a certified coal standard and precision is better than ±0.02 %.
TRIS fractions were determined by a two-step sequential extraction of
iron monosulfides and pyrite (Fossing and Jørgensen, 1989). The acid
volatile sulfur (AVS) fraction was extracted by the reaction with
1 MHCl for 1 h under a continuous stream of di-nitrogen gas. The H2S released
was quantitatively precipitated as ZnS and then determined
spectrophotometrically with a Specord 40 spectrophotometer following the
method of Cline (1969). Chromium-reducible sulfur (CRS; essentially pyrite; FeS2), was extracted with hot acidic Cr(II)chloride solution. For
δ34S analysis in different TRIS fractions the ZnS was converted
to Ag2S by addition of 0.1 MAgNO3 solution with subsequent
filtration, washing, and drying of the AgNO3 precipitate as described by
Böttcher and Lepland (2000). The non-specified solid S fraction,
resulting from the difference between TSsolid, CRS, and AVS, was
suggested to present primarily organic-bond S (orgS). The δ34S
composition of this residual fraction was measured from the washed and dried
solid residue after the Cr(II) extraction step via C-IRmMS following the
approach of Passier (1999). Reactive iron was extracted from freeze-dried
sediments by the reaction with a 1 MHCl solution for 1 h (e.g., Canfield,
1989).
Iron was determined as Fe2+ after reduction with hydroxylamine
hydrochloride via spectrophotometry using ferrozine as the complexing agent
(Stookey, 1970). Reactive iron here is considered to be the sum of those
iron fractions that may still react with dissolved sulfide. This fraction
includes iron(III)oxyhydroxides and acid volatile sulfide (AVS, essentially
FeS) as well as a very minor contribution from dissolved Fe2+ in the
pore water (Canfield, 1989).
Microbial community analysis
Genomic DNA of 0.2–0.3 g of sediment was extracted with the EURx soil DNA kit
(Roboklon, Berlin, Germany) according to manufactory protocols. DNA
concentrations were quantified with a Nanophotometer® P360
(Implen GmbH, Munich, DE) and Qubit® 2.0 fluorometer
(Thermo Fisher Scientific, Darmstadt, Germany) according to the manufactory
protocols.
The 16S rRNA gene for bacteria was amplified with the primer combination
S-D-Bact-0341-b-S-17 and S-D-Bact-0785-a-A-21 (Herlemann et al., 2011).
The 16S rRNA gene for archaea was amplified with the primer combination
S-D-Arch-0349-a-S-17 and S-D-Arch-0786-a-A-20 (Takai and Horikoshi,
2000). The primers were labeled with unique combinations of bar codes. The
PCR mix contained 1× PCR buffer (Tris ⋅ Cl, KCl,
(NH4)2SO4, 15 mM MgCl2; pH 8.7) (Qiagen, Hilden,
Germany), 0.5 µM of each primer (Biomers, Ulm, Germany), 0.2 mM of
each deoxynucleoside (Thermo Fisher Scientific, Darmstadt, Germany), and
0.025 U µL-1 hot start polymerase (Qiagen, Hilden, Germany). The
thermocycler conditions were 95 ∘C for 5 min (denaturation),
followed by 40 cycles of 95 ∘C for 1 min (denaturation),
56 ∘C for 45 s (annealing), and 72 ∘C for 1 min
and 30 s (elongation), concluded with a final elongation step at
72 ∘C for 10 min. PCR products were purified with a Hi
Yield® Gel/PCR DNA fragment extraction kit
(Süd-Laborbedarf, Gauting, Germany) according to the manufactory
protocol. PCR products of three individual runs per sample were combined.
PCR products of different samples were pooled in equimolar concentrations
and compressed to a final volume of 10 µL with a concentration of 200 ng µL-1 in a vacuum centrifuge concentrator plus (Eppendorf,
Hamburg, Germany). Individual samples were sequenced in duplicates.
The sequencing was performed on an Illumina MiSeq sequencer by the company
GATC. The library was prepared with the MiSeq Reagent Kit V3 for 2×300 bp
paired-end reads according to the manufactory protocols. For better
performance due to different sequencing length we used 15 % PhiX control
v3 library.
The quality of the sequences was checked using the fastqc tool (FastQC A
Quality Control tool for High Throughput Sequence Data;
http://www.bioinformatics.babraham.ac.uk/projects/fastqc/, last access: 29 June 2018; by Andrews, 2010).
Raw sequence reads were demultiplexed, and bar codes were removed with the
CutAdapt tool (Martin, 2011). The subsequent steps included merging of
reads using overlapping sequence regions (PEAR; Zhang et al., 2014),
standardizing the nucleotide sequence orientation, and trimming and
filtering of low-quality sequences (Trimmomatic) (Bolger et al., 2014).
After quality filtering, chimera were removed by the ChimeraSlayer tool of
the QIIME pipeline. Subsequently, sequences were clustered into operational
taxonomic units (OTUs) at a nucleotide cutoff level of 97 % similarity and
singletons were automatically deleted. To reduce noise in the dataset,
sequences with relative abundances below 0.1 % per sample were also
removed. All archaeal libraries contained at least > 18 500 sequences,
while bacterial libraries contained at least > 12 500 sequences. OTUs were taxonomically assigned employing the GreenGenes
database 13.05 (McDonald et al., 2012) using the QIIME pipeline
(Caporaso et al., 2010).
Representative sequences of OTUs were checked for correct taxonomical
classification by phylogenetic tree calculations in the environment.
Relative abundance of sequences related to known methanogens, anaerobic
methanotrophs (ANME), and sulfate reducers were used to project microbial
depth profiles. Sequences have been deposited at NCBI under the BioProject
PRJNA356778 with the sequence read archive accession numbers
SRR5118134-SRR5118155 for bacterial and SRR5119428-SRR5119449 for archaeal
sequences, respectively.
ResultsPore water geochemical patterns and pore water isotope composition
Substantial amounts of dissolved salts with EC maxima of up to 11.5 mS cm-1 occurred at
peat depths below 30 cm b.s.f. (centimeters below surface; Fig. 2a, Table A1) and corresponded with brackish pore water proportions of up to
60 % (based on Baltic Sea salinity reported by Feistel et al., 2010).
Only at spot 1, with the greatest distance to the coastline, did lower EC values
(max. 3.4 mS cm-1) indicate minor brackish pore water proportions
(5 %–6 %). At the other three spots, EC values were similar, i.e.,
exhibited no lateral salinity graduation along the remaining Baltic
Sea–freshwater transect.
Depth distribution of electrical conductivity (EC, a) and pore
water O isotope composition (b). Panel (c) depicts a scatter plot of pore
water O isotope composition and salinity. Grey transparent dots in
(c) represent a common positive δ18O–H2O vs. salinity
relationship derived from a sampling campaign of Baltic Sea surface water
(Westphal, unpublished).
Vertical trends in pore water stable O isotope composition were similar for
all spots and complementary to the salinity and EC patterns with an upwards
increase from 60 to 10 cm b.s.f. (Fig. 2b). The resulting salinity–δ18O relationship was negative (except for the low salinity gradient at
freshwater spot 1) and thus inverse to the common salinity–δ18O
trend characteristic for Baltic coastal waters (Fig. 2c). This suggests that
distribution patterns of salinity have formed independently from evaporative
fractionating effects observed in the top pore water layers.
The pore water geochemistry in the peatland was increasingly diversified
with depth: while the top 10 cm b.s.f. was comparatively homogenous across all
spots, specific patterns evolving from diagenetic differences emerged
primarily in deeper pore waters. Principal component analysis (Fig. 3)
revealed the pore water geochemical composition below 10 cm b.s.f. to be
constrained by two major components that evolved in opposed lateral
directions and, in concert, explained 90 % of the variation in pore water
composition. A distinct gradient associated with a depth increase in EC and
the associated conservative ions (Cl-, Na+, Br-) suggests a
persistent brackish impact at spots 2, 3, and 4 (first principal component,
explained 55 % of the total variation). Only at spot 1, farthest away from
the coastline, was the EC increase with depth minute. This EC gradient was
further negatively correlated with pH, indicating a general decrease in pH
with depth and the highest pH values around 7.0 at spot 1. A second distinct
lateral gradient was delineated by the concentrations of dissolved Fe, Mn,
DIC, and Ca, which occurred in higher abundances at spots 1 and 2 closest to
the upstream terrestrial catchment boundary (second principal component,
explained 35 % of the total variation). Such a lateral shift in pore water
geochemistry is probably related to the supply of mineral solutes from
terrestrial inflow. In this regard, the pore water composition of spot 2
united the elevated supply in mineral compounds from terrestrial inflow with
persisting remnants of former brackish impact.
Principal component biplot of pore water geochemical patterns
within the peatland. Different colors indicate different sampling locations
within the brackish–freshwater continuum with spot 1 closest to the
freshwater catchment and spot 4 closest to the Baltic Sea. The size of the
data points scales with sampling depth (smallest points indicate surface
patterns; largest points indicate pore water composition at 60 cm in depth).
Sulfur speciation, S isotope patterns, and sulfate reducing communities
We found distinct differences in the S biogeochemical patterns across spots indicating different sulfate supply and transformation processes along the
terrestrial–brackish continuum. In the following, we structured the results
spot-wise according to the specific S regime and address first spot 1 (low
solid sulfur and low sulfate), then spots 3 and 4 (high solid sulfur and low
sulfate), and finally spot 2 (high solid sulfur and partially high sulfate
concentrations).
Spot 1
Spot 1 characterized by low salinities and mineral inflow from the near
freshwater catchment exhibited the lowest sulfate concentrations of ≤0.3 mM. H2S concentrations hardly exceeded the detection limit
(∼1µM, Fig. 4). Sulfate made up only a small
proportion of the TSdiss pool, thereby indicating a higher abundance of
a non-specified dissolved S fraction, probably composed of dissolved organic
S, polysulfides, and S intermediates.
Speciation of dissolved (a) and solid (b) S compounds, S isotope
composition of solid S compounds (c), and average relative abundances of
sulfate-reducing bacteria (SRB, d). δ34S and δ18O
ratios of SO42- are displayed in Fig. 6a. The residual dissolved S
(ResS in a) refers to a non-specified S fraction resulting from the
difference between total dissolved S, H2S, and SO42-. ResS is
most likely composed of dissolved organic S, polysulfides, and S
intermediates. Solid S fractions (b) include iron monosulfide
operationally defined as acid volatile sulfur (AVS), pyrite extracted as
chromium-reducible sulfur (CRS), and a residual fraction suggested to
consist primarily of organic S (orgS). δ34S at AVS could only
be measured at spot 1 and the top of spot 2. SRB were extracted from two
replicates of the 16S rRNA bacterial community sequencing and are assigned to
the Deltaproteobacteria (Delta-SRB) and the Nitrospirae phyla (genus
Thermodesulfovibrionaceae – Thermo-SRB). Chloroflexi Dehalococcoides
(Chloroflexi) have not been assigned to SRB in the classical sense; however,
they could be potentially involved in S metabolism (Wasmund et al., 2016).
Note the different x axis scales.
In addition, the abundance of solid S was lowest at spot 1 (≤0.7 %dry weight (dwt) TSsolid). Among solid S compounds, organic-bond S constituted
the dominant solid S fraction (0.1 to 0.5 %dwt) with relatively stable
δ34S ratios (+8.1 ‰ and +9.8 ‰). Pyrite
contents (measured as CRS) were low despite abundant pore water Fe and
available solid iron (Fig. 5). Only at spot 1 did we find a low though
consistent abundance of iron monosulfides (0.1 %dwt, measured as AVS).
Biogeochemical turnover processes here might operate under sulfate-limited
conditions resulting in lower sedimentary S contents and accumulation of
iron monosulfides.
Mobile Fe species. Available solid iron was extracted as HCl
soluble iron from the sediment matrix and is composed of iron mono-sulfide
and non-sulfidized ferric Fe.
In correspondence with the low sulfate contents, no sulfate-reducing
bacteria occurred at spot 1.
Spots 3 and 4
Despite the persisting brackish impact found in the deeper pore waters of
spots 3 and 4 closest to the Baltic Sea, we found hardly any pore water
sulfate in the top 20 cm b.s.f. (≤0.1 mM) and only moderate
SO42- levels down to 30 cm b.s.f. (0.1–1 mM). H2S abundance was
essentially restricted to the depth at spot 3 (up to 347 µM).
Low pore water sulfate concentrations prevented δ34S
measurements at the majority of the data points. However, the single δ34S value of +86.4 ‰ measured at 60 cm b.s.f. of
spot 3 (Fig. 6a) indicated a remarkable 34S enrichment in relation to
Baltic Sea water SO42- (+21 ‰; Böttcher
et al., 2007). Sulfur isotope fractionation to this extent is likely to
result from a superposition of enzymatic kinetic fractionation associated
with a reservoir effect and constitutes striking isotopic evidence for the
exhaustion of the brackish sulfate pool by intense DSR (Hartmann and
Nielsen, 2012). Despite missing isotope measurements, it is likely that the
low sulfate concentrations at the remaining depth sections of spot 3 and
along the depth profile of spot 4 result from the same intense sulfate
reduction processes.
(a) S and O isotope composition of sulfate. Sufficient
SO42- for δ34S and δ18O ratio analysis
was only available at the bottom of spot 2 and spot 3 (here only δ34S). (b) Rayleigh plot for measured SO42- depletion at spot 2.
We measured high amounts of TSsolid (up to 3.5 %dwt) at the depth of
spot 3. In both, spots 3 and 4, organic-bond S constituted the dominant solid
S fraction (0.5 to 3.3 %dwt) but was completely missing at the depth of spot 4. Pyrite was less abundant (0.2–0.3 %dwt) and exhibited a wide range of
δ34S ratios (-15 ‰ to +11 ‰). As pyrite
δ34S ratios essentially reflect the isotopic signature of the
sulfide pool derived from DSR (Butler et al., 2004; Price and Shieh,
1979), the found variation in pyrite δ34S ratios reflected
different stages of a reservoir effect that varies in response to the
openness of the system (i.e., connectivity to the sea).
In correspondence with the exhaustion of the brackish sulfate pool, the
relative abundance of SRB was generally small (< 5 %) and most
likely substrate-limited. SRB were from the Deltaproteobacteria class and the
Thermodesulfovibrionaceae genus of the Nitrospirae phylum. With 40 % relative abundance, Chloroflexi of the class
Dehalococcoidetes represented the dominating bacterial group at the 1 mM SO42- concentration depth of spot 3.
Spot 2
At spot 2 – the interface between brackish impact and mineral inflow from
the freshwater catchment – we found a sharp rise in SO42-
concentration from ≤0.3 mM at the top 20 cm up to 32.8 mM at 60 cm b.s.f.
The latter exceeded the quantities expected from marine supply (Kwiecinski, 1965; Feistel
et al., 2010) by a factor of 8. The pronounced
concentration gradient at spot 2 was associated with a remarkable variation
in the stable isotope composition showing a downcore decrease in δ34S–SO42- from +82.9 ‰ to +22.7 ‰ and
a decrease in δ18O–SO42- from +30 ‰ to
+11 ‰ (Fig. 6a). δ34S values > +80 ‰ at 30 cm b.s.f. of spot 2 suggest the brackish
sulfate pool in the top pore waters to be microbially exhausted under the
same reservoir effect as in spots 3 and 4. The δ18O and δ34S ratios of excess SO42- in 60 cm b.s.f. (δ34S:
+22.7 ‰; δ18O:
+11.4 ‰) corresponded well with modern-day seawater
SO42- (δ34S: +21 ‰; δ18O: +9 ‰; Böttcher et al., 2007).
Altogether, the sharp sulfate concentration and isotope gradients at spot 2
could demonstrate the entire spectrum of sulfate speciation from the
persistence of a marine sulfate reservoir at 60 cm b.s.f. towards progressing
sulfate depletion in the upper peat layers.
To test this hypothesis, we applied a closed-system (Rayleigh-type) model
(Eq. , Mariotti et al., 1981) to the data from spot 2 and gained an
estimate for the δ34S ratios of the initial SO42- reservoir (δ34SSO4,initial2-) and the kinetic isotope
enrichment factor ε:
δ34SSO4,depth2--δ34SSO4,initial2-=εln(fSO4,depth2-).
Here δ34SSO4,depth2- represents the S isotope values
measured in specific depths of spot 2, and fSO4,depth2-
constitutes the fraction of remaining pore water SO42- in relation
to the initial sulfate reservoir (32.8 mM SO42-, measured in 60 cm b.s.f. at spot 2). The fit through four data points (R2: 0.99;
p > 0.05) revealed the δ34S ratios of the initial
SO42- reservoir (+24 ‰) to be close to the
34S signature of the Baltic Sea (Fig. 6b). The isotopic offset is
within the uncertainty of the estimate. The isotope enrichment factor
ε was estimated to be -27 ‰, which is within
the range reported for DSR in laboratory studies with pure cultures
(Kaplan and Rittenberg, 1964; Canfield, 2001; Sim et al., 2011) and in
the field (Habicht and Canfield, 1997; Böttcher et al., 1998).
The pronounced sulfate distribution patterns at spot 2 went along with the
highest amounts of pyrite (0.5–1.4 %dwt). Pyrite contents increased with
depth and partially exceeded the amounts of organic-bond S. The patterns in
pyrite δ34S ratios did not correspond with the vertical trend
in sulfate availability. Instead, δ34S values were lowest in 20 cm b.s.f. (-15 ‰) and stabilized around
+2 ‰ below.
Interestingly, at peak sulfate supply of spot 2, the relative abundance of
Deltaproteobacteria did not exceed 5 %. Instead, the SRB community at depth was dominated by the Thermodesulfovibrionaceae genus
that contributed up to 21 % of all bacterial 16S rRNA sequences. Likewise
with spot 3, Chloroflexi of the class Dehalococcoidetes also represented the dominating bacterial group
at the depth of spot 2.
Dissolved methane concentrations, isotopic signature, and methanogenic
communities
Measured pore water CH4 concentrations were up to 643 µM
with
equivocal vertical patterns across spots (Fig. 7a), reflecting the
methane-specific spatial variability that evolves from small-scale
heterogeneity in production and consumption processes and from ebullitive
release events (Chanton et al., 1989; Whalen, 2005). Here, we use the
isotope composition of CH4 (Fig. 7b) and DIC (Fig. 7c) to provide a
clearer (and probably more robust) indication for patterns of methanogenesis
and methanotrophy. Methanogenesis is a highly fractionating process: in
comparison to the starting organic material (δ13C∼-27 ‰ in this study), the
produced CH4 is distinctively 13C-depleted, whilst at the same
time, CO2 becomes considerably enriched in 13C (Whiticar et
al., 1986). In this respect, high δ13C-DIC ratios up to
+4.2 ‰ suggest intense methanogenic (i.e., 13C-DIC
fractionating) processes in 20–40 cm b.s.f., whereas DIC on top was
comparatively depleted in 13C as is characteristic for methane oxidation
in the aerated surface layers. δ13C-DIC ratios below 40 cm b.s.f.
converged towards the isotopic signature of bulk organic C
(-26 ‰).
Concentration patterns and isotope ratios for CH4(a, b) and
DIC (c), as well as average relative abundances of methanogens and
methanotrophs (d).
At spot 2, we found the most pronounced downward drop in δ13C-DIC ratios with a minimum of -23.9 ‰ in 60 cm b.s.f. This pattern coincided with a consistent downward decrease in δ13C-CH4 ratios from -57 ‰ to -68 ‰ and suggests
that methanogenesis operates under higher 13C fractionation associated
with thermodynamically less favorable conditions at the bottom of spot 2.
δD ratios of methane did not exhibit a concurrent increase but
varied unrelated to δ13C-CH4 ratios in a range between
-333 ‰ and -275 ‰. Based on the C and D isotopic ratio
threshold raised by Whiticar (1986), acetate fermentation revealed to be the
dominant methane production pathway at our study site (Fig. 8). A concurrent
rise in both δD-CH4 and δ13C-CH4 ratios at the depth of
spot 1 suggests a shift towards dominating CO2 reduction and/or an
increase in methanotrophy.
Projection of the CH4 stable isotope composition to
differentiate dominating methanogenic pathways and methanotrophy. Isotope
thresholds to confine methanogenic pathways are based on Whiticar et al. (1986).
The concurrent increase in δ13C-CH4 and δD-CH4 values at spot 1 suggests a downward shift towards increasing
CO2 reduction or CH4 oxidation rates at depth.
Together with high δ13C-DIC ratios in the upper parts of the
peat, 16S rRNA sequences related to methanogens (Fig. 7d) provided further
evidence for intensive methane production. At spot 2, we found the largest
divergence with 90 % methanogen-related sequences at the surface while in
deeper regions (10–50 cm b.s.f.) less than 7 % of the archaeal domain could
be attributed to methanogens. Surprisingly, at 60 cm b.s.f. of spot 2,
methanogen percentages increased abruptly up to 41 % despite high
relative abundances of SRB. Spot 1 exhibited the lowest methanogen
proportions, which decreased from 21 % at the top down to 1 % in 50 cm b.s.f.
The methanogen community was mostly dominated by Methanosaeta, an obligate acetotrophic
archaea genus that thrives in terrestrial organic-rich environments.
The Methanosaeta proportion usually scaled with the methanogen percentage and contributed
70 %–100 % to the methanogenic community. Whilst methanogenic pathways
derived from the isotopic composition of CH4 can be obscured by the
fractionating effect of methanotrophy, the phylogenetic structure of the
methanogenic community provided clear evidence for acetate fermentation as
the
prevailing methanogenic pathway in most of the peatland.
Sequences related to aerobic methanotrophs of the genus Methylosinus were only found at
30 cm b.s.f. in spot 4 representing approximately 1.5 % of all bacterial
sequences (data not shown). Aerobic methanotrophs were underrepresented in
our dataset.
Consistent with the concurrent depth increase in δ13C-CH4
and δD-CH4, spot 1 (Fig. 8), situated at the fringe of the
freshwater catchment, exhibited high abundances of anaerobic methanotrophs
of the ANME-2d clade that are so far implicated to use NO3-
(Raghoebarsing et al., 2006) and/or Fe(III) (Ettwig et al., 2016) as
electron acceptors.
DiscussionPore water biogeochemical patterns
Overall, the pore water geochemistry of the Hütelmoor was characterized
by two different aspects: a legacy effect delineated by the lateral
brackish–terrestrial continuum below 20 to 30 cm in depth and an overlying
recent layer representing the uniform freshwater regime induced by
rewetting.
Despite a continuous groundwater inflow from the forested catchment
(Miegel et al., 2016), relics of former brackish and mineral
terrestrial inflow are preserved in the deeper layers of the peat body. This
is exemplified by high pore water EC values that exceeded those reported
directly after the last brackish water intrusion event in 1995 (Bohne
and Bohne, 2008). In fact, discharge within the peatland is channeled
through rapid flow in the drainage ditches while water movement within the
interstitial peat body seems to be mostly restricted to vertical exchange
processes (evaporation, precipitation) with minor lateral flow (Selle et
al., 2016). Therefore, we assume that drainage-induced hydrological
alterations reinforced the segregation of the peat pore matrix from
subsurface lateral exchange. This would allow for the preservation of
residual signals in deeper pore waters and would further confine
contemporary biogeochemical transformation processes to the recycling of
autochthonous matter. The new top freshwater layer, established after
flooding in 2010, overprints lateral differences along the brackish–fresh
continuum and unifies the upper pore water geochemistry in the entire
peatland.
Sulfur transformation
Along the entire brackish–terrestrial transect, virtually no sulfate was
abundant in the newly developed fresh pore water layer at the top 20 cm.
However, distinct differences in sulfur speciation across spots were
preserved below 20 cm b.s.f. and seemed to reflect the gradual exposure to
former brackish intrusion and terrestrial inflow.
Spot 1 appeared to be virtually unaffected by any brackish impact with
biogeochemical turnover processes operating under sulfate-limited
conditions. Low sedimentary S contents and the accumulation of iron
monosulfides as representative for freshwater environments are strong points
for this conclusion.
Also at spots 3 and 4, contemporary biogeochemical processes essentially
operated under sulfate-limited conditions, although these areas had been
exposed to flooding from the nearby Baltic Sea. High sedimentary S
concentrations in conjunction with the 34S composition of the remaining
sulfate suggest that the brackish sulfate reservoir has been essentially
exhausted through DSR with the produced sulfide being either incorporated as
diagenetically derived S in organic compounds or precipitated as
34S-enriched pyrite minerals (Brown and MacQueen, 1985; Hartmann and
Nielsen, 2012). Hence, if diking of coastal wetlands prevents the
replenishment of the brackish sulfate reservoir, the latter can be almost
completely consumed through DSR as has been demonstrated by the Rayleigh
distillation model. The rapid exhaustion of the brackish sulfate reservoir
is likely to be reinforced in coastal peatlands where vast amounts of C
compounds constitute an extensive electron donor supply for DSR.
Prevalent sulfate limitation at spots 1, 3, and 4 was reflected by the
virtual absence of the sulfate-reducing microbial community. Interestingly,
minor remnants of the brackish sulfate pool (1 mM SO42-) at the depth
of spot 3 were associated with 40 % relative abundance of Chloroflexi of the class
Dehalococcoidetes. Genomes of this group in marine sediments have been shown to code for
dsrAB genes (Wasmund et al., 2016). Through their ability to reduce sulfite
they may be involved in S redox cycling. Indeed, further research is
required to better establish their function in the S cycle.
S geochemistry at spot 2, which unites the effects of brackish water
intrusion with mineral inflow of terrestrial origin, differed substantially
from the other spots with remarkably high sulfate concentrations (33 mM) at
depth. The mineral impact from terrestrial inflow was not only reflected by
high concentrations of dissolved constituents (Fe, DIC, Mg, Ca, Mn) but also
by high contents of labile iron minerals and dissolved ferrous iron.
Interactions with poorly ordered ferric hydroxides can supply Fe(III) as
a competitive electron acceptor next to sulfate (Postma and Jakobsen,
1996) and may, therefore, inhibit the efficient microbial reduction of the
brackish sulfate reservoir. Amorphous ferric hydroxides effectively
suppressed DSR in a recently rewetted Baltic coastal wetland (Virtanen
et al., 2014). In our study, high contents of labile iron minerals and
dissolved ferrous iron at the depth of spot 2 coincided with a high abundance
of Thermodesulfovibrionaceae and a concurrently minor occurrence of Deltaproteobacteria. Recent in vitro experiments suggest
Thermodesulfovibrionaceae can utilize ferric iron as an electron acceptor next to sulfate (Fortney
et al., 2016). Indeed, the demonstration of Fe(III) reduction by
Thermodesulfovibrionaceae under in situ conditions is currently still pending. Nevertheless, high
contents of labile iron minerals, the remarkable accumulation of pore water
iron, and the absence of typical iron reducers (Geobacteraceae, Peptococcaceae, Shewanellaceae, Desulfovibrionaceae, Pelobacteraceae) could suggest
Thermodesulfovibrionaceae prefer Fe(III) as an electron acceptor over sulfate. Thus, the unique
SO42- concentration patterns at spot 2 may be attributed to the
inhibited microbial consumption of the brackish sulfate reservoir caused by
the delivery of alternative electron acceptors from the nearby freshwater
catchment.
Altogether, our results demonstrate the potential fate of the brackish
sulfate reservoir in coastal wetlands under closed system conditions caused
by diking. Microbial transformation processes have decoupled the sulfate
distribution patterns from the relic brackish impact and have caused marked
differences in contemporary sulfate biogeochemistry: on the one hand, DSR
exhausted the brackish sulfate reservoir in wide parts of the peatlands,
whereas, on the other hand, the preferential consumption of competitive
electron acceptors from terrestrial origin allowed for the local
accumulation of large sulfate concentrations. Indeed, these relic signals of
brackish–terrestrial intermixing are constrained to the deeper pore water
regions below 30 cm b.s.f. as recent rewetting measures established a
homogeneous freshwater regime in the top layers of the entire peatland.
Methane production and consumption
δ13C-DIC ratios and a thriving methanogenic community indicate the
establishment of distinct methane production zones in the recently formed
freshwater layer across the entire peatland. In line with the prevalent
freshwater characteristics of the newly formed pore water layer, the
methanogen community was dominated by Methanosaeta, an obligate acetotrophic genus
typical of terrestrial organic-rich environments. Indeed, thermodynamically
favorable methanogenic conditions were confined to the top layers since
isotopic evidence and archaeal distribution patterns indicate a downward
shift towards non-fractionating metabolic processes (Barker, 1936;
Lapham et al., 1999) at the bottom. This vertical transition was most
pronounced at spot 2, probably indicating a potential suppression of
methanogenesis by high concentrations of sulfate and labile ferric iron
compounds at depth.
Surprisingly, we observed mutual coexistence of SRB (22 % of all bacterial
sequences) and methanogens (> 40 % of all archaeal sequences)
at high SO42 concentrations (32.8 mM) in 60 cm b.s.f. at spot 2.
Simultaneous methanogenesis and DSR have been reported under the abundance
of methanol, trimethylamine, or methionine as methanogenic precursors
(Oremland and Polcin, 1982). However, the concurrent high abundance of
Methanosaeta (30 %) at the depth of spot 2 suggests competitive consumption of acetate by
both SRB and methanogens. Although Liebner et al. (2015) emphasized the
relevance of community structure with regard to prevailing methanogenic
pathways, total abundance data could potentially yield more insight into this
issue.
Sequences related to aerobic methanotrophs of the genus Methylosinus were only found at
30 cm b.s.f. in spot 4, representing approximately 1.5 % of all bacterial
sequences (data not shown). The phenomenon of a lagged reestablishment of
methanotrophs in comparison to methanogens after rewetting in this
particular peatland is addressed in another publication (Wen et al.,
2018).
Despite the overlap of methane production zones anticipated from δ13C-DIC ratios with sulfate reduction zones, we could not find evidence
for the syntrophic consortium of anaerobic methanotrophs (ANME) and sulfate
reducers that is commonly associated with the anaerobic oxidation of methane
coupled to sulfate reduction (AOM-SR) in marine environments (Boetius et
al., 2000). However, we cannot exclude that AOM-SR is driven by archaea that
are so far not known for this function. One potential candidate phylum is
the Bathyarchaota that have been shown to encode an untypical version of the functional
gene for methane production and consumption (methyl coenyzme M reductase
subunit A, mcrA) (Evans et al., 2015). These archaea dominated spot 2 with
48 %–97 % relative sequence abundance of the archaeal community between 10
and 60 cm (data not shown).
While we cannot supply microbial evidence for AOM-SR, high abundances of
anaerobic methanotrophs of the ANME-2d clade at spot 1 suggest anaerobic
methane oxidation coupled to electron acceptors of terrestrial origin.
Methanotrophs of the ANME-2d clade are so far known to utilize
NO3- (Raghoebarsing et al., 2006) and ferric iron (Ettwig
et al., 2016) as electron acceptors, both of which were abundant at the
respective spot. This observation is further supported by the trend in
δ13C-CH4 and δD-CH4 that potentially
indicates a downward increase in methanotrophy at spot 1. The biogeochemical
characteristics at this very location result most likely from formerly drier
conditions due to slightly higher elevation in combination with prevalent
inflow from the nearby forest catchment.
Our results demonstrate how rewetting of a coastal peatland established a
distinct freshwater regime in the upper pore water layers, which, in
conjunction with prevalent anaerobic conditions and a vast stock of labile C
compounds, offers favorable conditions for intense methane production and
explains the high methane emissions reported in Hahn et al. (2015) and
Koebsch et al. (2015). As intense methane production was confined to the
upper pore water layers in the entire peatland, it did not interfere with
high sulfate concentrations locally preserved as the legacy of former brackish
impact in the bottom. Instead, isotopic and microbial evidence suggested
mineral compounds of terrestrial origin to constitute an electron acceptor
for anaerobic methane oxidation, which is an often neglected – though it is
an
important process in freshwater environments (Segarra et al., 2015). Our
results indicate that this process can also occur in disturbed coastal
peatlands. Indeed, the quantitative effects of anaerobic methane consumption
on methane emissions in coastal and/or rewetted peatlands need to be
addressed in future studies.
Conclusions
In this study, we investigated the biogeochemical and hydrological
mechanisms that turn disturbed and remediated coastal peatlands into strong
methane sources. Our study demonstrates how human intervention overrides the
sulfate-related processes that suppress methane production and thereby
suspends the natural mechanisms that mitigate greenhouse gas emissions from
coastal environments. Hence, the climate effect of disturbed and remediated
coastal wetlands cannot simply be derived by analogy with their natural
counterparts. Instead, human alterations form new transient systems where
relic brackish signals intermingle with recent freshwater impacts. The
evolving biogeochemical patterns overprint naturally established gradients
formed, for instance, by the distance to the coastline. In particular, the
decoupling of sulfate abundance from salinity is of high practical relevance
for greenhouse gas inventories that establish methane emission factors based
on the empirical relation to salinity as an easily accessible proxy for sulfate
concentrations.
Coastal environments are subject to particular pressure by high population
density while at the same time their potential as coastal buffer zones is
moving more and more into the focus of policy makers and land managers. From
a greenhouse gas perspective, the exposure of diked wetlands to natural
coastal dynamics would literally open the floodgates for a replenishment of
the marine sulfate pool and constitute an efficient measure to reduce
methane emissions. However, in practice, this option has to be weighed
against concurrent land use aspects.
Data availability
Geochemical data are represented within this paper in the appendix
(Table A1). Sequences have been deposited at NCBI under the BioProject
PRJNA356778 with the sequence read archive accession numbers
SRR5118134-SRR5118155 for bacterial and SRR5119428-SRR5119449 for archaeal
sequences.
Site parameters, pore water, and soil characteristics. Water level
and soil depth are given in centimeters above and centimeters below surface (cm a.s.f. and cm b.s.f.,
respectively).
FK and MEB have formulated the research question and planned the study
design. FK acquired funding. FK, GJ, MK, MW, and SK collected the samples.
MEB, SL, AS, MG, TS, and SK provided resources and lab instrumentation for
sample analysis. FK, AS, IS, MK, GJ, SK, and JW conducted the geochemical
analyses. MW, SL, and VU conducted the microbial sequencing analysis. BL
validated the results. FK visualized the data and prepared the original
draft with contributions from all coauthors.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
This work was supported by the DFG Research Training Group BALTIC TRANSCOAST (grant DFG GRK
2000) and the Helmholtz Terrestrial Environmental Observatories (TERENO) Network. This is BALTIC TRANSCOAST publication number
GRK2000/0023. Franziska Koebsch was supported by the Helmholtz Association of German
Research Centers through the Helmholtz Postdoc Program (grant PD-129) and
the Helmholtz Climate Initiative REKLIM (Regional Climate Change). Franziska Koebsch was further
supported by the European Social Fund (ESF) and the Ministry of Education,
Science and Culture of Mecklenburg-West Pomerania within the scope of the
project WETSCAPES (ESF/14-BM-A55-0030/16). Torsten Sachs and Susanne Liebner were each supported by
a Helmholtz Young Investigators Group (grants VH-NG-821 and VH-NG-919).
Biogeochemical and stable isotope work was supported by the Leibniz
Institute for Baltic Sea Research (IOW). We wish to express our gratitude to
Lisa Kretzschmann, Anke Saborowski, and Simon Strunk for their commitment to field
work under tough conditions. Bennet Juhls and Simon Strunk have helped with map
creation. The study would not have been possible without the laboratory and
bioinformatics support by Andrea Gottsche, Anke Saborowski, Lisa Kretzschmann, Anne Köhler, Birgit Plessen, Vera Winde, Fabian Horn, Xi Wen, and
Heiko Baschek.
We acknowledge financial support by Deutsche Forschungsgemeinschaft and Universität Rostock within the funding programme Open Access Publishing.
Review statement
This paper was edited by Jianming Xu and reviewed by two anonymous referees.
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