Introduction
The amount of organic carbon (OC) stored in the northern circumpolar
permafrost (PF) amounts to ∼1300 Pg OC of which
∼800 Pg OC is perennially frozen (the remaining 500 Pg is
non-permafrost, seasonally thawing active-layer permafrost or talik;
Hugelius et
al., 2014). Northern Hemisphere circumpolar soils thereby hold roughly half
of the global soil OC pool (Tamocai
et al., 2009). Modelled future climate scenarios predict continued amplified
warming in the Arctic for the coming 100 years
(IPCC, 2013). This will further destabilise
permafrost, leading to increased delivery of terrestrial OC to the Arctic
Ocean. The potential decomposition of this relict permafrost carbon (PF-C)
and its subsequent release to the atmosphere as CO2 or
CH4
constitutes a positive feedback to global warming
(IPCC,
2013; Koven et al., 2011; Schuur et al., 2015; Vonk and
Gustafsson, 2013). Considering the size of the Arctic PF-C pool it is
important to better understand the dynamics and extent of its vulnerability
to remobilisation in response to climate warming.
Many recent studies have focused on current carbon cycling in the Arctic
land–ocean continuum with possible linkages to climate change. Constraining
how this system responded to earlier climate warming may help us to better
predict the future response of PF-C and its climate couplings. The last
glacial–interglacial transition constituted a major climate rearrangement on
Earth. The increase in mean temperature coupled with sea level rise is
thought to have profoundly destabilised PF-C and further released CO2
to the atmosphere
(Ciais
et al., 2013; Crichton et al., 2016; Köhler et al., 2014; Tesi et
al.,
2016a). Several studies have suggested that there was a warming-coupled
translocation of terrestrial carbon during the climate warming that ended
the latest glacial period
(e.g.,
Bauch et al., 2001a; Ciais et al., 2013; Mueller-Lupp et al., 2000; Tesi et
al.,
2016a) similar to what is predicted to happen as a consequence of the
anthropogenic climate change (Barnhart
et al., 2014; Vonk and Gustafsson, 2013).
Many of the previous Holocene timescale studies in the East Siberian Arctic
Shelf (ESAS) have focused on the Laptev Sea
(e.g.,
Bauch et al., 2001b; Mueller-Lupp et al., 2000; Tesi et al., 2016a). This
study focuses on the East Siberian Sea (ESS), which has not yet been
extensively studied, especially for the historical reconstruction of PF-C
dynamics. The ESS receives terrestrial OC by coastal erosion, fluvial inflow
and possibly seabed erosion
(Karlsson
et al., 2016; Semiletov et al., 2005; Stein and Macdonald, 2004; Tesi et
al., 2014, 2016b; Vonk et al., 2010). The coast of the ESS is dominated by
carbon-rich Ice Complex deposits (ICD) consisting of old Pleistocene
material
(Schirrmeister
et al., 2011; Semiletov, 1999a, b; Vonk et al., 2012). These large ICD
bluffs are vulnerable to coastal erosion
(Semiletov et al., 2013;
Stein and Macdonald, 2004; Schirrmeister et al., 2011; Vonk et al.,
2012). Coastal erosion can be further intensified with warming-enhanced
processes like loss of sea ice cover, increasing frequency of storms,
degradation of ice-bonded coasts and sea level rise
(Barnhart et
al., 2014; Jones et al., 2009; Stein and Macdonald, 2004). The largest
rivers directly emptying into the ESS are Indigirka and Kolyma with
suspended matter discharge of 11.1×1012 gyr-1 and 123±19×109 gyr-1 (Gordeev,
2006; McClelland et al., 2016, respectively) and an input also from the
Lena River. The Lena River drains into the Laptev Sea, but its exported
terrestrial OC is also transferred to the ESS via the Siberian Coastal
Current
(e.g.,
Alling et al., 2012; Sánchez-García et al., 2011). However, studies
by
Vonk
et al. (2010, 2012) suggest that the contribution of ICD-PF erosion to the
ESS sediment OC dominates over river discharge (ranging from 36 to 76 %
in comparison to 5–35 %, respectively).
In this study we investigate land-to-ocean transfer and the fate of PF-C during
the latest state of the post-glacial eustatic sea level rise until the
present day. Our main objectives are to determine the sources and
remobilisation fluxes of terrestrial OC and the composition and
degradation status of the OC that was buried in ESS sediments during the
Holocene. We characterise the OC composition by quantifying lignin phenols,
cutin acids and other compounds yielded upon CuO oxidation to constrain the
sources and degradation status of PF-C and the contribution of
marine OC. Furthermore, we use a mixing model based on the isotopic
composition (Δ14C, δ13C) of the deposited OC to
quantify the contribution of three different sources: topsoil-PF from
active-layer deepening, ICD-PF and marine plankton. Additionally, we study
how OC deposition fluxes have changed over time in response to the sea level
rise and Holocene warming.
Map of the East Siberian Arctic Shelf showing the location of
the sampling site (station SWERUS C3-1-58; Schlitzer, 2015). Also shown in the map is the location of the
sediment core PC23 (station SWERUS C3-1-23;
Tesi et al., 2016a). The red line
marks the isobath (34 m of water depth), which is approximately where the coastline was at the beginning of the sediment archive (GC58) ∼9500 calyrs BP
(Lambeck et al.,
2014).
Materials and methods
Background and study area
The ESS is located off the northeastern Siberian coast between the Laptev Sea
and the Chukchi Sea (Fig. 1). The ESS is one of the largest shelf seas
(987 000 km2) in the Arctic Ocean and one of the shallowest
(mean depth 52 m; Jakobsson, 2002).
Thermokarst landscapes (i.e., thawing ice-rich permafrost) cover
∼20 % (3.6×106 km-2) of the northern
circumpolar permafrost region
(Olefeldt et al., 2016). Ice
Complex deposit and thermokarst landscapes cover 2400 km of the ESS
coastline (Grigoriev and Rachold, 2003). The modern average rate
of coastal retreat in the ESS and the adjacent Laptev Sea is 1–10 myr-1 (Grigoriev, 2010), though even higher
retreat rates (up to 24 and 30 myr-1) have been reported locally in the most
actively eroding parts
(Kanevskiy
et al., 2016; Romanovskii et al., 2004). The coastal erosion rates have
increased in the Arctic in recent decades
(Barnhart
et al., 2014; Günther et al., 2015; Jones et al., 2009). According to
recent studies
(Bröder
et al., 2016b; Semiletov et al., 2013; Tesi et al., 2016b; Vonk et al.,
2012), a large fraction of the remobilised PF-C is degraded during
cross-shelf transport and released back to the contemporary carbon cycle. To
better predict the consequences of the permafrost thaw, it is important to
understand both the amount of remobilised organic carbon and its
fate.
The shelf of the ESS contains terrestrial permafrost formed during the sea
level low of the last glacial maximum
(Jakobsson et al., 2014). During the
Pleistocene–Holocene transition the ESAS was flooded when the sea level rose
rapidly
(Lambeck
et al., 2014; Mueller-Lupp et al., 2000). This global marine transgression
started ∼20000 calyrs BP
(Lambeck et al.,
2014) and flooded the ESAS between ∼11000 and ∼7000 calyrs BP
(Bauch
et al., 2001a; Mueller-Lupp et al., 2000). The rate of the sea level rise was
on the order of 1 cmyr-1 or more (Cronin et al.,
2017; Stanford et al., 2011) in the early Holocene. The sampling site of the
sediment core investigated in this study was flooded around
11 000 calyrs BP (Lambeck et al.,
2014). The early Holocene temperatures in the Arctic regions were on average
1.6±0.8 ∘C higher than today
(Kaufman
et al., 2004) and the sea ice was at a low (Fisher et al.,
2006).
Post-glacial sea level rise with warming and wetting of the climate caused a
major relocation of permafrost carbon from land to the Arctic Ocean
(Bauch
et al., 2001a; Tesi et al., 2016a). Today the period when the ESS is only
partially covered with sea ice is on average 3 months per year, which is one
of the reasons why the area remains fairly unstudied
(Stein and Macdonald, 2004; Vetrov and
Romankevich, 2004).
Sampling
A gravity core (called GC58) was collected in the ESS at 54 m of water depth as
part of the international SWERUS-C3 research expedition on I/B Oden in August 2014.
The coring site (Leg 1, station 58; 74.4387∘ N,
166.0467∘ E) is located ∼500 km from the modern
shoreline (Fig. 1). An additional sediment core was collected at the same
site (MUC58) using a sediment multicorer (Oktopus GmbH, Germany), which is
specifically designed to preserve the sediment–water interface. The total
length of GC58 was 78 cm, while MUC58 was 32 cm long. The GC58 core was split
in half during the expedition and kept refrigerated (+4 ∘C). In
the laboratory at Stockholm University, one half was subsampled at
1 cm
intervals and kept frozen at -18 ∘C. The multicore was sliced
during the expedition at 1 cm intervals and then immediately frozen
(-18 ∘C). Prior to analyses, the samples were freeze-dried at the
Department of Environmental Science and Analytical Chemistry, Stockholm
University, Sweden.
210Pb dating
Radiogenic 210Pb was analysed with a gamma-ray spectrometer (GRS) at
the Department of Geology of the Swedish Museum of Natural History in
Stockholm, Sweden. The GRS determines the decay energy of radioisotopes in
counts per second by measuring the gamma emission of the sample at a known
energy level.
Radiocarbon (14C) ages of the mollusc shells retrieved from
the sediment core GC58. The 14C ages are shown in years BP with an age
error (yrs) and as calibrated 14C ages (calyrs BP) with 2 standard
deviations (±2σ) of the individual 14C dates. The
calibration curve used was Marine13 (Reimer et
al., 2013) and a ΔR value of 50±100 yrs
(Bauch et al., 2001a). Also shown are the
δ13C (‰) values of the mollusc shells.
Corrected
NOSAMS
Type
Age 14C
Age error
δ13C
Age 14C
Age 14C
2σ
depth*
Accession no.
(yrs BP)
(yrs)
(‰)
(Cal yrs BP)
(Cal yrs BP)
(cm)
median
mean
3.5
OS-119395
Mollusc shell, fragments
895
25
0.55
462
455
184
8.5
OS-120688
Mollusc shell, fragments
> Modern
–
1.70
39
45
64
34.5
OS-120689
Mollusc shell, fragments
2260
20
1.55
1807
1806
250
39.5
OS-120690
Mollusc shell, fragments
2210
15
1.55
1748
1746
244
47.5
OS-123161
Mollusc shell, fragments
7960
35
0.90
8372
8372
220
51.5
OS-119396
Mollusc shell, fragments
8010
25
1.06
8426
8429
224
54.5
OS-120691
Mollusc shell, fragments
8020
20
0.49
8437
8441
226
65.5
OS-119397
Mollusc shell, Macoma calcarea
8780
25
-2.46
9384
9372
234
72.5
OS-120692
Mollusc shell, fragments
8880
20
-0.91
9493
9499
244
78.5
OS-120693
Mollusc shell, fragments
8950
25
-0.79
9579
9595
264
* Corrected depth is the original depth +3 cm to account for core top loss
during sampling (Sect. 2.4).
Prior to the GRS analysis, a subsample of approximately 10 g was homogenised
and placed in a plastic container for at least 3 weeks to reach secular
equilibrium between the radioisotopes of lead and radium (210Pb and
226Ra, respectively). The samples were analysed for 210Pb (46.51 keV),
226Ra (186.05 keV) and 137Cs (661.66 keV) on an EG&G
Ortec® coaxial low-energy photon spectrometer containing a
high-purity germanium detector. The counting period for each sample lasted
1–3 days depending on the amount of 210Pb in the sample. An
externally calibrated U-series standard (pitchblende; Stackebo, Sweden) was
used to determine the relative efficiency of the gamma detector system. For
each sample a minimum of 350 counts was acquired. A blank (empty container)
sample was measured to correct for the background activity. The original
method is described in detail by
Elmquist et al. (2007).
Two different models were used for the 210Pb dating: a CRS (constant rate
of supply) model which assumes a constant rate of supply of excess
210Pb fallout and a CIC (constant initial concentration) model which
assumes a constant initial concentration of excess 210Pb (Appleby and
Oldfield, 1978).
Bayesian modelling of 14C ages for the chronology
For the age–depth model construction, mollusc shells retrieved (hand-picked)
from GC58 were analysed for their radiocarbon (14C) content at the
US NSF National Ocean Sciences Accelerator Mass Spectrometry (NOSAMS)
facility at the Woods Hole Oceanographic Institution (WHOI; MA, USA). Prior
to the analysis the mollusc shells were rinsed with MilliQ water and
sonicated. The analysis followed standard procedures of NOSAMS
(Pearson et al., 1998; Table 1).
To account for natural differences in the amount of 14C in the
atmosphere and differences between the marine environment and the
atmosphere (e.g., Stuiver and Braziunas, 1993), all
14C data were calibrated with the Marine13 calibration curve
(Reimer et al., 2013). The offset in the local
reservoir age was taken into account by using a ΔR of 50±100 years.
Since there are no ΔR values for the ESS in the literature,
this ΔR value was taken from a study in the Laptev Sea
(Bauch et al.,
2001a). The radiocarbon dates are reported in calendar years before present
(calyrs BP; Stuiver and Polach, 1977).
The age model of the core was built with the OxCal v4.2 program based on the
radiocarbon-dated mollusc shells and a depositional model (P_sequence, k=0.5; Ramsey,
2008; Ramsey and Lee, 2013). Also, the base of the adjacent multicore
dated with 210Pb was used in the model. The 210Pb date used was an
average age (50 yrs BP) from the two 210Pb dating models (CRS, CIC) for
the bottom layer (12.5 cm) of the multicore (Table S2 in the Supplement). The
age model of GC58 was constructed with a Bayesian statistics approach using
the reservoir age (ΔR) and the depth as a prior model and measured
radiocarbon dates as likelihoods. The posterior probability densities were
acquired with a Markov chain Monte Carlo procedure which calculates possible
distributions in order to date each sediment layer using the given prior
model and likelihoods (Ramsey, 2008).
Sampling with a heavy gravity corer often disturbs the sediment–water
interface and thereby causes losses of the surface sediments. The organic
carbon (OC) content of GC58 was therefore compared to the OC content of the
adjacent MUC58 to identify possible loss. According to the
comparison, the top 3 cm was likely lost in GC58 (Fig. S1 in the Supplement)
and thus corrected for.
Alkaline CuO oxidation
Microwave-assisted alkaline CuO oxidation was carried out using the method
by Goñi and Montgomery (2000). Each
homogenised subsample of around 300 mg was mixed with 300 mg of cupric oxide
(CuO) and 50 mg of ammonium iron (II) sulfate hexahydrate
((NH4)2Fe(SO4)2⚫6H2O). After thorough
mixing, nitrogen-purged 2M NaOH was added to each sample. Alkaline oxidation
was performed with an UltraWAVE Milestone 215 microwave digestion system at
150 ∘C for 90 min.
A known amount of internal recovery standards (ethyl-vanillin, cinnamic
acid) was added to the CuO reaction products and then acidified to pH 1 with
concentrated HCl (35 %). The CuO reaction products were repeatedly
extracted using ethyl acetate (EtOAc). Anhydrous sodium sulfate
(NaSO4) was added to remove the remaining water. The extracts were
dried in a CentriVap (Christ RVC 2-25) at 60 ∘C, redissolved in
pyridine and stored in a freezer (-18 ∘C) until further
analysis.
Finally, the samples were analysed with a gas chromatograph mass
spectrometer (GC-MS; Agilent 7820A) using a DB5-MS capillary column
(60 m × 250 µm, 0.25 µm stationary phase thickness; Agilent J&W) at
an initial temperature of 60 ∘C followed by a ramp of
5 ∘Cmin-1 until reaching 300 ∘C. Prior to the GC-MS
analysis, the extracts were derivatised with bis(trimethylsilyl)trifluoroacetamide (BSTFA) +1 % trimethylchlorosilane (TMCS) to
silylate exchangeable hydrogens. The quantification of the samples was based
on the comparison of the key ions to commercially available standards.
Concentrations of CuO oxidation products were normalised to the organic
carbon content of the sample and are reported as mgg-1 OC.
Bulk organic carbon and stable carbon isotope analyses
For the total organic carbon content (TOC), the total nitrogen content (TN)
and the stable carbon isotope analysis (δ13C) of TOC,
subsamples of 10–15 mg were homogenised and placed in silver capsules,
acidified with 1.5 M HCl to remove carbonates and then dried at 60 ∘C. The TOC, TN and δ13C-TOC were quantified with an elemental
analyser (Carlo Erba NC2500) connected via a split interface to a Finnigan MAT
Delta V mass spectrometer at the Stable Isotope Laboratory of the Department
of Geological Sciences at Stockholm University.
For radiocarbon (14C) analysis of the bulk organic carbon, subsamples
of sediment were acidified with 1.5 M HCl and sent to NOSAMS. To account for
the time between the deposition and the measurement, the 14C dates were
calibrated with Eq. (1) using the age data derived from the age model.
The bulk radiocarbon data are reported as Δ14C
(Stuiver and Polach, 1977):
Δ14C=(Fm×eλ(1950-Yc)-1)×1000,
where Fm is the fraction modern, λ is 1/mean life of
radiocarbon = 1/8267 and Yc is the year of collection derived from the age model
(Stuiver and Polach, 1977).
Source apportionment
The carbon isotope fingerprint of OC (Δ14C, δ13C)
can be used to quantitatively diagnose the relative contribution of
topsoil-PF, ICD-PF and marine OC assuming isotopic mass balance
(e.g., Vonk et al.,
2012). In other words, the carbon isotopic signatures may help to understand
whether the OC comes from coastal erosion as a result of the post-glacial
warming and sea level rise, active-layer deepening of permafrost carbon in
the watershed (as a response to the post-glacial warming) or sedimentation
of marine phytoplankton. These different sources have a natural variability
in their isotopic composition (endmembers). This variability needs to be
taken into account to correctly estimate the relative source contributions
and the associated uncertainties (e.g.,
Andersson, 2011). In previous studies a Bayesian Markov chain Monte Carlo
(MCMC) approach has been used to estimate the relative source
contributions for individual data points
(Andersson
et al., 2015; Tesi et al., 2016a). Here, we expand this approach to include
the time dependence of the down-core isotopic signatures, taking
advantage of the relatively small variability in the 78 δ13C
data points, whilst also using the 10 Δ14C points. The
time dependence of different proportions was taken into account by following
the approach of Parnell et al. (2013).
The method is described in detail in the “Supplementary methods” section in the Supplement.
The endmember values for the three source classes were taken from the
literature (ICD-PF and topsoil-PF values compiled in
Vonk et al., 2012; marine OC
from Smith et al., 2002) for topsoil-PF (Δ14C=-126±54 ‰, δ13C=-28.2±1.96 ‰; mean ± standard
deviation) representing thaw of the active layer of permafrost, marine
OC
(Δ14C=-60±60 ‰, δ13C=-21±1 ‰) resulting from primary
production of phytoplankton and ICD-PF (Δ14C=-940±84 ‰, δ13C=-26.3±0.63 ‰) resembling the old Pleistocene material from
coastal erosion. The endmember value for ICD-PF was corrected with Eq. (1)
to account for the age of the deposition.
Grain size analysis
Prior to the grain size analysis subsamples of sieved (500 µm)
sediments from GC58 were homogenised. The grain size analysis was done with
a Malvern Mastersizer 3000 laser diffraction particle size analyser, which
can measure particles between 10 nm and 3.5 mm. Sodium hexametaphosphate (10 %)
was used to disaggregate the particles suspended in deionised water.
To further aid the disaggregation, all samples were exposed to ultrasound
for 60 s and allowed to disperse in continuous flow for 3 min in total
(including 60 s of ultrasonication) prior to the measurements. To control
the concentration of the sample in the flow during the measurements, the
obscurity was kept between 5 and 15 %. High sample obscurity (i.e., high
concentration) would cause multiple light scatterings, thus distorting the
results. Each sample was analysed in five replicates. The measurements were
carried out at the Department of Geological Sciences at Stockholm
University, Sweden.
An age–depth model of the sediment core GC58 based on
radiocarbon-dated
(14C) mollusc shells (see Table 1) and 210Pb (base of a
multicore collected at the same location; see
Table S2 in the Supplement). All the modelled dates were calibrated with the Marine13 calibration
curve (Reimer et al., 2013). A ΔR value
of 50±100 yrs was used to account for the differences in the local
reservoir age based on Bauch et al.
(2001a). The core GC58 dates back to ∼9500 calyrs BP. The
calibrated age probability distributions are plotted for each radiocarbon
date in grey. Outliers are coloured red. The blue shading indicates the
modelled 2σ probability intervals for the entire depth range of the
core, and the tiny black curves indicate 2σ for the individual measurements.
Results and discussion
Age chronology of the core
The deepest part of the sediment core GC58 dates back to ∼9500 calyrs BP, i.e., to the early Holocene. The age–depth model shows an evident
hiatus in the middle of the core between 39.5 and 40.5 cm resulting in an
age gap of ∼ 6500 years (∼8200–1700 calyrs BP; Fig. 2).
In addition, there is a shorter gap in the chronology between
∼9300 and ∼8500 calyrs BP. In studies from
the adjacent Laptev Sea such age discrepancies have not been observed
(Bauch
et al., 2001a, b; Tesi et al., 2016a). It therefore seems
likely that there has been a local event causing the removal of sediment
layers. There might not have been accumulation during those periods, or the
age gap could be a condensed unit of sediment. An actual sediment transport
process giving rise to such a putative total halt in the sedimentation rate
is rather elusive and unlikely. Since the whole ESAS is a very shallow shelf
where sea ice is formed (Conlan et
al., 1998; Jakobsson, 2002), another explanation for an age gap is ice
scouring as observed in the Laptev Sea (Ananyev et al.,
2016), especially at ∼8500 calyrs BP when the sea level was
around 18 m lower
(Lambeck et al.,
2014) than today and the water depth at the coring site was around 32 m. At
the time of the second age gap (∼1700 calyrs BP), the water
depth at the coring site was approximately 52 m. An ice scouring event could
have formed a gouge at the sea bottom that was later refilled with sediment
(Barnes et al., 1984).
The accumulation rates of GC58 obtained from the 14C measurements vary
between 0.2 and 1.4 mmyr-1 (17.0–138.9 cmkyr-1) and mass
accumulation rates (MAR) spanned 0.02–0.1 gcm-2yr-1.
Bauch et al. (2001a) have reported similar sedimentation rates (0.1–2.6 mmyr-1)
from the outer shelf of the Laptev Sea around the same time period. The
linear sedimentation rate for the adjacent sediment core MUC58 derived from
210Pb dating is 1.3 mmyr1 with an average MAR of
0.03 gcm-2yr-1. Similar accumulation rates with 210Pb-dated sediment cores
have been reported in other studies from the ESS: 1.1–1.6 mmyr-1 (Vonk et al., 2012) and
1.4–1.5 mmyr-1 (Bröder et
al., 2016a). The slight difference in accumulation rates using 210Pb
chronology compared to 14C may be due to active biological mixing
giving higher accumulation rates for the shorter timescale of more
surficial sediments
(Baskaran
et al., 2017; Boudreau, 1994).
Sediment grain size, stable carbon isotopes and biomarker composition of
organic matter
Grain size can be used to describe the depositional environment. The
sediment core GC58 consists mostly of clay and silt with a fraction of sand
(Fig. S2 in the Supplement). The higher sand content that is observed at
∼8500 calyrs BP may reflect a higher-energy depositional
regime likely due to proceeding marine transgression and energetic coastal
dynamics. Bauch et al. (2001a)
have reported a shift from sandy silt to clayey silt around 7400 calyrs BP
from a sediment core collected in the eastern Laptev Sea. They attribute
this change to the end of the sea level rise and the establishment of more
stable conditions. The GC58 sediment core has a hiatus at that time period,
but it has a similar clayey silt composition at the top part of the core
(∼1700 calyrs BP until today). This may indicate comparably
similar stable conditions in the ESS in the last 1700 calyrs BP.
The total organic carbon (TOC) concentrations in GC58 vary from 0.5 to
1.1 % (Table S1 in the Supplement) with the highest TOC content in the surface
sediments. These data agree with average TOC contents reported for the ESS
(Semiletov
et al., 2005; Stein and Macdonald, 2004; Vetrov and Romankevich, 2004; Vonk
et al., 2012). The OC fluxes for GC58 calculated with the 14C age–model
(covering ∼9500 calyrs BP) range between 1.2 and 10.9 gm-2yr-1 (Fig. 3a). The OC fluxes for MUC58 calculated with the
210Pb chronology (covering the most recent ∼100 years) are
similar and vary from 0.4 to 6.1 gm-2yr-1 (Table S2 in the Supplement).
The OC fluxes show an increasing trend from the bottom of the core
toward the top in both cores. A similar trend has been reported by
Bröder et al. (2016a) from the
ESS using two 210Pb-dated sediment cores. For GC58, the high OC flux at
the very top of the core is likely related to the merging of the two dating
systems (14C and 210Pb), which causes a higher sediment
accumulation rate at the top of the core and thus higher fluxes.
Organic matter composition of the sediment core GC58. The x axis
has breaks due to gaps in the sediment chronology. (a) Organic carbon fluxes
(gm-2yr-1) were high at the bottom of the core. The high fluxes
at the top of the core are likely related to the merging of two dating
systems (210Pb and 14C; see Sect. 3.2). The sea level rose rapidly
in the early Holocene
(Lambeck et al.,
2014). (b) Both lignin and cutin fluxes (mgm-2yr-1) decrease
toward the core top. High fluxes at the top of the core are influenced by
the OC fluxes and likely do not show an actual increase in the fluxes of
lignin and cutin (see Sect. 3.2). (c) Low molecular weight fatty acids
(LMW-FA) show an influence of marine organic matter at the top of the core.
(d) The δ13C (‰) values illustrate a
gradual shift from terrestrial-dominated to more marine-dominated input of
organic matter towards the core top. The Δ14C
(‰) values (corrected for the time between the
deposition and the measurement) show that the bulk organic carbon is older
at the bottom of the core than at the core top. The drop in the Δ14C values ∼1700 calyrs BP is likely an artefact
caused by the age model used to correct for the Δ14C values.
Lignin phenols and cutin acids are useful proxies for tracing carbon of
terrestrial origin because both compounds are solely biosynthesised in
terrestrial plants. Lignin is an essential component in the cell walls of
vascular plants (Higuchi, 1971), while cutin is a lipid
polyester, which forms a protective wax layer on the epidermal cells of leaves
and needles with other lipids (e.g.,
Kunst and Samuels, 2003). These compounds have been demonstrated to be useful
in studying terrestrial OC in the Arctic
(e.g.,
Amon et al., 2012; Bröder et al., 2016a; Goñi et al., 2013; Tesi et
al., 2014). Both lignin and cutin fluxes show a similar trend with the
highest fluxes at the bottom of the core (∼9500 calyrs BP)
indicating a high proportion of terrestrial organic matter (Fig. 3b). The
large variability in the fluxes between ∼9500 and
∼8200 calyrs BP compared to the latest ∼1700 calyrs BP suggests that the system was more dynamic at that time. The
rapid decrease in both lignin and cutin fluxes indicates a change from
terrestrially dominated to marine-dominated input at ∼8400 calyrs BP in this part of the ESS.
Bauch et al. (2001b) suggested a
similar regime shift from terrestrial to marine in the Laptev Sea between
∼8900 and ∼8400 calyrs BP based on the
occurrence of bivalves and benthic foraminiferal species. The same process
affecting OC fluxes is likely also causing higher lignin and cutin fluxes at
the top of GC58. The overall decrease in lignin and cutin fluxes as well as
concentrations (Table S3 in the Supplement) in time is likely due to increasing
hydrodynamic sorting and degradation during transport as transport times
from the coast became longer because of the marine transgression (Fig. 3a).
Bröder et al. (2016b) have
observed a similarly strong decrease in the amount of terrestrial organic
carbon depositions with increasing distance from the coast in the Laptev
Sea. A recent study by Tesi et al. (2016b) shows that the largest particles,
rich in lignin (i.e., plant debris), tend to be preferentially buried close
to the shore with the cross-shelf transport of lignin occurring
overwhelmingly bound to fine particles (with low settling velocities; i.e.,
of the total lignin deposited to the marine environment, only a fraction of
∼4–5 % travels across the shelf).
Other useful indicators of the marine input in organic matter are CuO-oxidation-derived low molecular weight fatty acids (LMW-FA). They are mainly
found in phytoplankton but also in other organisms, such as bacteria and
algae (Goñi and Hedges, 1995). C16FA:1
together with C14FA and C16FA serve as especially good proxies for marine OC as they are
highly abundant in marine sediments and very low in concentrations in ICD-PF
and topsoil-PF
(Goñi and Hedges,
1995; Tesi et al., 2014). The highest fluxes of LMW-FA are observed for the
very top of the core (Fig. 3c), indicating a larger proportion of marine OC.
The values decrease rapidly down-core as marine FA are readily degraded
(e.g.,
Bröder et al., 2016a; Canuel and Martens, 1996). This trend may also be
influenced by the change in input from terrestrial- to marine-dominated
sources.
The stable isotopic composition of bulk OC (δ13C) may be used
to distinguish between marine and terrestrial organic matter (Fry
and Sherr, 1984). The δ13C values for C3-photosynthesised
terrestrial carbon are between -23 and -30 ‰, whereas
marine carbon has a less depleted δ13C signature between -18 and 24 ‰ (e.g., Fry
and Sherr, 1984). However, these endmember values may differ depending on
the region, especially in the Arctic where open water and sea ice
phytoplankton exhibit different isotopic fingerprints
(Kohlbach et al., 2016). The
δ13C values for GC58 range from -23 to -25 ‰ (Fig. 3d) with the most depleted values (i.e., most
terrestrial) between ∼9500 and ∼ 8200 calyrs BP
and the least depleted values (i.e., most marine) from ∼1700 calyrs BP until the modern time.
Mueller-Lupp
et al. (2000, and references within) have argued that δ13C
values in sediments of the Arctic Ocean can have a terrestrial overprint in
δ13C composition caused by the rapid degradation of planktonic
organic matter; i.e., the amount of marine organic matter in the total organic
matter pool in the Arctic is relatively low. Yet, the gradual change in
δ13C indicates that the contribution of marine organic matter
is greater at the top of the core where the δ13C values are
less depleted.
It is notable that the values for all the different parameters shown in Fig. 3
on both sides of the age gap (between ∼8200 and
∼1700 calyrs BP) are nearly continuous in spite of the
∼6500 year hiatus (except for the bulk Δ14C OC
values). Either the values actually are similar on both sides of the hiatus
or, alternatively, this could be explained by bioturbation mixing the older
part of the core with the newer deposits, thus resulting in an apparent
continuity in property values across the hiatus. The Δ14C
values suggest that there was more 14C-depleted material deposited
at
∼1600 calyrs BP, causing a drop in the Δ14C values. Though more likely, as the Δ14C values are
dependent on time, any uncertainty in the age model would have an effect on
the Δ14C values.
Degradation status of terrestrial organic matter
Lignin phenols provide insight into the degradation status of the deposited
terrestrial organic matter. The acid to aldehyde,
syringic acid to syringaldehyde (Sd / Sl) and vanillic acid to vanillin
(Vd / Vl) ratios of lignin phenols have been used to study the degradation of lignin
(e.g.,
Opsahl and Benner, 1995; Hedges et al., 1988). However,
Goñi et al. (2000) and Tesi et al. (2014) have
argued that the acid to aldehyde ratios of lignin phenols might not serve as
a good degradation proxy for Arctic Ocean sediments as the material entering
the marine environment might have experienced degradation prior to entering
the marine system. This is supported by our data as both the Sd / Sl and Vd / Vl
ratios show great variability throughout the core (Fig. S3 in the Supplement).
The ratio of 3,5-dihydrobenzoic acid to vanillyl phenols (3,5-Bd / V) is
another proxy used to constrain the degradation status of terrestrial
organic matter in sediments
(e.g.,
Hedges et al., 1988; Tesi et al., 2014, 2016a). Specifically, this
proxy is used to distinguish diagenetically altered mineral soil OC from
relatively fresh vascular plant debris
(Farella et
al., 2001; Louchouarn et al., 1999; Prahl et al., 1994). The only source of
3,5-Bd in the marine environment is brown algae, which are not common in
the study area
(Goñi and Hedges,
1995). The low 3,5-Bd / V ratio at the bottom of the core
(∼9500–8200 calyrs BP) implies that the organic matter
that was deposited in that period was relatively undegraded (Fig. S3 in the Supplement). The extent of degradation gradually increases toward the top of
the core. However, hydrodynamic sorting may affect the degradation values as
the largest particles of fresh vascular plant debris are likely buried close
to the coast (Tesi et al.,
2016b). The input of organic matter was higher before ∼8200 calyrs BP, presumably due to coastal erosion caused by the marine
transgression. When sediments are quickly buried they can serve as a more
effective sink for terrestrial organic matter
(Hilton et al., 2015). As the material is
less degraded and the sedimentation rates are high in GC58 between
∼9500 and ∼8200 calyrs BP, the input of
organic matter was likely high causing it to be quickly buried. Similarly high
input of terrestrial material has been observed in the Laptev Sea
∼11000 calyrs BP
(Tesi et al., 2016a).
The location of the study site is currently ∼500 km offshore
so transport time and thereby the oxygen exposure time of the organic matter
in the benthic compartment are now longer than in the earlier phase of the
Holocene. The longer distance from the coast allows more time for organic
matter to degrade before burial
(Bröder et al., 2016b).
Hartnett et al. (1998) have also shown that the
burial efficiency of organic carbon decreases as a function of oxygen
exposure time. The same trend can be seen in the fraction of remaining lignin
(flignin/terrOC), i.e., the amount of lignin as a ratio of the observed
and expected (assuming conservative mixing, i.e., no degradation)
concentrations of lignin and terrestrial OC (terrOC; see “Supplementary
methods” in the Supplement for details). In GC58 the flignin/terrOC decreases down-core
likely as a result of the proceeding marine transgression (Fig. S4 in the Supplement). This trend suggests that with longer transport time the lignin
degradation is more extensive due to the protracted oxygen exposure time and
hydrodynamic sorting
(Keil
et al., 2004; Tesi et al., 2016a). We estimated this lateral transport time
to be ∼1.4 kyr longer at modern times than at the beginning
of the Holocene for the station GC58 (Fig. S5 in the Supplement). To model the
lateral transport times, we used flignin/terrOC with individual
degradation rates for terrOC and lignin (Bröder et al., 2015;
see “Supplementary
methods” in the Supplement).
The dual-carbon-isotope-based (δ13C, Δ14C)
source apportionment of organic carbon (OC) illustrates fractions (%;
mean ± SD) of old Pleistocene permafrost (ICD-PF) in brown, thaw of
active-layer permafrost (topsoil-PF) in green and primary production (marine
OC) in blue of the sediment core GC58. The ICD-PF is the dominant fraction
throughout the core.
Dual-isotope-based source apportionment of OC
The source apportionment results show that most of the organic matter
originates from coastal erosion since ICD-PF material is the largest
fraction (41–91 %) throughout the core (Fig. 4). Earlier studies
demonstrated that the decay of fresh marine organic matter is more rapid
compared to the degradation of terrestrial organic matter
(Karlsson
et al., 2011, 2015; Salvadó et al., 2016; Vonk et al., 2010). This may
lead to the selective preservation of terrestrial organic matter in the
sediments of the ESAS
(Karlsson
et al., 2011, 2015; Vonk et al., 2010). The contribution of topsoil-PF is
fairly low throughout the core (3–23 %). This may be due to the location
of GC58 between two major rivers (Kolyma and Indigirka), resulting in
relatively low amounts of fluvial inflow depositing topsoil permafrost.
To further interpret our results within a larger context of PF-C
destabilisation during post-glacial warming, we compared our results with
another transgressive deposit collected in the Laptev Sea (PC23, Fig. 1;
Tesi et al., 2016a). For the Laptev Sea
(PC23), there was a predominant influence of watershed-sourced material via
river discharge during the onset of the Holocene followed by similar
contributions of marine OC and ICD-PF fractions (both sources varying
between 31 and 56 %) from ∼8300 calyrs BP to present.
For the ESS (GC58), the contribution of ICD-PF is more prominent for the
same time period, indicating a higher significance of coastal erosion for
the ESS compared to the Laptev Sea (Fig. 5), especially when compared to the
early Holocene signature. Topsoil-PF fractions in PC23 are slightly higher
(8–25 %) than in GC58 (3–23 %) from ∼8300 calyrs BP
to current day. The difference is likely caused by a strong influence of the
Lena River at the sampling location of PC23 and less fluvial inflow to GC58
due to its location farther away from the mouths of the Lena, Kolyma and
Indigirka rivers.
When the shoreline was farther seaward during the early Holocene, the
location of the core PC23 from the Laptev Sea experienced a large influence
by material derived from the Lena River (80–90 %;
Tesi et al., 2016a). This material
was supplied to the Laptev Sea in response to the deglaciation and
associated active-layer deepening in the watershed
(Tesi et al., 2016a). Although the
record of GC58 does not go back in time to the glacial–interglacial
transition at the very onset of the Holocene, we suggest that coastal
erosion was likely an important process affecting the permafrost carbon
supply and deposition at that time. This seems possible, especially
when considering the location of the core GC58 between the rivers and as
has been observed in modern day shallower sediments in the ESS
(Bröder
et al., 2016a; Vonk et al., 2012).
Dual-carbon isotope (δ13C, Δ14C)
composition of the sediment cores GC58 (this study) and PC23
(Tesi et al., 2016a). Topsoil-PF
refers to organic matter from the active layer of permafrost, ICD-PF to
relict Pleistocene Ice Complex deposit permafrost (yedoma) and marine OC to
organic matter from primary production. The endmember values for different
sources are taken from a
dataset compiled by Vonk et al. (2012) and a study by Smith
et al. (2002). The green arrow points to the direction from the bottom to
the top of the core (GC58).
Sources of terrestrial organic matter
The lignin fingerprint of organic matter sources in GC58 is consistent with
the dual-carbon isotope modelling. Here we focus on the ratios of cinnamyl to vanillyl
phenols and syringyl to vanillyl phenols (C / V and S / V, respectively).
The C / V ratio can be used to differentiate between woody (i.e., shrubs and
trees) and non-woody (i.e., leaves, needles, grasses) plant tissues as
origins
of terrestrial OC since cinnamyl phenols are produced only in non-woody
vascular plant tissues (Hedges et
al., 1988). Moreover, the S / V ratio differentiates between gymnosperms
(conifers) and angiosperms (flowering plants) as syringyl phenols are
produced solely in angiosperms
(Hedges et al., 1988). Thereby
higher S / V ratios mean more of a contribution from angiosperm plants.
Lignin composition of the sediment core GC58 (blue circles). The
ratio between cinnamyl and vanillyl phenols (C / V) is used as a proxy to
distinguish between soft and woody plant tissues. The ratio of syringyl to
vanillyl phenols (S / V) indicates the difference between gymnosperm and
angiosperm plants. The boxes indicate the typical values for S / V and C / V ratios
characterising different plant material (ranges from
Goñi and Montgomery, 2000). Measured S / V and C / V ratios for Ice Complex
deposit permafrost (ICD-PF) are shown with green triangles
(Winterfeld et al., 2015) and with a red square (±standard deviation; Tesi et al.,
2014). Measured S / V and C / V ratios for topsoil-PF (Lena River POC) are
illustrated with orange diamonds (Winterfeld et al., 2015).
Also shown is the lignin composition of the sediment core PC23 (black diamonds)
from the Laptev Sea (study by Tesi
et al., 2016a).
The S / V and C / V ratios in GC58 show that the terrestrial material
transported to the ESS originates mainly from soft tissue material (i.e.,
grasses and leaves) both from angiosperm and gymnosperm plants (Fig. 6). The
lignin fingerprint of old Pleistocene material (ICD-PF) is characterised by
high ratios of both C / V and S / V i.e., a high abundance of soft plant tissues
from tundra steppe vegetation (e.g., grass-like material;
Tesi et al., 2014;
Winterfeld et al., 2015). Observations from the Laptev Sea (sediment core
PC23, Fig. 1) reveal a much stronger influence from woody material
indicating a watershed source, likely from the Lena River, rather than from
coastal erosion (Fig. 6). It should be noted that the lignin phenols are
susceptible to degradation. Cinnamyl phenols in particular are known to
degrade fairly fast, which may lower the C / V ratios
(Opsahl and Benner, 1995). However, even
considering degradation effects, the relatively high C / V and S / V values that
characterise GC58, indicate grass-type material typical of tundra and steppe
biomes and ICD-PF deposits (Tesi
et al., 2014; Winterfeld et al., 2015).
Conclusions
This down-core study provides new insights into terrestrial carbon dynamics
in the ESS from the early Holocene warming period until the present. Our
results suggest a high input of terrestrial organic carbon to the ESS during
the last glacial–interglacial period caused by permafrost destabilisation.
This material was mainly characterised as relict Pleistocene permafrost
released via coastal erosion as a result of the sea level ingression.
The flux rates of both lignin and cutin compounds show a declining trend in
the early Holocene, suggesting a change from mainly terrestrial- to marine-dominated input. The same change can be seen in the stable carbon isotope
(δ13C) data, which imply a regime shift from terrestrial- to
more marine-dominated sediment input at ∼8400 calyrs BP.
The source apportionment data highlight the importance of coastal erosion as
a terrestrial carbon source to the ESS during the Holocene time periods of
∼9500–9300, ∼8500–8200 and ∼1700 calyrs BP to the modern day. This is
supported by the lignin composition, which suggests a deposition of
tundra and steppe vegetation (i.e., grasses) grown during the Pleistocene. Both
the biomarker and grain size data imply that the conditions have been more
stable in the ESS in the past ∼1700 calyrs BP compared to
the early Holocene.
The comparison of the source apportionment results (δ13C,
Δ14C) and the lignin fingerprint (C / V and S / V ratios) for the
sediment cores GC58 and PC23 shows a difference in the carbon sources
between the ESS and the adjacent Laptev Sea. The relict Pleistocene
permafrost, mostly originating from coastal erosion, may be more dominant in
the ESS than in the Laptev Sea. Data for the sediment core PC23 show that
the Laptev Sea instead had a relatively high input of terrestrial carbon
from the watershed, which is likely due to the influence of the Lena River.
The accelerating coastal erosion rates along the Siberian coast and
amplified warming in the Arctic predicted by many climate models are likely
to cause permafrost destabilisation and remobilisation of terrestrial carbon
to the marine environment, as observed at the beginning of the Holocene. To
better understand the consequences of the permafrost thawing processes, the
extent of degradation of terrestrial carbon in the marine environment should
be better constrained. Also, to improve the understanding of the processes
in the ESS and in the whole Arctic region, more historical down-core studies
are needed.