Permafrost degradation in the catchment of major Siberian rivers, combined
with higher precipitation in a warming climate, could increase the flux of
terrestrially derived dissolved organic matter (tDOM) into the Arctic Ocean
(AO). Each year, ∼ 7.9 Tg of dissolved organic carbon (DOC) is
discharged into the AO via the three largest rivers that flow into the
Laptev Sea (LS) and East Siberian Sea (ESS). A significant proportion of this
tDOM-rich river water undergoes at least one freeze–melt cycle in the
land-fast ice that forms along the coast of the Laptev and East Siberian seas
in winter. To better understand how growth and melting of land-fast ice
affect dissolved organic matter (DOM) dynamics in the LS and ESS, we
determined DOC concentrations and the optical properties of coloured
dissolved organic matter (CDOM) in sea ice, river water and seawater. The
data set, covering different seasons over a 9-year period (2010–2019), was
complemented by oceanographic measurements (T, S) and determination of the
oxygen isotope composition of the seawater.
Although removal of tDOM cannot be ruled out, our study suggests that
conservative mixing of high-tDOM river water and sea-ice meltwater with
low-tDOM seawater is the major factor controlling the surface distribution
of tDOM in the LS and ESS. A case study based on data from winter 2012 and
spring 2014 reveals that the mixing of about 273 km3 of
low-tDOM land-fast-ice meltwater (containing ∼ 0.3 Tg DOC)
with more than 200 km3 of high-tDOM Lena River water
discharged during the spring freshet (∼ 2.8 Tg DOC yr-1)
plays a dominant role in this respect. The mixing of the two low-salinity
surface water masses is possible because the meltwater and the river water
of the spring freshet flow into the southeastern LS at the same time every
year (May–July). In addition, budget calculations indicate that in the
course of the growth of land-fast ice in the southeastern LS,
∼ 1.2 Tg DOC yr-1 (± 0.54 Tg) can be expelled from
the growing ice in winter, together with brines. These DOC-rich brines can
then be transported across the shelves into the Arctic halocline and the
Transpolar Drift Current flowing from the Siberian Shelf towards Greenland.
The study of dissolved organic matter dynamics in the AO is important not
only to decipher the Arctic carbon cycle but also because CDOM regulates
physical processes such as radiative forcing in the upper ocean, which has
important effects on sea surface temperature, water column stratification,
biological productivity and UV penetration.
Introduction
The mean annual air temperature in the Arctic continues to rise (Overland et
al., 2019), resulting in a rapid decrease in summer sea-ice extent and
volume (Stroeve and Notz, 2018) and, consequently, a longer ice-free season.
Furthermore, the terrestrial permafrost temperature is increasing (Biskaborn et
al., 2019) with major implications for the Arctic carbon cycle. Due to the
accelerated degradation of terrestrial permafrost, an estimated 1035 ± 150 Pg of organic carbon stored in the upper 3 m of circumpolar
permafrost soils (Hugelius et al., 2014) can be either mineralised and
mobilised as terrestrial dissolved organic matter (tDOM) into the
hydrosphere or released as gaseous emissions into the atmosphere (Plaza et
al., 2019). The release of soil carbon into the hydrosphere in combination
with an increasing freshwater discharge from Arctic rivers (McClelland et
al., 2004; Rawlins et al. 2010; Haine et al., 2015) might thus increase the
flux of tDOM into the ocean (Frey and Smith, 2005; Guo et al., 2007;
Prokushkin et al., 2011; Tank et al., 2016). Changes in the land–ocean
fluxes of tDOM in the Arctic are of particular importance for the global
carbon cycle since the rivers at high northern latitudes export significant
quantities of tDOM to the Arctic Ocean (AO). Dissolved organic matter (DOM) is typically quantified via
its carbon content (dissolved organic carbon; DOC), which contributes
roughly half to the total mass of DOM. Currently, the annual riverine input
of DOC into the AO is about 25–36 Tg C yr-1 (Raymond et al., 2007;
Anderson and Amon, 2015), with the six largest Arctic rivers discharging
about 18–20 Tg C yr-1 (Stedmon et al., 2011; Amon et al., 2012). The
three major Siberian river systems (Ob, Yenisey and Lena) account for about
14 Tg C yr-1 (Holmes et al., 2012), with the Lena River alone
discharging 6.8 Tg C yr-1 DOC into the Siberian Laptev Sea (LS) (Juhls
et al., 2020). The total riverine DOC flux to the LS is about 8.3 Tg C yr-1 (Manizza et al., 2009), with the Lena drainage basin (2.61×106 km2) accounting for about 70 % of the total area
of the LS watershed. The LS additionally receives freshwater from the
outflow of the Kara Sea (KS), which transports river water from Ob and
Yenisey through the Vilkitsky Strait into the northwestern LS (Janout et
al., 2015). The combined DOC flux of the two major rivers discharging to the
East Siberian Sea (ESS; Kolyma and Indigirka) is ∼ 1.1 Tg DOC yr-1
(Opsahl et al., 1999), which is ∼ 50 % of the total annual
riverine DOC flux to the ESS (Manizza et al., 2009). The combined area of the
drainage basin of the rivers Kolyma and Indigirka (1.01×106 km2) accounts for ∼ 75 % of the total ESS
watershed.
The terrigenous input (particulate and dissolved) of carbon by coastal
erosion to the entire AO is estimated to be 15.4 Tg yr-1 (9.2–24.2 Tg yr-1 with 1 standard deviation; Terhaar et al., 2021). However, DOC
concentrations measured in Guba Buor-Khaya bay east of the Lena Delta, a region
known for the most rapid and extensive coastal erosion of the Siberian
Arctic coastline, suggest that direct DOC input from coastal erosion must be
small compared to the riverine DOC flux (Alling et al., 2010).
If the input of tDOM into the AO increases as a consequence of climate
change, it is crucial to achieve a better understanding of tDOM transport
dynamics and biogeochemical cycles. Previous studies in the Arctic marginal
seas have reported a strong negative linear relationship between salinity
and DOC, which implies that tDOM-rich river water mixes with DOM-poor marine
waters from the AO without significant losses and gains along the salinity
gradient (i.e. conservative mixing) (Kattner et al., 1999; Köhler et al.,
2003; Amon, 2004; Amon and Meon, 2004; Shin and Tanaka, 2004; Gueguen et al.,
2005; Matsuoka et al., 2012; Pugach and Pipko, 2013; Pavlov et al., 2016;
Tanaka et al., 2016; Pugach et al., 2018). In contrast, a number of studies
indicate significant degradation of tDOM in the Arctic Ocean (Belanger
et al., 2006; Alling et al., 2010; Stedmon et al., 2011; Letscher et al.,
2011; Kaiser et al., 2017a), the Fram Strait (Granskog et al., 2012) and
Hudson Bay (Granskog et al., 2009; Granskog, 2012). One reason for the
partly contradictory observations could be the extreme seasonality in the
discharge of Arctic rivers and the associated high variability in the tDOM
concentrations and composition. In the course of the peak spring discharge
(spring freshet), which has the highest annual tDOM concentrations, about
half of the annual tDOM is exported to the AO (Cauwet and Sidorov, 1996;
Stedmon et al., 2011; Holmes et al., 2012). In summer, autumn and winter,
tDOM concentrations in Arctic rivers are significantly lower (Stedmon et
al., 2011; Juhls et al., 2020). Several studies have shown that tDOM
discharged during the spring freshet displays a different chemical
composition and a higher biological and photochemical lability compared to
the summer discharge (Osburn et al., 2009; Amon et al., 2012; Mann et al.,
2012; Kaiser et al., 2017b). This could lead to tDOM degradation of 20 %–40 % within less than 1 month during the spring freshet (Holmes et al.,
2008). Therefore, seasonal changes in tDOM composition and lability are
another possible explanation for the observed variability in degradation
rates.
Considering the high seasonal variability in input and chemical composition
of tDOM and a predicted strong degradation of tDOM that is controlled by
multiple processes, it appears puzzling that most studies in the LS have
observed a conservative mixing of tDOM. This is especially true since the
sea-ice growth–melt cycle in the Siberian shelf seas, which has a
significant influence on the salinity–tDOM relationship (Anderson and Amon,
2015), should lead to much greater variability in tDOM concentrations along
the salinity gradient. During sea-ice growth, DOM is expelled from ice
crystals and concentrates in brines, which can then drain into the
underlying seawater (Müller et al., 2013). The DOM remaining in the ice
undergoes significant fractionation. The humic-like DOM fraction is most
prone to rejection from the sea ice along with the brine during freezing,
while the protein-like fraction is least prone to this process, which shifts
the molecular composition of DOM in the ice towards a higher proportion of
low molecular weight compounds (Müller et al., 2013; Granskog et al.,
2015a; Retelletti-Brogi et al., 2018, Zabłocka et al., 2020). On the other
hand, melting of CDOM-poor drifting pack ice (Kowalczuk et al., 2017) or
immobile land ice (Wegner et al., 2017) is also important because it
provides a low-salinity source that can dilute higher DOM concentrations in
the surrounding seawater (Amon, 2004; Mathis et al., 2005; Granskog et al.,
2015b; Logvinova et al., 2016; Tanaka et al., 2016).
At the end of winter, the LS shelf north of the land-fast ice (Fig. 1) is
covered by large areas of open water (polynyas) and thin (< 50 cm)
pack ice, which has almost completely retreated by the end of June (Itkin
and Krumpen, 2017). The retreat of the pack ice is predominantly controlled
by the prevailing atmospheric conditions in April and May (Janout et al.,
2016a), whereas land-fast ice that forms in winter in the LS and ESS between
115 and 170∘ E retreats only during the month of July (Selyushenok et al., 2015). Bauch et al. (2013) and Janout et al. (2016b) showed that the decline in fast ice is also influenced by the input
of warm river water during spring freshet. While the maximum extent of fast
ice in the LS (∼130×103 km2) shows only small annual variations (observation period 1999–2013; Selyushenok et al., 2015), the maximum area of fast ice in the East Siberian
Sea (140 to 170∘ E) varies annually from 130 to 200×103 km2 (2012–2021, data from
https://www.meereisportal.de, last access: 15 May 2021; Spreen et al., 2008). At the end of winter, the fast
ice also reaches its maximum thickness of ∼ 2 m (Kotchetov et
al., 1994; Li et al., 2020). Thus, the maximum total volume of fast ice in
the Laptev and East Siberian Sea is approximately between 520
and 660 km3 . Assuming an average
salinity of the fast ice of 4, the land-fast ice meltwater corresponds to a
maximum freshwater equivalent of up to ∼ 580 km3 (reference salinity 34.8).
The presence of tDOM-poor meltwater from the land-fast ice that contains a
high proportion of river water and tDOM-rich river water in the Arctic shelf
seas, together with the extreme seasonality of fluvial tDOM input, make it
difficult to study tDOM removal processes based on the interpretation of the
relationship between tDOM concentration and salinity (Granskog, 2012;
Anderson and Amon, 2015). Although degradation processes are obviously
important for DOM dynamics in the Arctic Ocean, our study focuses mainly on
the physical drivers controlling the distribution of tDOM in the LS and ESS.
In particular, we aim to understand whether sea-ice growth and melt have a
discernible influence on the distribution of tDOM in the Siberian shelf
seas. We attempt to show that degradation of tDOM is probably not the
determining factor for the observed tDOM concentration distributions in the
LS and ESS but that instead physical mixing of meltwater from land-fast ice
with river and seawater plays a key role.
This research is based on one of the most comprehensive bio-optical, DOC
and hydrographic data sets from the Siberian Arctic. We report on the first
multi-year study of the coloured fraction of dissolved organic matter
(CDOM), which includes observations in river runoff during spring freshet (2014), marine waters from the coastal area to the shelf edge of the LS
(2010–2019) and ESS (2019), and ice and water samples from the land-fast ice
of the southeastern LS (2012). Sampling was carried out in different seasons
of the year over a period of 9 years (2010–2019). Although it was not
sampled on each expedition (Table 1), we used the stable oxygen isotope data
of the water samples to calculate the river water and sea-ice meltwater (sim)
fractions.
Materials and methodsField sampling
Seven summer ship expeditions between 2010 and 2019 and one sea-ice-based
winter expedition in 2012 were carried out to perform oceanographic and
biogeochemical measurements in the LS (Table 1 and Fig. 1). During the
ship-based expeditions, we operated a CTD profiler (conductivity–temperature–depth; Seabird 19+) to sample the basic hydrographic
parameters. The CTD is operated on a carousel water sampler (SBE 32SC)
allowing for automated water sampling at pre-selected depths. During the winter
expedition (Ti12) a stand-alone CTD (Seabird 19+) was lowered through an
ice hole along with Nansen water samplers. In this study we used the
practical salinity scale to describe salinity (PSS-78). Ice cores (9 cm
diameter) were collected in 2012 at stations on the land-fast ice between
19 March and 24 April. Six ice cores taken at three different locations
(Fig. 1) in March/April 2012 were analysed (Ti12_ice, Table 1). Sub-samples were taken at 10 cm intervals to represent the different
periods of ice formation, with the youngest ice in the lower section and the
oldest ice from the first stages of ice formation in early winter in the
uppermost sections of the ice core.
Map of the Laptev (LS) and East Siberian (ESS) seas and station
distribution during the ship-based expeditions in summer (red dots). The
border between the LS and the ESS runs along 140∘ E. The LS was
sampled in 2010, 2011, 2013, 2014, 2018 and 2019. Only a few stations could
be repeatedly sampled in multiple years. The ESS was sampled in 2019. Maps
of the stations of each expedition with a detailed description of the
location and date are available under the respective data sets (Table 1;
Hölemann et al., 2020a–g; Kattner et al., 2010). In addition, samples
of land-fast ice (fast ice) and the water column were taken from March to
April 2012 (yellow stars). The extent of fast ice describes the situation on
1 May 2011. The grey line between A and B represents the oceanographic
transect shown in Fig. 6. Bathymetric data were provide by IBCAO V3
(Jakobsson et al., 2012).
Lena River data presented here were collected as part of the Lena 2014
expedition. Samples were taken close to the Samoylov Station (72∘22′ N, 126∘29′ E; green pentagon in Fig. 1) in the
central part of the Lena Delta from 21 May to 16 June 2014. The sample treatment and methods used are the same as for sea water and ice samples (see below).
Expedition data and the number of samples (#) used in this
study. The data are accessible on PANGAEA.
CampaignRegionYearMonth#CDOM#DOCReferenceδ18OReferenceLD10LS2010Aug98Kattner at al. (2010)NE10LS2010Sept2340Hölemann et al. (2020a)XBauch et al. (2018)YS11LS2011Aug/Sept8179Hölemann et al. (2020b)XBauch et al. (2018)Ti12LS2012Mar/Apr2825Hölemann et al. (2020c)XBauch and Thibodeau (2020)Ti12_iceLS2012Mar/Apr5247Hölemann et al. (2020c)XBauch and Thibodeau (2020)VB13LS2013Sept360Hölemann et al. (2020d)XBauch et al. (2018)Lena 2014Lena Delta2014May/June4444Eulenburg et al. (2019)VB14LS2014Sept112108Hölemann et al. (2020e)XBauch et al. (2018)AT18LS2018Aug/Sept102102Hölemann et al. (2020f)TA19_4LS & ESS2019Sept/Oct179162Hölemann et al. (2020g)Sample preparation and analytical methods
Water samples were immediately filtered through pre-combusted Whatman GF/F
glass microfiber filters (4.7 cm diameter) with a nominal pore size of
approx. 0.7 µm. The filtrate was filled into acid cleaned and
pre-rinsed high-density polyethylene bottles, stored dark and cold for
CDOM analysis, and frozen at -20 ∘C for DOC analysis. Filtered DOC
samples from the LENA 2014 expedition were filled into a pre-combusted 20 mL
glass vial, acidified with 25µL HCl (Merck, Suprapur grade, 10 M)
and stored in the dark at 4 ∘C. The ice cores were drilled using
an electromechanical ice corer (Kovacs Enterprise, USA) and subsequently
placed in polyethylene bags and transported frozen to the land-based
laboratory. Within a few hours after coring, the core was placed in
polyethylene containers in the dark at room temperature until they were
completely melted. Immediately after melting, the salinity of the meltwater
was determined with a conductivity metre (WTW 197i) with a standard
conductivity cell (WTW TetraCon 325), and samples were subjected to vacuum
filtration. Also samples for the determination of dissolved inorganic
nutrient concentrations, chlorophyll a, particulate matter and stable oxygen
isotopes were taken from the ice cores. Only Salinity, DOC and CDOM were
used in this study. In addition, under-ice water samples were collected.
Sample storage and analytical procedures were identical to those described
for water samples.
DOC was determined by high-temperature catalytic oxidation (HTCO) with a
Shimadzu TOC-VCPN analyser. In the autosampler, 6 mL of sample volume
(pre-combusted or thoroughly rinsed vials) were acidified with 0.12 mL HCl
(2 M) and sparged with oxygen (100 mL min-1) for 5 min to remove
inorganic carbon. A 50 µL sample volume was injected directly on the
catalyst (heated to 680 ∘C). Detection of the generated CO2
was performed with an infrared detector. Final DOC concentrations were
average values of triplicate measurements. If the standard variation or the
coefficient of variation exceeded 0.1 µmol L-1 C or 1 %,
respectively, up to two additional analyses were performed and outliers were
eliminated. Detection limit was 7 µmol L-1 C with an accuracy of
± 2 µmol L-1 C determined with low carbon water and
seawater reference material (DOC-CRM, Hansell Research Lab, University of
Miami, US). Quality control was assured by measuring one Milli-Q blank and
two standards after every six samples.
The optical properties of CDOM provide information on both the amount of
DOM present and its chemical properties (Coble, 2007). In order to retrieve
the absorption by CDOM, optical density (OD) spectra were analysed on a dual
beam spectrophotometer (Specord200, Jena Analytik) within 2 months after the
expedition. Spectra were measured from 200 to 750 nm using quartz cuvettes
with a path length of 5 cm or 10 cm, according to the expected absorption
intensity of the sample. OD of each sample was measured three times against
ultra-pure water. Napierian absorption per metre was calculated based on the
averaged OD value using 2.303 × OD /L, where L is the lengths of the cuvette.
In this study, we present the absorption coefficient (m-1) at a
wavelength of 350 nm: aCDOM(350). The wavelength was selected to make
the results comparable with previous studies in Arctic waters (Granskog et
al., 2012; Gonçalves-Araújo et al., 2015; Pavlov et al., 2016). The
spectral slope of the absorption spectra in the wavelength range between
275 and 295 nm (S275-295) and between 300 and 600 nm (S300-600) was
calculated by fitting with an exponential function aCDOM (λ) =aCDOM (λ0) ×e-S(λ-λ0).
The usefulness of S275-295 and other spectral slopes at shorter
wavelengths for determining the origin and diagenesis of CDOM was
demonstrated in the studies by Granskog (2012) and Fichot and Benner (2012).
Based on the relationship between S275-295 and DOC-normalised lignin
yield, Fichot and Benner (2012) showed that S275-295 is lower than 20 µm-1 when the proportion of terrestrial DOC is higher than
50 %. Steeper spectral slopes (i.e. higher S275-295) are typically
associated with a higher proportion of marine CDOM. Steeper S275-295
slopes have also been associated with photochemically induced shifts in
molecular weights (Helms et al., 2008). The spectral region between 275
and 295 nm lies at the short-wavelength edge of the natural solar spectrum.
Solar ultraviolet radiation shows significant degradation for CDOM in
natural aquatic ecosystems. In contrast to 295 nm, almost no photons are
present at 275 nm in the lower atmosphere. It is therefore assumed that
solar radiation absorbed by CDOM would always lead to a greater change in
aCDOM(295) than in aCDOM(275) and consequently to an increase in
S275-295 so that increased slope values are a good indicator of
photodegradation, while microbial degradation should have the opposite
effect (Helms et al., 2008). However, in marine systems such as the LS and
ESS, which are characterised by a high riverine input of terrestrial CDOM,
non-terrestrial sources and non-photochemical processes only become visible
once the absorption of terrestrial CDOM has been removed to a high degree
(Granskog, 2012).
Stable oxygen isotopes were analysed at the Stable Isotope Laboratory of
COAS at Oregon State University (Corvallis, USA), applying the CO2–water isotope equilibration technique, and were analysed by dual inlet mass spectrometry
(Thermo, DeltaPlus XL). The overall measurement precision for all δ18O analyses was ± 0.04 ‰. The
18O /16O ratios were calibrated with Vienna Standard Mean Ocean
Water (VSMOW) and reported in the usual δ notation (Craig, 1961).
For a quantitative interpretation of the oxygen isotope data, an exact match
of salinity and δ18O values is essential. Therefore, in
addition to CTD measurements, bottle salinity was determined directly within
the water samples taken for δ18O analysis using an AutoSal
8400A salinometer (Guildline Instruments) with a precision of ± 0.003 and an
accuracy greater than ± 0.005. The river water and sea-ice meltwater
(sim) contributions can be quantified with a mass-balance calculation, which
was previously applied in the Arctic Ocean basins (Bauch et al., 2011) and
shelf regions (Bauch et al., 2005). The basis for the mass-balance
calculation is the assumption that each sample is a mixture of marine water
(fmar), river-runoff (frw) and sea-ice meltwater (fsim).
The following equations determine the balance:
1fmar+frw+fsim=1,2fmar⋅Smar+frw⋅Srw+fsim⋅Ssim=Smeas,3fmar⋅Omar+frw⋅Orw+fsim⋅Osim=Omeas,
where fmar, frw and fSIM are the fractions of marine water,
river-runoff and sea-ice meltwater in a water parcel, and Smar,
Srw, Ssim, Omar, Orw and Osim are the corresponding
salinities and δ18O values. Smeas and Omeas are the
measured salinity and δ18O of the water samples. The analytical
errors from δ18O and salinity measurements add up to
approximately ± 0.3 % for each of the fractions. The additional
systematic error depends on the exact choice of endmember values. When
endmember values are varied within the estimated uncertainties (Bauch et
al., 2013), both fractions are shifted by up to ∼ 1 % in
absolute values, but relative results are qualitatively always conserved
even when extreme variations in endmember values are tested (Bauch et al.,
2011).
The endmembers are well known for the study area (Bauch et al., 2010, 2013).
The only exception is the sea-ice endmember that is based on an assumption
on the signature of the source water since sea ice and underlying water can
move independently from each other. Within the direct vicinity of the Lena
River, the summer surface layer is strongly influenced by summer discharge
of the Lena River, so the low δ18O summer surface signature in
this area is not a useful endmember for the sea ice formed during winter.
Therefore, the average surface value from the winter polynya region of -7 ‰ in δ18O is applied as source water for
sea-ice formation to all stations with a surface δ18O lower
than -7 ‰. Differences in calculated sea-ice
meltwater and river water fractions in the southern LS, when a constant
polynya value is used instead of each station's surface signature, are
generally small (for a detailed discussion see Bauch et al., 2010), and
calculated fractions remain stable relative to each other.
ResultsCDOM-DOC relationship in the Laptev Sea and East Siberian Sea
Samples collected in the Lena River, LS, ESS, and the adjacent Nansen and
Amundsen basins illustrate the range of variability in DOC concentrations
and aCDOM(350). The highest DOC concentrations (> 1200 µmol L-1) we observed in river water were measured during the
peak of the spring freshet of the Lena River between 29 May and 7 June 2014
(see Sect. 3.2). The lowest riverine DOC concentrations (384 µmol L-1) were recorded on 23 May 2014 before the onset of the spring
freshet. The coastal waters near the Lena Delta are characterised by DOC
concentrations of 185 to 853 µmol L-1 in
the surface mixed layer (0–10 m depth), with the highest values
(> 500 µmol L-1, salinity < 5) occurring near
the river mouth and in the mixed layer underneath the land-fast ice east of
the delta (300–520 µmol L-1, salinity 9–20). The marine water
masses of the outer shelf and in the basins with salinity above 20 show
lower DOC concentrations of 50–75 µmol L-1, with the lowest
value recorded in the Atlantic Intermediate Water on the continental slope
of the LS at a water depth of 300 m.
To characterise the relationship between the DOC concentration and the
optical property – aCDOM(350) – of DOM in water and ice samples, we
applied a non-linear regression analysis according to the method described in
Juhls et al. (2019) and Matsuoka et al. (2017). Statistical analyses of the
water samples from the LS and ESS show a strong relationship between
aCDOM(350) and the DOC concentration (r2= 0.99; RMSE = 0.8 m-1, n=527) (Fig. 2a). In contrast, ice samples from the
land-fast ice of the southeastern LS (Fig. 2a, blue diamonds) show a
significantly different relationship. The aCDOM(350) in the ice cores with
DOC concentrations > 300 µmol L-1 is lower compared
to water samples with comparable DOC concentrations. In particular, the
DOC-rich (> 400 µmol L-1) ice samples show
aCDOM(350) values that are up to 5 times lower than in seawater with
the same DOC concentration. All ice samples with DOC concentrations
> 200 µmol L-1 and aCDOM(350) < 2.5
were taken from the upper 70 cm of a 2 m long ice core from the coastal area
near one of the mouths of the Lena River. We also used the model developed
by Stedmon and Markager (2001) to differentiate between marine and
terrestrial organic matter (Fig. 2b). The model that suggests a dominance
of tDOM in the LS and ESS is based on the relationship between the spectral
slope coefficient (300–600 nm) and CDOM absorption at 375 nm. In the LS and
ESS most data points show high S300-600 values (> 17 µm-1) that stay relatively constant as aCDOM(375) decreases. The
results of these investigations are discussed in more detail in Sect. 4.1.
(a) The relationship between DOC and aCDOM(350) for Lena
River water, ice samples and seawater samples from the LS and ESS. The
regression and r2 was calculated only on the basis of river and seawater
samples. Panel (b) shows the model developed by Stedmon and Markager (2001) to
differentiate between marine and terrestrial organic matter. Data points
from the LS and ESS that lie within the model boundaries of the Stedmon and
Markager (2001) model (dotted red lines) are defined as marine CDOM, while
the rest are terrestrial in origin.
The tDOM characteristics in the Lena River during the spring freshet
The dominant feature of the hydrological cycle of the Lena is the spring
freshet in May and June. We sampled this event in the central part of the
Lena Delta from 21 May to 19 June 2014 (Lena 2014). The discharge data
measured ∼ 200 km upstream at Kyusyur were provided by
ArcticGRO (Shiklomanov et al., 2020). In their study based on ArcticGRO
discharge data, Juhls et al. (2020) give a more detailed description of the
annual variability in the tDOM input of the Lena River during the spring
freshet and throughout the year. We corrected the discharge data for the
distance between Kyusur and Samoylov Island as described in Juhls et al. (2020) assuming a mean flow propagation speed of 88 km d-1.
(a) Corrected daily Lena River discharge data, (b)aCDOM(350)
and S275-295, and (c) DOC concentration measured near Samoylov Island
in the Lena Delta during the spring freshet in May/June 2014.
The highest aCDOM(350) (Fig. 3b) and DOC concentration (Fig. 3c) were
measured during the peak river discharge on 31 May (Fig. 3a). During the
observation period 211 km3 of freshwater with a
flow-weighted average aCDOM(350) of 26.1 m-1 and a DOC load of
2.83 Tg entered the LS, which is one third of the annual DOC discharge of
the Lena.
The tDOM in the Laptev and East Siberian seas (ESS) and in
land-fast ice of the LS
Samples from the LS showed a significant negative linear relationship
between salinity and aCDOM(350) (r2= 0.90, n=659,
p< 0.01) (Fig. 4a) and between the fraction of river (meteoric)
water (frw) and aCDOM(350) (r2= 0.94, n=465,
p< 0.01) (Fig. 4b). The data points in the LS are well below the
theoretical mixing line resulting from a conservative mixing between the
Lena River water discharged after the spring freshet (post-freshet, “Summer”
in Fig. 4), which has an average aCDOM(350) of 14.4 m-1 (standard
deviation 3.4 m-1, n=26; Juhls at al., 2020), and the tDOM-poor seawater (salinity > 34) of the Nansen Basin. Looking at the
theoretical conservative mixing line of the flow-weighted annual river input
of CDOM (“Annual” in Fig. 4) that includes the spring freshet (April 2018 to April 2019, aCDOM(370) of 18.7 m-1, calculated from data published by Juhls
et al., 2020) and the seawater, the discrepancy becomes even more apparent.
The freshwater endmember calculated from the data in the LS on the basis of
salinity has an aCDOM(350) of 11.7 m-1 (interception of linear fit
in Fig. 4a). The aCDOM(350) of the river water endmember calculated
from frw (interception of linear fit in Fig. 4b) is ∼ 10 m-1, which is about 50 % lower than the annual flow-weighted average
aCDOM(350) of the river water. Whether this can be explained by the
dilution of low-tDOM meltwater from the land-fast ice of the southeastern LS
(Fig. 5b) that contains up to 90 % river water (Eicken et al., 2005) is
discussed in Sect. 4.2. The group of data points with frw between 40 % and 60 % that fall below the regression line are mostly from the
surface water of the southern LS and were sampled in September 2011 (Fig. 4,
YS11-data points outlined in red). The possible reason for the deviation in
this sample cluster is discussed in Sect. 4.3.
(a) Salinity, aCDOM(350) and the percentage of sea-ice
meltwater fraction (fsim %) measured in the LS in summer (NE10,
LD10, YS11, VB13, VB14, AT18 and TA19_4) and (b) percentage
of river water (frw %), aCDOM(350) and S275-295. The
solid black line represents the linear regression fit to the data in the LS.
The dashed black lines show the theoretical conservative mixing line between
Lena River water after spring freshet (Summer, aCDOM(350) of 14.4 m-1
calculated from Juhls et al., 2020) and the marine endmember (Nansen Basin, NB, seawater), as well as the theoretical conservative mixing line of
year-round flow-weighted average CDOM absorption of the Lena River (Annual,
aCDOM(350) of 18.7 m-1 based on year-round measurements in the Lena
Delta, April 2018–April 2019; Juhls et al., 2020) and the marine endmember.
Red outlined dots are water samples from the surface layer (< 20 m)
taken north and east of the Lena Delta in September 2011. Black dots in (a) represent samples on which stable oxygen isotope analysis was not performed
(LD10).
To describe the influence of the formation and melting of sea ice on the
distribution of tDOM in the LS, the stable oxygen isotope composition of the
water was studied. At salinities > 25, most water samples, which
are well below the regression line (Fig. 4a, linear fit) showed an increased
fraction of sea-ice meltwater (positive fsim). In contrast, many
samples above the regression line showed higher proportions of brine
influence (negative fsim). The addition of brines, which are expelled
from the sea ice during formation, was most apparent in a winter-water
sample from the northern edge of the land-fast ice taken at 10 m water depth
that showed a minimum fsim of -20 % at a salinity of 29.7 and a
comparatively high aCDOM(350) of ≈ 4 m-1. Shelf waters
during winter also showed higher DOC concentrations (mean 254 µmol L-1, salinity from 10 to 31) compared to summer shelf waters
of the same salinity (mean 199 µmol L-1).
Sampling in the ESS was carried out in August and September 2019. Across the
ESS, salinities were above 20 in 2019 (Fig. 5a). Even in coastal areas near
the Indigirka and Kolyma rivers surface-water salinities were above 20.
Moreover, aCDOM(350) in the ESS (Fig. 5a, coloured dots) was
significantly lower compared to LS water (Fig. 5a, black dots) at similar
salinities.
In the LS, S275-295 showed an inverse relationship with frw
(Fig. 4b), indicating on the one hand a fluvial source of CDOM and on the
other hand a more intense photobleaching of CDOM in the water masses on the
outer shelf and continental margin. When interpreting the slope data,
however, it must be kept in mind that in marine systems characterised by a
high input of terrestrial CDOM, non-terrestrial sources and
non-photochemical processes only become visible when the absorption of
terrestrial CDOM has been removed to a high degree (Granskog, 2012).
(a) Salinity and aCDOM(350) in the ESS (August and September 2019, coloured dots) and LS (black dots). The broken line shows the
theoretical conservative mixing line between the Lena River water that
enters the LS after the spring freshet (post-freshet) and the marine
endmember (Nansen Basin seawater, NB). The colour of the dots shows the
slope of the absorption spectra in the range between 275 and 295 nm
(S275-295). (b) Salinity, aCDOM(350) and S275-295 in
land-fast ice (squares) and sub-ice water samples (diamonds) in the LS
collected in (March/April 2012). The arrows indicate the direction of
theoretical conservative mixing lines between the land-fast-ice meltwater
and the Lena River water, as well as between the meltwater and the marine
endmember. The black dots represent the samples from the LS and ESS
(2010–2019).
The DOC concentration and optical properties of DOM in the land-fast ice of
the southeastern LS were determined in March and April 2012. The ice core
samples had a salinity of 3.6 and an aCDOM(350) of 0.99 m-1
(median of all cores and subsections; n=52), which was approximately 10 % of the aCDOM(350) observed in seawater with comparable salinity
(Fig. 5b). Unfortunately, δ18O measurements of LS land-fast ice
cannot be used directly as endmember for sea-ice meltwater as they represent
a mixture of seawater and river water contained within the ice. The median
of the DOC concentration in the ice was 96.2 µmol L-1. The highest
aCDOM(350) values of up to 4 m-1 were observed in the lowermost 20 cm of an ice core sampled near the northern border of the land-fast ice. The
high aCDOM(350) coincided with high chlorophyll a concentrations (Katya Abramova, personal communication, 2016). However, CDOM absorption spectra of algae-rich ice
samples exhibited no ultraviolet absorption shoulders that indicate
mycosporine-like amino acids from sympagic autotrophic organisms (Xie et
al., 2014; Granskog et al., 2015b). Similar to the marine waters with an
aCDOM(350) < 5 m-1 (Figs. 4a and 5b), the land-fast ice showed
high S275-295 values with a median slope of 20.7 µm-1. An
exception was the two chlorophyll-rich ice samples from the northern edge
of the land-fast ice with slopes below 18 µm-1, which
corresponds to the S275-295 slope of the tDOM in the winter runoff of
the Lena River but which could also be caused by a higher proportion of
marine DOM.
DiscussionDynamics of tDOM in the freshwater-influenced LS during spring and
summer
Our results show that aCDOM(350) correlates strongly with the DOC
concentration in riverine and marine LS and ESS waters (Fig. 2a). The high
correlation is clear evidence for the terrestrial origin of the DOM, which
is transported with the river water to the LS and ESS (Lara et al., 1998;
Dittmar and Kattner, 2003; Amon et al, 2012; Mann et al, 2016; Kaiser et al,
2017b; Juhls et al., 2019). The results of the data analysis using the model
developed by Stedmon and Markager (2001) also indicate a dominance of tDOM
in the LS and ESS. These results are consistent with those of
Gonçalves-Araujo et al. (2015), who investigated the composition and
concentration of CDOM in the southeastern LS using excitation emission
matrix spectroscopy and parallel factor analysis from a data set collected
in September 2013. They found that CDOM in the region of freshwater
influence (ROFI) of the LS was dominated by terrestrial humic-like material.
Kattner et al. (1999) were also able to show on the basis of chemical
analyses of DOM that the high input of tDOM not only shapes the ROFI of the
Lena but that 60 % of the DOC of the entire LS and the adjacent Eurasian
basin are also of terrestrial origin. One reason for the generally low DOM
contribution from marine algae, which would significantly alter the
dispersion in the CDOM vs. DOC dependence (Danhiez et al., 2017), could be
the timing of sampling, which was mainly conducted between late August and
mid-September when chlorophyll concentrations on the LS shelf are generally
below 1 mg m-3 (Janout et al., 2016a).
To describe the tDOM dynamics in the Siberian shelf seas we analysed the
optical properties of DOM along the salinity gradient in the LS, the ESS and
the adjacent continental slope. In water samples from the LS, we found a
statistically robust negative correlation between aCDOM(350) and
salinity (r2= 0.90, n=659, p< 0.01) (Fig. 4a) and between aCDOM(350) and the fraction of river (meteoric) water
(frw) and (r2= 0.94, n=465, p< 0.01)
(Fig. 4b). The strong linear relationship suggests that tDOM mixes
conservatively with seawater during the transit of river water across the
shelf. The calculated aCDOM(350) concentration of the river water
endmember of ≈ 10 m-1 (Fig. 5, linear fit of samples from the
LS) is approximately 30 % below the mean aCDOM(350) absorption
measured in the Lena River after the spring freshet (July–September, 14.4 m-1 with a standard deviation of 3.4 m-1; Juhls et al., 2020) and
nearly 50 % below the flow-weighted annual average aCDOM(350) of the
Lena River of 18.7 m-1 observed in 2018/2019 that also includes the
spring freshet (data from Juhls et al., 2020). Because aCDOM(350) in
river water may show intra-seasonal fluctuations between 10 and 22 m-1 in the post-freshet season (Juhls et al., 2020), the relatively
high aCDOM(350) of 13.6–16 m-1 at salinities < 6 observed
in the river plume near the Lena Delta in summer 2010 (Fig. 4a) were likely
caused by short-term fluctuations of aCDOM(350) in the river.
The dominance of riverine input of terrestrial DOM and the resulting close
relationship between salinity and aCDOM(350) is also evident from a
hydrographic transect across the LS shelf observed in September 2010 (Fig. 6; the area of the transect is indicated in Fig. 1), The oceanographic
transect runs from the mouth of the Lena River in the southeastern LS (A,
Fig. 6) to the Taymyr Peninsula in the northwestern LS (B, Fig. 6). While
the southeastern LS is dominated by the tDOM-rich river water plume of the
Lena River, the northwestern LS is mainly affected by the inflow of
tDOM-poor seawater from the Nansen Basin (Janout et al., 2013).
Salinity (white isoline) and aCDOM(350) (colour) on a
∼ 1000 km long oceanographic transect from the Lena Delta
(left) to the northwestern LS (September 2010; NE10). For the location of
the profile see Fig. 1. Small gray dots represent salinity measurements
(metre averaged), and black dots indicate water sampling for CDOM analysis. The
figure was prepared with Ocean Data View (ODV) using DIVA gridding
(Schlitzer, 2002). The black bars at the top of the panel indicate the
changes in the direction of the transect (see Fig. 1).
Assuming conservative mixing, the input of more than 200 km3 of tDOM-rich freshwater with a flow-weighted average
aCDOM(350) of 26.1 m-1 during the spring freshet in May and June
should result in shelf aCDOM(350) values that are above the theoretical
mixing line between the river water and the marine endmember (Fig. 4b).
However, the aCDOM(350) vs. frw distribution of all summer
(August/September) expeditions to the LS does not indicate the presence of
freshet-related river water with high aCDOM(350) values on the shelf
(Fig. 4b). Alling et al. (2010) suggested that the absence of tDOM-rich
waters from the freshet on the LS shelf in 2008 was caused by a rapid
wind-driven eastward transport into the adjacent ESS, but this is not
consistent with the observation that when winds are predominantly easterly,
as they were in the summer of 2008, the ROFI in the LS extends far north to
the outer shelf of the LS (Janout et al, 2020). The same atmospheric forcing
as in 2008 was also observed in 2011, when strong southeasterly summer
winds coincided with mixed layer salinities below 16 ∼ 300 km
north of the Lena Delta (Janout et al., 2020). The northern part of the ROFI
should therefore contain a significant amount of tDOM-rich freshwater from
the spring freshet, but this is not evident from the data (data from YS11,
Table 1).
Another explanation for the absence of freshet-related high aCDOM(350)
values within the ROFI could be the removal of tDOM, i.e. nonconservative
mixing. This would imply that ∼ 50 % of the tDOM discharged
into the southeastern LS in May and June had already been removed prior to
the August and September sampling period. In this context, photochemical
processes, flocculation of tDOM or rapid microbial degradation could play a
major role. A previous study in the Mackenzie River, Canada, showed that
photodegradation of tDOM was highest after spring freshet (Osburn et al.,
2009). On the other hand, Belanger et al. (2006) noted that under a closed
ice cover, photochemical reactions of tDOM carried by the Mackenzie River
into the Beaufort Sea do not play a significant role in the degradation of
tDOM. In the turbid nearshore waters of southeastern Hudson Bay, which have
a high proportion of river water (frw> 25 %),
self-shading could additionally slow photochemical reactions
(photobleaching) of tDOM (Granskog, 2012). Since during the Lena spring
freshet the turbid river plume mostly flows below the land-fast ice
(Hölemann et al., 2005) that covers the inner shelves from October to
mid-July (Barreis and Görgen, 2005; Selyuzhenok et al., 2015),
photodegradation of tDOM should not play an important role in the LS either.
Flocculation of DOM in the Lena River plume has been postulated as a
theoretically possible removal mechanism for DOM in summer
(Gonçalves-Araujo et al., 2015), which was, however, not directly
observed. An investigation of the under-ice river plume during the spring
freshet in 1996 (Hölemann et al., 2005) also gave no indication of
flocculation processes. Furthermore, Gonçalves-Araujo et al. (2015)
found that the degradation of tDOM by microbial processes plays only a minor
role in the Lena Delta region. Nevertheless, if we would assume that 50 %
of the tDOM were removed shortly after the spring freshet, the
aCDOM(350) values observed in the LS should be much lower because the
seawater is further diluted by CDOM-poor meltwater from the land-fast ice
(Fig. 5b) that melts in the southeastern LS in June/July. This process,
important to the LS and ESS, is discussed in more detail in the next two
chapters.
The importance of land-fast ice meltwater for the distribution of tDOM
in the LS
DOM is expelled from the ice matrix during sea-ice growth and is enriched in
the brine (Giannelli et al., 2001; Müller et al., 2013). Only about 10
to 40 % of the original DOM and other impurities remain in the sea ice
(Petrich and Eicken, 2010). Furthermore, Müller et al. (2013) were able
to show that a relatively higher proportion of DOM is incorporated into the
sea ice in relation to the dissolved inorganic substances. Most of the DOM
then flows off into the underlying water column along with the saline brine
(Amon, 2004; Anderson and Macdonald, 2015). This explains the observed low
aCDOM(350) of 0.99 m-1 (median) and DOC concentration of 96.2 µmol L-1 (median) in the LS land-fast ice that forms
from tDOM-rich surface waters sourced by the Lena River (Fig. 5b). Helms
et al. (2008) showed that as a result of this process, the molecular weights
of DOM in the ice and brine also decrease. Jørgensen et al. (2015)
further suggest that the transformations of DOM during sea-ice formation
also increase its bioavailability. We assume that the transformation
processes of the DOM during ice growth, which shifts the molecular
composition of DOM in the ice towards a higher proportion of low molecular
weight compounds (Müller et al., 2013; Granskog et al., 2015a;
Retelletti-Brogi et al., 2018; Zabłocka et al., 2020), are also responsible
for the high S275-290 values (> 19 µm-1) of CDOM in the ice (Fig. 5b).
The formation of the land-fast ice, which stores 30 %–50 % of the annual
freshwater discharge of the Lena River (Eicken et al., 2005), begins in
mid-November with phases of rapid development in January and February
(Selyuzhenok et al., 2015) during which the release of tDOM-rich brine is
highest. In the following summer, when the land-fast ice melts, the
tDOM-poor meltwater mixes with the ambient seawater and thus could dilute
the tDOM concentration of the surface mixed layer (Mathis et al., 2005;
Logvinova et al., 2016). North of the Lena Delta, sea-ice melt begins in
June (Janout et al., 2020) and progresses eastwards, while the seaward edge
of the land ice slowly moves southwards (Selyuzhenok et al., 2015). In
general, the ice in the southeastern LS retreats completely by the end of
July. The melting of sea-ice and land-fast ice in the southeastern LS thus
occurs at the same time and in the same region where the spring freshet
enters the LS, leading to a mixing of the two surface water masses (Bauch et
al, 2013).
To assess the influence of mixing on the tDOM concentration, we calculated
the first tDOM budget for the southeastern LS. During the course of the
freshet in 2014 the Lena discharged about 211 km3 of
freshwater with a flow-weighted average aCDOM(350) of 26.1 m-1 and
a DOC load of 2.83 Tg. The maximum land-fast ice area in the southeastern LS
is 134×103 km2 (Selyuzhenok et al., 2015)
with an average ice thickness of 2.04 m (Kotchetov et al., 1994; Barreis and
Görgen, 2005). With a salinity of 3.6 (median), the land-fast ice thus
holds a freshwater equivalent of 245 km3 (reference
salinity 34.8) with a median aCDOM(350) of 0.99 m-1 and a total
DOC load of 0.28 Tg. Our budget calculations reveal that the mixing product
(MIX) of land-fast-ice meltwater and the river water from the spring freshet
has an aCDOM(350) of 12.6 m-1 (Fig. 7). This is close to the
average aCDOM(350) of the post-freshet discharge of the Lena River of
14.4 m-1 (SD ± 3.4 m-1).
Salinity and aCDOM350 measured in land-fast ice from the LS
(squares), Lena River water during the spring freshet in May/June 2014
(stars), and marine waters from the LS (black dots) and ESS (red dots). The
dashed lines represent the theoretical conservative mixing lines between the
spring-freshet river water (flow-weighted CDOM absorption) and the Nansen
Basin (NB) seawater, as well as between the mixing product of the
spring-freshet river water with the land ice meltwater (MIX) and the seawater. The green squares indicate aCDOM350 endmember values of the
spring-freshet river water and land-fast ice meltwater (freshwater
equivalent), as well as the calculated endmember value for the mixing
product of the two sources.
Since aCDOM(350) is in the same range in MIX and in post-freshet river
water, the three main freshwater sources (spring-freshet river water,
post-freshet river water and meltwater from the land-fast ice) appear as a
single freshwater endmember in the aCDOM(350) vs. salinity (and
frw) distribution. Because the sampling season in the southeastern LS
ends when the land-fast ice begins to melt and does not start again until
August, it has not yet been possible to take samples during the important
phase of MIX water mass formation.
The proportions of meltwater and brine visible in the oxygen isotope data
from the LS at salinities > 25 are not clearly pronounced at
salinities < 25 (Fig. 4a). However, the observation that the oxygen
isotope data from the shelf water with salinities < 25 show no
dilution with meltwater but even indicate increased amounts of brine does
not mean that no meltwater was added. This seemingly contradictory
conclusion can be explained by the high export of sea ice in winter (Itkin
and Krumpen, 2017). Part of the brine formed in winter is still present in
the LS the following summer (Bauch et al., 2009a), while the newly formed
sea ice from which the brine originates is continuously exported to the
central AO. Because of the high sea-ice export, relatively young and thin
drifting pack ice (< 50 cm) covers most of the LS north of the
land-fast ice at the end of winter (Itkin and Krumpen, 2017). The melting of
this thin pack-ice cover and the land-fast ice leads to a weakening of the
isotopic brine signal (Bauch et al., 2013), but the isotopic composition
nevertheless remains dominated by the brine-enriched winter water. Hence,
the stable oxygen isotope signal of the sea-ice meltwater fraction is masked
by the high brine signal from the winter. When interpreting the isotope
data, it is important to note that the calculated sea-ice meltwater
fractions do not include the meltwater resulting from the high and variable
proportion of frozen river water that was part of the land-fast ice (Bauch
et al., 2010). Thus δ18O measurements of land-fast ice from the
LS cannot be used directly as an endmember for sea-ice meltwater.
Constructing the theoretical conservative mixing line (aCDOM(350) vs.
salinity) between MIX and the Nansen Basin seawater (Fig. 7) reveals
that the samples from the LS lie within the range of this mixing line.
However, the linear fit between aCDOM(350) and frw (Fig. 4b) shows a
freshwater endmember whose aCDOM(350) is about 30 % lower than in
MIX or in the post-freshet river water of the Lena. This could be an
indication for the photochemical or microbial removal of tDOM in the LS
shelf system. However, in this context it must also be taken into account
that the freezing of the river water in the fast ice – 30 %–50 % of the
annual freshwater discharge of the Lena River (Eicken et al., 2005) – also
leads to the extraction of tDOM from the river water. Although
biogeochemical- and/or photochemical-induced removal of tDOM cannot be ruled
out, our results suggest that the mixing of multiple water masses rather than
removal is the main factor controlling tDOM distribution in the LS. Similar
findings were discussed for the Chukchi Sea (Tanaka et al., 2016).
Sea-ice meltwater as possible cause for the <inline-formula><mml:math id="M337" display="inline"><mml:mrow><mml:msub><mml:mi>a</mml:mi><mml:mi mathvariant="normal">CDOM</mml:mi></mml:msub></mml:mrow></mml:math></inline-formula>(350) anomaly in
the LS in 2011 and the <inline-formula><mml:math id="M338" display="inline"><mml:mrow><mml:msub><mml:mi>a</mml:mi><mml:mi mathvariant="normal">CDOM</mml:mi></mml:msub></mml:mrow></mml:math></inline-formula>(350) distribution in the ESS in 2019
Previous studies indicate that approximately 30 % of the river-induced
DOC is already remineralised on the shelf (Cooper et al., 2005), with the
western ESS being a hotspot for organic matter degradation (Anderson et al.,
2011; Anderson and Amon, 2015; Semiletov et al., 2016). Thus, increased
removal of tDOM could also be the cause for the low-tDOM cluster of samples
that we observed in the LS in September 2011 (red outlined dots in Fig. 4a and b). This cluster represents water samples from the upper water column
(< 20 m) north of the Lena Delta that has a frw of 40 %–60 %
(Fig. 4b). The aCDOM(350) in the low-aCDOM(350) cluster was up to 50
% lower than in samples with the same frw percentage lying on the
mixing line of river water (summer discharge) and seawater. The general
hydrography of the LS in September 2011 was characterised by low surface
mixed layer salinities and sea surface temperatures (SSTs), which were
∼ 2 ∘C higher than the long-term average
(ERA5; Copernicus Climate Change Service, 2017). High SSTs caused by
solar warming support the assumption that photodegradation could be the
major factor leading to tDOM loss and a S275-295> 20 µm-1 (Fig. 4b). An alternative explanation to the removal of
tDOM in 2011 could be a change in the relative proportion of one of the
freshwater endmembers or a change in concentration in one of the tDOM
sources. In 2011, large areas of the southeastern LS were covered by melting
land-fast ice until the end of July. Exceptionally strong southeasterly
winds in August 2011 (Janout et al., 2020) steered the land-fast ice
meltwater plume to the northwest, resulting in additional mixing of
CDOM-poor meltwater into the shelves north of the Lena Delta and further
dilution of the shelf water, leading to low aCDOM(350).
The relationship between salinity and aCDOM(350) measured in the ESS in
summer 2019 was significantly different from the relationship observed in
the LS. The entire ESS shelf had salinities above 20, and aCDOM(350)
was about 50 % lower than in LS samples having the same salinity (Fig. 5a coloured dots and Fig. 7 red dots). During sampling in the last week of
September 2019, the ESS was ice-free. However, during the first week of July 2019, the ESS was almost completely covered with land-fast ice and drifting
pack ice that did not completely melt until the last week of July (data from
https://www.meereisportal.de, last access: 9 February 2021; Spreen et al., 2008).
Pugach et al. (2018) found that the distribution of tDOM concentration in
the ESS is mainly controlled by wind forcing. During our sampling in the ESS
in 2019, winds from the east (ERA5; Copernicus Climate Change
Service, 2017) pushed the inflow from the Pacific Ocean westward far onto
the ESS shelf. At the same time, southeasterly winds over the LS might have
blocked the Lena ROFI from extending into the western ESS (Anderson et al.,
2011; Janout et al., 2020). In addition, summer 2019 was characterised by
anomalously low river discharge into the LS and ESS (ArcticGRO version 2020-01-23; Shiklomanov et al., 2020). The wind forcing and low river
discharge led to a tDOM and salinity distribution in the ESS in 2019 that
was comparable to the distribution in 2008 and 2011, during which wind
forcing was similar (Alling et al., 2010; Pugach et al., 2018). In contrast,
in summers when the Lena River ROFI extends further eastward, the western
ESS is characterised by lower salinities and higher tDOM concentrations
(Pugach et al., 2018), with a salinity–tDOM relationship that corresponds to
the one we observed in the ROFI of the LS. Alling et al. (2010) assumed
that due to the long residence time of the water of the Lena River on the
ESS shelf, degradation of tDOM may be more advanced, leading to generally
low tDOM concentrations in the western ESS, as observed in August and
September 2008. This assumption is supported by the generally high
S275-295 values (Fig. 5a), which indicate that the CDOM in the ESS
has already been strongly transformed (e.g. by photochemical processes), but as in the LS, the addition of large amounts of DOM-poor meltwater from
the land-fast ice and meltwater from the pack ice, which may show even lower
CDOM absorptions (Kowalczuk et al., 2017), should result in a strong
dilution of the riverine tDOM. The ESS west of 170∘ E is
characterised by an extensive land-fast ice belt that varies in extent from
130 000 to 200 000 km2. The land-fast
ice, which contains a freshwater equivalent of 230–360 km3, usually melts back in the first 2 weeks of July, and drifting fields of
decaying ice might persist until August. We assume that the absence of the
Lena ROFI in the western ESS, the generally low river runoff and the
resulting higher fraction of low-tDOM sea-ice meltwater in the ESS led to
the comparatively low tDOM concentrations in the western ESS observed in
2008 and 2019. This assumption contradicts Alling et al. (2010), who
emphasise that the analysis of oxygen isotope data showed no evidence of
meltwater input in the western ESS in September 2008. However, as the
analysis of the ice distribution in 2008 (data from
https://seaice.uni-bremen.de/data, last access: 9 February 2021; Spreen et al., 2008) shows, the land-fast
ice still covered large areas of the southern ESS in mid-July. It is
therefore more likely that the difficulty in interpreting oxygen isotope
data from seasonally ice-covered shelves, which are characterised by high
sea-ice export rates and thus high brine fractions in the water column
(negative fsim values; see Sect. 4.2), has resulted in the masking of the
sea-ice meltwater signal (positive fsim value).
Transport dynamics of tDOM-rich brines in the LS
We used fast ice and seawater data from the southeastern LS to estimate how
much DOC might be expelled along with brine from growing land-fast ice in
the LS. The observed average DOC concentration in surface waters (0–5 m)
east of the Lena Delta (south of 73.4∘ N and east of
125.0∘ E) was 475 µmol L-1 (SD ± 165 µmol L-1). Thus, the observed median DOC concentration in the land-fast ice
of 96.2 µmol L-1 indicates that 379 µmol L-1 has
been removed during ice growth. This resulting value was multiplied by the
total volume of fast ice (273 km3; Kotchetov et al., 1994;
Barreis and Görgen, 2005; Selyuzhenok et al., 2005). Assuming that the
DOC is removed from the ice along with the brines, about 1.2 Tg yr-1
(± 0.54 Tg) DOC would thus be expelled from the growing land-fast ice.
The process of brine release is reflected in the water samples from the
polynya at the northern land-fast ice edge in the southeastern LS, which
show high brine fractions with simultaneously increased aCDOM(350)
(Ti12, winter 2012). This budget calculation, however, does not take into
account that the DOM in the brines has a higher bioavailability
(Jørgensen et al., 2015). Thus, the rapid removal of DOC in the brine or
the remineralisation of the remaining ice-bound DOC could also play an
important role.
Due to the low salinity of the surface mixed layer in the southeast LS at
the beginning of the freezing season and the continuous outflow of
freshwater from the Lena River in winter (Juhls et al., 2020), the loss of
buoyancy, which resulted from the addition of brine to the surface mixed
layer, is not high enough to completely erode the pycnocline by convective
mixing processes. This leads to permanently stratified waters under the
land-fast ice of the southeastern LS (Bauch et al., 2009a), as well as in the
polynya north of the Lena Delta, even during phases of high ice production
(Krumpen et al., 2011). The strong stratification results in the trapping of
tDOM-rich brines in the upper water column followed by advection, together
with the winter river discharge (Macdonald et al., 1995). How fast the
brines are advected across the shelf depends on the ocean currents in
winter. Eicken et al. (2005) described under-ice spreading velocities of
1–2.7 cm s-1. Moored velocity records from below the land-fast ice
(1998–1999, Hölemann unpub. data) showed a residual northeastward flow
of ∼ 2 cm s-1. These velocities are on the same order of
magnitude as those measured north of the Lena Delta in winter (Bauch et al.,
2010; Janout et al., 2013). Further north the geostrophic velocity fields
show eastward velocities above 5 cm s-1 with northward directions on
the northeastern LS shelf in March/April (Kwok and Morrison, 2011). At the
continental margin residual eastward velocities can reach more than 20 cm s-1 (Janout et al., 2015). The tDOM-rich brine formed in the southeast
LS in December could thus be transported over several hundred kilometres in
a northeasterly direction until the end of June, when the land-fast ice
melts. This is in accordance with Bauch et al. (2009b) who showed that the
residence time of the brines on the shelf could be as short as 1 year. A
significant proportion of the surface waters (0–20 m) of the LS and western
ESS leaves the Siberian shelf north of the New Siberian Islands (Morison and
Kwok, 2012) and supplies the Transpolar Drift Stream with tDOM-rich water
masses (Charette et al., 2020), while denser bottom waters flow north of the
New Siberian Island further to the east and leave the shelf in the western
ESS (Anderson et al., 2017). The transport of tDOM in the water masses at the
upper halocline is confirmed by investigations in the East Greenland Current
where a higher CDOM absorption occurring between 30 and 120 m water depth at
salinities between 32 and 33 is explained by a high fraction of brine and
river water from the Siberian shelves (Granskog et al., 2012).
During the northward transport of the brine-enriched water masses across the
shelf, an increase in density occurs due to the further influx of brines
from ice formation in leads and coastal polynyas of the western LS (Janout
et al., 2017). Erosion of the density stratification at the end of winter
and mixing with denser water masses from the western shelf of the LS could
dilute the tDOM-rich brines from the southeastern LS but at the same time
increase the density and bring them to greater water depths where they then
flow into the Arctic halocline. The seasonal evolution of the density
stratification in the LS is described in more detail in Janout et al. (2020). The water depth at which the brine is transported across the LS
shelf and into the Arctic halocline depends mainly on the density structure
of the water column in winter, which in turn depends on the position of the
ROFI in the previous summer (Bauch et al., 2012). An eastward-directed
transport south of the New Siberian Islands towards the ESS at a depth greater
than ∼ 20 m is inhibited by the shallow water depths of the
Dmitry Laptev (10 m) and Sannikov straits (18 m).
Conclusions
The eastern LS and western ESS together represent a region where much of the
freshwater-influenced waters of the Siberian shelves discharge into the
transpolar drift current and the Arctic halocline, which carries the shelf
waters towards Greenland and the Nordic seas (Morison et al., 2012;
Timmermans and Marshall, 2020). This region is thus a key region for a
better understanding of the Arctic marine carbon cycle. Analysis of the
extensive data set in this study, spanning several years and different
seasons, illustrates that the distribution of tDOM in the LS and ESS is
mainly determined by the physical mixing of river water, sea-ice meltwater
and seawater. Our study highlights that the formation and melting of
land-fast ice in the Laptev and East Siberian seas significantly shapes the
concentration distribution of tDOM in the area of the Siberian shelf seas.
In this context, the observed concentration distribution on the shelves
indicates a conservative mixing behaviour of tDOM introduced by the river
rather than a removal or an additional DOM supply from other sources. The
reason for this could be the short residence time of the river water on the
Siberian shelves, as well as a reduced photochemical degradation of the tDOM
due to the ice cover of the southern shelves lasting until July. The
extremely high CDOM absorption and suspended matter concentrations within
the ROFI (Juhls et al., 2019; Heim et al., 2014), which lead to a
self-shadowing effect, may also reduce photodegradation. The results of the
study also illustrate that the growth of fast ice and the associated
formation of CDOM-rich brines have an important role in the transport
pathways of tDOM across the Siberian shelves and the AO.
The duration of the land-fast-ice season in the LS is reduced by 2.8 d per
year (observational period 1999 to 2013; Seyushenok et al., 2015), while in
the second decade of the 21st century the onset of the spring freshet of the
Lena River happens about 6 d earlier than at the beginning of the
observational period in 1940 (Shiklomanov et al., 2020). Further changes of
the ice regime and the timing of spring freshet will certainly have an
impact on the dynamics of tDOM in the AO. In addition, the decline of Arctic
sea ice and the associated longer ice-free season will lead to changes in
wind forcing in the shelf systems of the Arctic and to an increased input of
solar radiation into the water column. This will significantly change
freshwater transport pathways, the heat content of the water column, and
stratification in the LS and ESS. The heat content in turn influences the
formation of new ice and thus also the production of DOM-rich brines.
Because density stratification also controls where and at which depth the
DOM-rich brine leaves the shelf, changes in stratification also impact the
future transport pathways of tDOM in the AO. Studying these processes is
important not only to decipher the Arctic carbon cycle but also because it
regulates physical processes such as radiative forcing in the upper ocean,
which has important effects on, for example, sea surface temperature, water
column stratification and UV penetration (Pegau, 2002; Hill, 2008; Granskog
et al., 2015b; Gnanadeskian et al., 2019; Soppa et al., 2019).
Data availability
The data sets used for this study (Table 1) can be downloaded from the
Pangaea website (Bauch and Thibodeau, 2020; Bauch et al., 2018; Eulenburg et al., 2019; Hölemann et al., 2020a, b, c, d, e, f, g; Kattner et al., 2010).
Author contributions
JAH was responsible for sample collection, data
analysis, and the organisation and writing of the manuscript. BJ, MJ and DB
contributed with discussions and data collection at sea. BPK analysed the DOC
samples and contributed with discussions. BH contributed with discussions
and data records. All authors contributed to the article and approved the
submitted version.
Competing interests
The authors declare that they have no conflict of interest.
Disclaimer
This article reflects only the authors' views; the
funding agencies and their executive agencies are not responsible for any
use that may be made of the information that the article contains.
Publisher's note: Copernicus Publications remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
Acknowledgements
We thank the crews and colleagues on board the
research vessels involved in sampling. Our thanks also go to the helicopter
pilots, the staff of the airport in Tiksi and the colleagues who made the
winter expedition possible. We also thank the colleagues of the
Russian–German Otto Schmidt Laboratory in St. Petersburg for support and
accessibility of laboratory instruments for sample analysis. We are grateful
to Piotr Kowalczuk and an anonymous reviewer for the constructive comments on
the manuscript. Our special thanks go to Heidemarie Kassens who headed all
summer and winter expeditions and made sure that the scientific programme
could be carried out as planned even under the most difficult conditions. We
greatly acknowledge Leonid A. Timokhov, Heidemarie Kassens and Vladimir V. Ivanov for coordinating the German–Russian Laptev Sea System project.
Financial support
This research has been supported by
the Bundesministerium für Bildung und Forschung (grant nos. 03G0833 and 03F0776). Bennet Juhls was supported by Geo. X, the Research Network for Geosciences in Berlin and Potsdam (grant no. SO_087_GeoX).
The article processing charges for this open-access publication were covered by the Alfred Wegener Institute, Helmholtz Centre for Polar and Marine Research (AWI).
Review statement
This paper was edited by Anja Engel and reviewed by Piotr Kowalczuk and one anonymous referee.
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