OSOcean ScienceOSOcean Sci.1812-0792Copernicus PublicationsGöttingen, Germany10.5194/os-13-349-2017Shelf–Basin interaction along the East Siberian SeaAndersonLeif G.leif.anderson@marine.gu.seBjörkGöranHolbyOlaJutterströmSaraMörthCarl MagnusO'ReganMatthttps://orcid.org/0000-0002-6046-1488PearceChristofhttps://orcid.org/0000-0002-4866-3204SemiletovIgorStranneChristianhttps://orcid.org/0000-0003-1004-5213StövenTimTanhuaTosteUlfsboAdamJakobssonMartinhttps://orcid.org/0000-0002-9033-3559Department of Marine Sciences, University of Gothenburg, P.O. Box 461, 40530 Gothenburg, SwedenDepartment of Environmental and Energy Systems, Karlstad University, 651 88 Karlstad, SwedenIVL Swedish Environmental Research Institute, Box 530 21, 400 14 Gothenburg, SwedenDepartment of Geological Sciences, Stockholm University, 106 91 Stockholm, SwedenDepartment of Geoscience, Aarhus University, Aarhus, DenmarkInternational Arctic Research Center, University Alaska Fairbanks, Fairbanks, AK 99775, USAPacific Oceanological Institute, Russian Academy of Sciences Far Eastern Branch, Vladivostok 690041, RussiaThe National Research Tomsk Polytechnic University, Tomsk, RussiaHelmholtz Centre for Ocean Research Kiel, GEOMAR, Kiel, GermanyCenter for Coastal and Ocean Mapping/Joint Hydrographic Center, Durham, NH 03824, USADivision of Earth and Ocean Sciences, Nicholas School of the Environment, Duke University, Durham, NC 27704, USALeif G. Anderson (leif.anderson@marine.gu.se)27April20171323493635December201616December20162April20173April2017This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://os.copernicus.org/articles/13/349/2017/os-13-349-2017.htmlThe full text article is available as a PDF file from https://os.copernicus.org/articles/13/349/2017/os-13-349-2017.pdf
Extensive biogeochemical transformation of organic matter
takes place in the shallow continental shelf seas of Siberia. This, in
combination with brine production from sea-ice formation, results in cold
bottom waters with relatively high salinity and nutrient concentrations, as
well as low oxygen and pH levels. Data from the SWERUS-C3 expedition with
icebreaker Oden, from July to September 2014, show the distribution of such
nutrient-rich, cold bottom waters along the continental margin from about
140 to 180∘ E. The water with maximum nutrient concentration,
classically named the upper halocline, is absent over the Lomonosov Ridge at
140∘ E, while it appears in the Makarov Basin at 150∘ E and
intensifies further eastwards. At the intercept between the Mendeleev Ridge
and the East Siberian continental shelf slope, the nutrient maximum is still
intense, but distributed across a larger depth interval. The nutrient-rich
water is found here at salinities of up to ∼ 34.5, i.e. in the
water classically named lower halocline. East of 170∘ E transient
tracers show significantly less ventilated waters below about 150 m water
depth. This likely results from a local isolation of waters over the Chukchi
Abyssal Plain as the boundary current from the west is steered away from
this area by the bathymetry of the Mendeleev Ridge. The water with
salinities of ∼ 34.5 has high nutrients and low oxygen
concentrations as well as low pH, typically indicating decay of organic
matter. A deficit in nitrate relative to phosphate suggests that this
process partly occurs under hypoxia. We conclude that the high nutrient
water with salinity ∼ 34.5 are formed on the shelf slope in
the Mendeleev Ridge region from interior basin water that is trapped for
enough time to attain its signature through interaction with the sediment.
Introduction
The extensive, flat, and shallow shelf areas of the Laptev and East Siberian
seas are particularly influenced by the changing climate in the Arctic.
Coastal erosion from wave action becomes widespread when the summer sea-ice
cover shrinks and river discharge increases in warmer humid conditions, both
affecting organic matter and nutrient supply (Charkin et al., 2011). At the
same time, the decrease in summer sea-ice coverage changes the dynamics of
the ocean by increasing vertical mixing and brine production in the fall
when sea ice again starts to form over areas that in the past used to be
sea-ice covered. The changes may impact shelf basin exchange (e.g. Dethleff,
2010; Nishino et al., 2013).
Map of the Arctic Ocean with general currents at intermediate depths
over the deep basins and exchange with the surrounding oceans (a).
The black frame indicates the investigated area that is illustrated in
panel (b) with the hydrographic station positions of sections A to F
as white points and those of sediment cores in yellow stars. The Arctic Ocean
Section 1994 stations are in green and the orange points show the positions
of the ACSYS 96 stations, which are used as historic references. The yellow
frame borders the area where the historic sea-ice coverage has been
evaluated; see Fig. 10. Abbreviations: Fram Strait (FS), Bering Strait (BS)
Nansen Basin (NB), Amundsen Basin (AB), Makarov Basin (MB), Canada Basin
(CB), Chukchi Plateau (CP), and Chukchi Abyssal Plain (CAP).
Here we assess data collected in 2014 along the continental shelf break of
northern Siberia. Acquired oceanographic and bottom sediment data add to our
understanding of water mass modification in the central Arctic Ocean basin.
The objectives are to describe the spreading of shelf waters, including
those richest in nutrients, from the East Siberian Sea and assess their
sources, as well as to evaluate potential effects of diminishing sea-ice
coverage under a warmer climate.
The Arctic Ocean has an area of about 9.5×1012 m2 of
which more than half is comprised of shallow continental shelf seas
(Jakobsson, 2002). The deep central part consists of several basins; the
Nansen and Amundsen basins are together denoted the Eurasian Basin, and the
Canada and Makarov basins constitute the Amerasian Basin. The Lomonosov
Ridge stretches from the continental slope of the Laptev Sea to the slope
off northern Greenland and separates the Eurasian Basin from the Amerasian
Basin (Fig. 1). The deep waters of the Arctic Ocean are supplied from the
Atlantic Ocean, entering either through the eastern Fram Strait (Fram Strait
Branch, FSB) or over the Barents Sea (Barents Sea Branch, BSB). The latter
water flows into the Kara Sea before exiting through the St Anna Trough
along the continental margin where it covers a depth range down to about
1500 m (e.g. Schauer et al., 2002). Both branches flow to the east and
follow the bathymetry in a cyclonic pattern around the basins (Rudels et
al., 1994, Fig. 1), the difference being that the FSB takes the inner turn
and is largely restricted to the Eurasian Basin. It is mainly the BSB that
flows over the Lomonosov Ridge into the Makarov Basin north of the Laptev
Sea.
The upper waters are entering from both the Pacific and Atlantic oceans,
where the latter either pass over the Barents shelf or through Fram Strait.
The upper waters have classically been divided into a surface mixed layer
(SML) that varies seasonally, an upper halocline of mainly Pacific origin,
and a lower halocline of Atlantic origin (e.g. Jones and Anderson, 1986;
Rudels et al., 1996). The flow pattern of these waters differs. The lower
halocline primarily follows the underlying Atlantic layer, while the upper
halocline, and even more so the surface mixed layer circulation, is much
impacted by the dominating wind field (e.g. Jones et al., 2008). The flow of
the surface water is dominated by transport from the Laptev Sea towards Fram
Strait, the Transpolar Drift, and one cyclonic circulation in the Canada
Basin, the Beaufort Gyre. The size of the latter is determined by the
atmospheric pressure field, where a negative Arctic Oscillation results in a
larger Beaufort Gyre compared to a positive Arctic Oscillation (Proshutinsky
et al., 2009).
The properties of the surface mixed layer and the upper halocline are
modified over the shelves, and for the SML also in the central Arctic Ocean
by, e.g. mixing with river runoff, sea-ice melt, and brine from sea-ice
formation. Biogeochemical processes also modify the chemical signature,
e.g. lowering the nutrient concentration of the SML through primary
production and increasing the nutrient concentration in the upper halocline
through remineralization of organic matter (e.g. Jones and Anderson, 1986).
The latter process has been reported to occur in the Chukchi Sea (Bates,
2006; Pipko et al., 2002), East Siberian Sea (Nishino et al., 2009; Anderson
et al., 2011), and Laptev Sea (Semiletov et al., 2013, 2016).
One of the most pronounced signatures of the upper halocline of the central
Arctic Ocean is a silicate maximum, which was first reported in 1968–1969 from
observations made from the drifting T-3 ice island in the Canada Basin
(Kinney et al., 1970). In 1979 the silicate maximum was observed during the
LOREX study over the Lomonosov Ridge and into the fringe of the Amundsen
Basin (Moore et al., 1983). In 1994 no silicate maximum was observed in the
Makarov Basin along a section from the Chukchi Sea to the North Pole (Swift
et al., 1997). It is clear that the distribution of the upper halocline with
its prominent silicate signature has varied much in the past and with
changing sea-ice coverage it might vary even more in the future. In this
contribution we give some indications of the latter.
Furthermore, recently high silicate concentrations were found at salinities
∼ 34.5 along the continental slope of the eastern East
Siberian Sea (Nishino et al., 2009; Anderson et al., 2013). Nishino et al. (2009)
suggested that this silicate maximum was produced by decomposition of
opal-shelled organisms along the continental margin. Based on δ18O data collected in 2008, Anderson et al. (2013) reported a brine
content of at least 4 % and a small temperature minimum signature
associated with the high silicate concentration. The present study will
expand on the formation process of this water
Methods
Water column data in this study were obtained along six oceanographic sections
across the shelf break (A–F; Fig. 1) during the SWERUS-C3 (Swedish–Russian–US
Arctic Ocean Investigation of Climate–Cryosphere–Carbon
Interactions) expedition in 2014 with Swedish icebreaker Oden. SWERUS-C3 is a
multi-disciplinary international program focusing on investigating the
functioning of the Climate–Cryosphere–Carbon (C3) system of the East
Siberian Arctic Ocean. The expedition consisted of two legs with the icebreaker
Oden. Leg 1 started 5 July in Tromsø, Norway, and followed the Siberian
continental shelf to end in Barrow, Alaska, on 21 August. Leg 2 took the return
route from Barrow and ended in Tromsø on 3 October after concentrating the
field program to the continental shelf break, slope, and the adjacent deep
Arctic Ocean basin. Data from leg 2 focusing on the shelf break are
discussed in this study.
Water samples were collected using a rosette system equipped with 24 bottles
of the Niskin type, each having a volume of 7 L. The bottles were closed during
the return of the CTD rosette package from the bottom to the surface and
water samples for all constituents were drawn soon after the rosette was
secured in the sampling container.
The following constituents are used here: bottle practical salinity,
dissolved inorganic carbon (DIC), total alkalinity (TA), pH, oxygen,
nutrients (NO3-+ NO2-, PO43-, SiO2),
and the transient tracer sulfur hexafluoride (SF6). The order of
sampling was determined by the risk of contamination meaning that transient
tracer samples were collected first followed by oxygen, the carbon system
parameters, nutrients, and salinity.
Salinity and temperature data were collected using a SeaBird 911+ CTD with
dual SeaBird temperature (SBE 3), conductivity (SBE 04C), and oxygen sensors
(SBE 43) attached to a 24 bottle rosette for water sampling. Salinity data
were calibrated against deep water samples analysed onboard using an Autosal
8400B lab salinometer. The salinometer was calibrated using one standard seawater ampule (IAPSO standard seawater from OSIL Environmental Instruments
and Systems) before each batch of 24 samples. The accuracy of the Autosal
salinities and CTD salinities should both be within ±0.003 and the
accuracy for temperature ±0.002 ∘C. Water samples for salinity
were analysed for more than 90 % of the depth and when no data were
available the CTD salinity was used in the evaluation.
The water samples for determination of the transient tracer SF6 were
directly drawn from the Niskin bottles using 250 mL glass syringes. The
samples were stored in a cooling bath that was continuously rinsed with cold
surface water to prevent outgassing of the tracers. Measurements were
directly performed on board, using a purge and trap GC-ECD system similar to
the “PT3” set-up described in Stöven and Tanhua (2014). The column
composition was as follows: the trap consisted of a 1/16” column packed
with 70 cm Heysep D, the 1/8” precolumn was packed with 30 cm Porasil C
and 60 cm Molsieve 5Å and the 1/8” main column with 200 cm Carbograph
1AC and 20 cm Molsieve 5Å. The precision for onboard measurements was
±0.02 fmol kg-1 for SF6. In the evaluation of the data,
SF6 is given in partial pressure normalized to 1 atm, which is equal to
the mixing ratio on the volume scale in parts per trillion
(ppt). The advantage
of using partial pressure instead of concentrations for dissolved gases in
the ocean is that the partial pressure is not influence by temperature and
salinity effects. The precision is given in concentration since it is related
to the absolute value per kg seawater. Age modelling based on these
transient tracers is complicated and erroneous at high latitudes due to
ambiguous reasons (Stöven et al., 2015, 2016). Hence, we do not provide
any statements about the ventilation timescale but rather the ventilation
states of the water masses in the Arctic Ocean based on the concentration
distribution.
An automated Winkler titration system was used for the oxygen measurements
with a precision of ∼ 1 µmol kg-1. The accuracy
was set by titrating known amounts of KIO3 salts that were dissolved in
sulfuric acid. As the amount was known to better than 0.1 % the accuracy
should be significantly less than the precision.
DIC was determined by a coulometric titration method based on Johnson et
al. (1987), having a precision of 2.0 µmol kg-1, from
duplicate sample analyses, with the accuracy set by calibration against
certified reference materials (CRM; Batch nos. 123 and 136), supplied by A.
Dickson, Scripps Institution of Oceanography (USA). TA was determined by an
automated open-cell potentiometric titration (Haraldsson et al., 1997), with
a precision better than 2.0 µmol kg-1 and the accuracy
ensured in the same way as for DIC. pH was determined by a spectrophotometric
method, based on the absorption ratio of the sulfonephtalien dye, m-cresol
purple (mCP) (Clayton and Byrne, 1993). Purified mCP was purchased from the
laboratory of Robert H. Byrne, University of South Florida, USA. The accuracy
was estimated to 0.006 from internal consistency calculations of analysed CRM
samples and the precision, defined as the absolute mean difference of
duplicate samples, was 0.001 pH units. The seawater pH is reported on the
total scale and in situ temperature.
Geographic information of sediment cores and their biogenic silica
content.
The partial pressure of carbon dioxide (pCO2) was calculated from the
combination of pH and TA, and pH and DIC, using CO2SYS version 1.1 (van
Heuven et al., 2011) with the stoichiometric dissociation constants of
carbonic acid (K1∗ and K2∗) and bisulfate
(KHSO4∗) given by Millero (2010) and Dickson (1990),
respectively. Input data included salinity, temperature, PO4, and
SiO2. The reported values are the average of the two calculated for each
sample. The uncertainty was computed using a Monte Carlo approach (Legge et
al., 2015) and is, expressed as double standard deviation, about 2.5 %.
Sections of salinity (left) and silicate in µmol L-1
(right) of the upper 300 m of sections A to F; see Fig. 1 for location of
sections. Sections drawn using Ocean Data View (Schlitzer, 2017).
Besides the extensive sampling and measuring of the water column, analyses
were also performed on sediments. Sediment samples from six coring stations
along the SWERUS leg 2 cruise track (Fig. 1) were taken from four different
depths in the upper 16 cm (Table 1). Two different types of coring devices
were used: a gravity corer (GC) and multi-corer (MC). These 24 samples were
analysed for biogenic silica (BSi) content, with the aim of investigating a
possible sedimentary source of the silicate maximum observed in the water
column. Biogenic silica was measured using a wet alkaline extraction
technique (Conley and Schelske, 2001). Samples were
freeze dried and approximately 30 mg of homogenized sediment was placed in a
mild alkaline solution (1 % Na2CO3) at 85 ∘C and
aliquots were taken at 3, 4, and 5 h during this leaching process. For each
of these subsamples, dissolved Si was measured by Inductively Coupled Plasma
Spectrometry, using a Thermo ICAP 6500 DUO. All BSi is assumed to have
dissolved after 2 h leaching, after which only Si from minerals is being
released. Based on this principle, the zero-hour intercept of the slope from
the 3, 4, and 5 h Si concentrations is used to calculate the biogenic
fraction. This method was validated by including blanks, and standards from a
previous inter-laboratory comparison exercise (Conley, 1998). The relative
uncertainties associated with this method are estimated to be ±20 %
of the measured value and precision of the ICP is from certified standard
measurements better than 5 %.
Results
The salinity distribution along the continental margin from the Lomonosov
Ridge to the Chukchi Sea shows a similar general pattern, but with some
significant variations especially in the top 50 m (Fig. 2). The
thinnest layer of low-salinity surface water is found at the Lomonosov Ridge
(section A), which increases in thickness eastward in the study area. In
section B we find the lowest salinity of 24.55 at 10 m water depth, followed
by a very sharp halocline with the salinity increasing from about 32 at 50 m
to 34 at 100 m depth. Further to the east the halocline is less sharp with,
e.g. the 34 salinity isoline deepening to a depth of more than 200 m. Here,
also the > 33 isolines deepens from the shelf towards the deep
basin, especially in sections D and E.
The silicate distribution is variable between the sections (Fig. 2). Over
the Lomonosov Ridge (section A) the highest silicate concentration, reaching
15 µmol L-1, is found in the surface. In section B the maximum
is instead found at about 50 m depth and varies horizontally, with the
highest concentration exceeding 30 µmol L-1. At this depth the
salinity is around 33. Further to the east at section C, the concentration
in the silicate maximum is higher and is found somewhat deeper and also at a
larger salinity range. It extends horizontally all over the shelf and slope,
although with concentrations decreasing some 100–150 km seaward from the
shelf break. At the station farthest out in the deep basin the concentration
is close to the maximum in section B.
Properties in the upper 1000 m along section D of Fig. 1. Sections
drawn using Ocean Data View (Schlitzer, 2017).
Sections D and E are fairly close to each other and both show a similar
pattern. The maximum silicate concentration, above 50 µmol L-1,
is close to the bottom at 100–150 m depth (Fig. 2). From here the
concentration decreases gradually away from the shelf break, to the lowest
maximum at the outer station, around 30 µmol L-1. Another
specific characteristic of the silicate distribution at these sections are
the wide depth range of concentrations more than 15 µmol L-1.
Here it spans the range of about 50 to 250 m, whereas in section C it only spans
50 to 150 m. To some degree this is attributed to the more gradual increase
in salinity with depth, but there are also high concentrations at salinities
above 34.5. In section F the silicate concentration is lower and also spans
a narrower depth range. However, this section starts further away from the
shelf break and may be difficult to compare with the other sections.
Biogenic silicate (BSi) concentrations (percent dry weight) in the
upper 16 cm of the sediment (a) at coring sites shown in
panel (b). The symbols of the coring sites are colour coded after
measured BSi concentration from the lowest concentrations in blue to the
highest in red. The black bars represent the depth layer of the sediment that
is analysed.
The waters of high silicate concentration have other distinct characters
such as high concentrations of the other nutrients, phosphate and nitrate,
high apparent oxygen utilization (AOU) and pCO2, and low pH (Fig. 3).
The top 100–150 m is colder than 0 ∘C and the nutrient maximum as
represented by phosphate is largely confined to the coldest water. There are
some small differences in the exact pattern of the different parameters,
e.g. the AOU maximum is located slightly deeper than that of phosphate
farthest out in the deep basin.
pSF6 (ppt per volume) profiles from surface to 550 m depth,
coloured by AOU (µmol kg-1) of the stations in sections A to
F.
Biogenic silica concentrations in the analysed sediments varied widely
between the different sites. The full names of the cores include the prefix
SWERUS-L2, which henceforth is omitted. The most western sites (coring
stations 21MC1, 25MC1, 27MC1 ∼ water column sections A, B, C)
had BSi levels of less than 0.5 % (Fig. 4). Values increased slightly to
the east, reaching up to 1 % BSi in station 18MC1 (∼ water
column section D) and up to 2 % in station 7GC1 (∼ section F).
Concentrations in the most eastern station 5GC1 located on the western
flank of Herald Canyon are, however, much higher and reach up to 13.5 %.
The surface generally contains the highest concentration of BSi at all
sites, except for station 5GC1 where the concentration increases down core.
These subtler differences should however be treated with caution due to the
large uncertainties associated with the measurement method.
The mean mixed layer partial pressure of SF6 along all sections is
∼ 8.1 ppt (Fig. 5), which is slightly below the contemporaneous
atmospheric value of 8.4 ppt. At all sections except A, a SF6 minimum
is associated with the maximum in AOU. Close to the shelf in section B this
SF6 minimum is 6.4 ppt at 80 m and shoals polewards to 50 m with
increasing partial pressure to the range 6.7–7 ppt. The elevated AOU
values are 75–138 µmol kg-1 at these depths. The SF6
minimum becomes more significant at section C with partial pressures between
4.5–5.1 ppt at 95–130 m (135–183 µmol kg-1 AOU). The maximum
deepens eastwards to about 200 m at sections D, E, and F with partial
pressures between 2.5–3.4 ppt and 90–118 µmol kg-1 AOU.
The SF6 partial pressure in the AW (Atlantic Water)
layer between 250 and 600 m is homogeneously distributed with a mean value
of about 6 ppt at sections B and C (Fig. 5). In contrast, sections D, E, and
F show significant lower mean partial pressures of 4.1–3.4 ppt in the same
depth interval with the lowest values at section F. Note that the deep
SF6 partial pressures at sections E and F are close to the values in the
overlying minimum at 200 m and the minimum can thus not be defined by
SF6 data only. However, the minima can clearly be separated by the AOU
values since the warm AW layer shows constant low values of about
50 µmol kg-1 along all sections.
The bottom water partial pressure of SF6 has a general trend of
decreasing values at a specific isobaths from the west to the east (Fig. 6).
The highest partial pressures of 6.1–6.9 ppt can be found at section B
between 100 and 500 m. Section C shows increasing partial pressures with
depth from 4.6 ppt at 100 m to 6.5 ppt at 500 m, with decreasing values
deeper. A similar gradient of 4.4 to 5.7 ppt can be found at sections D, E, and F at the same depth
range. Below ∼ 500 m the partial pressure decreases with increasing
depth, reaching the detection limit at 1900–2000 m in the Makarov Basin
(Fig. 6).
Discussion
In section A, the most western that is located over the Lomonosov Ridge, the
highest silicate concentrations are found in the surface. Hence, no
sub-surface maximum typical of the upper halocline water is present here.
The surface silicate maximum is associated with low salinity, which is a
typical signature of runoff from the Lena River. Thus, this surface water
has a substantial fraction of freshwater from river, even if there also is a
contribution of sea-ice melt. The high silicate surface water is also seen
in section B and partly in section C but not further to the east, indicating
the limit of the river plume to this part of the deep Arctic Ocean.
SF6 partial pressure (ppt) in the sample collected closest to
the bottom, typically 5–10 m above.
Depth profiles of temperature (a), salinity (b),
silicate (c), and N∗∗(d) in the upper 300 m of
all stations at sections D and E; see Fig. 1 for locations.
Temperature (a), oxygen (b), silicate (c)
and N∗∗(d) versus salinity for the stations at
sections D and E. The panels (a) and (b) are from the CTD
output, while (c) and (d) are from water samples analysed.
The highest silicate concentrations are found at the shelf slope of sections D and E
and this is also where the salinity isolines shoal (Fig. 2). The
increase in salinity along the shelf slope at bottom depths less than 250 m
is accompanied by an increase in temperature as illustrated in the depth panels
(Fig. 7a and b) with no indication of mixing with another water mass, as all
data from these sections have the same shape in a T-S panel (Fig. 8a) for
salinity > 32.5. The slope of the isolines infers a near-bottom
increase of the current due to geostrophic shear, which is superimposed on
the typical overall eastward current (e.g. Rudels, 2012). Although it is not
possible to determine the absolute current velocity from just a geostrophic
calculation, our data together with the known direction of the mean flow
suggest that we have a bottom intensified flow in the eastward direction.
The magnitude of this increase is about 3 cm s-1 (based on the density
difference and distance between the two slope stations located at 164 and
241 m water depth in section E) over the depth range 100–150 m, which is not
negligible.
At sections D and E is the maximum observed silicate concentration about
56 µmol L-1 and found at ∼ 120 m (Fig. 7c) but
elevated concentrations are found down to nearly 250 m depth. Plotting
silicate concentrations against salinity (Fig. 8c) shows a clear pattern
with a shallow maximum around 33 and a deep maximum at 34.5. These maxima
are also evident in the section panels in Fig. 2. When the nutrient maximum
is accompanied by an oxygen minimum (Fig. 3) it suggests organic matter
mineralization, and if this occurs at low oxygen levels, nitrate is lost as
electron acceptor via either denitrification or anammox. Such conditions can
only be met close to, or in, the sediments of the shelves within the Arctic
Ocean as all other waters observed in this region are well oxygenated. A
deficit in nitrate is seen when computing the property N∗∗= 0.87 × [NO3] - 16 × [PO4] + 2.9 (Codispoti et
al., 2005), which gives a constant value if the classical
Redfield–Ketchum–Richard N : P ratio (Redfield et al., 1963) is followed. A
low value indicates denitrification. The N∗∗ profiles (Fig. 7d) show a broad
minimum focused at depths around 100 m, strongly indicating that this water
has had its signature influenced by hypoxic conditions, i.e. loss of
nitrate when used as electron acceptor during mineralization of organic
matter at low oxygen concentration.
More information on the formation history of the high salinity silicate
maximum water can be obtained from property versus salinity panels (Fig. 8).
The T-S curve show a typical shape for the halocline, with a warmer low-salinity
water at S≈31, followed by a temperature minimum at 32 < S < 33
and then increasing temperature with salinity to a
maximum in the Atlantic Layer, followed by colder water towards the highest
salinity in the deep water (Fig. 8a). The temperature minimum has
historically been attributed to winter water, often with a signature of
brine contribution (e.g. Aagaard et al., 1981; Anderson et al., 2013). This
brine enhanced water follows the shelf bottom and gets enriched in organic
matter decay products during its flow towards the deep basin. A nearly
strait mixing line can be seen in the salinity range from about S=34 to
that of the temperature maximum (Fig. 8a), i.e. no Tmin at the high
salinity silicate maximum water. Oxygen profiles, on the other hand, show a
clear minimum at S≈34.5 (Fig. 8b) indicating organic matter
remineralization. Comparing the oxygen signature with those of silicate and
N∗∗ (Fig. 8c and d) reveals interesting information. The broad silicate
maximum around S≈33 has a minimum in N∗∗ but no minimum in oxygen
even if the concentration is some 100 µmol kg-1 below
saturation, while the silicate maximum at S≈34.5 has a clear
oxygen minimum but with only a slight minimum in N∗∗. The most plausible
explanation for this pattern is that the nutrient maximum at low salinity
had a higher oxygen concentration before exposure to organic matter decay at
the sediment surface. The waters with S > 34 at some stations
with lower oxygen and higher silicate concentrations also have lower N∗∗
(Fig. 8b, c and d), indicating less mixing and thus potentially more
recent contact with the sediment surface.
The conclusion is that both nutrient maxima are formed in contact with
hypoxic sediments, with one maxima at salinity around S≈33 mainly
being formed on the shelf where the preformed water is well oxygenated by
interaction with the atmosphere during ice-free periods and ice formation
periods with cooling and convection, while the nutrient maximum at S≈34.5 is formed at the shelf break of more than 100 m depth. Such
a scenario is consistent with the SF6 partial pressure of the silicate
maximum at S≈34.5 being close to those in the deep basin, while
that around S≈33 has a significantly higher level of around 7 ppt
(Fig. 5). At section B the maximum AOU is associated with S≈33
and found at about 50 m depth (Fig. 5b) and at C it is found at around 100 m
depth (Fig. 5c). At the latter section the maximum AOU is found at S≈34.5 associated with the SF6 partial pressure minimum of
around 5 ppt. At the same salinity there is also a weaker minimum in section B
at about 75 m depth. In sections D, E, and F the SF6 minimum is also
found at S≈34.5 but at a deeper depth of 200 m, all associated
with the AOU maximum. However, at these sections the AOU maximum has a
SF6 partial pressure close to that of the water deeper, indicating that
the basin water in the Chukchi Abyssal Plain is the source of this high
salinity nutrient maximum water. The presence of the high salinity SF6
minimum at section C, and to a lesser degree at section B, points to the
existence of a westward penetration of water at the shelf break. However,
this does not need to be a persistent flow, but can be something that occurs
intermittently. Strengthening of these concussions are seen in Fig. 9 where
the pSF6 interpolated to a salinity of 34.5 has a strong gradient with
increasing partial pressure towards the west and significant higher silicate
concentrations at lower pSF6.
The formation of the silicate maximum at S≈34.5 on the shelf
break is in line with Nishino et al. (2009), who suggested that the silicate
maximum at this high salinity was produced by decomposition of opal-shelled
organisms along the continental margin. Anderson et al. (2013) showed that
this high salinity silicate maximum had a brine content of at least 4 %
and that the CTD record had a small temperature minimum signature. This was
not the situation in 2014, illustrating that the conditions likely are not
constant with time. This water with a salinity of just over 34 has
historically been named lower halocline water (Jones and Anderson, 1986),
without a signature of any nutrient maximum. Hence, this more recent finding
of high silicate concentrations along the continental margin of the East
Siberian Sea is either a local or a new phenomenon. The sediment record (Fig. 4)
clearly shows that the BSi content is low in the slope off the western
part of East Siberian Sea (sections B and C; coring stations 27MC1, 25MC1,
21MC1), making local decomposition in this region unlikely. Further to the
east the BSi content increases slightly to the location of sections D and F,
with a large increase at the most eastern station in the Herald Canyon (13.5 % BSi at site 5GC1)
where opal-shelled organism, primarily diatoms, are
abundant in the bottom sediments. This strongly supports a Pacific Ocean
source of silicate, but does not exclude that some of the silicate-rich
water enters the eastern East Siberian Sea before transformation and escape
to the slope and deep central basins. Such a scenario is consistent with the
decreasing silicate concentrations to the west in the salinity range 34.3 to
34.7 (Fig. 9b). However, it is not possible to fully elucidate the transport
and transformation of silicate from these few sediment profiles, especially
since they are also from variable bottom depths (Table 1). Nevertheless,
these sediment observations do not contradict occasional westward flow along
the shelf break, as suggested by the SF6 signature.
SF6 partial pressure (ppt) interpolated to the salinity
34.5 (a) and silicate versus pSF6 in the salinity range 34.3 to
34.7 colour coded by longitude (b).
Variability is also seen in a comparison with historic data. Our section F
is on the border to the Chukchi Abyssal Plain, where the Arctic Ocean
Section hydrographic program collected a section of data in 1994. In Fig. 10
we compare these two data sets and it is clear that the silicate maximum at
S≈34.5 was more or less absent in 1994. However, at stations with
bottom depths ∼ 180 m the silicate concentration was close to
20 µmol L-1 towards the seafloor, and at the station with bottom
depth ∼ 250 m, the concentration decreased to 18 µmol L-1
towards the seafloor. These are the stations where the salinity
does not reach the maximum salinity in Fig. 10a. Hence, there seems to be a
signal from the shelf slope that did not penetrate deep into the Chukchi
Abyssal Plain. Also N∗∗ had relative to the deepest data higher values at S≈34 except for at the shallowest stations with elevated silicate
concentrations (Fig. 10b). Centred at S≈33 the silicate maximum
is higher and the N∗∗ minimum is lower in 1994, indicating a stronger
contribution of organic matter decay at low oxygen levels from the shelf.
There is also an indication of a wider salinity interval of the silicate
maximum in 2014 compared to 1994, especially towards the high salinity end
(Fig. 10a).
Silicate (a) and N∗∗(b) versus
salinity in the Chukchi Abyssal Plain area; data from the Arctic Ocean
Section in 1994 in green (stations marked by green points in Fig. 1b) and
from SWERUS-C3 in red (section F in Fig. 1b).
Historically the extent of the nutrient maximum has varied but few studies
along the Siberian continental margin have been reported. Data from east of
175∘ E for years between 2001 and 2010 were compiled by Nishino et al. (2013)
showing the presence of the nutrient maximum but with some
variability in both the maximum concentration and the vertical extent
between the years. Our 2014 data show a clear maximum in the Makarov Basin,
at section B some 150 nautical miles east of the Lomonosov Ridge at about
longitude 153∘ E, with higher concentrations in the sections further to
the east (Fig. 2). As the concentration of silicate generally increases
towards the shelf in sections B and C and also increases from section B to C
(Fig. 2) it is likely that the source is the local shelf area. Data obtained
from RV Polarstern in the same area as section B (see Fig. 1b for station
locations) during the summer of 1996 (ACSYS 96) did not show any elevated
silicate concentrations in the halocline (Fig. 11). This could be an effect
of either that in 1996 the nutrient-rich water was confined closer to the
shelf on its transport to the east, or that the production site of this
water has extended further to the west since 1996. We find it most plausible
that the latter is the cause as the sea-ice climate has changed
significantly over the shelf south of these sections during the last 20 to
30 years (Fig. 12) that potentially have moved the production areas further
to the west. Up to about the year 2000 most summers had more than 50 %
sea-ice coverage, with a few years with less than 10 %. During the last
10 years the typical conditions for the month of September is more than 90 %
open water. All through the record the area is more or less ice covered in
November, a situation that lasts until April. Consequently, there has been
more sea-ice formation and, thus, brine production in this region during the
last 10 years compared to the 1980s and 1990s. When this sea-ice formation
is further away from the coast line the initial salinity is probably also
higher and thus also the resulting brine.
Silicate versus depth for data collected in 1996 (green) and from
section B in 2014 (red). The positions of the stations in 1996 are shown by
orange points in Fig. 1b.
Percentage of ice-free area in the region:
latitude 76 to 80∘ N, longitude 140 to 150∘ E (framed
yellow in Fig. 1b), for each month from 1980 to 2014, evaluated from the
passive microwave data of NSIDC (Cavalieri et al., 1996).
Salinity (a), silicate (b), and
SF6(c) as a function of distance from the bottom for all
stations deeper than 400 m in sections D and F. Series 1 to 11 represent the
bottom depth, increasing from 483 to 1120 m.
Indications of shelf plumes penetrating all the way to the deep basin are
seen in sections D and F where salinity, silicate, and SF6 levels
increases towards the bottom, a signature that generally fades away down the
shelf slope (Fig. 13). This is not observed in the more western sections and
is consistent with shelf plumes penetrating down into this eastern region. A
rough computation of the fraction of shelf water can be done as follows. With
an increase in SF6 partial pressure from the intermediate to the bottom
water of about 0.5 ppt, and the intermediate water and shelf water partial
pressures being 1.5 and 8 ppt, respectively, a little less than 10 % of
shelf water is needed. This is quite substantial but not unrealistic if
matched with the other properties. The shelf water silicate concentration
should be ∼ 25 µmol L-1 in order to achieve the
observed 2 µmol L-1 increase, and the shelf water salinity
would be 36 to get an increase of 0.15. These computations completely ignore
mixing and entrainment, but provide some indication that a shelf water
contribution to the deep water of the Chukchi Abyssal Plain is realistic. The
realism of the shelf source concentrations is supported by observations and
modelling. For instance, silicate concentrations in the range 40 to
60 µmol L-1 was observed on the western flank of Herald
Canyon in the summer of 2008, even if the salinity was well below 34
(Anderson et al., 2013). Windsor and Björk (2000) used a polynya model
driven by atmospheric forcing to compute ice, salt, and dense water production
in different regions of the Arctic Ocean over 39 winter seasons from 1959 to
1997. Two regions were east and west of the Wrangle Island where mean
salinities of 37.0 and 35.9 were produced, respectively, well within the
range needed.
Conclusions
We have showed that this region of the Arctic Ocean is much more dynamic and
variable than previously reported. Our data collected in the summer of 2014
are consistent with a shelf–basin exchange scenario as summarized in Fig. 14.
A boundary current of Atlantic Layer water follows the shelf break from
the west to the east, where some of the water crosses the Lomonosov Ridge
into the Makarov Basin. This boundary current follows the shelf break to the
Alpha Ridge where it turns towards north at its western flank. The water at
the corresponding depth in the Chukchi Abyssal Plain has a substantially
lower partial pressure of SF6, consistent with a more isolated
circulation in this region.
Summary of deduced circulation. Green arrow shows the runoff
spreading in the surface out north of the New Siberian Islands. The light
brown illustrates the export of nutrient-rich water from the shelf into the
deep basin at a salinity of around 33 and the dark brown interrupted line the
nutrient-rich water of a salinity around 34.5. The dark blue arrows in the
deep basin show the intermediate deep (500–1500 m depth) boundary currents.
The yellow dotted line illustrates the deep water plumes off the shelf break.
Map drawn using Ocean Data View (Schlitzer, 2017).
Surface water with substantial input of river water exits the shelf north of
the New Siberian Islands to follow the Lomonosov Ridge out into the central
basins. High nutrient water with salinity centred at 33 exits the East
Siberian Sea from its western end and contributes to the cold halocline of
the central Arctic Ocean. Compared to historic data the high nutrient water
is found outside the shelf break further to the west in 2014, which is
associated with a lower degree of ice cover during the summer north of the
New Siberian Islands. Where the Mendeleev Ridge connects to the shelf slope
a water body with salinity around 34.5, elevated nutrient concentrations and
low SF6 partial pressure hugs the shelf slope. Water of such property
is also found further to the west. As the source of the low SF6 partial
pressure most likely is in the Chukchi Abyssal Plain, at least an occasional
flow to the west follows, a conclusion that is supported by the surface
sediment biogenic silicate (BSi) content. In the eastern study region plumes
of high salinity, silicate, and SF6 levels flow off the shelf into the
deep basin.
Data are available upon request to the corresponding
author.
The authors declare that they have no conflict of
interest.
Acknowledgements
We thank the supporting crew and Master of I/B Oden as well as the support
of the Swedish Polar Secretariat.
This research was supported by grants from the Swedish Research Council
(contract 621-2013-5105); the Swedish Research Council Formas (project
reference 214-2008-1383, A.U. 214-2014-1165); the Swedish Knut and Alice
Wallenberg Foundation (KAW); the European Union FP7 project CarboChange
(under grant agreement no. 264879). I. Semiletov acknowledge support from the
Russian Government (grant no. 14.Z50.31.0012/03/19.2014), the Far Eastern
Branch of the Russian Academy of Sciences and the ICE-ARC-EU FP7 project.
The transient tracer measurements were supported by the Deutsche
Forschungsgemeinschaft in the framework of the “Antarctic Research with
comparative investigations in Arctic ice areas” priority program by a grant
to T. Tanhua and M. Hoppema; Carbon and transient tracers dynamics: a
bi-polar view on Southern Ocean eddies and the changing Arctic Ocean (TA 317 = 5, HO 4680 = 1).
Edited by: T. Tesi
Reviewed by: S. Nishino and two anonymous referees
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