The flow of Pacific water to the North Atlantic exerts a globally
significant control on nutrient balances between the two ocean basins and
strongly influences biological productivity in the northwest Atlantic.
Nutrient ratios of nitrate (NO3-) versus phosphate
(PO43-) have previously been used to complement salinity
characteristics in tracing the distribution of Pacific water in the North
Atlantic. We expand on this premise and demonstrate that the fraction of
Pacific water as determined by NO3- : PO43- ratios can be
quantitatively predicted from the isotopic composition of sub-euphotic
nitrate in the northwest Atlantic. Our linear model thus provides a
critically important framework for interpreting δ15N signatures
incorporated into both modern marine biomass and organic material in
historical and paleoceanographic archives along the northwest Atlantic
margin.
Introduction
Pacific water from the Bering Strait constitutes a major fraction of the
polar outflow to the northwest Atlantic Ocean (McLaughlin et al., 1996;
Jones et al., 2003; Aksenov et al., 2010). Besides redistributing heat and
freshwater (Tang et al., 2004; Carmack et al., 2016), it also plays a
critical role in the transport of nutrients between the two ocean basins
(Tremblay et al., 2015; Lehmann et al., 2019). Pacific water has relatively
high nutrient concentrations (Macdonald et al., 2010). These nutrients
support high productivity on the Bering and Chukchi shelves (Arrigo and van
Dijken, 2011), which in turn fuels high rates of sedimentary denitrification
both in the shelf regions and along the Bering continental slope (Devol et
al., 1997; Lehmann et al., 2005, 2007; Chang and Devol, 2009; Granger et
al., 2011; Brown et al., 2015). The resulting excess in silicate
(Si(OH)4) and phosphate (PO43-) (relative to nitrate
(NO3-)) is a significant source of these nutrients to the Atlantic
(Torres-Valdés et al., 2013). In particular, the excess PO43-
supports N2 fixation in the Atlantic, thereby helping to balance the
global oceanic nitrogen budget (Yamamoto-Kawai et al., 2006).
Pacific-derived nutrients also influence biological productivity along the
northwest Atlantic shelf complex. The NO3- deficit in Pacific
water sets an upper limit on productivity, which otherwise would be higher
in the presence of more NO3--enriched Atlantic water (Harrison and
Li, 2008). The Si(OH)4 and PO43- excess influences plankton
community composition (Harrison et al., 2013; Fragoso et al., 2017).
Interannual and decadal-scale variability in the circulation of Pacific
water into the northwest Atlantic may help to explain recent observed
changes in the magnitude and composition of primary productivity with
potential bottom-up effects on ecosystem functioning (Drinkwater et al.,
2003; Greene et al., 2013; Townsend et al., 2015).
Given its importance to downstream circulation, nutrient budgets, and
productivity, it is useful to track the distribution of Pacific water using
chemical tracers. Jones et al. (1998) characterized
NO3- : PO43- relationships for “pure” Pacific and
Atlantic endmember waters. They further demonstrated that the concentrations
of NO3- and PO43- in a water sample relative to the
endmember relationships may be used to quantify the contribution of Pacific
water (i.e., “fraction Pacific water”, or fPW). With this approach, the
spatial and depth distributions of fPW were used to map the flow of Pacific
water through the Arctic and North Atlantic oceans (Jones et al., 1998,
2003). The same approach has also been used to deconvolute fluxes of
freshwater originating from Pacific water from those of sea ice meltwater and
meteoric water (Yamamoto-Kawai et al., 2008; Sutherland et al., 2009;
Azetsu-Scott et al., 2012; Benetti et al., 2016). In another study,
time-series nutrient data were used to track fPW and thereby infer changes
in circulation patterns over a 30-year period in Disko Bay, Greenland
(M. O. Hanson et al., 2012). The use of NO3- : PO43- as a proxy
for Pacific water has, however, important limitations. For example, the
approach requires an assumption of constant stoichiometry associated with
the uptake and recycling of nutrients, which may not hold in all regions
(Michel et al., 2002; Mills et al., 2015). Moreover, sensitivity to
NO3- source and sink processes such as N2 fixation and
denitrification may lead to an under- or overestimation of fPW,
respectively. Finally, seawater NO3- : PO43- ratios are
not preserved in organisms or sedimentary archives, thus limiting their use
in establishing changing baselines in an ecological or paleoceanographic
context.
The nitrogen (15N : 14N) and oxygen (18O : 16O) isotope
ratios in NO3- (expressed as δ15NNO3 and δ18ONO3) represent a complementary tool to trace the distribution
and modification of Pacific water, possibly addressing shortcomings related
to the use of stoichiometric nutrient tracers. Coupled N and O isotope
ratios provide insights into the internal cycling of NO3-, as well
as into input and removal processes. The preferential reaction of the lighter
14N and 16O during both phytoplankton uptake and denitrification
results in an enrichment of δ15N and δ18O of the
dissolved NO3- pool with a ratio of ∼ 1 (Casciotti
et al., 2002; Granger et al., 2004, 2008; Sigman et al., 2005). Conversely,
the recycling or regeneration of NO3- via nitrification (the
oxidation of ammonium (NH4+) to nitrite (NO2-) and
NO3-) leads to a decoupling of the N and O isotopic signature of
NO3- (Sigman et al., 2005; Lehmann et al., 2005; Granger and
Wankel, 2016). The δ15N of newly nitrified NO3-
depends on the N isotopic composition of its source substrate and hence
mirrors the isotopic signature of the organic matter exported from the
surface. In contrast, the δ18O of newly nitrified
NO3- remains independent from its N source and approaches the
δ18O signature of seawater (δ18OH2O+1.1 ‰; Casciotti et al., 2008; Sigman et al., 2009;
Buchwald et al., 2012). The resulting NO3- isotope fingerprints of
particular water masses have led to their increasing use as unique water
mass tracers (e.g., Granger et al., 2018; Lehmann et al., 2018)
The goal of this paper is to establish the use of δ15NNO3
and δ18ONO3 as a new chemical oceanographic tracer for
tracking the distribution of Pacific water to the northwest Atlantic. We
present new data from Baffin Bay, the Davis Strait, and the Labrador Sea (Fig. 1), highlighting differences in NO3- isotopic ratios among the
different water masses found in those regions. We evaluate the preservation
of δ15NNO3 and δ18ONO3 signatures during
southward advection of Pacific water from the Arctic Archipelago to the
Labrador Shelf and present a linear relationship between δ15NNO3 and fPW for the northwest Atlantic margin. Lastly, we
entertain implications of our findings for regional isotope ecological and
paleoceanographic studies.
Materials and methodsSample collection and nutrient measurements
Seawater samples were collected opportunistically during four different
expeditions that sampled 25 stations along the NW Atlantic margin from the
mid-Labrador Shelf to northern Baffin Bay between the years 2005 and 2016
(Fig. 1). New and previously published data are presented. New data were
collected during (1) expedition MSM45 of the Maria S. Merian in August 2015 and (2) an
ArcticNet expedition (AMD-2016-002a) of the Canadian Coast Guard Ship (CCGS)
Amundsen from July through September 2016. Previously published data are from (3) expedition HUD-2005-016 of CCGS Hudson in June 2005 and (4) a GEOTRACES (GN02)
expedition aboard CCGS Amundsen in July and August 2015. Stations associated with
each expedition are indicated in the Fig. 1 station legend. Sample
collection and analytical protocols for the MSM45 and AMD-2016-002a
expeditions are given below. Protocols for the HUD-2005-016 expedition are
provided in Sherwood et al. (2011), and for the GEOTRACES expedition they are provided in
Lehmann et al. (2019). Samples were collected under ice-free conditions
during all expeditions. Station and bottle data are provided as a
supplementary data file.
Samples from the MSM45 and AMD-2016-002a expeditions were collected with a
rosette water sampler holding 24×10 L Niskin bottles mounted with a
conductivity–temperature–depth (CTD) profiler equipped with sensors for
dissolved oxygen and fluorescence. During the MSM45 expedition, samples for
nutrient and NO3- isotope analysis were collected from the Niskin
bottles into separate, triple-rinsed, 60 mL high-density polyethylene (HDPE)
bottles with pre-filtration through 0.45 µm
surfactant-free cellulose acetate (SFCA) membrane filters and stored at -20 ∘C. Concentrations of NO3-, NH4+, and
Si(OH)4 were measured post-cruise at Dalhousie University according to
standard protocols (Grasshoff, 1969) using a Bran+Luebbe autoanalyzer III. During the AMD-2016-002a expedition, samples for nutrient analysis were
pre-filtered (0.2 µm) into 15 mL acid-rinsed centrifuge tubes. The
samples were analyzed on board the ship for NO3-, NO2-,
PO43-, and Si(OH)4 concentrations using a Bran+Luebbe
autoanalyzer III following standard protocols (Grasshoff, 1969). Samples for
NO3- isotope analysis were collected into acid-cleaned and
triple-rinsed 60 mL HDPE bottles without pre-filtration and stored at -20 ∘C.
Map and inset detail of study region with World Ocean Atlas
(WOA18) climatological mean temperature at 100 m water depth (colour
shading), bathymetry (black contours), major surface currents (arrows), and
sampling stations. Abbreviations: West Greenland Current (WGC), Baffin
Island Current (BIC), Labrador Current (LC), Nares Strait (NS), Jones Sound
(JS), Lancaster Sound (LS), Fury and Hecla Strait (FHS), Hudson Strait
(HS), Hatton Basin (HBn), Saglek Bank (SB), Hamilton Bank (HB).
NO3- isotope analyses
Seawater samples were prepared for the measurement of dual N and O isotope
ratios in NO3- following the denitrifier method (Sigman et al.,
2001; Casciotti et al., 2002). This method quantitatively converts
NO3- present in the water samples to nitrous oxide (N2O) by
introducing cultured denitrifying bacteria (Pseudomonas chlororaphis f. sp. aureofaciens, ATCC no. 13985) that
lack N2O reductase activity. The resulting N2O gas is then
analyzed by isotope ratio mass spectrometry. Isotopic ratios are
reported in delta notation following Eq. (1):
δ15Norδ18O=[Rsample/Rstandard-1]×1000,
where R represents either 15N : 14N or 18O : 16O, the
standard is the N2 in the atmosphere (air) or the oxygen in Vienna
Standard Mean Ocean Water (VSMOW), and the units are reported as per mille
(‰) deviation from the standard ratios. Sample δ18ONO3 data were not corrected for the δ18O of
seawater because the latter varies minimally, from -2.2 ‰ to 0.2 ‰ (Lehmann et al., 2019).
Samples from the MSM45 expedition were analyzed at Dalhousie University
using a Thermo Scientific Delta V Plus isotope ratio mass spectrometer (IRMS) interfaced with a Thermo GasBench
inlet. Data were calibrated using seawater-based reference material USGS32
(δ15N =+180 ‰ vs. air, δ18O =+25.7 ‰ vs. VSMOW), USGS34 (δ15N =-1.8 ‰ vs. air, δ18O =-27.9 ‰ vs. VSMOW), and IAEA-N3 (δ15N =+4.7 ‰ vs. air, δ18O =+25.6 ‰ vs. VSMOW) (Böhlke et al., 2003; Gonfiantini,
1984). NO2- concentrations were below the detection limit of 0.2 µM, so no prior NO2- removal was performed. The blank size
constituted < 2 % of the overall sample size for the standard 20 nmol target N2O concentrations and < 5 % for the low-concentration samples from the biologically productive zone. Analytical
reproducibility based on replicate measurements averaged 0.2 ‰ for δ15N and 0.4 ‰
for δ18O.
Samples from the AMD-2016-002a expedition were analyzed at the University of
Basel using a Thermo Scientific Delta V Plus IRMS with a customized purge-and-trap system (modified after McIlvin and Casciotti, 2010, 2011).
NO3- samples were analyzed in parallel to isotope reference
material USGS34 and IAEA-N3. NO2- concentrations were < 0.36 µM, which is < 4 % of the corresponding
NO3- concentrations, so no prior NO2- removal was
performed. Blanks constituted < 1 % of the overall sample size.
Analytical reproducibility based on replicate measurements averaged 0.2 ‰ for δ15N and 0.3 ‰
for δ18O.
Definitions and calculations
The seawater potential temperature (θ) and potential density anomaly
referenced to surface pressure (σθ) were calculated from
CTD data using the “oce” package in the R computing platform (Kelley and
Richards, 2017). Water masses (Table 1) were operationally defined on the
basis of θ and/or σθ thresholds following
published conventions (Stramma et al., 2004; Fratantoni and Pickart, 2007;
Azetsu-Scott et al., 2012). Note that Halocline Water (HW) and Labrador
Shelf Water (LShW) are distinct water masses, despite overlapping θ
and σθ characteristics. The latter is formed in the
Hudson Strait area from mixing of HW with Irminger Water (IW) and Hudson Bay
outflow water (Sutcliffe et al., 1983; Straneo and Saucier, 2008). The base
of the biologically productive zone (zP) was determined from CTD
profiles of chlorophyll fluorescence as the shallowest depth below the
subsurface chlorophyll maximum where values < 0.1 mg m-3 were
encountered. Apparent oxygen utilization (AOU) was calculated from CTD in situ
dissolved oxygen profiles using equations in Weiss (1970). The N*
parameter, quantifying NO3- to PO43- imbalances relative
to Redfield stoichiometry, was calculated following Eq. (2):
N*=NO3--16PO43-+2.95,
where the constant 2.95 forces a global mean N* of zero (Gruber and
Sarmiento, 1997; Deutsch et al., 2001). Regenerated PO4 (Preg) was
calculated from AOU and the stoichiometric constant of Anderson and
Sarmiento (1994) following Eq. (3):
Preg=AOU/170.
The ratio of regenerated to measured PO43- is reported as
Preg/meas.
Fraction Pacific water (fPW) was calculated from N and P concentration data
in relation to N : P relationships for Pacific and Atlantic endmembers,
following Jones et al. (1998). For the purposes of data representativity,
accessibility, and propagation of error calculations, we derived new
equations for Atlantic and Pacific waters using public domain data from the
2019 version of the Global Ocean Data Analysis Project (GLODAP; Olsen et
al., 2019). The N : P relationships are sensitive to the choice of dissolved
inorganic nitrogen (DIN) species (NO3-, NO2-,
NH4+), particularly in highly productive shelf regions where a
significant fraction of the total dissolved N may not be fully nitrified
(Yamamoto-Kawai et al., 2008; Mills et al., 2015). The equations reported
here use only NO3- concentrations, as the inclusion of
NO2- and NH4+ was found to have a negligible impact on
fPW calculations (less than calculated uncertainties; see below).
Pacific endmember data were selected from a region encompassing the Canada
Basin of the Beaufort Sea (Fig. S1a). The data were filtered (bottom depths
> 500 m) to exclude shelf waters where NH4+ accounts
for a significant fraction of the DIN. Data were further filtered (S≤ 33.5; NO3-> 2.5 µM) to exclude waters of
Atlantic origin, as well as data from below a kink in the NO3--vs.-PO43- relationship where NO3- is exhausted before
PO43- (Yamamoto-Kawai et al., 2008). The resulting relationship
for Pacific water (PW) was calculated following Eq. (4):
NO3PW=14.07±0.09×PO4PW-11.53±0.15r2=0.85,n=4109.
This relationship is within the error of the one reported in Yamamoto-Kawai et
al. (2008), which was based on total DIN-vs.-PO43- data for a
region also encompassing the Chukchi Sea (Fig. S1b). This indicates that the
NH4+ that accumulates on the Bering and Chukchi shelves in summer
is largely nitrified by the time the Pacific-origin shelf waters reach the
Canada Basin and Amundsen Gulf (Brown et al., 2015; Granger et al., 2018). In
other words, the use of total DIN (instead of just NO3-) to define
the Pacific N : P relationship would have a negligible effect on the derived
Pacific endmember relationship in Eq. (4).
Atlantic endmember data were obtained from the Irminger Sea (Fig. S1a),
which, based on drifter trajectories, represents the source region for
waters entering the Labrador Sea via the Irminger Current (Cuny et al.,
2002; Jakobsen et al., 2003). Data from the region were filtered (S≥ 35; NO3-> 2.5 µM) to exclude polar waters
entering the area through the Fram Strait (Sutherland et al., 2009), and waters
affected by nutrient drawdown. The resulting relationship for Atlantic water
(AW) entering the Labrador Sea was calculated following Eq. (5):
NO3AW=15.54±0.10×PO4AW-0.26±0.10r2=0.89,n=2669.
For any given sample PO43- concentration, Eqs. (4) and (5)
define theoretical endmember NO3PW and NO3AW
concentrations, respectively. The fPW was then calculated from the sample
NO3- concentration in relation to NO3PW and
NO3AW, following Eq. (6):
fPW=NO3sample-NO3AW/NO3PW-NO3AW.
Negative values were considered devoid of Pacific water and were set
to zero. Analytical error in NO3- and PO43- measurements
averaged 1 % and 2 %, respectively. Errors propagated through Eqs. (4), (5), and (6) with statistical bootstrapping (n=10 000, with uniform
distributions) resulted in uncertainties of ±3 % (at the 95 %
confidence level) in fPW estimates.
Water mass definitions, depth ranges, and statistical summaries
(mean ± 1 SD (or absolute difference where n< 3) and
number of samples in parentheses) of physical and chemical properties by
water mass. HW: Halocline Water; BBW: Baffin Bay Water; LShW: Labrador Shelf
Water; IW: Irminger Water; LSW: Labrador Sea Water; NEADW: Northeast
Atlantic Deep Water; DSOW: Denmark Strait Overflow Water. All properties
were calculated for waters below the biologically productive zone (zp).
* It is not possible to calculate fPW for BBW. See Sect. 3.2.4 for
details. NA: not available.
ResultsHydrographic summary
Pacific water propagates as a halocline layer through the Canadian Arctic
Archipelago, then through Baffin Bay, and into the Labrador Sea via the Davis
and Hudson straits (Fig. 1; Tang et al., 2004; McLaughlin et al., 2004;
Steele et al., 2004). The 25 stations that were
sampled for this study are distributed along this net transport pathway of
Pacific water and are therefore ideally situated for investigating the
distribution of NO3- isotopic ratios with respect to fPW. The
stations are grouped into five hydrographic regimes on the basis of common
water column properties. A summary of hydrographic properties by regime
follows below. To help with visualization, the data are colour-coded by
hydrographic regime consistently throughout the subsequent figures. A
diagram of θ–S data is shown in Fig. 2, with delineations for the
different water masses. Depth profiles are shown for all stations in Fig. 3
and separately for each hydrographic regime in Figs. S1–S5.
Temperature–salinity diagram for individual stations. For each
profile, the 30 m depth level is indicated by open symbols; the bottom level
is indicated by shaded symbols. Data from < 30 m omitted for
clarity. Bold black lines indicate temperature and σT limits
for water masses discussed in text: HW, Halocline Water; LShW, Labrador
Shelf Water; BBW, Baffin Bay Water; IW, Irminger Water; LSW, Labrador Sea
Water; NEADW, Northeast Atlantic Deep Water; DSOW, Denmark Strait Overflow
Water.
The Baffin Bay regime was represented by three stations (BB2, BB3, CAA3).
Station BB2 was located inside the central Baffin gyre at a depth of 2300 m.
BB3 was located along the path of the Baffin Island Current at a depth of
1243 m. CAA3 was located at the southern side of Lancaster Sound at a depth
of 690 m. The three stations displayed similar θ and S profiles,
characteristic of Baffin Bay more broadly (Fig. 3; Tang et al., 2004). A
surface layer, formed by summer warming and melting extended down to 20–30 m. Below this, a layer of almost isothermal, cold (θ<-1.5 ∘C) water, increasing in salinity from < 32 to 34,
extended down to 200 m. This layer represents the Pacific-sourced HW, which
is formed in the Beaufort Sea and adjacent shelves and then modified by
regional winter cooling and sea ice formation in northern Baffin Bay and
along the northwestern Greenland coast (Bourke et al., 1989; Münchow et
al., 2015; Rysgaard et al., 2020). Below the HW, a warmer (θ
approaching +2 ∘C) and saltier (S> 34) layer
extended down to 700 m. This layer represents the diluted remnants of
Atlantic-sourced IW, often referred to as West Greenland Intermediate Water,
which flows northward via the West Greenland Current and spreads throughout
the entire Baffin Bay (Tang et al., 2004; Münchow et al., 2015). The
θ and S data over this depth interval at CAA3 were more variable,
reflecting the interleaving of different water masses by the complex tidal
currents in Lancaster Sound (Fig. S3; Prinsenberg and Hamilton, 2005). At
stations BB2 and BB3, the waters below 700 m form a distinct tail on the
θ–S diagram (Fig. 2). We refer to this as “Baffin Bay Water” (BBW;
θ< 2 ∘C; 27.5 >σθ≤ 27.8), which, for convenience, groups waters generally referred to as
Baffin Bay Deep Water for 1200 <z< 1800 m and Baffin Bay
Bottom Water for z> 1800 m (Tang et al., 2004), as well as the
shallower waters from 700 <z< 1200 m. The BBW is also
distinguished by the rapid increase in AOU with depth (Fig. 3c).
The northern Davis Strait regime was represented by four stations (177, 179,
BB1, ROV7). Station 177 was located within 2 km of coastal Baffin Island at
a depth of 376 m. Despite the coastal location of station 177, it is
hydrographically connected to more open water via a deep, northeast-trending
cross-shelf trough (Broughton Trough). Station 179 was located on the Baffin
shelf break at 186 m. Station BB1 was located on the northern flank of the
Davis Strait sill at a depth of 1042 m. Station ROV7 was located over the
Greenland slope (Disko Fan) at a depth of 932 m. Hydrographic profiles at
these four stations were similar to those of the Baffin Bay regime, with the
characteristic HW and IW layers (Fig. 3). A seemingly thicker surface layer
extending down to > 50 m at station 177 is a result of the later
sampling date (late September) than at the other three stations, which were
sampled in late July or early August. The HW layer was thicker at stations 177
and 179, which are located in the path of the Baffin Island Current, and
thinned out toward the more centrally located BB1 and ROV7, also evident from
the shallowing isopycnals (Fig. S4 σθ profiles; Tang et
al., 2004; Azetsu-Scott et al., 2012). Stations BB1 and ROV7 sampled the IW
(θ> 2 ∘C; S> 34.4) from 300–500 m and BBW below about 700 m.
Depth profiles for (a) potential temperature, (b) salinity, (c)
apparent oxygen utilization (AOU), (d) PO43-, (e) NO3-,
(f) N*, (g)δ15N of NO3-, and (h)δ18O of
NO3-. Points not connected by lines in (h) are suspected
analytical outliers.
The Labrador Shelf regime comprised seven stations (009, 018, 024, 030, 154,
147, 143). Station 009 was located over a > 900 m deep basin in
the main channel of the Hudson Strait but had a similar hydrographic profile
to the other stations on the Labrador Shelf. Stations 018, 024, and 030 were
located on an along-shelf transect, located at depths of 200, 534, and
535 m, respectively. Stations 154, 147, and 143 were located along an outer
cross-shelf transect of Hamilton Bank at depths of 202, 245, and 344 m,
respectively. A surface layer extended down to about 30 m at all stations,
underlain by the remnants of the HW, modified by tidal mixing and warming
southward of the Davis Strait (Fig. 3; Tang et al., 2004). This layer is often
called the “Cold Intermediate Layer” (Colbourne et al., 2016) but herein
is referred to as “Labrador Shelf Water” (LShW; θ< 2 ∘C, S< 34.2) for ease of reference. LShW extended
down to between 150–300 m and was underlain by IW where the bottom depth
exceeded 300 m. The influence of IW increased from west to east, as becomes
apparent from the cross-shelf increase in θ and S at stations 030,
154, 147, and 143 (Fig. S5; Fratantoni and Pickart, 2007).
The outer Hudson Strait regime comprised five stations (ROV1, ROV2, ROV3,
ROV5, ROV6) concentrated around an area seaward of the Hudson Strait, around
the shelf break (Fig. 1 inset). Stations ROV1 and ROV5 were located on the
sill of an outer-shelf bathymetric depression (Hatton Basin) at
an approximately 500 m water depth. Stations ROV2 and ROV3 were located along
the northern flank of Saglek Bank at 279 and 436 m, respectively. Station
ROV6 was located further north, at a 456 m depth, but had similar hydrography
to the other four stations (Fig. 3). The surface and bottom currents in
these areas are quite strong, up to 0.60 m s-1 at station ROV3, and
generally flowing NW to SE but with a strong tidal influence linked to the
macrotidal oscillation in Frobisher Bay (Zedel et al., unpublished bottom
current meter data from NE Saglek Bank). The surface layer extended down to
about 30 m at all stations. Nutrient data, specifically PO43-
concentrations and N* values, discussed below, clearly distinguished the
surface waters of this regime from the other regimes presented above. The
water mass structure was overall similar to that of the Labrador Shelf
regime but with a thinner, warmer, and saltier layer of LShW underlain by
warmer and saltier IW (Fig. S6).
Finally, the Labrador Basin regime comprised six stations (006, 013, 016,
033, LS2, K1) located in the deep waters of the Labrador Sea, at depths of
1280 m (station 013) to 3292 m (station 033). Hydrographic profiles at these
stations (Fig. 3) reflect the well-known water mass structure in the
Labrador Sea (e.g., Yashayaev and Loder, 2016). Doming of isopycnals leads to
the thinning and shoaling of the IW layer from the margins (e.g., station 013, IW, 70–500 m) to the center of the basin (e.g., station K1, IW, 30–150 m) (Fig. S7). Below the IW, a thick layer of Labrador Sea Water (LSW, 27.68 >σθ≤ 27.80, θ> 2 ∘C) extended down to 1500–2000 m, underlain by Northeast
Atlantic Deep Water (NEADW, 27.80 >σθ> 27.88) to 2400–2700 m and then Denmark Strait Overflow Water
(DSOW, σθ> 27.88).
NutrientsNutrient concentrations in the biologically productive zone
Sampling was conducted in the months of June–August, which follows the
spring bloom throughout most of the study region (Tremblay et al., 2006;
Frajka-Williams et al., 2010). Complete or near-complete utilization of
NO3-, PO43-, and Si(OH)4 was observed in the upper
30 m of the water column at all sites, with evidence of partial nutrient
utilization to < 120 m. Minima in NO3- in the surface
waters averaged < 1 µM and did not vary by hydrographic
region (Fig. 3e). Minima in PO43-, by contrast, exhibited a
striking bimodal distribution with respect to region, with concentrations
< 0.1 µM for most of the outer Hudson Strait and Labrador
Basin stations and > 0.4 µM for all of the Baffin Bay,
Davis Strait, and Labrador Shelf stations (Fig. 3d). Minima in Si(OH)4
exhibited a similar bimodality (< 1 µM and > 5 µM) with respect to hydrographic regions (Figs. S1–S5). Thus,
NO3- was relatively more limiting to primary production than
either PO43- or Si(OH)4 in the colder and fresher hydrographic
regions, as observed previously (Tremblay et al., 2006; Harrison and Li,
2008; Martin et al., 2010; Ferland et al., 2011; Fragoso et al., 2017).
Nutrient concentrations below the biologically productive zone
Nutrient concentrations generally stabilized below the biologically
productive zone (depth >zP), with the exception of BBW, in
which concentrations increased rapidly with depth (Fig. 3d, e). The elevated
concentrations result from in situ nutrient regeneration in Baffin Bay Deep Water
and Baffin Bay Bottom Water (Jones et al., 1984; Tremblay et al., 2002; Lehmann et al.,
2019). Baffin Bay is a 2300 m deep basin enclosed by < 700 m deep
sills. The enclosed bathymetry and permanent halocline restrict circulation,
thereby trapping particulate organic matter (POM) and remineralized
nutrients. More precisely, given the long residence time of the deep and
bottom waters (77–1450 years; Top et al., 1980; Wallace et al., 1985), high
fluxes of POM originating from the productive northern Baffin Bay (Klein et
al., 2002; Tremblay et al., 2002; Lalande et al., 2009) accumulate at depth.
The subsequent in situ remineralization of this sinking POM leads to the observed
increase in nutrients, seen as an increase in Preg/meas, and drawdown
of oxygen (increase in AOU) in the deep basin (Fig. 3c–e). While O2
concentrations remain too high to support denitrification in the water
column, dissimilatory NO3- consumption in the sediments is
supported by the low oxygen concentrations in the water above and acts as a
potential sink for dissolved NO3- in the lower water column
(Lehmann et al., 2019). Indeed, BBW had PO43- concentrations almost 2-fold higher (1.4 ± 0.3 µM) and Si(OH)4 concentrations 3-fold higher (41 ± 25 µM) than any of the other water masses (Table 1) but only
somewhat higher NO3- (19 ± 3 µM). For water masses
other than BBW, there were significant differences in NO3- and
PO43- but not Si(OH)4 (one-way ANOVA). HW and LShW had lower
NO3- concentrations (∼ 10 µM) than IW, LSW, and
NEADW (∼ 15 µM). The distribution of PO43- by
water mass was similar, except that the concentration in HW (0.98 µM)
was closer to that of IW, LSW, and NEADW (> 1 µM) than that of
LShW (0.88 µM).
NO3- vs. PO43- data for samples from all water
depths in this study, with lines representing empirically derived Atlantic
and Pacific water endmember N : P relationships. Endmember lines are enclosed
by 95 % confidence intervals. Red- to blue-coloured lines between Atlantic
and Pacific endmembers represent lines of constant fraction Pacific water
(fPW), in increments of 0.2. Also shown are dashed grey lines of constant
N*.
Nutrient ratios
NO3--to-PO43- stoichiometry is expressed in profiles of
N*, which were available for 19 of the 25 stations with paired
NO3- and PO43- concentration data (Fig. 3f). Positive N*
reflects a water mass history of excess fixed NO3-, e.g., by net
N2 fixation; negative N* reflects a NO3- deficit, relative to
the mean global ocean (Gruber and Sarmiento, 1997), induced by
denitrification in the broadest sense (i.e., including other modes of
suboxic DIN transformations to N such as anammox). N* signatures can be
imported from other ocean regions or can be generated within a given water
mass or region, depending on biogeochemical conditions. Positive N* occurred
throughout most of the outer Hudson Strait and Labrador Basin profiles. The
water masses LSW, NEADW, and DSOW (z> 200 m) all showed mean N*> 2 µM (Table 1), consistent with an Atlantic
origin (Gruber and Sarmiento, 1997; Jenkins et al., 2015). The deflections
to lower N* at z∼ 100 m (Fig. 3f) correspond to IW, which has
lower N* due to mixing at the shelf–slope front (Cuny et al., 2002;
Fratantoni and Pickart, 2007). Negative N* occurred through most of the
Baffin Bay, Davis Strait, and Labrador Shelf profiles (Fig. 3f). As noted
earlier, HW partially originates from the Bering and Chukchi shelf areas,
where sedimentary denitrification fuelled by high water column productivity
acts as a sink for dissolved NO3-. The resulting pronounced
minimum in N* (<-10 µM; Yamamoto-Kawai et al., 2008; Mills
et al., 2015) propagates via HW through the Canadian Arctic Archipelago and
into Baffin Bay (Carmack and McLaughlin, 2011; Tremblay et al., 2015). This
import of fixed-N-deficient waters explains the lowest N* values in HW
(Table 1), with minima < 6 µM at stations CAA3, BB3, BB2,
BB1, 177, and 179 (Fig. 3f). BBW had the next most negative N*, likely due
to the upward propagation of partially denitrified bottom water nutrients
(Tremblay et al., 2002; Lehmann et al., 2019). LShW had the most variable N*
signatures, resulting from the mixing of HW and IW. The effect of this
mixing is clearly evident in the cross-shelf increase in N* on the Labrador
Shelf, from <-3 µM at station 154 to > 2 µM at station 143 (Fig. S5).
Fraction Pacific water
A cross plot of NO3- vs. PO43- concentrations helps
to visualize the fraction of Pacific water among individual water samples
relative to lines representing pure Atlantic and Pacific waters (Fig. 4).
Note that the Atlantic line (fPW = 0) coincides with a line of constant N*= 2 µM (Fig. 4), which is also the average N* value for North
Atlantic intermediate waters (Gruber and Sarmiento, 1997). The Pacific line
(fPW = 1) falls between the lines where N* is -10 to -12 µM. Thus,
depth profiles of fPW mirror those of N* (Figs. S3–S7). Data from the
Labrador Basin regime fall on or close to the Atlantic line because the
water masses at these stations (IW, LSW, NEADW, DSOW) are mostly
Atlantic-sourced. Data from the other regimes plot increasingly toward the
Pacific line in the order of the Hudson Strait, the Labrador Shelf, the Davis Strait,
and Baffin Bay. Maxima in fPW (> 0.6) were found within the core
of HW sampled at Lancaster Sound (CAA3) and along the path of the Baffin
Island Current (BB2, BB3, BB1, 177, 179) (Figs. S3, S4).
Water samples with PO43-> 1.25 µM and
NO3-> 17.5 µM correspond to BBW (Fig. 4). The
data fall along a N : P trajectory with a slope of 9.6 ± 0.3. This slope
is considerably lower than the slopes of either Atlantic or Pacific
endmember waters. It arises from in situ remineralization of POM, as indicated by
Preg/meas values > 0.5 (Figs. S3, S4), with a partial loss
of NO3- via sedimentary denitrification (Lehmann et al., 2019).
The denitrification generates N* values as low as -4.3 µM. Thus, the
process that leads to low N* in BBW is separate and distinct from the
processes that generate low N* in HW. As a result, it is not possible to
calculate fPW for BBW, because remineralization and denitrification
overprint the preformed N : P signatures (Jones et al., 2003).
Another complication with fPW estimates, as noted in Sect. 2.3, is that
elevated concentrations of NO2- and NH4+ may alter
apparent N : P ratios with respect to the derived endmember relationships.
Within the overall study region, NO2- and NH4+
concentrations below the euphotic zone are generally < 1 µM
(Harrison and Li, 2008; Martin et al., 2010; Azetsu-Scott et al., 2012).
Where measured in the present study, NO2- concentrations were
< 0.36 µM and NH4+ concentrations were < 1 µM, except for six samples from the Labrador Basin regime with
NH4+ concentrations up to 2.5 µM (Figs. S3–S7). Thus, with
the exception of those few samples, the overall low NO2- and
NH4+ concentrations should have little impact on fPW estimates.
δ18ONO3 vs. δ15NNO3 by station and
NO3- concentration. The dense cluster of data centered at bottom
left represents deep waters (>zp). Diagonal lines represent
1:1 isotopic fraction of 18O and 15N. Arrows denote isotopic
fractionation associated with NO3- assimilation in the
biologically productive zone and differences in water mass N-cycling
histories.
NO3- isotope ratio variability
Isotope ratios of NO3- were measured at all 25 stations for
δ15NNO3 and all but the three stations from the
HUD-2005-016 expedition for δ18ONO3. Patterns of isotopic
variability are presented separately for waters in and below the base of the
biologically productive zone (zP) in the following sub-sections.
Isotope ratios of NO3- in the biologically productive zone
For waters above zP, δ15NNO3 increased from values of
around 5 ‰–6 ‰ to maxima of 12 ‰
toward the surface (Fig. 3g). The δ18ONO3 similarly
increased from < 2 ‰–11 ‰
(Fig. 3h). In a cross plot of δ18ONO3 vs. δ15NNO3, these isotopic enrichments extend approximately along
lines of 1:1 (Fig. 5). The increase in isotopic ratios coincides with a
decrease in NO3- concentrations and an increase in chlorophyll,
as interpreted from fluorescence profiles (Figs. S1–S5). Together, these
patterns are consistent with coupled (identical) fractionation of 15N
and 18O during NO3- assimilation (Granger et al., 2004;
Sigman et al., 2005).
To further demonstrate the effect of NO3- assimilation on isotopic
ratios, δ15NNO3 and δ18ONO3 are plotted
against the natural logarithm of NO3- concentrations, where the
“kinks” in the relationships represent the base of the NO3-
assimilation zone (Fig. 6). To the left of the kinks, both δ15NNO3 and δ18ONO3 increase with decreasing
NO3-, again consistent with coupled fractionation of 15N and
18O during NO3- assimilation. Moreover, assuming a mainly
vertical supply of nutrients to the euphotic zone, the isotopic composition
of the NO3- used in assimilation may be approximated by the minima
in δ15N and δ18O at the kinks (Fig. 6; Rafter and
Sigman, 2016; Peters et al., 2018). In this respect, the δ15N
of the assimilated NO3- increases from its lowest values at the
Labrador Basin stations to its highest values at the Baffin Bay stations. The
minima in δ18ONO3 data show the opposite trend, with the
lowest values at the Baffin Bay stations and highest values in the Labrador
Basin.
Isotope ratios of NO3- below the biologically productive
zone
Below zP, δ15NNO3 ranged from 4.1 ‰–6.5 ‰ (Fig. 3) and varied significantly by water mass (one-way
ANOVA; Table 1). The Atlantic-derived water masses (IW, LSW, NEADW, DSOW)
sampled in the Hudson Strait and Labrador Basin regimes had the lowest mean
δ15NNO3 (4.8 ± 0.3‰). This
value is identical to the δ15NNO3 of North Atlantic
intermediate-depth waters; it represents the basin-scale N isotopic mass
balance between relatively 15N-depleted NO3- in Atlantic subtropical
thermocline water and Mediterranean Overflow Water and relatively
15N-enriched NO3- in Antarctic Intermediate Water (Marconi et
al., 2015). The Pacific-influenced HW, as well as BBW sampled in Baffin Bay
and the Davis Strait, displayed the highest mean δ15NNO3 (> 6 ‰). The elevated δ15NNO3 in HW reflects its predominant origin in the western
Arctic. At the entrance to the western Arctic, Pacific-origin NO3-
propagating onto the Bering Shelf has an already-high δ15NNO3 (6.3 ‰; Lehmann et al., 2005). As
Pacific waters flow across the productive Bering and Chukchi shelves,
NO3- becomes further isotopically enriched due to benthic coupled
nitrification–denitrification (CPND), which results in the removal of
isotopically light NH4+ from the system and the efflux of heavy
NH4+ into the overlying water column (Granger et al., 2011; Brown
et al., 2015). Subsequent water column nitrification leads to the
characteristically high δ15NNO3 signature of the western
Arctic upper halocline (∼ 8.0 ‰; Brown et
al., 2015; Granger et al., 2018; Fripiat et al., 2018). The δ15NNO3 signature in BBW, on the other hand, is consistent with
in situ remineralization in deep Baffin Bay, as indicated by high AOU and nutrient
concentrations. The high δ15NNO3 (> 7.0 ‰; Fig. 3g) indicates that the POM exported to the deep
Baffin Bay is largely fuelled by Pacific-derived nutrients in northern Baffin
Bay (Lehmann et al., 2019), given that the N isotopic composition of newly
nitrified NO3- largely reflects the signature of its source
substrate. The LShW sampled on the Labrador Shelf exhibited intermediate and
more variable δ15NNO3 signatures (5.4 ± 0.7 ‰), which, as with the corresponding N* data, was
consistent with mixing of HW and IW across the Labrador Shelf.
Values of δ18ONO3 also varied significantly by water mass
(one-way ANOVA) but as a mirror image of δ15NNO3 (Table 1).
The LSW, NEADW, and DSOW exhibited higher δ18ONO3
(> 1.8 ‰); HW and BBW had lower values
(> 0.3 ‰). The low δ18ONO3
in HW is within the range of values previously reported for the western
Arctic upper halocline layer (Brown et al., 2015; Granger et al., 2018;
Fripiat et al., 2018), where values close to 0 ‰ are
indicative of the highly remineralized NO3- pool, as nitrification
introduces a low δ18O close to a value of ambient seawater
(+1.1 ‰; Casciotti et al., 2008; Sigman et al., 2009;
Buchwald et al., 2012). Low δ18ONO3 values associated with
BBW similarly reflect the high proportion of remineralized nutrients in the deep
Baffin Bay (Lehmann et al., 2019). The significantly higher δ18ONO3 in the Labrador Sea subsurface layer reflects the remote
signal of partial assimilation in the Southern Ocean, as well as a higher
δ18O of water oxygen atoms that are incorporated during
remineralization in transit in the Atlantic versus the Arctic (Marconi et
al., 2015; Granger et al., 2018). The differential NO3- isotope
tagging of the various sub-euphotic water masses, which is a function of
their different origin and N-cycling history, holds great potential to trace
the distribution of these water masses in the northwest Atlantic and thus to
assess the contribution from Pacific sources.
To explore the biogeochemical drivers of δ15NNO3 and
δ18ONO3 in the different water masses further, a
correlation matrix of physical and chemical variables was constructed (Fig. S8). The strongest covariates of δ15NNO3 were fPW (r=0.89) and N* (r=-0.86), followed by θ (r=-0.75) and then
variables associated with diatom and POM remineralization in BBW:
Si(OH)4 (r=0.65), AOU (r=0.64), and PO43- (0.61).
Upon recalculating the correlation matrix without BBW (see below), the correlations
with fPW (r=0.91) and N* (r=-0.89) became even stronger, followed by
salinity (r=-0.82), θ (r=-0.81), and NO3- (r=-0.65). All five of these parameters exhibit multicollinearity; that is,
waters with high fPW also have low N* and are colder, are fresher, and have
less preformed NO3- than waters with low fPW. The same parameters
were also correlated with δ18ONO3 but opposite in sign
(Fig. S9).
(a)δ15NNO3 and (b)δ18ONO3 plotted against
the natural logarithm of NO3- concentrations. The kink
represents the base of the NO3- assimilation zone. For clarity,
only some of the station data are connected by lines.
In a plot of δ15NNO3 versus N*, the regression line
through the main group of data is highly significant (p≪ 0.001) with an r2 value of 0.78 (Fig. 7a). Note that BBW data plot
above and to the right of the rest of the data. We hypothesize that this
shift arises from remineralization of PO43- and NO3-,
followed by loss of the NO3- via sedimentary denitrification
(Lehmann et al., 2019). A conceptual model of this two-step process is shown
in the Fig. 7a inset. The source of preformed nutrients in the deep Baffin
Bay is still debated (Tang et al., 2004), but, assuming a dominantly
Atlantic source (Azetsu-Scott et al., 2012), the preformed nutrients would
plot near the other Atlantic waters with N* values > 0 µM
and δ15NNO3∼+5 ‰
(e.g., Marconi et al., 2015; Granger et al., 2018; Fripiat et al., 2018).
Under a simplifying assumption of near-Redfield stoichiometry, the N* would
remain unchanged during remineralization. (We note, however, that lower-than-Redfield N : P uptake has been documented in Baffin Bay – Harrison et al.,
1982 – which would shift the N* to lower values.) The POM originates in the
overlying 15NNO3-enriched HW, which would generate remineralized
NO3- with relatively high δ15N in BBW. Subsequent
sedimentary denitrification would shift the N* to lower values because the
process acts as a sink for NO3- but not for PO43-. This
process has a negligible effect on water column δ15NNO3, as
the NO3- is completely reduced in the sediments such that there is
effectively no isotopically modified NO3- to diffuse back into the
overlying water (Brandes and Devol, 1997; Lehmann et al., 2005, 2007). This
again highlights that the processes affecting N* and δ15NNO3 in BBW are separate and distinct from those influencing
HW or the other water masses (Lehmann et al., 2019).
The relationship between δ15NNO3 and fPW is also highly
significant (p<< 0.001) with an r2 of 0.80 (Fig. 7b).
The intercept, corresponding to 100 % Atlantic water, is 4.8 ± 0.04 ‰. This value coincides exactly with previous estimates
of the mean δ15NNO3 (4.8 ‰) in North
Atlantic intermediate and deep waters (Marconi et al., 2015). This is not
surprising given that samples of 100 % Atlantic water are represented in the data distribution
(Fig. 4). The regression also predicts the δ15NNO3 for 100 % Pacific water at 8.3 ± 0.2 ‰. This value
likewise matches with previous measurements of the Pacific halocline water
measured in the eastern Beaufort Sea, downstream of the centers of CPND
(δ15NNO3=8.0± 0.1 ‰;
Brown et al., 2015; Granger et al., 2018). Such accurate prediction of the
Pacific endmember is remarkable, considering the degree of extrapolation
beyond the limit (fPW ≤ 0.6) of sample data.
(a)δ15NNO3 vs. N*. Regression line excludes BBW
because these waters are affected by remineralization and denitrification as
indicated by the inset conceptual schematic. See text for explanation. (b)δ15NNO3 vs. fPW, with BBW excluded because it is not possible to
calculate fPW when NO3- and PO43- ratios are not
conserved. Data in both plots are for depths >zp to avoid
the effects of NO3- assimilation. Regression lines in both plots
are bounded by 95 % confidence intervals.
Relationships for δ18ONO3 vs. N* and δ18ONO3 vs. fPW (Fig. S9) were also highly significant (p≪ 0.001), although weaker, with r2 values of 0.38 and
0.35, respectively. The δ18ONO3 vs. fPW relationship
predicts δ18ONO3= 1.9 ± 0.1 ‰ for 100 % Atlantic water and δ18ONO3=-0.4± 0.3 ‰ for 100 %
Pacific water, close to previous direct measurements of Atlantic and Pacific
water in their respective source regions (Marconi et al., 2015; Brown et
al., 2015; Granger et al., 2018).
DiscussionPreservation of NO3- isotope signatures
One of the objectives of this paper is to assess the preservation of
NO3- isotope signatures during transit of Pacific water from the
Canadian Arctic Archipelago southward into the northwest Atlantic. Accurate
prediction of δ15NNO3 and δ18ONO3 in
Atlantic and Pacific source waters based on the relationships in Figs. 7b
and S9b is perhaps the strongest indication that the signatures are well
preserved. This preservation is likely facilitated by the oxic conditions in
all water masses, as well as by the extreme vertical density gradient, which
isolates the HW from vertical mixing as it propagates downstream from the
Arctic (Tremblay et al., 2015). We also consider that N2 fixation would
act to increase N* while decreasing δ15NNO3, and while
there is some evidence of diazotrophy in northern waters (Blais et al.,
2012; Sipler et al., 2017; Harding et al., 2018), reported rates are small to
negligible and are therefore unlikely to impact water column δ15NNO3 signatures, as suggested by the data in Fig. 7a.
Prediction of fraction Pacific water from δ15NNO3
Another objective of this study is to evaluate the use of δ15NNO3, specifically, as a water mass tracer because it is
potentially reflected in the δ15N of living and detrital
biomass. Based on the preceding results, sub-euphotic zone δ15NNO3 may be considered the product of two-endmember mixing
of Atlantic and Pacific water (Fig. 7b). Thus, by inversion of the linear
regression in Fig. 7b, fPW may be derived. To achieve normality of model
residuals, it was necessary to remove data where fPW = 0, as well as three
outliers identified in quantile–quantile plots. The resulting final model
follows Eq. (7):
fPW=0.23±0.02×δ15NNO3-1.06±0.08,r2=0.83,p≪0.001,n=46.
The 95 % confidence intervals of fPW predictions range from 0.02 to 0.1,
which is roughly equal to the error associated with estimating fPW from
NO3- vs. PO43- data (Sect. 3.3; see also Jones et al.,
2003). In other words, δ15NNO3 may be used to estimate fPW
about as well as nutrient data. Of course, it is less labour intensive to
analyze and use N : P concentrations to estimate fPW, but there are scenarios
in which a proxy δ15NNO3 approach is complementary or even
essential. For example, δ15NNO3 of a water mass integrates
the overall history of nutrient cycling, thereby smoothing out short-lived
contingencies inherent to seawater chemistry data. N : P concentration data
alone also cannot identify and/or distinguish between the various processes
(i.e., remineralization, nitrification, denitrification, diazotrophy)
affecting nutrient concentrations and stoichiometric ratios. As a case in
point, the two-step process of remineralization followed by sedimentary
denitrification in BBW would not be obvious without paired N : P and δ15NNO3 data (Fig. 7a) and invites further investigations to help
clarify their proportional effects on bottom water N : P ratios (Lehmann et
al., 2019). We also suggest that isotope data can be used to screen samples
that deviate from a two-endmember mixing model for the calculation of fPW.
Finally, the relationship in Eq. (7) provides, for the first time, a
coherent framework for interpreting δ15N signatures
incorporated into living and paleo-organic materials in the hydrographically
complex northwest Atlantic marine ecosystem.
Incorporation of δ15NNO3 into baseline δ15N for food web and paleoceanographic studies
In isotope ecology and paleoceanography contexts, the term “baseline”
δ15N usually refers to the δ15N of primary
producer (phytoplankton) biomass. This baseline δ15N signature
is propagated to organisms higher up in the food web, overprinted by trophic
fractionation, which is often assumed to be about +3.4 ‰ per trophic level but is in fact widely variable
(Minagawa and Wada, 1984; Vander Zanden and Rasmussen, 2001). The baseline
signature may also be altered by bacterial degradation during sinking and
sedimentation of particulate organic material (Lehmann et al., 2002;
Robinson et al., 2012). In either instance, it is critical to know, or be
able to approximate, baseline δ15N in order to interpret the
environmental significance of δ15N variability recorded in
organisms or sediments.
Phytoplankton fractionate against the heavier isotopes of N and O during
growth on NO3-. Thus, under open-system conditions, the
δ15N of the phytoplankton will be lower than that of the
NO3-. The isotopic fractionation varies from about 2 ‰–10 ‰, depending on phytoplankton species and growth
conditions (Needoba et al., 2003). Under the semi-closed conditions of
NO3- drawdown in the ocean euphotic zone, the δ15N of both
NO3- and phytoplankton increase, following Rayleigh
fractionation kinetics (Fig. 6). If the NO3- is exhausted, the
δ15N of phytoplankton will, by isotope mass balance, converge
on that of the original, unassimilated NO3-. Hence, baseline
δ15N reflects the combined influences of δ15NNO3 and the degree of NO3- utilization (Altabet et
al., 1999; Trull et al., 2008). It is difficult to distinguish between these
influences unless δ15N is measured on paired samples of
phytoplankton and NO3-. This has not been performed anywhere in our
study region, apart from in studies located in more inshore areas (Ostrom et
al., 1997). Nevertheless, it is possible to make general inferences about
nutrient drawdown and its effect on baseline δ15N from a
consideration of regional nutrient–plankton bloom dynamics. In the Labrador
Sea and Baffin Bay, light is the principal limiting factor to phytoplankton
growth for most of the year; however, during the peak summer growth period,
NO3- becomes co-limiting or limiting as concentrations within the
mixed layer are depleted (Harrison and Li, 2008). This applies even in the
more light limited Arctic, where productivity is tightly coupled to
NO3- availability (Tremblay et al., 2006; Martin et al., 2010).
Therefore, the δ15N of the accumulated phytoplankton biomass
should approach that of the pre-assimilated NO3-, as identified by
the kink in Fig. 6. This was confirmed in a study of spring bloom
dynamics in the North Water Polynya in northern Baffin Bay, where the
δ15N of phytoplankton converged on the δ15NNO3 of Arctic halocline water (8.3 ‰) as
the fraction of unassimilated NO3- was drawn down to < 10 % of the pre-bloom concentrations (Tremblay et al., 2006). Considering
that northern Baffin Bay is located at a latitude of maximum light
limitation, we would predict that the patterns observed there also apply to
Pacific-influenced waters of the more southerly Baffin Bay and continental
shelves of eastern Canada, except perhaps to inshore and upwelling regions,
where NO3- would be less limiting. For Atlantic-influenced waters,
NO3- is already relatively less limiting than P and Si (Fig. 3).
Under these conditions, phytoplankton will be more likely to fractionate
against 15N (becoming “lighter”), thereby amplifying the existing
differences in δ15NNO3 between the Atlantic- and Pacific-derived water masses (Fig. 6). Additional studies are needed to determine
the effective fractionation, if any, over seasonal and longer timescales.
Implications for isotope ecology
Results presented here may help to explain previously documented spatial
variability in organism δ15N in the northwest Atlantic and
Baffin Bay regions. For example, Sherwood and Rose (2005) examined bulk
δ15N in invertebrates and fish in waters off Newfoundland and
Labrador. Organism δ15N within each feeding guild was
consistently higher, by up to 2.7 ‰, at coastal sites
compared to shelf break sites. Part of this offset may be explained by the
cross-shelf gradient in fPW, which increases from < 0.02 at the
shelf break to > 0.5 at the coast (Pepin et al., 2013; see also
the fPW and δ15NNO3 profiles for stations 143, 147, and
154 – Fig. S5), and corresponds to an increase of 2.3 ‰
in δ15NNO3 based on Eq. (7). Similarly, in studies located
off western Greenland, J. Hansen et al. (2012) and Hedeholm et al. (2012)
reported that δ15N in primary consumers increased by 2 ‰–3 ‰ over a latitudinal gradient from 60–72∘ N.
Subsurface fPW off southern Greenland is essentially zero (Sutherland et
al., 2009; Azetsu-Scott et al., 2012) and increases northward as IW and
West Greenland Shelf Water mixes with HW, reaching values > 0.4
based on sparse nearby data (e.g., station ROV7 profile, Fig. S4; J. Hansen et
al., 2012). The corresponding increase in δ15NNO3 is
> 2 ‰. This suggests that, in the examples
above, the spatial variability in organism δ15N may be
attributed largely to the differential water mass partitioning, rather than
to spatial variations in the degree of NO3- utilization directly at
the respective study sites. Finally, Sherwood et al. (2008) examined the
bulk δ15N in the tissues of deep-sea corals collected along the
continental slope from the Hudson Strait (62∘ N) to the southwest
Grand Banks of Newfoundland (43∘ N). They found no overall change in
δ15N with respect to latitude, but this is consistent with the
minimal latitudinal change in fPW (< 0.1) along the path of the
slope component of the Labrador Current (Jones et al., 2003). Overall, these
examples reiterate the fundamental importance of accounting for variability
in baseline δ15N in isotope ecology studies (e.g., de la Vega
et al., 2021). It is not always feasible to measure δ15N in
NO3- or primary producers directly; thus we suggest that baseline
δ15N for the Canadian Arctic and northwest Atlantic region may be
approximated, to a first degree, from nutrient concentrations and either of
the N* or fPW relationships presented in Fig. 7.
Implications for paleoceanography
Our results also have important implications for regional paleoceanographic
interpretations of δ15N as recorded in sedimentary organic
matter and in long-lived biological archives. With respect to sediments,
δ15N is confounded by site-to-site differences in sedimentation
rates and diagenetic effects (Robinson et al., 2012). Nevertheless, known
spatial patterns track the expected distribution of fPW, with lower values
of 4 ‰–6 ‰ in the central Labrador Sea and Southwest
Greenland margin and higher values of 6 ‰–9 ‰ on the
Labrador Shelf and in northern Baffin Bay (Muzuka and Hillaire-Marcel, 2000;
Cormier et al., 2016; Kienast et al., 2020; Limoges et al., 2020). Thus, by
extension, downcore trends in δ15N should reflect
advection-related temporal changes in fPW. Based on arguments in Sect. 3.6, this advection influence is likely to exceed the influence of surface
water NO3- utilization, particularly where NO3- is
limiting. This may help to explain why downcore variations in δ15N are positively correlated with other biomarker and
micropaleontological proxies for Arctic throughflow to Baffin Bay (Cormier et
al., 2016; Limoges et al., 2020), confirming the potential of sedimentary
δ15N to quantitatively reconstruct changes in fPW in the past,
at least in areas where local changes in nutrient utilization did not play a
greater role. This also applies to records of δ15N recorded in
biological archives such as deep-sea corals, which have been shown to track
changes in the southward advection of the Labrador Current over the 20th century (Sherwood et al., 2011). We note that, as the N-cycling regimes in
the source region and/or in the North Atlantic may have shifted in the past,
long-term changes in downcore or archival δ15N may also be
influenced by variability in endmember δ15NNO3 signatures
(i.e., NO3- “inventory-altering” effects; Galbraith et al., 2013),
particularly for the Pacific water endmember which is sensitive to primary
productivity via sedimentary CPND on the western Arctic shelves. Thus, long-term variability in δ15N should be carefully interpreted in the
context of all three influences – nutrient utilization, advection, and
changes to endmember δ15NNO3 signatures – together with
other lines of paleoceanographic evidence.
Conclusions
The flow of Pacific water through the Canadian Arctic Archipelago and into
the northwest Atlantic plays a key role in global thermohaline circulation
and biogeochemical cycling. The isotopic composition of NO3-
presents a new way to track this influence, expanding on the existing N : P
stoichiometry approach. Isotopically distinct Pacific water (δ15NNO3=8.3 ‰; δ18ONO3= 0 ‰) travels as a subsurface halocline layer
through the Canadian Arctic Archipelago and onward to Labrador Shelf, with
little apparent alteration other than mixing with Atlantic water (δ15NNO3=4.8 ‰; δ18ONO3= 2.0 ‰). The resulting two-endmember mixing of
Pacific and Atlantic water is described by a new empirical relationship that
may be used to estimate the fraction of Pacific water from δ15NNO3. The deep waters of Baffin Bay are distinctly different,
with nutrient inventories showing an imprint of both in situ remineralization and
sedimentary denitrification. These deep waters are isolated below 500 m and
therefore do not influence baseline δ15N incorporated into
primary producer biomass. Rather, baseline δ15N throughout the
Labrador–Baffin region should primarily reflect the fraction of Pacific
water, particularly where NO3- is the limiting nutrient. Overall,
these results provide a new framework for interpreting spatial and temporal
patterns of δ15N in isotope ecology and paleoceanography
contexts. In particular they highlight the potential of δ15N
recorded in sedimentary and organic paleo-archives to quantitatively
reconstruct changes in the fraction of Pacific water in the past.
Code availability
R code used for data processing is available upon request from the corresponding author.
Data availability
Data presented in this article may be accessed via the following link:
10.5281/zenodo.5129246 (Sherwood et al., 2021).
The supplement related to this article is available online at: https://doi.org/10.5194/bg-18-4491-2021-supplement.
Author contributions
OAS, SHD and MK conceptualized the research with input from all authors.
OAS, SHD, and MK collected the samples. SHD and NL carried out nitrate isotope
analyses. OAS, SHD, NL, CB, and MK analyzed the data. OAS, SHD, and NL prepared
the manuscript with contributions from all authors.
Competing interests
The authors declare that they have no conflict of interest.
Disclaimer
Publisher’s note: Copernicus Publications remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
Acknowledgements
We thank the captain, crew, and science staff of R/V Maria S. Merian MSM45 and CCGS
Amundsen AMD-2016 expeditions. We also thank ArcticNet (a Canadian Network of Centres
of Excellence) and Amundsen Science for their in-kind contributions to
expedition logistics and scientific equipment. McKenzie Mandich (Dalhousie
University) analyzed nutrient concentrations for the MSM45 samples.
Jean-Éric Tremblay facilitated sample collection and nutrient analysis
during the AMD-2016 expedition. Thomas Kuhn (University of Basel) analyzed
nitrate isotopes for the AMD-2016 samples. Claude Hillaire-Marcel provided
input on earlier drafts of the manuscript.
Financial support
Funding for this project was provided by the Canadian Foundation for Innovation and by the Natural Sciences and Engineering Research Council (NSERC) of Canada through a Climate Change and Atmospheric Research grant to Paul Myers (RGPCC433898); a Ship Time Allocations Committee grant to Evan N. Edinger (515528-18); a Strategic Partnerships grant to Markus Kienast and Owen A. Sherwood (521427-18); and Discovery grants to Owen A. Sherwood (RGPIN-2018-05590) Carolyn Buchwald (RGPIN-2018-05568), Evan N. Edinger (RGPIN-2014-04826), Markus Kienast (RGPIN-2016-04885), and Claude Hillaire-Marcel (RGPIN-2018-05031).
Review statement
This paper was edited by Marcel van der Meer and reviewed by two anonymous referees.
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