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.

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): 1δ15Norδ18O=[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.

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): 2N*=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): 3Preg=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): 4NO3PW=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): 5NO3AW=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): 6fPW=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.

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.

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.

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).

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.

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.

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): 7fPW=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.

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.

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.

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.