Using Radiocarbon Measurements of Dissolved Inorganic Carbon to Determine a Revised Residence Time for Deep Baffin Bay

The Canadian Arctic is warming at three times the rate of the rest of the planet and the effects of climate change on the Arctic marine carbon cycle remains unconstrained. Baffin Bay is a semi-enclosed, Arctic basin that connects the Arctic Ocean to the north to the Labrador Sea to the south. While the physical oceanography of surface Baffin Bay is well characterized, less is known about deep water formation mechanisms within the Basin. Only a few residence-time estimates for Baffin Bay Deep Water (BBDW) exist and range from 20 to 1450 years. Better residence time estimates are needed to understand the oceanographic significance of Baffin Bay. Here we report stable carbon (δ¹³C) and radiocarbon (Δ¹⁴C) values of dissolved inorganic carbon (DIC) collected aboard the CCGS Amundsen in 2019. DIC δ¹³C and Δ¹⁴C values between ranged between -0.7‰ to +1.9‰ and -90.0‰ to +29.8‰, respectively. Surface DIC δ¹³C values were between +0.7‰ to +1.9‰, while deep (>100m) values were 0.0 to -0.7‰. Surface DIC Δ¹⁴C values ranged between -5.4‰ to +22.9‰, while deep DIC (>1400m) DIC Δ¹⁴C averaged -82.2 ± 8.5‰ (n = 9). To constrain natural DIC Δ¹⁴C values, we quantified the amount of atmospheric “bomb” ¹⁴C in DIC (Δ¹⁴C_bomb; using the potential alkalinity method; P_alk) and anthropogenic DIC (DIC_anth; using the ΔC^* method). Both proxies indicate an absence of Δ¹⁴C_bomb and DIC_anth below 1000m. Using two previously proposed deep water formation mechanisms and our corrected DIC Δ¹⁴C_natural values, we estimated a ¹⁴C-based residence time of 360-690 years for BBDW. Based on these residence times, we infer carbon is likely stored for centuries in deep Baffin Bay.

Introduction

Approximately half of the carbon dioxide (CO₂) emitted into the atmosphere from fossil fuels has been absorbed by the oceans. It has been suggested the oceans comprise the only true sink of atmospheric CO₂ over the past 200 years, storing carbon in the deep ocean on timescales of ~1000 years (Sabine et al., 2004). Most of this oceanic carbon is stored as dissolved inorganic carbon (DIC; ~38,000 GtC), the largest marine carbon reservoir (Heinze, 2014). The Arctic Ocean is an important region for the marine carbon cycle (Hamilton and Wu, 2013) and is warming at three times the rate of the rest of the planet (Arctic Monitoring and Assessment Programme Arctic Climate Change Update, 2021). The effect of climate change on the marine carbon cycle in the Arctic remains unconstrained.

Many Arctic regions act as sinks of anthropogenic carbon below the mixed layer, through the export of particulate organic carbon (POC) and downwelling of surface DIC. Previous studies have used DIC concentrations in Arctic gateways coupled with mass transport models to establish an Arctic Ocean DIC budget (e.g., MacGilchrist et al., 2014; Olsen et al., 2015). The Arctic Ocean absorbs roughly 225 ± 49 Tg C yr^-1 via air-sea gas exchange during the summer months (MacGilchrist et al., 2014). These studies suggest the Arctic Ocean is a carbon sink for most of the year. Open water biological activity and air-sea gas exchange account for most CO₂ uptake. The sink is assumed to be minimal during the winter due to sea ice cover (Shadwick et al., 2011). While DIC and total alkalinity measurements have played a critical role in our understanding of Arctic carbon cycling and ocean acidification, these studies require repetition of high spatial resolution field measurements to estimate carbon fluxes. Most Arctic sampling campaigns are limited to the summer/fall, limiting air-sea gas exchange estimates (Bates and Mathis, 2009).

Marine DIC radiocarbon (Δ¹⁴C) and stable carbon (δ¹³C) measurements are powerful tracers of water mass sources, ages, and advection. Additionally, the decrease in atmospheric δ¹³CO₂ (i.e., the Suess effect) has been used as key evidence for anthropogenic carbon sequestration in the form of marine DIC (Körtzinger et al., 2003). DIC δ¹³C values serve as valuable endmembers for understanding DIC cycling in dynamic Arctic Ocean environments. For example, surface DIC δ¹³C values in the Arctic Ocean range between -0.6 and +2.2‰, with deep DIC δ¹³C ranging from -1.0 to +2.0‰, in both the Canadian and Eurasian Basins (Griffith et al., 2012; Bauch et al., 2015). DIC δ¹³C values in the Beaufort Sea, which feeds Pacific water to the Canadian Arctic Archipelago (CAA), range between -0.5 to +2.2‰ (Mol, 2017). The North Atlantic Ocean, which receives Arctic water through Davis and Fram Strait, consists of DIC δ¹³C between -0.1 to +1.9‰ (Humphreys et al., 2016).

Radiocarbon is naturally produced in the stratosphere through the capture of slow cosmic ray neutrons by atmospheric ¹⁴N nuclei (Trumbore et al., 2016). In addition, the infiltration of atmospheric “bomb” CO₂ into the global oceans since the 1960s has made ¹⁴C a useful physical oceanographic tracer. Northern hemisphere atmospheric Δ¹⁴C values reached +1000‰ with surface marine DIC Δ¹⁴C peaking at roughly +200‰ a decade later. Atmospheric CO₂ Δ¹⁴C values have declined to 0‰ today, due to dilution of this signal via increased fossil fuel output and the incorporation of atmospheric CO₂ into the oceans and terrestrial biosphere (Guilderson et al., 2000; Brown and Reimer, 2004). Similarly, the influx of fossil (Δ¹⁴C = -1000‰) CO₂ into the marine DIC reservoir (i.e., the Suess Effect) has proven a useful tool for quantifying anthropogenic impacts on the ocean (Chen and Millero, 1979). However, to date, only a handful of Arctic DIC Δ¹⁴C measurements have been published. DIC Δ¹⁴C values of -60‰ were found in the deep Eurasian Basin, and values as low as -105‰ to -112‰ in the deep Canada Basin (Östlund et al., 1987). Surface DIC Δ¹⁴C values range between -35‰ to +19‰ in the Beaufort Sea, and a value of +52‰ was measured at a single station in the Amundsen Basin (Druffel et al., 2017). Deep DIC Δ¹⁴C values in the Beaufort Sea were as low as -111‰ (Griffith et al., 2012).

In this study, we present the first DIC Δ¹⁴C and new δ¹³C data for Baffin Bay. We use Δ¹⁴C and δ¹³C values of DIC to trace water mass characteristics, DIC sources, the penetration of bomb vs. anthropogenic (fossil) Δ¹⁴C and to estimate the first ¹⁴C-based residence time of deep Baffin Bay.

Methods

Sample Collection

DIC samples for isotopic analysis were collected from 11 stations throughout Baffin Bay in July 2019 aboard the icebreaker CCGS Amundsen (Figures 1A–C). Water samples were collected throughout the water column for each station using a rosette equipped with the SBE 911plus conductivity, temperature, depth probe (CTD) and 24 × 12L Niskin bottles. Prior to DIC Δ¹⁴C and δ¹³C sample collection, 250 mL Pyrex bottles (Corning #1395-250), caps and 45 mm clear pouring rings were cleaned by soaking in 10% HCl for 1 hour, rinsed with copious amounts of 18.2 MΩ Milli-Q water (TOC <4ppb) and dried overnight. The bottles were then baked at 540°C for 2 hours. Samples for DIC and Total Alkalinity (A_T) concentrations were collected into 250 mL borosilicate glass bottles, overflowed three times, poisoned with 100μL of saturated HgCl₂, sealed with Apezion M grease on ground glass stoppers and secured with electrical tape. Samples for DIC Δ¹⁴C followed a similar process but were re-opened briefly after collection in the wet lab, poisoned, re-capped and stored in the dark at room temperature. Nutrient samples were collected in 20 mL polyethylene flasks and measured immediately at sea.

FIGURE 1

Figure 1 Area and station maps with temperature and salinity in Baffin Bay. (A) broader area map indicates major currents (white arrows) land masses and Baffin Bay gateways (grey boxes). Gateways include Davis Strait (DS; 640m), Nares Strait (NS; 250m), Jones Sound (JS; 120m), and Lancaster Sound (LS; 300-800m). (B) Data from all stations visited in the 2019 research cruise were used to create the interpolated surface distributions of sea surface temperature (°C). Stations sampled for DIC are indicated by large black diamonds and labeled. (C) Data from all stations visited in the 2019 research cruise were used to create the interpolated surface distributions of salinity from 2-3 metres depth are reported in Practical Salinity Units (PSU). Plots b and c are Ocean Data View interpolations resulting from all CTD casts on the cruise (small black dots; n=52 stations) and are used to indicate the presence of the Baffin Island Current (BIC) and Western Greenland Currents (WGC) during the time of sampling.

DIC, Total Alkalinity, Salinity, and Nutrient Analysis

DIC concentrations were analyzed by coulometry (Johnson et al., 1993) and A_T by open cell potentiometric titration after (Dickson et al., 2007). DIC and A_T samples were analyzed within 15 months of collection at both the Bedford Institute of Oceanography (BIO; stations 193, 196, and 224) and the Institute of Ocean Sciences (IOS; stations BB15, BB18, 227, 224, BB2, 204, 210 108, and 323). Analytical precision was determined by repeated analysis of a bulk seawater sample. For samples analyzed at Bedford Institute of Oceanography, precision for DIC and A_T was ±0.03%. ± 0.05%, respectively. Results for DIC and A_T at IOS were calibrated against certified reference material (CRM) supplied provided by Dr. Andrew Dickson (Scripps Oceanographic Institution, San Diego, USA), and the analytical precision, based on the average difference between replicates, was ±1.1 µmol kg^-1 for DIC and ±2.0 µmol kg^-1 for A_T. An inter-lab comparison of DIC and A_T concentrations yielded errors of ±7.76 µmol kg^-1 and ±14.27 µmol kg^-1 (n = 24), respectively. Bottle salinity data were collected from the rosette system and analyzed onboard using an AUTOSAL 8400B salinometer and reported in the Practical Salinity Scale units (PSU). Nutrient concentrations were measured by colorimetry following Aminot and Kérouel (2007) at Université Laval with an analytical precision of ±1 µmol kg^-1.

Stable Isotope (δ¹³C) Analysis

For each sample, 1 mL of seawater was transferred into a pre-cleaned 12 mL Exetainer vials, which had been flushed with UHP He gas within a glove bag to remove of any atmospheric CO₂. 100 µL of 85% H₃PO₄ was then added to each sample, which was shaken and left to equilibrate for at least 12 hours at room temperature. This allowed the DIC to equilibrate with CO₂ in the headspace of the vials (Olack et al., 2018). Isotopic standards were prepared using external calcite and limestone standards (IAEA NBS-18, δ¹³C = -5.01‰; IAEA NBS-19, δ¹³C = +1.95). Internal bicarbonate standards for calibration correction were also used (VL-1 δ¹³C = -3.21‰; VL-2 δ¹³C = -15.80‰; VL-3 δ¹³C = -13.89‰ from Ján Veizer Stable Isotope Laboratory; OXG δ¹³C = -7.5‰ from Oxford University). DIC δ¹³C values were measured using a Thermo-Finnigan Gas Bench coupled to a Thermo-Finnigan Delta V-Plus Continuous Flow Isotope-Ratio Mass Spectrometer (IRMS) at the University of Ottawa, Ján Veizer Stable Isotope Laboratory. Results are reported in standard per mil (‰) notation and relative to Vienna Pee Dee Belemnite (V-PDB) for δ¹³C. Average external precision based on sample replicates (n = 100) was ±0.1‰.

Natural Abundance Radiocarbon (Δ¹⁴C) Analysis

Seawater DIC was extracted using the headspace CO₂ method for Δ¹⁴C following Gao and co-workers (2014). Approximately 40 mL of seawater was transferred to pre-cleaned and pre-weighed 60 mL vials (Fisher Scientific #5719398) by decanting in a glove bag flushed multiple times with ultra-high purity (UHP) N₂ gas. Vials were filled with ~40mL sample, were then sealed with Viton and silicone septa, secured with caps, and re-weighed to calculate seawater volume. Each vial was then injected with 0.5 mL of 85% H₃PO₄ (Fisher Scientific #A260500) using a 5mL gas-tight glass syringe (Hamilton #81520) and a 23-gauge side-port needle (Fisher Scientific #14815488). Acidified samples were gently shaken and heated in a block at 75°C for 2 hours.

Isotopic standards were prepared using in-house carbonate standards (Fm = 0.9444 ± 0.0018; Coral STD, UC Irvine), IAEA-C2 travertine (Fm = 0.4114 ± 0.0003), and ¹⁴C-free NaHCO₃ (Fm = 0, uOttawa). Standards were weighed and transferred in a glove bag into pre-cleaned 60 mL I-Chem vials, along with 40 mL of stripped standard water (100 mL of Milli-Q water stripped with UHP N₂ gas for 15 minutes) and sealed. Standards were then acidified with 0.5 mL of 85% H₃PO₄ and heated at 75°C for 2 hours.

Combusted OX-1 (Fm = 1.0398), OX-2 (Fm = 1.3407) standards and in-house Glycine (Fm = 0) were used as modern normalization standards and dead process blank standards during AMS sample analysis. These standards were weighed into pre-baked (540°C/2 hrs) 6 mm quartz tubes with ~150 mg of CuO, sealed under vacuum and combusted (850°C/3 hrs). Combusted standards were then expanded into a vacuum line to cryogenically purify the resultant CO₂ that was then graphitized using the sealed-tube Zn method (Xu et al., 2007).

DIC samples and standard headspace was extracted using a pre-flushed 60 mL BD Luer-Lock syringe equipped with a BD 2.5 cm needle and inserted into the vacuum line system through a septum. The CO₂ was then cryogenically purified with liquid nitrogen and quantified manometrically. Purified CO₂ was graphitized using the sealed-tube Zn method (Xu et al., 2007), pressed into targets at uOttawa and analyzed for Δ¹⁴C at the Keck Carbon Cycle AMS facility at the University of California Irvine. DIC radiocarbon data is reported as Fraction Modern (Fm), Δ¹⁴C and apparent ¹⁴C age following the conventions set forth by (Stuiver and Polach, 1977) and corrected for year of sample collection and for modern vs. dead C blanks associated with sample preparation and extraction.

Estimation of Anthropogenic (DIC_anth) and “Bomb” DIC ¹⁴C

Our measured DIC Δ¹⁴C values include contributions of both atmospheric “bomb” ¹⁴C (from 1950s above-ground thermonuclear weapons testing) and fossil (anthropogenic) CO₂ that are incorporated into the DIC reservoir through air-sea gas exchange. This DIC mixture is expressed in Equation 1.

$\begin{array}{ll}DI{C}_{measured}{\text{Δ}}^{14}{C}_{measured}=DI{C}_{natural}{\text{Δ}}^{14}{C}_{natural}+DI{C}_{bomb}{\text{Δ}}^{14}{C}_{bomb}+DI{C}_{anth}{\text{Δ}}^{14}{C}_{anth}& (1)\end{array}$

We use two methods to correct DIC Δ¹⁴C_measured values to Δ¹⁴C_natural values which can then be used to estimate a ¹⁴C-residence time for BBDW. The potential alkalinity method (P_alk) proposed by Rubin and Key (2002) was used to separate the amount of Δ¹⁴C_bomb from Δ¹⁴C_natural. P_alk is first calculated using measured total alkalinity (A_T), then corrected for biological activity using nitrate concentrations and normalized to salinity (Equation 2). Δ¹⁴C_{bomb corr} was then calculated based on the regression statistics between global Δ¹⁴C versus potential alkalinity measurements (Equation 3) corrected for a global surface potential alkalinity value of 2320, known as P_alk0. Δ¹⁴C_bomb was determined by calculating the difference between Δ¹⁴C_measured and Δ¹⁴C_{bomb corr} for each sample.

$\begin{array}{ll}{\text{P}}_{\text{alk}}=({\mathrm{A}}_{\text{T}}+\text{Nitrate})\times 35/\text{Salinity}& (2)\end{array}$

$\begin{array}{ll}{\text{Δ}}^{14}{C}_{bombcorr}=-59.0-0.962({\text{P}}_{\text{alk}}-2320)& (3)\end{array}$

The amount of anthropogenic carbon present in DIC (DIC_anth) was calculated by using the quasi-conservative tracer (ΔC^*) method following Lee et al. (2003). ΔC^* is calculated using measured DIC (DIC_measured) and alkalinity concentrations (A_T) in Equation 4 and is reported in µmol kg^-1. O_{2 measured} and O_{2 eq} represent measured oxygen concentrations and oxygen saturation levels at a given temperature and salinity, respectively. Preformed alkalinity concentrations $({\mathrm{A}}_{\text{T}}^0)$ are determined using Equation 5. Here, NO represents the relationship between oxygen and nitrate concentrations (Equation 6) and R_C:O and R_N:O represent the stoichiometric ratios which relate inorganic carbon (C), nitrate (N), and dissolved oxygen (O). R_C:O and R_N:O are based on the Redfield Ratio proposed by (Anderson and Sarmiento, 1994). DIC_eq is the DIC in equilibrium with pre-industrial atmospheric CO₂ levels with a fugacity of 280 μatm (Equation 7).

$\begin{array}{ll}\text{Δ}{\text{C}}^{*}=DI{C}_{measured}+{\text{ R}}_{\text{C}:\text{O}}\left({\text{O}}_{2eq}^{}-{\text{ O}}_{2measured}^{ }\right)-0.5\left[\left({\text{A}}_{\text{T }measured}^{ }-{\text{ A}}_{\text{T}}^0\right)-{\text{ R}}_{\text{N}:\text{O}}\left({\text{O}}_{2eq}^{}-{\text{ O}}_{2measured}^{ }\right)\right]-DI{C}_{eq}\left(fC{O}_2 =280\mu atm,A_T^0,\theta ,S\right)& (4)\end{array}$

$\begin{array}{ll}{\mathrm{A}}_{\text{T}}^0\left({\text{μmol kg}}^{-1}\right)=335.7+55.80\times \text{ S }+0.08924\times \text{ NO}& (5)\end{array}$

$\begin{array}{ll}\text{NO }={\text{ O}}_2-{\text{ R}}_{\text{O}:\text{N}}\times \text{ N}& (6)\end{array}$

$\begin{array}{ll}DI{C}_{eq}=2077-8.517\left(\text{θ }-9\right)+3.523\left(\text{S }-35\right)+0.6399\left({\mathrm{A}}_{\text{T}}^0-2320\right)& (7)\end{array}$

Final DIC_anth concentrations are calculated using Equation 8, with ΔDIC_diseq values from Lee et al. (2003) based on seawater potential density. In our study, DIC_anth values <5 µmol kg^-1 were assumed to contain zero DIC_anth (see Supplementary Table 2).

$\begin{array}{ll}DI{C}_{anth}=\text{ Δ}{\mathrm{C}}^{*}-\text{ Δ}DI{C}_{diseq}& (8)\end{array}$

In comparing North Atlantic vs. Baffin Bay DIC_anth and Δ¹⁴C_bomb depth profiles, we find that the P_alk proxy grossly overestimated Δ¹⁴C_bomb values for depths shallower than 400m. We found estimated DIC_anth values via the ΔC* proxy gave realistic values, albeit lower than in the North Atlantic. Both proxies were parameterized for the subtropical North Atlantic and have recognized limitations at high latitudes (Rubin and Key, 2002; Lee et al., 2003).

To correct our Δ¹⁴C_measured for DIC_anth and DIC_bomb Δ¹⁴C contributions, a two-endmember mixing model was used to determine the Δ¹⁴C per mil equivalence of DIC_anth, that accounts for DIC_anth, DIC_measured, and a Δ¹⁴C_anth value of -1000‰. For stations above 500m, only the per mil equivalence DIC_anth was subtracted from Δ¹⁴C_measured. For stations below 500m, both Δ¹⁴C_bomb and the per mil equivalence of DIC_anth were subtracted from Δ¹⁴C_measured, to determine Δ¹⁴C_natural. DIC Δ¹⁴C_natural values were converted to a fraction modern (Fm) value and then to an apparent ¹⁴C-age based on a half-life of 5730 years. To calculate a residence time of the region, we used DIC Δ¹⁴C and other relevant measurements from the A16N line in the North Atlantic in 2013 (https://cchdo.ucsd.edu/cruise/33RO20130803). DIC_anth and Δ¹⁴C_bomb corrections made to Δ¹⁴C_natural values and a corrected apparent ¹⁴C-age.

Results

DIC δ¹³C Values

DIC δ¹³C values ranged from -0.7‰ to +1.9‰ across Baffin Bay (Figure 2A). Enriched DIC δ¹³C values were observed at surface depths (+0.7‰ to +1.9‰) and depleted DIC δ¹³C were observed at depths greater than 100m (-0.0‰ to -0.7‰). Surface DIC δ¹³C values (less than 20m) were enriched along the coast of Greenland and in the Northern gateways (stations BB18, 227, 210, 108 and 323; δ¹³C = +1.3 ± 0.4‰, n = 12), while more depleted values were found along Davis Strait and central Baffin Bay (stations 193, 196, BB2, and 224; δ¹³C = +0.9 ± 0.3‰, n = 8). Mid-depth (200 – 700m) DIC δ¹³C ranged from -0.4‰ to +0.2‰, with slightly depleted values found at stations BB18 and 227 (δ¹³C = -0.30 ± 0.07‰; n = 4). Below 750m in Baffin Bay, DIC δ¹³C values had a narrow range (-0.28 ± 0.16‰; n = 19). Deep water (>1200m) at Station 224 was more depleted (-0.51 ± 0.11‰; n = 4) than DIC δ¹³C values at station BB2 (-0.27 ± 0.07‰; n = 4).

FIGURE 2

Figure 2 Depth profiles of DIC δ¹³C and Δ¹⁴C values in Baffin Bay. (A) measured DIC δ¹³C values, ± 0.1‰ (See Stable Isotope (δ¹³C) Analysis). Different sampled stations (Figure 1) indicated by legend, which cover the extent of the Baffin Island Current (BIC) in central Baffin Bay (circles), the Western Greenland Current (WGC) along the coast of Greenland, (diamonds). and the Northern gateways of Nares Strait and Lancaster Sound (triangles). (B) DIC Δ¹⁴C signatures, with errors ranging between ±1.2 to 2.0‰ (See “Natural Abundance Radiocarbon (Δ¹⁴C) Analysis” and “Estimation of Anthropogenic (DIC_anth) and “Bomb” DIC ¹⁴C”).

DIC Δ¹⁴C Values

Our measured Baffin Bay DIC Δ¹⁴C values had a ~120‰ range (from –90.0‰ to +29.8‰; Figure 2B). Surface (less than 20m) DIC Δ¹⁴C values ranged between -5.4‰ to +22.9‰. Positive surface DIC Δ¹⁴C were found along coastal Greenland (stations 196, BB15, BB18, 227 204, and 210; Δ¹⁴C = +17.6 ± 2.5‰, n = 12). In contrast, more negative DIC Δ¹⁴C surface values were found in the Northern gateways (stations 108 and 323; Δ¹⁴C = -1.7 ± 3.3‰, n = 3). Surface DIC Δ¹⁴C values at central stations (193, BB2, 224) fell between these two end members (Δ¹⁴C = +8.5 ± 3.7‰; n = 6). Mid-depth waters (200 – 700m) had DIC Δ¹⁴C between -1.11‰ and +29.8‰, with positive Δ¹⁴C values found in Northern Baffin Bay (stations BB2, 108, and 323; Δ¹⁴C = +23.8 ± 3.9‰, n = 7). Below 750m, DIC Δ¹⁴C values gradually decreased from +7.7‰ to –90.0‰, with the deepest water (1400 – 2400m) having the most negative Δ¹⁴C values (-82.2 ± 8.5‰; n = 9).

Discussion

Baffin Bay – Hydrographic Setting

Wide ranges in Baffin Bay surface water salinity and temperature were observed (Figures 1B, C; Figures 3A, B). Surface water (above 50m) on the western side of Baffin Bay was cooler and fresher, while surface water on the eastern side was warmer and more saline (Fox and Walker (2022), accepted). We note that the Ocean Data View interpolation of n=52 stations from Baffin Bay in Figure 1 are meant for visualizing these differences and do not reflect the true spatial structure of temperature or salinity. These surface water summer conditions are influenced by surface processes such as solar insulation, wind-driven mixing, and ice melt (Fissel et al., 1982). Surface water variance is also influenced by the current systems within Baffin Bay. Baffin Bay is a marginal sea ice zone experiencing full sea-ice coverage from December to April. Due to the relatively warm West Greenland Current, ice cover decreases earlier along the Greenland coast than on the western side of Baffin Bay . By late summer, Baffin Bay is typically clear of extensive sea-ice, although icebergs are common, especially in proximity to fjords. Sources of freshwater in Baffin Bay include local precipitation, river runoff, sea ice, ice bergs and glacial meltwater, and the Arctic Ocean outflow. It is estimated Baffin Bay has a total river inflow of ~100 km³ y^-1 (Shiklomanov et al., 2020). Work by Azetsu-Scott et al. (2012) suggests that glacial melt on the Greenland Shelf comprises ~6% total freshwater and that Arctic outflow water from the CAA and Northern Gateways dominates freshwater flux in Baffin Bay (~60% in Western Davis Strait). The Arctic outflow water from the CAA includes Pacific (31.5 < S < 33), sea ice meltwater and meteoric (Mackenzie R. and CAA tributaries) water endmembers. A small portion of meteoric water enters Baffin Bay via Nares Strait from Russian rivers (Mungall et al., 2017; Burgers et al., 2017).

FIGURE 3

Figure 3 Potential temperature-salinity diagram with DIC δ¹³C values throughout Baffin Bay. (A) Bottle salinity and potential temperature values plotted for each collected station. Different sampled stations (Figure 1) indicated by legend; Baffin Island Current (circles), Western Greenland Current (diamonds), Northern gateways (triangles). Dashed grey lines represent potential density anomaly contours. (B) Potential temperature versus salinity with DIC δ¹³C (color bar). Dashed boxes indicate designated currents and water masses; Arctic Water (AW), western Greenland shelf water (WGSW), Baffin Island current (BIC), West Greenland Irminger water (WGIW), and Baffin Bay deep and bottom water (BBDW; BBBW). Note: fewer data points on plot b are since not all Niskin bottle depths were sampled for DIC δ¹³C.

Baffin Bay water masses are indicated on Figure 3B. Two main surface currents dominate Baffin Bay: the Baffin Island Current (BIC) and the West Greenland Current (WGC). The BIC is a southerly flowing current containing Arctic Water (AW) to depths of 300m and has water temperatures <-1.6°C and salinities <33.8 (Münchow et al., 2015). AW enters Baffin Bay via the Northern Gateways (Nares Strait and Lancaster Sound) and is modified through glacial and sea ice discharge (Curry et al., 2011). In this study, AW was observed at both central stations (BB2, 224) between 100-300m. Western Greenland Shelf Water (WGSW) was evident along the coast of Greenland and was both warm (7°C) and saline (34.1), with a density anomaly (σ_θ = ρ_θ – 1000, where ρ is density and θ indicates potential temperature) ranging between 25.5 and 27.3 kg m^-3. WGSW is composed of waters from the East Greenland Current (EGC), which originate from AW exiting through Fram Strait and later enter Baffin Bay via the WGC (Cuny et al., 2002). As the WGSW progresses northward along the WGC extension, it becomes fresher and cooler, due to mixing with glacial meltwater (Münchow et al., 2015).

At depth, several water masses are present in Baffin Bay. West Greenland Irminger Water (WGIW) is defined by salinities >34.1, potential temperatures >2°C, and a density anomaly range of 27.3 to 28 kg m^-3 (Curry et al., 2011). WGIW originates from warm, saline water found in the Irminger Sea. Transitional Water (TrW) found below 250m, is modified AW that has mixed with WGIW (Curry et al., 2011). The TrW typically has temperatures >2°C, salinities <33.7, and σ_θ of 25.5 – 27.3kg m^-3 (Curry et al., 2011). Baffin Bay Deep Water (BBDW) has a temperature and salinity of 0°C and ~34.5, respectively, and is found between 1200 and 1800m (Curry et al., 2011). Baffin Bay Bottom Water (BBBW) has a temperature and salinity of -0.4°C and ~34.5, respectively. BBBW is more difficult to distinguish using temperature, salinity, and water mass density alone (Lehmann et al., 2019).

DIC δ¹³C as a Surface Water Mass Tracer

Surface DIC δ¹³C values can provide useful constraints on DIC source contributions. In Baffin Bay there are several possible DIC sources in the BIC and WGC, including: Pacific, North Atlantic, Arctic water, regional sea ice melt and glacial melt water. The many surface DIC sources in Baffin Bay results in a diversity of isotopic endmember values that could also contribute to deep-water DIC δ¹³C values. Understanding the processes shaping surface DIC δ¹³C distributions are therefore important to understanding the DIC cycle within Baffin Bay. Here we discuss surface and intermediate water DIC δ¹³C values as water mass tracers.

Our observed depth trends in DIC δ¹³C values are broadly consistent with the open ocean. Surface DIC δ¹³C values are enriched, whereas deep DIC δ¹³C values are depleted (Figures 2A, 3B). These trends in DIC δ¹³C are attributed to air-sea gas exchange with some isotopic fractionation of CO₂ across the air-liquid boundary and to a much lesser extent the fractionation of organic carbon through photosynthesis (Zhang et al., 1995). Depleted DIC δ¹³C values at depth can be attributed to the preferential oxidation of ¹³C-depleted sinking particulate and dissolved organic matter (POM; DOM) via heterotrophic respiration. Older aged water masses typically contain depleted DIC δ¹³C values, since there is more time for remineralization (Kroopnick, 1985; Emerson and Hedges, 2008).

Surface (above 20m) DIC δ¹³C is variable in Baffin Bay. The warm WGSW along the coast of Greenland contains enriched DIC δ¹³C values (>+1.03‰). These values are consistent with previously reported DIC δ¹³C from the North Atlantic along the southern tip of Greenland (+1.09 to +1.90‰; Humphreys et al., 2016). Within WGSW, we observe an increase in surface DIC δ¹³C values from station 196 to 210 (+1.03 to +1.90‰, respectively). One possible explanation for this trend is high nutrient loading from melting Greenland glaciers, leading to enhanced primary production and enriched DIC δ¹³C values as the WGC flows north. For example, station BB18, located in the Viagat Fjord behind Disko Island, had the highest surface DIC δ¹³C (+1.90‰). This station also had high POM and chlorophyll-a concentrations due to an abundance of diatoms (Fox and Walker (2022), accepted; Lovejoy pers. comm.). This marginal sea ice region and the inflow of saline Atlantic water have been previously shown to promote primary production (Krawczyk et al., 2021).

Northern Gateway stations (108 and 323) had enriched surface DIC δ¹³C values (+0.71 to +1.42‰). Surface water in Smith Sound (Station 108) is a mixture of both the inflow of AW from Nares Strait and WGSW (Hamilton and Wu, 2013). Thus, Station 108 does not represent a pure AW endmember. Surface water in Lancaster Sound (Station 323) contains Pacific Water entering Baffin Bay via the CAA. Enriched surface DIC δ¹³C values in Lancaster Sound may be indicative of this Pacific influence, which has been observed to have DIC δ¹³C values up to +2.20% in the Amundsen Gulf (Mol, 2017). Lancaster Sound is also a region of seasonal sea ice cover, high nutrient, primary production, and diatom species richness (Krawczyk et al., 2021). These factors contribute to the enriched surface DIC δ¹³C values at Station 323. Variability in surface DIC δ¹³C in Baffin Bay can also be attributed to sea-ice and glacial melt, as evident in large ranges in observed surface salinity as shown in Figure 3A.

Central Baffin Bay stations BB2 and 224 had more depleted surface DIC δ¹³C values (+1.06 to +1.50‰) than Eastern Baffin Bay stations along coastal Greenland (Figure 3). These lower DIC δ¹³C values coincide with AW within the BIC. The Arctic Ocean is covered in sea ice for most of the year, inhibiting air-sea gas exchange and increasing the time in which heterotrophic respiration can result in depleted DIC δ¹³C values (Morée et al., 2018). DIC δ¹³C values within the BIC (20 – 200m; +0.07 to +0.37‰) are significantly higher than that of WGIW (-0.39 to +0.07‰), reflecting more recent air-sea gas exchange in the BIC vs. older, intermediate water (Münchow et al., 2015).

Overall, the trends we observe in DIC δ¹³C values in temperature-salinity space (Figure 3B) suggest that DIC δ¹³C can be used as an effective tool for distinguishing carbon sources within water masses. A simple least squares regression of DIC δ¹³C (y) and potential density (x) resulted in statistically significant correlations at all stations in Baffin Bay; in particular for coastal Greenland (R^{2 =} 0.78-0.98) and Northern Gateways stations (R^{2 =} 0.85-0.96; see Table S3 and Figures S1 and S3). Surface DIC δ¹³C variability is strongly affected by sea-ice melt (more positive DIC δ¹³C) and the inflow of saline Atlantic water (more negative δ¹³C). This is especially true for several coastal Greenland stations adjacent to large fjord regions (Stations BB18, 227, 210; Figure S4 and Table S5). In contrast, deep DIC δ¹³C variability is driven by slow heterotrophic respiration at depth. This is indicated by depleted DIC δ¹³C values >1000m (σ_θ = ~27.7 kg/m³) at stations 224 and BB2 (-0.27 to -0.68‰ at 1000 to 2000m and -0.16 to -0.36‰ at 1200 to 2300m, respectively).

DIC Δ¹⁴C Distributions in Baffin Bay

We observe a wide range in DIC Δ¹⁴C values within Baffin Bay (-90.0‰ to +29.8‰; Figure 2B). DIC Δ¹⁴C values can reflect water mass aging via radioactive decay, varying degrees of recent air-sea gas exchange or allochthonous DIC Δ¹⁴C sources. For example, the advection of waters with different ¹⁴C reservoir ages (i.e., surface DIC Δ¹⁴C vs. atmospheric Δ¹⁴C offsets). With deep-water formation, modern surface DIC Δ¹⁴C can also contribute to observed deep ocean DIC Δ¹⁴C values (e.g., in the North Atlantic Ocean; Broecker et al., 1960).

Surface (<20m) DIC Δ¹⁴C ranged between -5.4‰ to +22.9‰. Surface DIC Δ¹⁴C values at Greenland stations (196, BB15, 227, 204, 210) had the most positive DIC Δ¹⁴C values (+14.3‰ to +22.9‰). These values represent positive DIC Δ¹⁴C endmembers from the North Atlantic at depths <150m (+30.2‰ to +62.3‰; Bullister and Barigner, 2020), indicating the WGSW has a large impact on DIC Δ¹⁴C values in Eastern Baffin Bay. Previous studies have noted summer upwelling of warmer, saline WGSW along Greenland (Melling et al., 2010) which would clearly contribute DIC with positive Δ¹⁴C values.

Surface water (above 20m) at the Northern gateway stations (108, 323) had far lower DIC Δ¹⁴C values (-5.4‰ to +0.9‰). These stations contain Arctic and CAA contributions to the Smith and Lancaster Sounds, respectively. Surface DIC Δ¹⁴C values in the Amundsen Gulf and Beaufort Sea range between -35‰ to +19‰, whereas a value of +52‰ was measured in the Eurasian Basin (Druffel et al., 2017). Thus, the lower DIC Δ¹⁴C values we observe in Lancaster Sound (station 323) are consistent with advection of Pacific water with low DIC Δ¹⁴C to Baffin Bay from the CAA. Pacific water nutrient contributions to Baffin Bay were recently reported (Lehmann et al., 2019), again suggesting this is a plausible explanation for our low Station 323 DIC Δ¹⁴C values. The low DIC Δ¹⁴C values we observe for Smith Sound (Station 108) are likely impacted by both AW entering from the Lincoln Sea and WGC water. While no DIC Δ¹⁴C data exist for the Lincoln Sea, this station experiences extended periods of sea ice cover – inhibiting air-sea gas exchange and likely resulting in lower surface DIC Δ¹⁴C values.

Surface DIC Δ¹⁴C values in central Baffin Bay are modern (BB2, 224; +5.8‰ to +15.7‰) and possibly comprise a mixture of low AW DIC Δ¹⁴C with positive DIC Δ¹⁴C from WGSW and the influence of seasonal sea ice cover. At these stations, large DIC Δ¹⁴C gradients (-90.0 to +20.6‰) were present in deep (>500m) water of similar densities (σ_θ = 27.56 to 27.72 kg m^-3). Below 1200m, DIC Δ¹⁴C values gradually decrease from -35.4‰ to -90.0‰ (or 225 to 675 ¹⁴C-years) towards the bottom (σ_θ = 27.70 ± 0.02 kg m^-3, BBDW). Unlike DIC δ¹³C, no significant relationships were found between potential density and DIC Δ¹⁴C (Table S4; Figure S2). This is because similar, modern DIC Δ¹⁴C values were measured across large surface salinity (and density) gradients. Conversely, at Stations BB2, 224, large gradients in DIC Δ¹⁴C values were present in mid-depth (500-1200 m) waters of relatively invariant σ_θ (e.g., 27.7 kg m^-3; Figures 2B; S2F, G).

While the distributions of DIC Δ¹⁴C appear consistent with the physical oceanography of Baffin Bay, measured DIC Δ¹⁴C values can also include atmospheric “bomb” carbon (Δ¹⁴C = -1000‰; Mahadevan, 2001; Sweeney et al., 2007) or anthropogenic fossil fuel carbon (DIC_anth; Carter et al., 2019). The contribution of DIC_anth and Δ¹⁴C_bomb are ever present in the modern global surface ocean, but their relative contributions and depth of infiltration in the water column vary widely. Since our goal is to derive a ¹⁴C-based residence age of Baffin Bay, we next quantify these contributions to deep Baffin Bay water.

Estimation of DIC Δ¹⁴C_bomb Contributions

Our estimates of Δ¹⁴C_bomb using the P_alk method (see Estimation of Anthropogenic (DIC_anth) and “Bomb” DIC ¹⁴C) for Station BB2 are shown in Figure 4. Both DIC Δ¹⁴C_bomb and Δ¹⁴C_{bomb corr} values are highly unconstrained above 400m, with very negative Δ¹⁴C_{bomb corr} values and very positive Δ¹⁴C_bomb values beyond what would be reasonably expected for seawater DIC Δ¹⁴C in 2019. This is a known limitation of the P_alk proxy which was parameterized for the subtropical and tropical Atlantic (<60°N/S) with salinities averaging 36.31 ± 0.35, and surface alkalinity averaging 2377 ± 22 µmol kg^-1 (Jiang et al., 2014). In comparison, surface Baffin Bay salinity and total alkalinity values average 32.2 ± 1.2 and 2194 ± 59 µmol kg^-1, respectively. Lower surface salinity and alkalinity values, combined with seasonal sea ice cover in Baffin Bay, limits the application of the P_alk method in the Arctic. The overestimation of bomb Δ¹⁴C above 400m in Baffin Bay is particularly telling. At their peak, global surface DIC Δ¹⁴C values reached +200‰ in the late 1960s (Rodgers et al., 2000). We note that Rubin and Key (2002) recognized the P_alk proxy predicts lower than expected Δ¹⁴C_{bomb corr} values north of 58°N and suggested a more reliable calibration value (P_alk0) of 2330 µmol kg^-1 for high latitudes. However, we find surface Δ¹⁴C_{bomb corr} values to remain unconstrained even using this P_alk0 value and that more work is needed to refine the P_alk proxy in the Arctic.

FIGURE 4

Figure 4 Calculated bomb Δ¹⁴C and natural Δ¹⁴C values for Baffin Bay deep water. Measured DIC Δ¹⁴C values for deepest station collected in Baffin Bay, BB2 (Figure 1). Calculated Δ¹⁴C_bomb and Δ¹⁴C_{bomb corr} values derived using the Potential Alkalinity (P_alk) method (Rubin and Key, 2002) method (see Estimation of Anthropogenic (DIC_anth) and “Bomb” DIC ¹⁴C). Dashed lines and open circle indicate overestimated bomb and underestimated natural Δ¹⁴C values above 400m.

Despite the unconstrained surface estimates of Δ¹⁴C_bomb and Δ¹⁴C_{bomb corr}, these parameters provide realistic values at depth (Figure 4). The confluence of the corrected and measured values suggests the P_alk proxy can accurately predict Δ¹⁴C_{bomb corr} at depth, and an absence of Δ¹⁴C_bomb in BBDW and BBBW. We believe the reliability of the P_alk proxy below 400m can be attributed to the similar salinity, nitrate, and alkalinity values in North Atlantic Deep Water (NADW) and deep Baffin Bay. For example, NADW below 1200m in the tropics (35°N to Equator) has average salinity and alkalinity values of 35.4 ± 1.1 and 2371 ± 37 µmol kg^-1, respectively (Takahashi et al., 1981). Average salinity and total alkalinity values below 1000m in BBDW are 34.39 ± 0.01 and 2290 ± 13 µmol kg^-1, respectively. While these deep ocean values are not identical, it appears that they fall within a range for which the P_alk proxy was developed (i.e., closer to deep sub-tropical Atlantic values). Because deep water alkalinity is globally less variable than surface waters, it is also possible the P_alk method is simply more robust in deep waters than in surface waters. Ultimately, since we can find no evidence of Δ¹⁴C_bomb below 1400m, there is no need to correct our deepest measured DIC Δ¹⁴C values for Δ¹⁴C_bomb when determining a ¹⁴C-based residence time.

Anthropogenic DIC Contributions

Calculated DIC_anth concentrations suggest an influx of DIC_anth via Davis Strait (Figure 5). At Station 193, DIC_anth = 25 - 40 µmol kg^-1 between 100 and 400 m and at Station 196, DIC_anth = 20 - 33 µmol kg^-1 at 0-100m (Figures 5A, B; Table S2). DIC_anth enters Davis Strait via the WGC and advects north within WGSW into central Baffin Bay. Davis Strait DIC_anth concentrations are only slightly lower than those previously reported for the North Atlantic (Lee et al., 2003). Thus, we conclude the North Atlantic supplies significant DIC_anth to Baffin Bay. DIC_anth contributions were lower at our Northern Gateway stations 108 and 323 (9.4 to 41.5 µmol kg^-1), and no DIC_anth was found below 100m at these stations. Between 150-300m the southward BIC also contains little DIC_anth. We note that the Ocean Data View interpolation in Figure 5 is from a limited number of stations, and thus meant for illustrative purposes only. The Ocean Data View distributions does not necessarily reflect the fine-scale spatial structure of DIC_anth in Baffin Bay.

FIGURE 5

Figure 5 Calculated DIC_anth concentrations found throughout central and Eastern Baffin Bay. Distributions of DIC_anth concentrations plotted from two transects along (A) the west coast of Greenland and (B) central Baffin Bay, as indicated by the inset on b). DIC_anth concentrations calculated based on methods by Lee et al. (2003). DIC_anth values <5µmol kg^-1 are assumed to contain zero DIC_anth (See Estimation of Anthropogenic (DIC_anth) and “Bomb” DIC ¹⁴C). Sample depths and locations are indicated by black dots, with station names labelled at the top of the plots. Dashed white contours are density anomaly (σ_θ in kg m^-3) isopycnals. We note that Ocean Data View interpolations of DIC_anth (from n=11 stations) are meant for interpretation and do not reflect an accurate spatial distribution of DIC_anth.

Although we are the first to apply the ΔC^* proxy to Baffin Bay, Earth System models estimate low DIC_anth inventories in the Arctic due to seasonal sea ice and limited air-sea gas exchange (Tjiputra et al., 2010). Lee and co-workers (2003) also found low DIC_anth inventories in surface waters at higher Atlantic Ocean latitudes (~65°N). Low DIC_anth at higher latitudes may also be attributed to mixing of intermediate waters with deeper “DIC_anth free” water. Short residence times (~2 years) of surface water within Baffin Bay may also limit the region’s surface DIC_anth inventories (Rudels, 1986; Poisson and Chen, 1987; Hoppema et al., 2001). Large volumes of water are annually exported south through Davis Strait (2.3 ± 0.7 Sv; Curry et al., 2011) perhaps also precluding the BIC from containing significant DIC_anth (Figure 5).

Above 50m, DIC_anth is variable in Baffin Bay. High DIC_anth is found in surface water along the coast of Greenland (29-46 µmol kg^-1) although surface DIC_anth was not found in the Vaigat Fjord (Station BB18) and on the west side of Davis Strait (193). Such wide-ranging DIC_anth in surface waters directly in contact with the atmosphere demonstrates limitations in the ΔC^* proxy at low temperatures and salinities. Like the P_alk method, the ΔC^* proxy was also developed for the mid-latitude open ocean. As a result, many of our surface DIC_anth values had high errors and/or negative values. The ΔC^* proxy however appears to provide more reasonable estimates of anthropogenic carbon in surface water than the P_alk method does for Δ¹⁴C_bomb. With the limitation of the P_alk method throughout Baffin Bay surface water, we assume the absence of Δ¹⁴C_bomb when DIC_anth is undetectable. As mentioned in section 2.5, DIC_anth values <5 µmol kg^-1 (including negative values) were considered free of DIC_anth (0 µmol kg^-1). Despite these DIC_anth proxy limitations, we do not observe the presence of DIC_anth in Baffin Bay below 800m. We believe this to be reasonable given the absence of Δ¹⁴C_bomb in BBDW. Thus, our deep water measured DIC Δ¹⁴C values do not require further correction for the addition of fossil C.

Estimating a ¹⁴C-Based Residence Age for Deep Baffin Bay

The formation of BBDW and BBBW is a topic of ongoing debate. Two main formation mechanisms have been proposed: mixing of cold, saline, brine-enriched surface shelf water with either Arctic Intermediate Water in Nares Strait and Smith Sound (Bailey, 1956; Collin, 1965; Bourke et al., 1989; Bourke and Paquette, 1991) or with Labrador Sea intermediate water entering via Davis Strait (Tang et al., 2004). Some research has argued that the depth of convective mixing from freezing would be too shallow (<100m) to produce the bottom water in Northern Baffin Bay (Muench, 1970), but this would not necessarily preclude downslope transport of dense water from northern shelves (Tang et al., 2004).

To constrain a ¹⁴C-residence age for deep Baffin Bay, it is important to consider each of these formation mechanisms. The addition of winter brine to AW at intermediate depths in Smith Sound has been previously supported by temperature, salinity and δ¹⁸O isotopic data (Bailey, 1956; Tan and Strain, 1980). Nares Strait is a relatively shallow region (~200m in Kane Basin). The potential density anomaly at 225m in Station 108 was 27.27 kg m^-3, which would not be dense enough to contribute directly to BBDW (σ_θ = 27.72 kg/m³). However, Bourke et al. (1989) suggested that mixing brine-enriched winter waters produced in the North Water Polynya in Smith Sound, with Arctic Intermediate Water could produce deep water densities. If we use measured DIC Δ¹⁴C values from 200-300 m at Station 108 (where C_ath was not detected and presumably little additional bomb Δ¹⁴C is present; Table S2) as a deep-water source endmember (Δ¹⁴C = +26.3‰; n = 2), then the ¹⁴C residence time of BBDW (σ_θ > 27.7 kg/m³, Δ¹⁴C = -90.7‰) would be roughly approximate to its apparent ¹⁴C age (690 ± 35 years).

The second plausible source of BBDW is water entering Davis Strait from the Labrador Sea over a 640m depth sill. Sverdrup et al. (1942) suggested the cooling of Labrador Sea deep water with winter surface water brine as a formation mechanism. The WGC is cooler and fresher than the EGC (3.0°C < θ < 5.5°C and 34.4 < S < 35.0) or the Irminger Current (3.5°C < θ < 5.5°C, S ~ 35.0; Lobb, 2004). In the northern Labrador Sea, the Irminger Current lies below the WGC (Rykova et al., 2009). As it flows northward, the upper portion of the Irminger Current mixes with the cooler and fresher WGC (Cuny et al., 2002). Inflow of WGIW via Davis Strait is seasonal and extends to depth of up to 1000m during the fall (Curry et al., 2014).

We use North Atlantic DIC Δ¹⁴C values (A16N GO-SHIP, Station 4; Bullister and Barigner, 2020) from the Davis Strait sill depth (640m) with a density anomaly (σ_θ = 27.53 to 27.67 kg m^-3) as a second source endmember from which to estimate a ¹⁴C residence time for Baffin Bay. We applied both the P_alk and ΔC^* proxies to correct for Δ¹⁴C_bomb and DIC_anth from measured North Atlantic DIC Δ¹⁴C to determine a Δ¹⁴C_natural depth profile (Figure 6). Both proxies provided robust estimates of Δ¹⁴C_bomb and DIC_anth for the North Atlantic station, consistent with previously published values (Rubin and Key, 2002; Lee et al., 2003). This North Atlantic BBDW source end-member was determined to have a ¹⁴C-age of 330 ± 35 years. Subtracting this age from the apparent ¹⁴C age of BBDW (690 ± 35 years) results in an estimated ¹⁴C residence time of 360 ± 35 years. Since both bomb Δ¹⁴C and DIC_anth are present in the North Atlantic, this BBDW formation mechanism requires extremely slow deep water formation rates.

FIGURE 6

Figure 6 DIC Δ¹⁴C_natural depth profiles, corrected for both Δ¹⁴C_bomb and DIC_anth contributions for Station BB2 and A16N. DIC Δ¹⁴C_natural values from A16N (Station 4; 62.75°N, -20.00°W just south of Iceland) were corrected for DIC_anth and Δ¹⁴C_bomb. Apparent ¹⁴C-ages are calculated based on Δ¹⁴C_natural values (see Natural Abundance Radiocarbon (Δ¹⁴C) Analysis) for depths with potential densities close to BBBW values (θ_σ = 27.7 kg m^-3). Dashed horizontal grey line indicates depth of the Davis Strait sill, at 640m. Dashed line with open circles for BB2 profile indicate unconstrained Δ¹⁴C_natural values based on limitation of the P_alk proxy. Red circles indicate depths at which apparent ¹⁴C-ages were calculated.

Our two estimated ¹⁴C residence times (360-690 years) for deep Baffin Bay provide new constraints on the ventilation age of this ocean basin. Our residence times fall between those previously estimated for BBDW of 20 to 1,450 years (Sadler, 1976; Top et al., 1980; Wallace, 1985). Both deep water formation mechanisms require cooling and brine rejection during sea ice formation that would increase the local DIC concentration of underlying waters (Rysgaard et al., 2007; Moreau et al., 2016; König et al., 2018). This increase in DIC would also selectively preconcentrate DIC with modern Δ¹⁴C (atmospheric), Δ¹⁴C_bomb, and fossil (C_anth) Δ¹⁴C values. Since we do not detect these Δ¹⁴C endmembers in BBDW, significant dilution must occur during such a deep-water formation mechanism (e.g., “shaving” of the deep-water plume; Bourke et al., 1989). This would also suggest that our estimated ¹⁴C residence times are minimum age estimates. Nevertheless, with a minimum residence time of 360 ± 35years, BBDW has the potential to store carbon on centennial timescales.

Summary and Implications

Seawater DIC δ¹³C and Δ¹⁴C values place new constraints on the sources and cycling of carbon in the Arctic. Baffin Bay is a dynamic region, with Arctic, Pacific, and Atlantic carbon source endmembers, as well as glacial inputs along the coast of Greenland. Δ¹⁴C_bomb and DIC_anth estimates through the ΔC^* and P_alk proxies are limited in their scope of application in the Arctic but highlight important DIC source constraints to Baffin Bay water masses. We observe DIC_anth-replete water entering from the North Atlantic Ocean, DIC_anth-free AW in Northern Baffin Bay, and no DIC_anth or bomb Δ¹⁴C in BBDW. Using two possible deep water formation mechanisms for BBDW, we determine a Baffin Bay ¹⁴C residence time of 360-690 years, suggesting deep Baffin Bay stores carbon for several centuries.

Baffin Bay exports cold, dense water to the Labrador Sea which is critical for North Atlantic Deep Water (NADW) formation (Goosse et al., 1997). Global climate change has resulted in warming and freshening of Arctic surface water exported through Baffin Bay and increased stratification in the North Atlantic Ocean. This stratification is now reducing NADW formation (Caesar et al., 2021). We now observe warming temperatures in the Arctic Ocean acting to reduce sea ice extent but also CO₂ uptake (MacGilchrist et al., 2014). With a residence time of 360-690 years, Baffin Bay may partially moderate Arctic climate change through storage of deep carbon on centennial timescales.

Data Availability Statement

The original contributions presented in the study are included in the article/Supplementary Material. Further inquiries can be directed to the corresponding author.

Author Contributions

BW conceived the hypothesis and the objectives of this investigation. SZ and BW drafted the manuscript with input from JW, BE, LM, and KA-S. SZ, BW, and JW performed radiocarbon and stable isotopic analysis of DIC samples. LM and KA-S contributed total alkalinity and DIC concentration data. All authors read, edited, and approved the final manuscript.

Funding

This work was supported by the Natural Sciences and Engineering Research Council (NSERC) of Canada through a Discovery Grant, Accelerator and Launch Supplements (RGPIN-2020-06501, RGPAS-2020-00071, DGECR-2020-00256; to BW), a Discovery Grant (RGPIN-2015-04780 to BE), the Canada Research Chairs program (BW), the NSERC Alexander Graham Bell Canadian Graduate Scholarship (SZ), the Ontario Graduate Scholarship (SZ), and Fisheries and Oceans Canada. This work is a contribution to ArcticNet, a Network of Centres of Excellence Canada.

Conflict of Interest

The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Publisher’s Note

All claims expressed in this article are solely those of the authors and do not necessarily represent those of their affiliated organizations, or those of the publisher, the editors and the reviewers. Any product that may be evaluated in this article, or claim that may be made by its manufacturer, is not guaranteed or endorsed by the publisher.

Acknowledgments

We gratefully acknowledge Chief Scientist Alexandre Forest, Anissa Merzouk and the staff of Amundsen Science and the crew of the CCGS Amundsen for the opportunity to participate in the 2019 research cruise. We also would like to thank Dr. Cara Manning for her help in coordinating the sampling and logistics of the Biogeochemistry groups aboard the CCGS Amundsen. Tonya Burgers (University of Manitoba) and Shawn Marriott (University of Calgary) collected DIC and A_T samples. Danielle Caleb and Marty Davelaar (Institute of Ocean Sciences) performed DIC and A_T analyses. We thank Drs. Xiaomei Xu and John Southon of the UC Irvine Keck Carbon Cycle AMS lab for their expertise and help with ¹⁴C analysis, Paul Middlestead, and the staff of the Ján Veizer Stable Isotope Laboratory at the University of Ottawa for aiding in δ¹³C analysis, Jonathan Gagnon, and Jean-Éric Tremblay (Université Laval) for nutrient measurements and Aislinn Fox for ODV analysis and providing site map figures. Finally, we acknowledge the reviewers, whose constructive comments greatly improved the paper.

Supplementary Material

The Supplementary Material for this article can be found online at: https://www.frontiersin.org/articles/10.3389/fmars.2022.845536/full#supplementary-material

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