The recent state and variability of the carbonate system of the Canadian Arctic Archipelago and adjacent basins in the context of ocean acidification

10 Ocean acidification driven by the uptake of anthropogenic CO2 by the surface oceans constitutes a potential threat to the health of marine ecosystems around the globe. The Arctic Ocean is particularly vulnerable to acidification and, thus, is an ideal region to study the progression and effects of acidification before they become globally widespread. The appearance of undersaturated surface waters with respect to the carbonate mineral aragonite (ΩA < 1), an important threshold 15 beyond which the calcification and growth of some marine organisms might be hindered, has recently been documented in the Canada Basin and adjacent Canadian Arctic Archipelago, a dynamic region with an inherently strong variability in biogeochemical processes. Nonetheless, few of these observations were made in the last five years and the spatial coverage in the latter region is poor. We use a dataset of carbonate system parameters measured in the Canadian Arctic 20 Archipelago (CAA) and its adjacent basins (Canada Basin and Baffin Bay) from 2003 to 2016 to describe the recent state of these parameters across the Canadian Arctic and investigate the amplitude and sources of the system’s variability over more than a decade. Our findings reveal that, in the summers of 2014 to 2016, the ocean surface across our study area served as a net CO2 sink and was partly undersaturated with respect to aragonite in the Canada Basin and the Queen 25 Maud Gulf, the latter region exhibiting undersaturation over its entire water column at certain locations. We estimate, using measurements made across several years, that approximately a third of the interannual variability in surface DIC in the CAA results from fluctuations in biological activity. In consideration of the system’s variability resulting from these fluctuations, we derive times of emergence of the anthropogenic ocean acidification signal for carbonate system 30 parameters in the study area.


Introduction
Ocean acidification and its repercussions on marine ecosystems constitute an important consequence of the ongoing rise in atmospheric carbon dioxide (CO2) concentrations. The world's oceans absorbed approximately one third of the anthropogenic CO2 released to the atmosphere 35 over the last two centuries of industrial activity (Sabine et al., 2004) and are currently a sink for ~24% of global annual anthropogenic carbon emissions (Le Quéré et al., 2018). Atmospheric carbon dioxide uptake by the surface oceans has well defined impacts on seawater chemistry, including a decrease of pH resulting from the dissociation of carbonic acid (H2CO3), the product of the reaction between water and dissolved CO2. A large fraction of the hydrogen ions released 40 by this reaction is neutralized by carbonate ions (CO3 2-), leading to a decrease of their concentration and, concomitantly, the saturation state of seawater with respect to the carbonate minerals calcite and aragonite. The saturation state is defined by: (1) ΩC,A= [Ca 2+ ][CO3 2-] / K * SP where square brackets denote concentrations and K * SP is the stoichiometric solubility product of 45 calcite or aragonite, the two most common marine CaCO3 polymorphs, at a given temperature, pressure and salinity.
The combination of these chemical reactions is most often referred to as ocean acidification (OA). As it proceeds, the dissolved inorganic carbon concentration (DIC; the sum of [H2CO3*], 50 [HCO3 -] and [CO3 2-]) in the surface ocean is expected to increase relative to the total alkalinity (TA; the capacity of a solution to neutralize protons), as the latter is nearly conservative in the surface ocean (Wolf-Gladrow et al., 2007). The global mean surface ocean pH currently sits ~0.1 units below its preindustrial value (Orr et al., 2005) and, according to Earth System models, under the IPCC's "business as usual" RCP8.5 emission scenario, is predicted to decrease by an additional 55 0.3 units by the end of this century (Bopp et al., 2013).
Marine calcifying organisms, many of which are important primary producers (e.g., coccolithophores), extract the constituents of their calcitic or aragonitic tests (shells) from seawater. In most cases, their ability to do so is directly dependent on the saturation state of the surrounding water. Supersaturated seawater (Ω > 1) will favor carbonate precipitation whereas undersaturated seawater (Ω < 1) favors carbonate dissolution. The majority of calcifying organisms, such as planktonic foraminifera and coccolithophores, undergo dissolution or exhibit substantially hindered growth when exposed to undersaturated seawater (e.g., Mostofa et al., 2016). Calcifying organisms found in the Arctic Ocean are subject to rapid environmental changes, as this polar ocean is warming more rapidly than others (Serreze et al., 2011) and is particularly vulnerable to acidification due to the low alkalinity and correspondingly weak buffer capacity of its cold waters (Shadwick et al., 2013). Atmospheric CO2 uptake by Arctic surface waters is further promoted by the rapidly melting seasonal ice cover (e.g., Stroeve et al., 2012), which exposes a gradually larger area of the ocean to gas exchange with the atmosphere and whose melt-water 70 dilutes calcium concentrations, alkalinity and carbonate ion concentrations, further decreasing Ω.
The Canadian Arctic Archipelago (CAA) and its adjacent deep basins, the Canada Basin (CB) and Baffin Bay (BB, Fig. 1), are part of the region projected to undergo the largest reduction in ice cover and, consequently, the largest decrease in surface pH (~0.6) and Ω (1 and 0.7 for ΩC 75 and ΩA, respectively) over the 21 st century (Popova et al., 2014). Recent observations (e.g. Yamamoto-Kawai et al., 2009b;Robbins et al., 2013;Qi et al., 2017) hint at a significant decrease of the aragonite saturation state of surface waters, notably near the continental shelves (Chierici and Fransson, 2009;Bates et al., 2009) as well as a rapid expansion of the undersaturated area in the Canada Basin. Aragonite saturation states remain above saturation but relatively low (ΩA = 1.5 80 -2.5) in the eastern CAA and Baffin Bay (Azetsu-Scott et al., 2010). Most of the time series describing carbonate mineral saturation states predate 2010 and/or do not extend geographically to the CAA and Baffin Bay.
In this study, we use a large observational dataset for this part of the Arctic to 1) describe 85 the recent state of the carbonate chemistry and its spatial variability in the Canadian Arctic Archipelago and adjacent basins, 2) investigate the interannual variability in carbonate system parameters and identify detectable temporal trends using time series spanning from 2003 to 2016 and 3) estimate the contribution of the temporal change in biological activity to the observed variability of surface DIC.

Canada Basin
The Canada Basin (CB), Canadian Arctic Archipelago and Baffin Bay accommodate the flow of surface waters from the North Pacific to the North Atlantic (Stigebrandt, 1984), as well as circulation of Atlantic waters at greater depths. The water mass structure of the southern Canada 95 Basin is representative of these broad circulation patterns and can be summarized as follows MacDonald et al., 1989;Lansard et al., 2012): a relatively cold and fresh surface layer that contains significant fractions of meteoric water (river discharge and precipitation) and sea-ice melt in the summer and becomes homogeneous in winter; an intermediate layer (~50-200m) of advected Pacific waters, often divided into summer and winter 100 varieties, the latter being distinctively rich in nutrients and metabolic CO2 and recognizable by a temperature minimum in the upper halocline; a layer of warm (~0.5°C) and saline (SP>34) Atlantic water; a cold bottom layer with practical salinities (SP) reaching 34.85. The main surface circulation feature in this area, the clockwise Beaufort Gyre, is the largest freshwater reservoir in the northern oceans, formed through Ekman pumping (Proshutinsky et al., 2009). This feature is 105 reversed at depth. The main source of freshwater to the Beaufort Sea (the southwest portion of the Canada Basin) is the Mackenzie River (Carmack and MacDonald, 2002), although the contribution of sea-ice melt is increasing significantly along with the accelerating reduction in ice cover (Yamamoto-Kawai et al., 2009a). The supply of freshwater at the surface, combined with the advection of pre-acidified waters from the Pacific (100-200 m) and the Atlantic (below 400 m; 110 Luo et al., 2016) Oceans result in the presence of three distinct and expanding ΩA undersaturation horizons in the Canada Basin (Wynn et al., 2016).

Canadian Arctic Archipelago
The CAA is a series of islands on the Canadian continental shelf, through which complex Strait, located centrally in the archipelago, inhibits the eastward flow of Atlantic waters, so that 120 only surface and Pacific-origin waters reach Baffin Bay (Bidleman et al., 2007). The properties of these water masses are substantially modified during this transit (McLaughlin et al., 2004)

Baffin Bay
The oceanographic regime of Baffin Bay is distinct from that of the CAA and CB, as it receives multiple inputs from both the Arctic and Atlantic Oceans. Cold and relatively fresh Arctic and Pacific-derived waters enter this 2300-m deep semi-enclosed basin through the Nares Strait as well as Jones and Lancaster Sounds (Muench, 1971;Jones et al., 1998Jones et al., , 2003. Warmer and more 135 saline Atlantic Ocean waters are transported from the Labrador Sea by the West Greenland Current (WGC) into Baffin Bay through the eastern side of Davis Strait, circulate cyclonically, i.e., in an anti-clockwise direction, before joining the southward Baffin Island Current (BIC) which exits Baffin Bay through the western Davis Strait (Bourke et al., 1989, Munchow et al., 2015. Atlantic Ocean waters are modified as they mix with Arctic inflows in Northern Baffin Bay, near the North 140 Water Polynya (Melling et al., 2001). The resulting water mass structure is described by Tang et al. (2004) as: 1) a cold (T<0°C) and relatively fresh (SP<33.7) surface layer, representing the mixed Arctic inputs, 2) a warm (T>0°C) and saline (SP>34) Atlantic Ocean water layer found at depths of ~300 to 800 m, and 3) a deeper layer of nearly constant salinity (SP = 34.5).

Methods
The dataset used in this study comprises data from 420 stations visited during various research cruises carried out aboard the Canadian Coast Guard Ship (CCGS) Amundsen between 150 2003 and 2016. Table 1 summarizes the timeframe and relevant data acquired during each cruise; figure 2 shows the position of each sampling station. Although ice conditions restricted most observations to the summer months, two winter time-series, acquired in 2003(CASES, Miller et al., 2011(CFL, Shadwick et al., 2011 are included in the dataset.

Sampling and measurements
Seawater was sampled separately for each measured parameter from Niskin bottles mounted on a Rosette system equipped with a Seabird SBE 911plus Conductivity-Temperature-Depth (CTD) sensor, which recorded in situ practical salinity (SP) and temperature data throughout the water column. The conductivity/salinity probe was calibrated post-cruise against 160 measurements carried out on discrete seawater samples using a Guildline Autosal 8400 salinometer (accuracy of ± 0.002 or less), itself calibrated with IAPSO standard seawater. Samples used for pH determination were drawn directly from the Niskin bottles into 125 mL low density polyethylene (LDPE) bottles with no headspace to avoid gas exchange with surrounding air and left to thermally equilibrate in a temperature bath set at 25.0 (±0.1) °C. pH, on the total proton 165 scale (pHT), was then measured spectrophometrically on a Hewlett-Packard 8453 UV-visible diode array spectrophotometer using m-Cresol purple (Clayton and Byrne, 1993) and Phenol red (Robert-Baldo et al., 1985) indicators at 434 and 578 nm or 433 and 558 nm, respectively, in a 5cm quartz cell. Daily calibrations were performed with TRIS buffer solutions at practical salinities of 25 and/or 35, depending on the salinity range of the samples. The reproducibility was found to 170 be ±0.005 pH units or better, based on duplicate measurements of the same samples with the same or both indicators. Samples destined for total alkalinity (TA) and dissolved inorganic carbon (DIC) analyses were drawn directly from the Niskin bottles into 250 or 500 mL glass bottles with groundglass stoppers, poisoned with solid mercuric chloride (HgCl2) to halt biological activity, and sealed with Apiezon M grease. TA and DIC from the GEOTRACES 2009 and all post-2010 cruises were measured onboard or at Dalhousie University on a Marianda VINDTA 3C instrument, following the protocol described by Dickson et al. (2007) and calibrated with Certified Reference Materials (CRM) provided by A.G. Dickson (Scripps Institute of Oceanography). The precision of the instrument 180 was found to be ± 2-3 µmol kg -1 based on repeated CRM analyses. The remaining DIC and TA analyses were performed respectively on a SOMMA instrument (Johnson et al., 1993) (Lewis and Wallace, 1998), using the carbonic acid dissociation constants determined by Mehrbach et al. (1973), refit by Dickson and Millero, (1987), the HSO4dissociation constants of Dickson (1990) and the total boron concentration (BT) from Uppström (1974). pCO2

Quality control
In order to assess the robustness of the computed DIC values, we calculated DIC from TA and pHT(25°C) and compared the results with the measured DIC values. The resulting coefficient of determination of the linear fit to the measured and calculated DIC values, R 2 , is 0.989, while the mean difference between calculated and measured DIC values is ~2 µmol kg -1 . We excluded 200 30 measurements that differed from the calculated values by more than 50 µmol kg -1 (2.5% of the mean DIC).
Questionable TA measurements, excluded from the dataset, were identified as those outside a range of 3 standard deviations from the mean salinity-normalized TA for individual 205 regions (CB, CAA, BB) characterized by internally consistent water mass assemblages. TA measurements obtained from the two instrumental methods (VINDTA and Radiometer Titrilab 865) used in 2015 and 2016 were also compared to ensure that data originating from both methods could be used interchangeably in the calculation of additional parameters and conjointly in time series. The resulting coefficient of determination between both datasets (R 2 ) is 0.988, the mean of 210 the non-systematic discrepancy between values is 6 µmol kg -1 and its maximum is 36 µmol kg -1 , respectively corresponding to 0.3% and 1.7% of the mean TA. The degree to which the results of this test are representative of the entire dataset is uncertain, but they constitute the best possible estimate of the uncertainty associated with the use of the two analytical methods used to measure TA. When TA measurements obtained from both methods deviated significantly ( >10 µmol kguniform salinity (Millero, 2005;p.268), was used to determine which data to discard. The deviation from TA values calculated from DIC and pHT(25°C) was used to complement the first method, especially at the surface, although the validity of DIC measurements was previously assessed using TA.

Error estimation
In order to quantify the error associated with the calculated carbonate system parameters reported in this study, we used the CO2SYS program modified by Orr et al. (2018), which applies error propagation to instrumental and constant-related uncertainties. For simplicity, we report the 225 mean uncertainty for each parameter (see Table 2), as the variance is minimal within our dataset.
We found the additional uncertainty associated with the unavailability of nutrient concentrations (P and Si) as input parameters in CO2SYS to be negligible (up to 0.0006 pH units, 1.5 µatm pCO2 and 0.006 Ω units, as determined using nutrient data where available). and Baffin Bay (BB) are presented in Table 3. It is important to note that the mean regional values 240 we report for the Canada Basin may be skewed by the higher density of stations located along the Mackenzie Shelf, and that our sample size for Baffin Bay consists of only 6 stations. Practical salinities considerably below 25 were mostly observed near the mouth of the Mackenzie River, with some in the Queen Maud Gulf (QMG). The discrepancy between SP and TA values observed at the surface and in most of the water column of the CB/CAA and Baffin Bay clearly illustrates 245 the change in water mass regime west of Lancaster Sound (see Sect. 2), while DIC, which is strongly affected by biological activity, shows a less prominent spatial pattern. In all regions, surface-water pCO2, of which we only consider data acquired over the top 5 m of the water column in order to render it more indicative of gas exchange potential, was largely undersaturated with respect to the atmosphere, by as much as 150 µatm (Fig. 3). This suggests that the region as a blurs the saturation threshold in such a way that ΩA values marginally below 1 might in reality represent supersaturated conditions, and vice-versa. It is important to note that, even without the 290 influence of climate change, areas of high riverine discharge naturally harbor lower carbonate mineral saturation states. Thus, undersaturated conditions in the QMG and elsewhere do not solely result from the recent increase of freshwater inputs described above, or air-sea gas exchange.
Nonetheless, changes in the amount of freshwater-driven dilution and atmospheric CO2 uptake affect the degree of this undersaturation as well as its spatial and temporal extent.

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Surface waters throughout the study area are supersaturated with respect to calcite, with ΩC ranges (mean) of 1.34 to 3.25 (2.04), 1.21 to 3.29 (1.97) and 2.38 to 2.70 (2.52) in the CB, CAA and BB, respectively. Uncertainties on ΩC values are on the order of 0.25-0.30, almost twice as large as those of ΩA, due to the larger uncertainty of the calcite stoichiometric solubility product 300 (Mucci, 1983).

Water column observations
Depth profiles of pHT, pCO2 and ΩA grouped by region (equivalent to water-mass regime) are presented in Fig. 7. We divided the broad Canadian Arctic Archipelago region into four sub- The most prominent feature in profiles of carbonate system parameters in the Canada Basin is the Upper Halocline Layer (UHL), a layer of water originating from the Pacific Ocean with a relatively lower pH due to its high metabolic CO2 content (Shadwick et al., 2011). In 2014-2016, the UHL was characterized by a pHT minimum of 7.82 ± 0.03, a pCO2 maximum of ~652 ± 6 µatm (both calculated at in situ temperature) and a ΩA minimum of 0.75 ± 0.16 in the central CB. This 315 pHT minimum migrates upwards from ~180 to ~ 140 m as the UHL encounters the continental shelf west of M'Clure Strait but maintains its amplitude. The presence of such an acidified layer exacerbates the vulnerability of the planktonic communities in this area, as, in addition to the aragonite undersaturation found at the surface, ΩA drops below one at depths of 100 to 125 m or even shallower waters in the Canada Basin. As CO2 naturally diffuses or mixes from the UHL to the overlying waters and the combination of gas exchange and freshening continues to generate undersaturated conditions at the surface, the entire photic zone (where ΩA < 1.5) may acidify and become undersaturated with respect to aragonite at a much faster rate than that of other oceans.
The shallowest subsurface aragonite saturation horizon we observe in the central Canada Basin was found at ~85 m in 2014. Our data do not corroborate the interpretation of Wynn et al. (2016) 325 that the upper boundary of the UHL is migrating downwards due to an expansion of the overlying Polar Mixed Layer (PML), at least not during the period of our observations. The ΩA crosses the saturation threshold back to supersaturation between 200 and 250 m, where Atlantic waters become predominant, as evidenced by a +0.8 °C temperature maximum at depths of 400 to 500 m (in contrast to a temperature minimum of -1.5 °C in the UHL). ΩA and pHT remain, respectively,  The Amundsen Gulf and the western portion of the Parry Channel (Fig. 7b, f, j) exhibit a similar water mass structure and carbonate system chemistry as the Canada Basin, as the dominant circulation pattern pushes water eastward from the CB to the CAA. Undersaturation with respect to aragonite does not occur at the surface in these areas, owing to higher salinities. Although the amplitudes of the ΩA, pHT and pCO2 excursions are slightly smaller than those found in the CB, 340 the UHL is considerably shallower in the western CAA. Consequently, ΩA falls below one at depths of 50 to 70 m, and the upper portion of the water column in those parts of the CAA might become undersaturated with respect to aragonite even more rapidly than in the CB. As reflected by the blue lines in Fig. 7 (b, f, j), the UHL becomes progressively less discernable on depth profiles as it undergoes modification and mixing during its transit from the CB to Lancaster Sound.

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Atlantic waters are found at the bottom of the water column in the Amundsen Gulf and Parry Channel. The saturation maxima at ~400 m (ΩA ~ 1.1 to 1.4) are significantly lower in these areas than at similar depths in the CB.
the layer of Pacific water (UHL) and the overlying surface water. ΩA is close to 1 at the surface and reaches a maximum of ~1.45 (± 0.16) at ~30 m, before dipping back below 1 between 50 and 75 m, like in other parts of the CAA. With the exception of one profile that exhibits a strong pCO2 maximum of 685 (± 6.49) µatm as well as pHT and ΩA minima of 7.80 (± 0.03) and 0.76 (±0.16), respectively, at a depth 125 m, this UHL feature becomes less prominent in the central CAA.

355
Degradation of settling organic matter (remineralization) might explain this exceptional peak of greater amplitude than those observed in the Canada Basin at Station 310 in September 2016.
Stratification becomes significantly weaker in the shallow waters (20-100 m) of the Queen Maud Gulf (Fig. 7c, g, k), despite freshwater addition theoretically strengthening stratification, as it isone of the areas with the strongest tidal mixing in the CAA (McLaughlin et al., 2004). Practical salinity 360 profiles in this area do not show values exceeding 30 at the bottom of the water column, significantly lower than the values above 32 observed at the same depths in the central CAA. The residence time of waters in this area might also be relatively high, due to its geographical isolation from the main channels of the CAA, possibly allowing more mixing to take place. The combined effects of low salinities from freshwater accumulation (mostly from river discharge) and the 365 efficient redistribution of CO2 through mixing result in low carbonate mineral saturation states throughout the water column. The QMG is the only region in our study area where, at some locations, the entire water column is undersaturated with respect to aragonite, making it an ideal location to study the effects of such conditions on aragonitic organisms. through Nares Strait, the warm and saline Atlantic water inflow from the Labrador Sea dominates the water-mass structure in the region and accounts for the high alkalinity of these waters relative to the CB and CAA (Münchow et al., 2015). In Baffin Bay, waters become 380 undersaturated with respect to aragonite and calcite at depths of ~600 m and ~1400 m, respectively. The pCO2 and pHT increase and decrease proportionally to each other with depth.

385
Of the 420 stations that make up our dataset, twenty-four were visited at least on two different years and match our comparability criteria for time-series. These criteria are: 1) the stations were sampled within 31 calendar days of each other (this criterion is not ideal since seasonality is highly variable and driven by complex sea-ice processes, including ice break-up) and 2) the stations are located within a 5 km radius. The mean time difference and distance between To quantify near-surface change, we averaged data from the top 25 m of the water column.
Across this depth interval, between 2007 and 2016, the temperature-normalized pCO2, DIC and 405 DIC/TA at site LS1 rose by 36 ± 9 µatm, 37 ± 10 µmol kg -1 and 0.008 ± 0.005, while pHT and ΩA decreased by 0.042 ± 0.037 and 0.12 ± 0.23, respectively. Given its uncertainty, the change in ΩA is insignificant. The same trend is visible between the same years at station CAA1, at a greater magnitude (+ 78 µatm pCO2, -0.088 pHT unit, -0.37 ΩA unit), except for a DIC decrease of 58 µmol kg -1 caused by a decrease in salinity of 1.76. The DIC/TA, which effectively normalizes DIC 410 against salinity, shows an increase of 0.021, proportional to the change in the other parameters.
Data from the nearby station CAA2 also displays a similar trend between 2009 and 2015 (+ 29 µatm pCO2, -0.041 pHT unit, +0.001 DIC/TA), albeit a strong positive change in salinity (+1.81) contributed to an increase of 0.06 ΩA unit (again, insignificant given its uncertainty of 0.23).

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Four stations located on the transect extending from Cape Bathurst to Banks Island complete the seven time-series of at least five years (Fig. 11). Three of those time series (AM1, AM2, AM4) span 2003-2009 or 2004-2009. Thus, given their spatial and temporal proximity, we expect a certain consistency in the trends they exhibit. Surprisingly, salinity increased at each of the stations over the study period, by 1.46, 0.70 and 0.13, respectively. Despite the consistent 420 salinity trend, stations AM2 and AM4 show opposite trends in carbonate system parameters, the former exhibiting a positive change in pCO2 (+ 49 ± 9 µatm) and DIC/TA (+ 0.010 ± 0.005) as well as a negative change in pHT (-0.089 ± 0.037) and ΩA (-0.20 ± 0.23), while the latter displays negative but somewhat smaller changes in pCO2 (-35 ± 9 µatm) and DIC/TA (-0.007 ± 0.005) as well as a positive change in pHT (+0.051 ± 0.037) and ΩA (+0.13 ± 0.23), mostly apparent near the 425 surface (<10 m). The two profiles recorded at site AM2 exhibit a clear difference in stratification; a sharp peak in carbonate system parameters (associated with a temperature maximum) appears at for which the mechanism is unclear, the latter year displaying considerably less acidic conditions (+ 0.25 pHT unit), while salinity and temperature profiles are nearly identical for the two years, suggesting no change in water masses. Total alkalinity measurements at this depth range in 2008 455 are substantially higher than those made in 2004, for similar salinities, but neither of these data appear disputable.
As previously stated, these time series are snapshots in time and cannot be assumed to represent the continuous evolution of the carbonate chemistry in the Canadian Arctic. Nonetheless, 460 even with a small sample size, we can confidently state that the temporal evolution of carbonate system parameters in the region does not display a systematic trend on sub-decadal timescales.
Moreover, most of the significant changes that our time series exhibit are associated to variations in the physical oceanography of the region (water-mass distribution and circulation) or surface processes (melting of sea-ice). Given the well-documented rapid melting of the sea-ice cover in 465 the region (e.g., Tivy et al., 2011), we did not expect to observe increases in summer surface salinity over time intervals of 5 to 9 years. Our time series, therefore, offer proof of the strong interannual variability of this highly dynamic system. Discerning the ocean acidification signal amid the various physical and biological sources of change would require continuous time series over a longer period of time. We estimated this period to 23 to 35 years for pH, 25 to 37 years for 470 pCO2, 31 to 46 for ΩA and 118 to 177 years for DIC using calculations of Time of Emergence, or the time required for the effects of a process to emerge from the natural variability of a system (see Appendix A). Time of emergence calculations are usually performed with large, continuous datasets from climate models; we therefore do not consider these results to be statistically robust.
Nonetheless, this exercise shows the relative variabilities of the carbonate system parameters and 475 highlights the particularly strong variability of DIC, which is the object of the next section. Its results also imply that without accurate measurements of the effects of biological activity and seaice processes (both major drivers of natural variability), direct detection of the ocean acidification signal will require at least 20-25 years of observations. Gathering data in the CAA in the next few years is therefore critical, as regular ship-based observational campaigns in the region started in 480 the early 2000's (Giesbrecht et al., 2014).

The biological contribution to interannual DIC variations: ΔDIC Bio
We define ΔDICBio as the change in the contribution of in situ biological activity (photosynthesis and respiration) to the DIC of a parcel of water over a given period of time.

485
ΔDICBio is calculated for each sampled depth at recurrently visited stations according to: where DICRef is computed in CO2SYS using the temperature-normalized seawater pCO2 calculated at a reference time, adjusted to the time of interest assuming a constant air-sea pCO2 gradient, and the TA measured at this time of interest. The change in global mean atmospheric 490 CO2 concentrations between the reference year and the year of interest is used to correct pCO2 to account for gas exchange (data from Dlugokencky and Tans, NOAA/ESRL, www.esrl.noaa.gov/gmd/ccgg/trends/). This approximation rests on the assumption that the yearly increase in surface water pCO2 follows that of the atmosphere (given stable biological production), as observations from global monitoring stations demonstrate (e.g., González-Dávila et al., 2010), 495 although the validity of this claim is weakened on short spatial and temporal (sub-decadal) scales (Fay et al., 2013, Wanninkhof et al., 2013. This also restricts our calculations to the upper portion of the water column (25 m) that is in direct contact with the atmosphere. Under the additional assumptions that DIC is only affected by gas exchange, biological activity and mixing, and that TA is not significantly affected by biological activity (Zeebe and Wolf-Gladrow, 2001), DICRef 500 represents the DIC of a parcel of water if its in situ biological component remained unchanged relative to a reference year (i.e., identical contribution, negative or positive, from the balance between photosynthesis and respiration). Because the reference pCO2 is calculated in part from TA, changes in water masses should not affect the results of this analysis, given the salinity range of the data subset used in the calculation of ΔDICBio (25.6 < SP < 33.7). Thus, ΔDICBio can provide 505 insights into the interannual variability of biological activity in the Canadian Arctic, without direct measurements of parameters such as chlorophyll or biomass. The uncertainty on ΔDICBio was calculated by applying standard error propagation to the procedure described above. It is important to note that this calculated uncertainty is purely mathematical and does not include the uncertainty associated with the assumptions made to calculate ΔDICBio, such as the constancy of the air-sea 510 ΔpCO2, which could be considerably affected by processes like changes in sea-ice cover. The extremely weak (r = 0.08) correlation between ΔDICBio and the time interval over which it applies provides additional evidence of the absence of a trend in the balance between 540 photosynthesis and respiration in the surface waters of the Canadian Arctic. The variability in this balance is driven by many interconnected, often localized processes. For instance, short-lived episodes of upwelling of halocline waters not only directly change the chemical properties at the surface, but also provide nutrients that stimulate biological activity .
Primary production in the Arctic is also closely linked to the seasonal cycle of sea-ice (e.g., Arrigo As previously mentioned, variations in water mass composition cannot directly explain variations 550 in ΔDICBio. Nonetheless, mixing is likely accompanied by changing nutrient concentrations, which influence the balance of photosynthesis and respiration (Tremblay et al., 2015).

Conclusions
Field observations of carbonate system parameters made between 2014 and 2016 in the Canadian Arctic reveal that surface waters of the region serve as a net CO2 sink in the summer and are generally close to saturation with respect to aragonite (1 < ΩA < 1.5). Surface undersaturation (ΩA < 1.0) is found predominantly in the central Canada Basin, as documented in previous years (Yamamoto-Kawai et al., 2011;Robbins et al., 2013), and in the freshwater-influenced Queen Time series of carbonate system parameters, although relatively short (<10 years) and incomplete, illustrate the strong interannual variability of the region, due in part to complex circulation patterns and varied water mass assemblages. Our estimates of ΔDICBIO, the change in the contribution of biological activity to DIC, suggest that variations in biological activity (the 570 balance between photosynthesis and respiration) account for approximately a third of the interannual variability of DIC measurements. Additional work must also be carried out to extend this estimate to other carbonate system parameters (pH, pCO2, Ω). In order to test the validity of the ΔDICBIO concept and its underlying assumptions, our results should be compared to direct measurements of biological productivity (e.g. biomass, chlorophyll-a) during the same time period.

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Without the latter, the quantification of the progression of ocean acidification in the surface waters of the Canadian Arctic will require longer and more continuous time series, the length of which can be estimated using the concept of Time of Emergence. Future work on ocean acidification in this region should focus on obtaining continuous time series of carbonate system parameters, especially in areas where surface waters might soon become undersaturated with respect to 580 aragonite, as well as bridging the gap between observations of carbonate mineral saturation and markers of ecosystem health.

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The time of emergence (ToE) of a process affecting a natural system is the time required for the measurable effects of this process to emerge from the natural variability of the system. The concept is predominantly applied in global climate change modeling studies, for which the results are either "years of emergence" based on a pre-industrial steady-state (e.g., Friedrich et al., 2012) or time intervals over which observations must be made in order to distinguish an anthropogenic 590 signal from its natural variability. Few of these studies have used observations (e.g., Sutton et al. 2016), and, to our knowledge, none of them have focused specifically on the Arctic.
We define the time of emergence according to the following equation:  (Table A1) are in general agreement with the values reported by Bates et al. (2014). C is a constant that sets the threshold of emergence at either 2 or 3 standard deviations (N), i.e., when the acidification signal becomes significant beyond natural variability as it emerges from 95% or 99.7% of the observed annual mean values, assuming the data (or the 605 naturally occurring values of the parameter they represent) are normally distributed. Only data collected from June to October, inclusively, are used to minimize the effect of seasonal variability.
Although the assumption of relative equilibration with atmospheric pCO2 might be less applicable at these depths, we applied the same technique using data from depths of 100 m and 300 m in order 610 to estimate the time of emergence below the surface, where the interannual variability should be smaller than at the surface, if no important changes in water masses occur. Significant acidification at these depths might be due in large part to advection (e.g., Luo et al., 2016) rather than gas exchange.

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The results of this analysis are presented in Table A2. Despite their similarity, our calculated times of emergence are consistently longer than those reported in modelling studies (Keller et al., 2014;Rodgers et al., 2015). This is consistent with the fact that coastal waters, that comprise a large portion of our dataset, exhibit a much higher 635 variability in pH (and other carbonate system parameters) than open oceans (Duarte et al. 2013).
Furthermore, direct observations are likely to integrate variability on temporal and spatial scales that are too small to be resolved by models. It is also important to note that distinct measurement

Estimation of the freshwater sources in the Queen Maud Gulf
In order to estimate the relative fractions of sea-ice melt and meteoric water (mostly river water) used by Lansard et al. (2012), the fractions of meteoric (river) water and sea-ice melt would be, respectively, 98% and 2%. A potential source of error affecting this estimate is the use of the δ 18 O 670 of Mackenzie River water as the riverine end-member, which might differ significantly from the oxygen isotope signature of the rivers discharging in the Queen Maud Gulf.

Data availability
The raw data collected as a part of the ArcticNet program, on which most of the observations 675 presented in this paper are based can be accessed through the Polar Data Catalogue (Mucci, 2017).
Complementary datasets, some of which are part of larger databases, are also available on various online repositories (François et al., 2012;Chierici et al., 2013;Giesbrecht et al., 2014;Papakyriakou et al., 2017). We would like to thank the Captains and crew of the CCGS Amundsen without whom, over the years, this project would not have been possible. This project was funded through View Software (Schlitzer, 2016).