The North Atlantic subpolar gyre (NASPG) is a major site for formation of intermediate and deep waters and thus plays a key role in the ocean meridional overturning circulation. The upper ocean circulation brings to its southern and eastern parts relatively warm and salty water of subtropical origin. This water is then cooled by significant heat loss to the atmosphere and freshened by local excess precipitation and by inputs of fresher water from the Arctic, as well as from continental and ice cap origin (Boyer et al., 2007). Part of this upper water then flows into the Nordic Seas, whereas the other part recirculates cyclonically in the gyre, steered by topography, such as around the Reykjanes Ridge (Fig. 1). This water is progressively transformed by winter mixing into intermediate waters in the Labrador and Irminger Sea or further entrained in the dense outflows of the Nordic Seas to form Atlantic deep waters (Mercier et al., 2015; Daniault et al., 2016; Rossby et al., 2017).

This region is the only large part of the world ocean which has experienced a surface cooling trend over the last century (Rahmstorf et al., 2015). The cooling has been related to changes in the meridional overturning circulation and to an observed overall surface freshening (Friedman et al., 2017). It also experiences very large decadal to multi-decadal variability (Yashayaev and Loder, 2016; Reverdin, 2010; Frajka-Williams et al., 2017; Robson et al., 2016) associated with Atlantic multi-decadal variability. This might result from atmospheric variability, as well as from changes in the strength of the meridional overturning circulation (Häkkinen and Rhines, 2004; Häkkinen et al., 2011; Hátún et al., 2005; Reverdin, 2010; Chafik et al., 2016). The most recent trend has been a large cooling and freshening since 2005, which reversed a previous warming and freshening since the mid-1990s (Robson et al., 2016). It was associated in 2014–2015 with particularly strong positive North Atlantic Oscillation (NAO) atmospheric conditions inducing large vertical mixing and deep convection in the Labrador Sea and the Irminger Sea (Yashayaev and Loder, 2016; Piron et al., 2017; de Jong and de Steur, 2016; Fröb et al., 2016).

The North Atlantic contributes substantially to the global oceanic uptake of CO2. This is mainly due to extensive biological activity during summer and considerable heat loss during winter, as well as due to the export of surface waters to the deep ocean by the ocean circulation and vertical mixing. As a result, a large anthropogenic carbon inventory is evaluated in this region (e.g., Khatiwala et al., 2013; Zunino et al., 2014). Takahashi et al. (2009) estimate that the annual mean air–sea CO2 flux in the North Atlantic, north of 50∘ N (representing only 5 % of the ocean surface), is 0.27 Pg yr-1, i.e., almost 20 % of the global flux. Although the mean annual carbon flux is a robust result for the North Atlantic (Takahashi et al., 2002, 2009; Watson et al., 2009; Schuster et al., 2013) there is still disagreement in the magnitude of seasonal, interannual to decadal variability depending on the method used to evaluate air–sea CO2 fluxes (Schuster et al., 2013). Compared to other basins, the air–sea CO2 fluxes' interannual variability in the North Atlantic appears relatively small (Rödenbeck et al., 2015; Landschützer et al., 2016). However, during some periods, significant variability has been recognized at a regional scale in the North Atlantic subpolar gyre (NASPG). It is related to changes in ocean fCO2 relative to atmospheric concentrations induced by either warming or deep convection (Corbière et al., 2007; Metzl et al., 2010; Rödenbeck et al., 2014). The variations of winds also impact air–sea CO2 fluxes (Wanninkhof and Trinanes, 2017). Based on a synthesis of pCO2 observations for years 1972–2006, Takahashi et al. (2009) evaluate a mean rate of 1.8 µatm yr-1 (±0.4) in the North Atlantic, i.e., close to atmospheric increase. This result was revisited and confirmed by McKinley et al. (2011) for the period 1981–2009. An analysis of recent data (2005–2014) across the NASPG near 59–60∘ N also suggests surface fCO2 trends that are near the atmospheric increase (Fröb et al., 2018). Interestingly, this study also illustrates a spatial variety of the mechanisms for these trends. For example, a large contribution of alkalinity changes related to a salinity decrease was found only in the Iceland Basin in surveys by VOS Nuka Arctica (Friederike Fröb and Are Olsen, personal communication, 2018).

The uptake of CO2 through air–sea gas exchange affects the seawater CO2 chemistry and leads to ocean acidification (Gattuso and Hansson, 2011). Over the past two decades, pH in North Atlantic surface waters has declined at a similar rate to global ocean pH (-0.0018 yr-1) (Lauvset et al., 2015; García-Ibáñez et al., 2016). However, in a similar way as for fCO2, the pH interannual variability could be significant at a regional scale. As an extreme case, based on the shorter 2001–2008 SURATLANT (SURveillance ATLANTique) winter data, Metzl et al. (2010) reported a pH rate of -0.0069 yr-1 associated with a particularly fast rise of oceanic fCO2 (up to 7.2 µatm yr-1).

Finally, the large uptake of anthropogenic CO2 in the ocean leads to a strong change in the isotopic composition of dissolved inorganic carbon (DIC), thus reducing its 13C/12C ratio (noted δ13CDIC hereafter). The δ13CDIC decrease in response to the human-induced perturbation is less than -0.007 ‰ yr-1 in polar surface regions (McNeil et al., 2001; Olsen et al., 2006), whereas it reaches -0.025 ‰ yr-1 in subtropical regions (Gruber et al., 1999). These observations have been used to validate oceanic models for these regions (Tagliabue and Bopp, 2008; Sonnerup and Quay, 2012). This decrease is generally masked by the seasonal cycle due to physical and biological mechanisms, which is as large as 1 ‰ in the NASPG (Gruber et al., 1999; Racapé et al., 2014). Although δ13CDIC provides additional information allowing us to further understand mechanisms of ocean region CO2 uptake in key regions such as NASPG, its interannual to decadal variability is still poorly documented.

In order to unravel the surface NASPG variability, it is necessary to revisit the region studied by Corbière et al. (2007), Metzl et al. (2010), and Racapé et al. (2014) with a more comprehensive data set. The data set encompasses temperature (T), salinity (S), water stable isotopes (δ18O and δD), dissolved inorganic carbon, total alkalinity (At), nutrients, and δ13CDIC. The 24 years of physical and geochemical data presented here cover the recent multi-decadal sea saw of warming–cooling. Hence, they might also provide an alternate view on the decadal variability in fCO2 and pH in the central part of the NASPG. fCO2 and pH are thus also computed from these data.

The data have been binned in time and in latitude bands for later use. The data and the binning method are first presented (Sect. 2). To illustrate the properties of the data set, we will then present the average seasonal cycle (Sect. 3). Finally, we will present first-order estimates of trends over the whole period (Sect. 4).

In this section, we will present the average seasonal cycle portrayed in the 2001–2017 data (Fig. 3). For T (not presented), one finds a seasonal cycle range increasing from north (4 ∘C) to south (12.3 ∘C in yellow box over the shelves and 6.4 ∘C offshore in red box; corresponding boxes are on the top panel of Fig. 3). This is close to the ranges portrayed in SST (sea surface temperature) climatologies, such as HadSST3 (Kennedy et al., 2011). The seasonal cycle in S (Fig. 3) is also found as described at the surface in the World Ocean Database 2013 (WOD13; Boyer et al., 2013) with a maximum in February–May (February–April on shelf) followed by a gradual decline until a minimum in September (October, north of 62∘ N). The amplitude also increases from north to south.

For the other parameters, the seasonal cycle presented has been normalized as described in Sect. 2.2.1 and reported for S=35.

The DIC seasonal cycle presents (Fig. 3) a maximum in March and a minimum in July (north) transitioning to August (south). The amplitude is at a maximum on the Newfoundland shelf (73 µmol kg-1) and north of 62∘ N (70 µmol kg-1), and it is smallest for 58–62∘ N (48 µmol kg-1). There is a slight decrease from March to April, indicating the beginning of carbon consumption during blooms, and a very steep decline in April to May.

At (Fig. 3) has much smaller seasonal variability than DIC (here expressed as deviation from average relationship At–S), with estimated uncertainties of several µmol kg-1, which is not small compared with the seasonal cycle amplitude (uncertainties are larger on the shelf in January–May). After a minimum in February, there is a peak in April in the interior and later (May–June) north of 62∘ N and on the Newfoundland shelf, followed by a decrease until August north of 54∘ N (and later further south). The At increase could be associated with late winter or early spring blooms reducing NO3 and PO4 for the formation of organic matter and decreasing DIC. The latter At decrease could be associated with calcifying organisms such as coccolithophores, which are known to produce large late spring or early summer blooms in this part of the Atlantic, usually well past the large diatom-dominated blooms (Signorini et al., 2012)

The δ13CDIC seasonal cycle (Fig. 3) updates Racapé et al. (2014). It mirrors the cycle in DIC, with a maximum δ13CDIC in July in the north shifting to August further south (and even early September on the shelf). This increase of 0.7 ‰ or more compared to late winter is associated with production of organic matter and the associated fractionation. The decrease later in the year is associated with remineralization and vertical mixing with depleted subsurface water (Racapé et al., 2013). Notice also a small spatial gradient in winter with increasing values from north to south (and decreasing salinity).

The three nutrients present a March maximum (as for DIC) associated with maximum entrainment of subsurface water in the mixed layer. Then, they present a slight decrease until April and a larger decrease until a July minimum for Si and an August minimum for PO4 and NO3. This shift in the time of the minima between the three nutrients suggests a dominance of non-siliceous organisms in the later portion of the bloom, when Si levels have been strongly depressed. For PO4 and NO3 there is a north to south decrease in the nutrients (but not so much for Si) in all seasons. The southward decrease is even stronger for NO3 on the Newfoundland shelf, a sign of the contribution of fresher water from the Pacific Ocean and western Arctic having experienced denitrification on the shelves (McTigue et al., 2016). In particular, summer NO3 levels are very low south of 54∘ N. The seasonal cycle of the three nutrients is consistent with climatologies available from WOD13 and also with the time series station in the Irminger Sea close to Iceland (Olafsson et al., 2010).

δ18O and δD (not shown) present no significant seasonal deviations from their average relationship with S north of 58∘ N. At 54–58∘ N, there is an early winter minimum and a maximum in April–May, during a period with overall small salinity seasonal variability. South of 50∘ N, in the shelf region, one finds a late summer maximum as described near the shelf break in Benetti et al. (2016), where it was related to sea ice meltwater. The 50–54∘ N latitude range presents an intermediate situation between the seasonal cycles in these two regions with a spring to September maximum.

Finally, the fCO2 and pH seasonal cycles are estimated from monthly mean DIC and At, as well as SST, SSS (sea surface salinity), and nutrients (Fig. 4) Here, we do it without normalizing to S=35, in order to compare fCO2, pH, and DIC seasonal cycles with the climatology constructed by Takahashi et al. (2014). Except for the southern region (46–50∘ N), fCO2 presents a pronounced maximum in February–March associated with vertical mixing and entrainment of remineralized DIC in the surface layer. It presents a minimum in June associated with the carbon use during the spring bloom. Similarly to DIC described above, the fCO2 seasonality is most pronounced in the most northern box (amplitude 90 µatm) and in the most southern box (80 µatm from January to May). In both regions, the fCO2 seasonal cycle exhibits a secondary maximum (August) and minimum (November), as in the climatology. In the central regions (54–58 and 58–62∘ N) the seasonal fCO2 amplitude is on the order of 40 µatm. The fCO2 values in these two boxes are very close for each month, as was also found for DIC and nutrients (Fig. 3). The oceanic fCO2 values are close to (in December–March in the north) or well below the atmospheric level (Fig. 4a). The largest ocean CO2 sink is observed in the southern region in May (fCO2ocean-fCO2atm=-110 µatm) as found in the climatology (Takahashi et al., 2009, 2014) and regularly observed in recent years (5 May cruises in 2004–2015, Wanninkhof and Pierrot, in Bakker et al., 2016). The pH seasonal cycle (Fig. 4b) mostly mirrors (with reverse sign) the seasonal cycle in fCO2 (Fig. 4a), and its amplitude ranges between 0.04 (in the gyre) and more than 0.1 (in the south and north).

For fCO2, At, DIC and pH seasonal cycles, we present (Fig. 5) comparisons to the climatological seasonal cycle (Takahashi et al., 2014) for the two boxes in the central gyre (54–58 and 58–62∘ N) where we observed homogeneous properties (Fig. 3). There are a few methodological differences between the two products. The climatology (Takahashi et al., 2014) is reported for reference year 2005, whereas the SURATLANT seasonal cycle constructed with 2001–2017 data is referenced to year 2010 (Sect. 2.2.1). The climatology for DIC and pH was calculated from fCO2 observations and reconstructed alkalinity, whereas for SURATLANT, it is fCO2 and pH that are computed from observed DIC and At. At 56∘ N (green lines), the fCO2 (pH) climatology is low (high) compared to SURATLANT, but with rather similar seasonality. At 60∘ N (blue lines), fCO2 and pH seasonality are stronger in the climatology with a more pronounced fCO2 minimum and pH maximum in June–July, and results are similar in other seasons (August–May). For both regions, the At seasonal amplitude is most pronounced in the SURATLANT data (about 10–15 µmol kg-1, Fig. 5b), with the largest difference in August. Despite these differences, the DIC cycles derived from independent observations and methods are very similar (Fig. 5c), which is typical of all regions.

We show the trends over the full period (1993–2016), using all data or restricted to the latitude range 50–63∘ N, and separately in summer (June–September) or winter (January–March). In this latitude range, At does not show any trend (Fig. 6), as well as salinity or nutrients (not shown). In contrast, we find positive trends for DIC (∼+0.79, 0.76 µmol kg-1 yr-1) respectively for winter and summer (notice no winter data included in 2016 or 2017) (Fig. 7). This trend is about half the one reported by Olafsson et al. (2009), based on winter observations in the northern Iceland Sea for years 1985–2008 (+1.4 µmol kg-1 yr-1), but close to the trend observed over 1990–2015 in the Irminger Sea upper ocean waters (+0.63 µmol kg-1 yr-1), a signal mainly explained by anthropogenic carbon uptake (Fröb et al., 2018). As also suggested by δ13CDIC data (described below), the positive DIC trend derived from SURATLANT data (Fig. 7) is likely due to anthropogenic CO2. Also, interestingly, in both seasons, the SURATLANT data present almost no trend until 2005, which corresponds to SST (and AMO) maximum. This lack of a DIC trend in the early part of the record was also reported in Metzl et al. (2010). This first part of the record also corresponds usually to a period of decreasing winter winds (decrease in frequency of the NAO+ situation) and thus an expected decrease in winter mixed layer depths (and also a decrease in subpolar gyre circulation).

For fCO2 positive trends are found as expected (Fig. 8). In the band 50–63∘ N, the summer (+2.1 µatm yr-1) and winter (+1.76 µatm yr-1) trends are close to the atmospheric increase. This is in the range of the long-term trend (25–30 years) estimated in the North Atlantic by Takahashi et al. (2009) and McKinley et al. (2011), but much lower than values near or above +3 µatm yr-1 previously reported for years 1993–2006 (Corbière et al., 2007; Schuster et al., 2009) or 2001–2008 (Metzl et al., 2010). This is also much larger than the +1.1 µatm yr-1 trend estimated by Lauvset et al. (2015) for years 1991–2011 in the NASPG. This highlights that the fCO2 trend is quite sensitive to the period (and data) selected (Fay and McKinley, 2013). Also, when using all data (all seasons and regions) the fCO2 trend is +1.9 µatm yr-1 (Fig. 9a) so that the δ-fCO2 (difference between ocean and atmospheric fugacities) presents no significant trend (Fig. 9b), suggesting that in this region the air–sea CO2 fluxes driven by δ-fCO2 (and winds) would also have presented a small trend. Although temperature interannual variations (up to +1.5 ∘C in 2005, or -1.5 ∘C in 2015; see Fig. B1) could explain rapid fCO2 changes for some periods (Corbière et al., 2007), over the longer term the fCO2 trends presented here are mainly explained by DIC (with At being relatively constant). The same is true for pH (Fig. 10) and its negative trend of -0.0021 yr-1 mirrors the fCO2 trend. Similarly to the fCO2 trend, this pH trend for the NASPG is close to the mean global ocean estimate of -0.0018 yr-1 (Lauvset et al., 2015). It is also comparable to other trends evaluated in the North Atlantic polar waters, ranging between -0.0017 and -0.0026 yr-1 depending on the periods, seasons, and regions (Bates et al., 2014; Lauvset and Gruber, 2014; Lauvset et al., 2015; Olafsson et al., 2009). Compared to the Irminger Sea, the difference of pH trends is mainly explained by observed DIC trends (+0.8 µmol kg-1 yr-1 for NASPG against +1.6 µmol kg-1 yr-1 for Irminger Sea, Bates et al., 2014).

For δ13CDIC, we find significant trends in all seasons for the 2005–2017 period for which samples are available. The winter trend (-0.014 ‰ yr-1) is smaller than the summer trend (-0.042 ‰ yr-1) (Fig. 11). Both are small compared to the seasonal cycle and have large uncertainties, due to the small number of years sampled. Notice also that the summer season presents large seasonal changes (Fig. 3). Furthermore, the trend estimates strongly depend on the adjustments that we apply separately to the data before and after 2010 (see Appendix A3). The trends have a similar magnitude before and after 2010, although the summer trend is sensitive to the large positive deviations of August 2010 data (see Racapé et al., 2014), with lower trends in 2010–2017 when not including it (-0.0040 vs. -0.042 ‰ yr-1). Altogether, the surface winter trend deduced from the adjusted data set is lower than the expected Suess effect in the atmosphere (-0.025 ‰ yr-1, based on current change rates in atmospheric δ13CO2, such as from the Alert station, White et al., 2015). It is also comparable with model estimates in the NASPG for other periods from Sonnerup and Quay (2012; -0.12 ‰ decade-1 over the period 1970–1995) and from Tagliabue and Bopp (2008; -0.10 ‰ decade-1 between 1970 and 2005).