Sea–air CO2 flux in the North Atlantic subtropical gyre: Role and influence of Sub-Tropical Mode Water formation

Deep-Sea Research II 91 (2013) 57–70

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Sea–air CO2 flux in the North Atlantic subtropical gyre: Role and influence of Sub-Tropical Mode Water formation Andreas J. Andersson a,b,n, Lilian A. Krug a,c, Nicholas R. Bates a, Scott C. Doney d a

Bermuda Institute of Ocean Sciences, Ferry Reach, St. George’s GE01, Bermuda Scripps Institution of Oceanography, University of California San Diego, La Jolla, CA 92093, USA c University of Algarve, Faro, Portugal d Woods Hole Oceanographic Institution, Woods Hole, MA, USA b

a r t i c l e i n f o

abstract

Available online 21 February 2013

The uptake of atmospheric carbon dioxide (CO2) into the mid-latitudes of the North Atlantic Ocean through the production of wintertime Sub-Tropical Mode Water (STMW) also known as Eighteen Degree Water (EDW) is poorly quantified and constrained. Nonetheless, it has been proposed that the EDW could serve as an important short-term sink of anthropogenic CO2. The objective of the present investigation was to determine sea–air CO2 gas exchange rates and seawater CO2 dynamics during wintertime formation of EDW in the North Atlantic Ocean. During 2006 and 2007, several research cruises were undertaken as part of the CLIMODE project across the northwest Atlantic Ocean with the intent to study the pre-conditioning, formation, and the evolution of EDW. Sea–air CO2 exchange rates were calculated based on measurements of atmospheric pCO2, surface seawater pCO2 and wind speed with positive values denoting a net flux from the surface ocean to the atmosphere. Average sea–air CO2 flux calculated along cruise tracks in the formation region equaled  18 76 mmol CO2 m  2 d  1 and  14 7 9 mmol CO2 m  2 d  1 in January of 2006 and March of 2007, respectively. Average sea–air CO2 flux in newly formed outcropping EDW in February and March of 2007 equaled  287 10 mmol CO2 m  2 d  1. These estimates exceeded previous flux estimates in this region by 40–185%. The magnitude of CO2 flux was mainly controlled by the observed variability in wind speed and DpCO2 with smaller changes owing to variability in sea surface temperature. Small but statistically significant difference (4.17 2.6 mmol kg  1) in dissolved inorganic carbon (DIC) was observed in two occurrences of newly formed EDW in February and March of 2007. This difference was explained either by differences in the relative contribution from different water masses involved in the initial formation process of EDW or temporal changes owing to sea–air CO2 exchange ( 25%) and vertical and/or lateral mixing ( 75%) with water masses high in DIC from the cold side of the Gulf Stream and/or from below the permanent thermocline. Based on the present estimate of sea–air CO2 flux in newly formed EDW and a formation rate of 9.3 Sv y (Sverdrup year¼ 106 m3 s  1 flow sustained for 1 year), CO2 uptake by newly formed EDW may constitute 3–6% of the total North Atlantic CO2 sink. However, advection of surface waters that carry an elevated burden of anthropogenic CO2 that are transported to the formation region and transformed to mode water may contribute additional CO2 to the total net uptake and sequestration of anthropogenic CO2 to the ocean interior. & 2013 Elsevier Ltd. All rights reserved.

Keywords: Carbon dioxide Sea–air CO2 flux Gas exchange Sub-Tropical Mode Water STMW Eighteen degree water EDW North Atlantic Ocean

1. Introduction In order to improve predictions on the future concentration of carbon dioxide (CO2) in the atmosphere, and its impact on climate over the next century, it is important to accurately identify both anthropogenic and natural sources and sinks of CO2. At present, the global ocean is a sink for anthropogenic CO2 from the atmosphere with uptake rates estimated at 1.4 to 2.3 Pg C y  1

n Corresponding author at: Scripps Institution of Oceanography, University of California San Diego, La Jolla, CA 92093-0202, USA. Tel.: þ 1 858 822 2486; fax: þ 1 858 534 7313. E-mail address: [email protected] (A.J. Andersson).

0967-0645/$ - see front matter & 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.dsr2.2013.02.022

(Pg¼1015 g C; Le Que´re´ et al., 2009; Takahashi et al., 2009) or roughly 15–26% of the anthropogenic CO2 released to the atmosphere. Despite major research efforts to characterize temporal and spatial variability of sea–air CO2 flux across the world oceans (Doney et al., 2009a,b; Takahashi et al., 2002, 2009), there are still many potentially important areas and regions that have not been adequately described and many relevant physical/biological mechanisms that are not fully understood. At present, the North Atlantic Ocean is believed to act as a significant net sink of CO2 with cold temperatures and strong winds during wintertime favoring a substantial ocean uptake of atmospheric CO2 (Bates, 2007; Bates et al., 2002; Takahashi et al., 2002, 2009; Wanninkhof et al., 2010; Watson et al., 2009).

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Analysis of repeat observations between 1993 and 2003 suggests that the uptake of anthropogenic CO2 in the North Atlantic Ocean for this time period was approximately 0.19 Pg C y  1 (Wanninkhof et al., 2010). Sea–air CO2 flux estimates based on measurements of surface seawater pCO2 from volunteer observing ships (VOS), which include both anthropogenic and natural background components, yielded a net influx of 0.2570.05 Pg C y  1 between 101N and 651N in 2005 (Watson et al., 2009). Over the past few decades, much attention has focused on the ocean uptake of CO2 during the wintertime formation of deep and intermediate waters in the sub-polar gyre of the North Atlantic Ocean (e.g.,  et al., 2007; Friis et al., 2005; Lefe vre et al., 2004; Olsen Corbiere et al., 2006; Omar and Olsen, 2006; Pe´rez et al., 2008, 2010; Schuster and Watson, 2007). Less attention has been given to the CO2 uptake associated with formation of Sub-Tropical Mode Water (STMW) in the mid-latitudes of the North Atlantic Ocean just south of the Gulf Stream (Bates et al., 2002; Gruber et al., 2002; Levine et al., 2011). The rates of formation of North Atlantic STMW and uptake of CO2 are poorly quantified and constrained (Bates et al., 2002; Marshall et al., 2009). During wintertime, pulses of cold air passing over the Gulf Stream (GS) and the northwest North Atlantic subtropical gyre causes the largest net heat loss observed anywhere in the world ocean, sometimes exceeding 1000 W m  2 (Marshall et al., 2009). The loss of heat from the surface ocean to the atmosphere and subsequent cooling creates gravitational instability, which combined with vertical and lateral shears from the GS (Joyce et al., 2009) result in a deepening of the mixed layer to depths greater than 400 m with subsequent isopycnal advection and mixing that emplaces STMW throughout the subsurface ( 250 to 450 m deep) of the subtropical gyre. Characteristically this homogenous water mass has a potential density (sy) in the range 26.45– 26.55 kg m  3 and a modal temperature close to 18 1C, and it is commonly referred to as Eighteen Degree Water (EDW; Worthington, 1959). The cooling and high winds associated with the EDW formation region may facilitate a significant ocean uptake of CO2 by increasing the gas solubility of CO2 in seawater and increasing the sea–air CO2 gas transfer rate, respectively. The formation of EDW may also entrain and subduct water that has taken up anthropogenic CO2 both within and outside of the formation region (Levine et al., 2011), which depends on the active modes of different water masses involved in the formation process and whether mixing is mainly vertical and/or lateral driven by gravitational or shear instabilities, respectively (Joyce et al., 2009). During springtime, the EDW is restratified and capped by a seasonal thermocline that prevents any further modification of the total dissolved inorganic carbon (DIC) concentration through entrainment with surface waters and gas exchange with the atmosphere (Bates et al., 2002). The formation and dissipation of STMW within the subtropical gyre of the North Atlantic Ocean are of significant climatic importance because of the large heat exchange associated with its formation (Marshall et al., 2009). STMW is also apparently important in regulating biogeochemical processes in the subtropical gyre of the North Atlantic Ocean by setting the nutrient balance (Bates and Hansell, 2004; Jenkins and Doney, 2003; Palter et al., 2005, 2008). Furthermore, Bates et al. (2002) proposed that STMW (i.e., EDW in the North Atlantic Ocean) could act as an important short-term sink for storing anthropogenic CO2 in the relatively shallow depths of the ocean. Based on observations from the Bermuda Atlantic Time-series Study (BATS) site off Bermuda in the 1990s, Bates et al. (2002) reported that seawater DIC content in the EDW increased twice as fast as the rate of DIC increase in surface seawater, which was similar to the expected rate required to maintain equilibrium with increasing atmospheric CO2 concentration. Extrapolating the observed extra

increase in DIC to global STMW resulted in a storage term for CO2 that was equivalent to as much as 3–10% of the annual oceanic net uptake of CO2. Based on a mass balance approach, the annual increase in EDW DIC of 2.5570.25 mmol kg  1 y  1 was attributed to three causes; (1) uptake of 0.9 mmol kg  1 y  1 of anthropogenic CO2 from the atmosphere; (2) uptake of 0.5570.17 mmol kg  1 y  1 owing to biological processes including remineralization of organic matter and CaCO3 dissolution (Gruber and Sarmiento 2002), and; (3) uptake of 1.1070.25 mmol kg  1 y  1 owing to gas exchange. Bates et al. (2002) estimated that sea–air CO2 flux of  5 to  10 mmol CO2 m  2 d  1 at the site of EDW formation in wintertime for a 3 months period could account for an increase of 1–2 mmol kg  1 y  1 in the EDW DIC inventory. However, the actual rates of CO2 gas flux during the formation of North Atlantic Ocean EDW had at the time of their analysis never been measured or calculated from direct measurements of relevant surface seawater CO2 parameters. The Climate Variability and Predictability (CLIVAR) Mode Water Dynamic Experiment (CLIMODE) project was a multidisciplinary research project with the objective of investigating atmospheric and oceanic interaction in the western North Atlantic Ocean, and rates and mechanisms controlling the formation of EDW in 2006 and 2007 (Marshall et al., 2009). With the CLIMODE project providing a physical framework, investigation of seawater CO2 uptake and dynamics associated with EDW formation was the subject of the present investigation. Some of the key questions motivating this investigation include: (1) what is the sea–air CO2 flux in the EDW formation region during wintertime, (2) what is the relative contribution from sea–air CO2 gas exchange to the DIC content of EDW; (3) what is the main process(es) setting the initial DIC content of EDW before it is subducted into the subtropical gyre; and (4) how much atmospheric CO2 is taken up by EDW every year? The results presented here mainly focus on the sea–air CO2 exchange observed in the wintertime of 2006 and 2007, but we also discuss the DIC inventory of newly formed EDW formed in the winter of 2007. As a context for the present study, sea–air CO2 gas exchange rates measured in wintertime in the region of EDW formation are compared with North Atlantic subtropical gyre conditions in summertime and time-series observations collected at the BATS site over the past decade.

2. Methods 2.1. Data collection During 2006 and 2007, several research cruises were undertaken as part of the CLIMODE field campaign across the northwest subtropical North Atlantic Ocean (Fig. 1) with the intent to study the pre-conditioning, formation rates, and the evolution of EDW as this water mass dispersed throughout the subtropical gyre. Multiple hydrocasts were conducted and sampled for physical and biogeochemical parameters at a number of locations and along several transects crossing the Gulf Stream. In order to study sea–air CO2 gas exchange across the region of wintertime EDW formation, a number of different approaches and instrumentations were used to measure the surface seawater partial pressure of CO2 (pCO2sw). During the winter of 2006 (19–31 January), the R/V Atlantis was equipped with a Submersible Autonomous Moored Instrument (SAMI, Sunburst sensors, MT), which continuously measured the pCO2 in seawater (73 matm) collected via the ship’s clean seawater flow-through system. The SAMI measurements were validated and corrected by discrete measurements of DIC and total alkalinity (TA) combined with temperature (T) and salinity (S) from the ship’s underway system. Regrettably, during the winter of 2007 (R/V Knorr, 4–28 February and 2–21 March), the SAMI failed and surface seawater pCO2 was only calculated from discrete

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Fig. 1. Cruise tracks and locations where sea–air CO2 flux was calculated from surface seawater CO2 parameters during CLIMODE cruises in 2006 and 2007 in the northwest Atlantic Ocean. Surface seawater pCO2 was directly measured using underway pCO2 systems during the winter of 2006 (blue line) and summer of 2007 (red line). In the winter of 2007 surface seawater pCO2 was calculated from discrete samples of DIC and TA collected every 3 h from the ship’s underway seawater system (black symbols). A few selected dates are shown next to the cruise tracks for reference. Stations 30 (Feb 23, 2007), 48 (Mar 11, 2007) and 8 (June 5, 2007) are highlighted because newly formed EDW was observed at these locations. The location of the Bermuda Atlantic Time-series Study (BATS) is also indicated. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

measurements of DIC and TA (and T and S) collected from the ship’s seawater flow-through seawater system approximately every 3 h during the second leg of the cruise (i.e., when it became apparent that the SAMI did not work). Seawater samples were collected in 200 ml Kimax brand glass bottles using Tygon tubing. The bottles were rinsed extensively before allowing the bottles to overflow by at least one and a half bottle volume (Dickson et al., 2007). Special care was taken to avoid turbulence and sampling of any air bubbles. Each sample was fixed with 100 ml saturated solution of mercuric chloride (HgCl2) in order to preserve the sample and prevent production of CO2 from biological activity. The same sampling procedure was exercised taking samples from Niskin bottles on the CTD rosette. The samples were subsequently shipped and analyzed at the Bermuda Institute of Ocean Sciences (BIOS). Seawater samples were analyzed for DIC using a VINDTA 3C (Marianda Inc.) and a UIC 5012 CO2 coloumetric system (UIC Inc.) while TA was based on potentiometric acid titrations (0.1 N HCl) using a VINDTA 2 S (Marianda Inc.) titration system. The performance and accuracy of both instruments were evaluated using Certified Reference Materials (CRM) provided by A.G. Dickson at the Scripps Institution of Oceanography. The standard deviation of replicate analysis of CRMs conducted during the analysis of underway samples collected during the March cruise for DIC was 71.9 mmol kg  1 (n ¼6) and 76.6 mmol kg  1 (n¼ 6) for TA. The standard deviations of CRMs analyzed in conjunction with analysis of samples from hydrocasts were 72.4 mmol kg  1 (n ¼19) for DIC and 73.0 mmol kg  1 (n ¼23) for TA. Surface seawater pCO2 was calculated using CO2SYS (Lewis and Wallace, 1998) with carbonic acid stoichiometric dissociation constants originally defined by Mehrbach et al. (1973) and refit by Dickson and Millero (1987). Propagation of the uncertainties associated with underway samples in the calculation of seawater pCO2 resulted in an error of approximately 5.4–5.8% or about 20 matm (Dickson, 2010; Dickson and Riley, 1978). During the summer of 2007 (2–14 June, R/V Atlantic Explorer), direct measurements of surface water pCO2 were obtained with

an automated underway system (General Oceanics model 8050 pCO2 Measuring System) based on a showerhead equilibrator and infrared (IR) detector. The system was regularly calibrated every 4 h with four CO2-in-air standard gases containing 0 (pure N2), 250, 370, and 450 ppm CO2. A LiCOR non-dispersive infrared gas analyzer (model 6252) was used to measure the CO2 concentration in the equilibrated air, and converted to pCO2 after correction for the gas equilibration and water pressure. CO2-in-air standards were provided by the Climate Monitoring and Diagnostic Laboratory National Oceanographic and Atmospheric Administration (CMDL/NOAA) and calibrated with World Meteorological Organization (WMO) standards. The precision for each measurement cycle of equilibrated headspace air was 70.4 matm. For all cruises, underway systems onboard measured temperature and salinity while anemometers and barometers collected wind and barometric pressure data. In order to place the 2006 and 2007 CO2 flux data into context, surface seawater pCO2 at BATS from August 1999 to December 2009 was calculated based on monthly measurements of temperature, salinity, DIC and TA (e.g., Bates, 2007) using CO2SYS (Lewis and Wallace, 1998) as previously described. Reanalysis data of daily averaged barometric pressure at sea level (pbaro) were obtained at 2.51 spatial resolution from the National Climate and Environment Program/National Center for Atmospheric Research (NCEP/NCAR) website (http://www.esrl.noaa.gov/psd/data/gridded/ data.ncep.reanalysis2.pressure.html). Windspeed data at 10 m above sea surface were based on daily or 3-day averaged QuickSCAT scatterometer data with a spatial resolution of 0.25 degree (ftp:// ftp.ssmi.com/qscat/). Since the QuickSCAT program was launched in August 1999, the calculations for the CO2 flux at BATS were conducted from this starting date. The error of wind speeds measured by QuickSCAT is typically on the order of approximately 1 m s  1 (e.g., Ebuchi et al., 2002). Atmospheric CO2 concentrations (pCO2atm) in the study region were obtained from Globalview, a product provided by the Cooperative Atmospheric Data Integration Project maintained by

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the National Oceanic and Atmospheric Administration Earth System Research Laboratory (NOAA/ESRL; http://www.esrl.noaa. gov/gmd/ccgg/globalview/co2/co2_description.html). Globalview is the data repository of the NOAA global atmospheric observational network of stations that collect weekly measurements of atmospheric CO2. For this study, atmospheric CO2 data from Tudor Hill tower located on the island of Bermuda (321290 N 0 64152 W) were used for the closest days matched up with cruises and BATS measurements. Comparisons of atmospheric CO2 were also made with two other stations adjacent to the study area; aircraft measurements in Massachusetts, USA (42.951 N, 70.631 W) and a tower in Grifton, North Carolina, USA (35.351 N, 77.381 W). The maximum difference in average concentrations of atmospheric CO2 during January 2006 between the three stations was 1.29 mmol mol  1. However, larger spatial and temporal gradients in atmospheric CO2 may be associated with variations in the anthropogenic signature of air masses coming off the North American continent. Based on measurements and numerical modeling of atmospheric CO2 at Grifton, NC, the synoptic variability at this location during wintertime is approximately 74 mmol mol  1 (71 standard deviation relative to the monthly average; Geels et al., 2004). Measurements of atmospheric CO2 every 3 h on the rim reef of Bermuda in January and February of 2011 reveal occasional excursion of 76 mmol mol  1, but the monthly standard deviation is less than 72 mmol mol  1 (http://www.pmel.noaa. gov/co2/story/Hog%2BReef). Hence, based on these data it is reasonable to assume that the average uncertainty associated with the atmospheric CO2 concentration adopted from Globalview in the study region may be in the order of 74 mmol mol  1.

2.2. Calculation of sea–air CO2 gas exchange rates The flux (F) of CO2 (mmol m  2 d  1) across the ocean– atmosphere interface where a positive value denotes flux from sea to air was calculated by the equation: F ¼ kUaUDpCO2

ð1Þ

where k is the CO2 transfer velocity coefficient, a is the CO2 solubility in seawater and DpCO2 is the difference between surface seawater and atmosphere partial pressure of CO2:

DpCO2 ¼ pCO2sw pCO2atm

ð2Þ

Wanninkhof (1992) calculated the relationship between wind and gas transfer velocity for instantaneous wind measurements from shipboard anemometers or scatterometer data as: k ¼ 0:31u2



 Sc 0:5 660

ð3Þ

where k is in cm h  1, 0.31 is the gas transfer scaling factor (G), u is wind velocity at 10 m (m s  1) and Sc is the Schmidt Number (dimensionless), which is the ratio of the molecular viscosity for seawater over the molecular diffusivity of that specific gas in seawater and is strongly dependent on the sea surface temperature (for CO2, Sc ¼2073.1  125.62Tþ3.6276T2  0.043219T3). The solubility of CO2 (mol l  1 atm  1) in seawater (a) was given by Weiss (1974) according to Zeebe and Wolf-Gladrow (2001) as:  9345:17 T 60:2409þ 23:3585 ln a ¼ exp T 100 "   #! T 2 þ S 0:0235170:00023656T þ 0:00047036 ð4Þ 100 where T is sea surface temperature (in Kelvin) and S is salinity. To obtain pCO2atm, assuming the air right above sea surface has 100%

humidity (Zeebe and Wolf-Gladrow, 2001): pCO2atm ¼ ½CO2 atm Uðpbaro pH2 o Þ

ð5Þ

where [CO2]atm is the atmospheric concentration of CO2, pbaro is the barometric pressure at sea level and pH20 is the saturation vapor pressure of water (Zeebe and Wolf-Gladrow, 2001),   6745:09 T 4:8489 ln 0:000544S ð6Þ PH2 o ¼ exp 24:4543 T 100 where T is sea surface temperature (in Kelvin) and S is salinity. The error associated with the flux calculations in the present study varies for each cruise due to the different approaches used to measure seawater pCO2. For daily average fluxes, the largest error was associated with the DpCO2 term arising from the uncertainty due to seawater and atmospheric pCO2 estimates. This error ([(error pCO2sw)2 þ(error pCO2atm)2]1/2) was large for the March cruise ( 720 matm) due to the fact that pCO2 was calculated from DIC and TA and smaller for the January (75 matm) and June ( 74 matm) cruises where surface seawater pCO2 was directly measured. These errors in DpCO2 translate into a random error in the sea–air CO2 flux of 52%, 13%, and 10%, respectively. The error associated with the sea–air CO2 transfer coefficient (k) is a function of (wind speed)2, but also systematic errors associated with the parameterization of the scaling factor (G) (e.g., Sweeney et al., 2007; Takahashi et al., 2009; Wanninkhof et al., 1992). The error associated with the scaling factor alone may contribute to an error of 730% in the sea–air CO2 flux (Stanley et al. 2009; Sweeney et al., 2007; Takahashi et al., 2009). Although errors in the scaling factor may be correlated to systematic errors in wind speed, the wind and scaling factor errors of the present study were considered uncorrelated following the reasoning of Takahashi et al. (2009). The fractional error associated with wind speed translates into a fractional error in sea–air CO2 flux multiplied by 2. Hence, an error in wind speed of 5% translates into an error in the CO2 flux of 10%. The errors associated with the ship-based anemometers are unknown, but are likely smaller than estimates based on QuickSCAT ( 1 m s  1). Hence, we assume an error of 10% in the sea–air CO2 flux associated with random errors in wind speed estimates. In order to estimate the total error (u{total}) associated with sea–air CO2 flux in the present study, the individual percentage errors discussed here were calculated for any particular flux estimate and combined according to: 2

2

2 2

uftotalg ¼ ðufDpCO2 g þ ufscaling factorg þ ufwind g Þ1=2

3. Results 3.1. CO2 gas flux in the North Atlantic subtropical gyre The most dominant feature of the sea–air CO2 fluxes observed in the EDW formation region and across the North Atlantic subtropical gyre in 2006 and 2007 was the general influx of CO2 from the atmosphere to the ocean (Fig. 2). The largest mean uptake of CO2 along cruise tracks was measured in January followed by March and then June (Table 1; Fig. 2). However, the average sea–air CO2 flux measured in June across a latitudinal gradient between 261N and 371N was significantly lower compared to the average fluxes measured in the EDW formation region during January and March. The largest influx of CO2 was 104734 mmol CO2 m  2 d  1 observed in January. However, this magnitude influx was unusual and only persisted for a few hours. It was more than twice as large as the flux measured at any other time (Fig. 2). The large influx was driven by low seawater pCO2 and consequently large DpCO2 (o 150 matm) during a period of intermediate wind strengths (15–18 m s  1; Fig. 3).

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Fig. 2. Calculated CO2 flux (mmol m  2 d  1) along the cruise tracks shown in Fig. 1 in (A) January 2006, (B) March 2007, and (C) June 2007. Gray color indicates CO2 fluxes in the proximity of the Gulf Stream. Positive values indicate efflux and negative values indicate influx. See Section 2 for a detailed description on how the relevant parameters required to calculate the CO2 flux (e.g., pCO2, T, S, u) were measured or calculated and the error associated with these estimates. The vertical black lines in the lower right corner of each panel indicate the average uncertainty in CO2 flux for each cruise. Table 1 Average and maximum ocean-atmosphere CO2 fluxes during CLIMODE cruises and at BATS in 2006 and 2007. The 10-year average flux measured at BATS is also presented for comparison. Cruises

CO2 sea–air flux (mmol m  2d  1) Average

Maximum influx

Maximum efflux

January 06 March 07 June 07 4351N o351N

 17.6 7 6  14.2 7 9  3.2 7 1  7.4 7 2  1.7 7 0.6

 103.77 34  46.17 28  21.97 7  21.97 7  16.47 6

þ 10.07 3

þ 0.47 0.1

BATS

Average 10-year climatology

Influx 2006, 2007

Date of measurement

January March June

 13.2 7 7  9.9 7 5  0.27 0.1

In comparison, the maximum influxes observed in March and June equaled  46728 mmol CO2 m  2 d  1 and  2277 mmol CO2 m  2 d  1, respectively. However, episodes with influxes exceeding  40 mmol CO2 m  2 d  1 were regularly observed in both January and March. These large CO2 fluxes were for the most part associated with strong wind events. Fig. 3 illustrates how the CO2 gas flux responded to variations in the independent parameters temperature, wind speed squared, and DpCO2 during the different cruises. This analysis demonstrated that the magnitude gas flux in 2006 and 2007 was most strongly controlled by the observed variations in wind and DpCO2 with little or weak influence from the observed variations in temperature (Fig. 3). The spatial distribution of surface seawater pCO2 in January of 2006 appeared in general to be higher on the cold side (i.e., north) of the Gulf Stream relative to the warm side (Fig. 4). Surface seawater pCO2 was in general lower than the atmosphere with DpCO2 typically ranging from close to 0 to 70 matm, although there were instances

 17.27 9  8.07 4  0.47 0.2

þ 0.47 0.1

29-January-06 6-March-07 20-June-07

when DpCO2 was lower than 150 matm. Sea–air CO2 fluxes were spatially variable, but the largest CO2 fluxes were observed on the warm side or in the core of the GS (Fig. 4). In March of 2007, surface seawater pCO2, DpCO2 and sea–air CO2 flux were also spatially variable (Fig. 5). Seawater pCO2 was in general lower relative to the atmosphere with DpCO2 ranging from just above 0 to 120 matm in the GS region and the northern Sargasso Sea. Both strong and weak influxes of CO2 were measured within very short spatial and temporal scales of kilometers and hours, respectively (Fig. 5). This was a result of large variations in wind speed during this time of the year linked to pulses of cold-air and low-pressure systems moving across the region. The largest CO2 fluxes were again observed within or on the warm-side of the GS, and also in the northern extent of the Sargasso Sea (Fig. 5). In June of 2007, surface seawater pCO2 and sea–air CO2 influx followed a distinct geographic pattern, the two properties decreasing and increasing, respectively, with increasing latitude and

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Fig. 3. Calculated sea–air CO2 flux (mmol m  2 d  1) during each cruise in January 2006 (top), March 2007 (middle), and June 2007 (bottom), as a function of sea surface temperature (SST 1C), wind speed (m s  1), and sea–air DpCO2 (matm). Least squares linear regressions indicate that there exists no or weak correlation between CO2 flux and SST (r2 ¼0.03, 0.02, 0.29), intermediate to strong correlation between CO2 flux and (wind speed)2 (r2 ¼ 0.47, 0.51, 0.81), and intermediate or weak correlation between CO2 flux and DpCO2 (r2 ¼0.37; 0.14; 0.46).

decreasing temperature (Fig. 6). Some variability from this pattern was observed with respect to pCO2 over relatively short distances (Fig. 6). DpCO2 ranged from small positive values just above 0 to about  45 matm. CO2 fluxes ranged from intermediate negative fluxes at high latitudes and weak negative fluxes or even positive fluxes at low latitudes (Table 1). 3.2. CO2 gas flux at BATS In general, the sea–air CO2 flux measured at BATS varied seasonally from a relatively small efflux of CO2 during summer and early fall (June to October), with maximum fluxes on the order of þ271 to þ5 72.5 mmol CO2 m  2 d  1, to a larger influx during winter and spring (November to May), with maximum fluxes well exceeding 10 mmol CO2 m  2 d  1 (Fig. 7). Based on the 10-year time series of CO2 flux at BATS, the average sea–air CO2 flux in the summer and early fall (June to October) ranged from 0 to þ2.571.3 mmol CO2 m  2 d  1 while the flux for the remainder of the year ranged from  371.5 to 13 77 mmol

CO2 m  2 d  1 (Fig. 7). The greatest CO2 influx was observed in January and the greatest CO2 efflux in August. Integrated annually, the location of the BATS and the ocean region this location represents acted as significant sinks of CO2. Compared to the average CO2 fluxes observed along the cruise tracks of the 2006 and 2007 CLIMODE cruises, the climatological flux averaged over a 10-year period at BATS was generally lower (Table 1). The flux recorded at BATS in January 2006 was similar to the average flux recorded during the January CLIMODE cruise. In contrast, the fluxes at BATS in March and June were significantly lower compared to the average fluxes recorded during these months in the GS region during the CLIMODE cruises (Table 1). 3.3. Formation and capping of EDW During February of 2007 formation of EDW was observed at the end of the first leg on the R/V Knorr (Joyce et al., 2009; Marshall et al., 2009). The formation event occurred following a cold air outbreak in the region with maximum heat fluxes

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Fig. 5. Spatial distribution of (A) surface seawater pCO2 (matm), (B) DpCO2 (matm), and (C) sea–air CO2 flux (mmol m  2 d  1) in the Gulf Stream region between 341N and 401N in March of 2007. Contour lines indicate sea surface temperature (1C). The approximate location of the northern edge of the Gulf Stream is highlighted by the gray line coincident with the most northern 18 1C isotherm. Fig. 4. Spatial distribution of (A) surface seawater pCO2 (matm), (B) DpCO2 (matm), and (C) sea–air CO2 flux (mmol m  2 d  1) in the Gulf Stream region between 361N and 401N in January of 2006. Contour lines indicate sea surface temperature (1C). The approximate location of the northern edge of the Gulf Stream is highlighted by the gray line coincident with the 18 1C isotherm.

of 1000 W m  2 on February 20 (Joyce et al., 2009). A transect perpendicular to the GS showed no evidence of surface outcropping of EDW in its southern extent on February 19 (stn 21– 25; Fig. 8), but evidence of outcropping and newly formed EDW was observed in the northern part of the transect on February 21 and 23 (stn 26–28, 30; Fig. 8). Note that the stations in the southern and the northern section of the transect were separated in time because of SeaSoar surveys conducted during the formation process in order to unravel the physical fine-scale details of this process (Joyce et al., 2009). Evidence of formation of EDW was clearly evident at station 30 (N39.25951 W54.09401) on February 23 with EDW extending to a depth of 500 m (Fig. 9;

Table 2). Nevertheless, the surface layer down to this depth still appeared weakly stratified (Figs. 9 and 10). On March 11, a vertical profile with properties similar to those found at Stn. 30 on February 23 was observed at N38.36471 W55.95371 approximately 100 nautical miles southwest of the previous location (Fig. 9). EDW extended to a depth close to 500 m. Compared to February 23, temperature was somewhat colder, salinity slightly lower, and potential density a little bit higher (Table 2). Essentially no density gradient was observed from the surface to 500 m depth at this location (Figs. 9 and 10). Average TA remained the same as on February 23, but both DIC and DO were slightly higher. The observed difference in means of DIC (4.172.7 mmol kg  1) was statistically significant (KruskalWallis non-parametric 1-way analysis of variance; p o0.01). The region of EDW formation was revisited in June of 2007. At this time, the EDW layer had been capped by a seasonal thermocline extending to a depth of about 200 m (Fig. 9). The EDW layer was observed just below the thermocline and extending to

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Fig. 6. Spatial distribution of (A) surface seawater pCO2 (matm), (B) DpCO2 (matm), and (C) sea–air CO2 flux (mmol m  2 d  1) in the North Atlantic subtropical gyre between 251N and 391N in June of 2007. Contour lines indicate sea surface temperature (1C). The approximate location of the northern edge of the Gulf Stream is highlighted by the gray line coincident with the most northern 21 1C isotherm.

Fig. 7. Sea–air CO2 flux (positive values indicate efflux; negative values indicate influx) at the Bermuda Atlantic Time-series Study (BATS) between 1999 and 2009 based on monthly measurements of surface seawater DIC and TA. The gray line indicates the average flux for each month between 1999 and 2009 and the black line the flux measured for each month.

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Fig. 8. (A) Location of transect across the Gulf Stream overlain on sea surface temperature obtained from satellite data in February of 2007 that coincided with a cold air outbreak and large heat fluxes (  1000 W m  2) on February 20 and formation of EDW (Joyce et al., 2009). Stations 21–25 were sampled on February 19, stations 26–28 on February 21, and station 30 on February 23 (station 21 is the most southern station). Because the transect was interrupted in time and space, no data interpolation has been conducted between stations 25 and 30. Latitude-depth cross sections for the transect for (B) temperature (1C), (C) DIC (mmol kg  1), and (D) dissolved oxygen (DO;mmol kg  1). Isopycnals defining the EDW (26.45 kg m  3 o sy o 26.55 kg m  3) are shown in each section by the dashed lines. Red numbers at the top of panels (B), (C) and (D) indicate station numbers.

Fig. 9. Seawater physical–chemical properties of newly formed Eighteen Degree Water (EDW) in February (green) and March (blue) of 2007, and also in June of the same year (red) after the EDW has been capped by the seasonal thermocline. (A) temperature (1C), (B) salinity, (C) potential density (kg m  3), (D) DIC (mmol kg  1), (E) TA (mmol kg  1), and (F) dissolved oxygen (DO;mmol kg  1).

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Table 2 Average EDW (sy ¼26.45–26.55 kg m  3) properties (7 1 std) in February, March and June of 2007.

Depth range T(1C) S ry (kg m  3) DIC (mmol kg  1) TA (mmol kg  1) DO (mmol kg  1)

02/23/07

03/11/07

06/06/07

0–499 m 17.9 70.1 36.560 70.015 26.49 70.02 2093.9 73.2 2395.87 2.8 218.5 72.6

0–449 m 17.7 7 0.02 36.5357 0.002 26.53 7 0.005 2098.0 7 3.7 2397.27 3.5 222.5 7 0.9

252–605 m 17.97 0.2 36.5657 0.017 26.51 70.04 2098.8 72.2 2400.67 5.1 214.6 76.5

Fig. 10. T–S plot from station 30 (circles) and 48 (squares) on February 23 and March 11, 2007, respectively. The insert represents the data range enclosed by the box in the upper right corner that encompasses EDW.

a depth greater than 600 m. The top boundary of the EDW layer was not as distinct as in the wintertime and subtle erosion was evident at the top of the pycnostad. Most notable was that the whole pycnostad had heaved downward in the water column. Relative to the vertical profile of DIC observed in March, the surface layer was significantly depleted in DIC. However, a corresponding increase in DIC was observed at the approximate depth of the seasonal thermocline (Fig. 9). This observed increase in DIC was accompanied by a decrease in DO. Notably, both physical and chemical properties of the EDW in June were very similar to observations in March and February (Fig. 9; Table 2). 3.4. CO2 dynamics in newly formed EDW The average DIC of the newly formed EDW in February and March of 2007 (Table 2) was somewhat lower than the average DIC of old EDW (2105.6711.0 mmol kg  1) observed upstream in the GS and capped by the seasonal pycnocline. Consequently, old EDW was reset during the formation process due to mixing with low carbon surface seawater while anthropogenic CO2 was taken up via gas exchange and entrainment of water masses with an elevated anthropogenic signature (Levine et al., 2011). We did not explicitly track the spatial path and Lagrangian temporal evolution of the EDW water column sampled in February, and therefore can not say definitively if the two observed occurrences of newly formed EDW were formed independently in different areas of the Gulf Stream, formed simultaneously in the same location or even if they were the same water mass observed at different times. For this latter scenario to be the case, a minimum advective transport at 0.13 m s  1 in a southwesterly direction would have been required (Fig. 1), which cannot be completely ruled out. We start the analysis for the case where the two occurrences of newly formed EDW formed independently from each other and followed different trajectories. In such a scenario, the relatively small variability observed in physical and chemical properties between

the two observations suggest that a similar suite of processes and active modes of mixing broadly controlled EDW formation though with some local-scale variations in the magnitude and relative contributions of different processes and/or water masses. Alternatively, if in fact the two observations of newly formed EDW were formed simultaneously or were the same water mass, the EDW observed on February 23 (or a water mass with similar properties) could have acted as the precursor to the EDW observed on March 11. In such a scenario, the observed temporal changes in seawater physical and chemical properties between the two observations would be due to processes occurring along its Lagrangian pathway rather than during the initial formation process. Thus, to account for this possibility, in the subsequent exercise, it was assumed that the EDW of February 23 served as the starting point (t¼ 0) and that any changes observed in the EDW at a later date represented time dependent changes. Regardless whether the water masses in reality were formed independently, this assumption provides an opportunity to evaluate the potential importance of how different processes (e.g., gas exchange, mixing, biological processes including primary production and respiration) could modify the DIC inventory of the EDW while it remained in contact with the atmosphere. 3.4.1. Increasing DIC Integrating the observed average increase in DIC of 4.172.7 mmol kg  1 between February 23 and March 11 between the surface and the base of the mixed layer corresponded to a vertical inventory increase of 2.171.4 mol of C m  2 or 124782 mmol C m  2 d  1. Hence, despite a small increase in the seawater concentration of DIC, this increase translates to a substantial increase in the total DIC inventory of the EDW because of the large depth of the mixed layer (  500 m). Conversely, a small error in the measurement of DIC produces a large error in the estimate of the total inventory. Nevertheless, the observed difference in mean DIC between the two time periods was statistically significant at the 99% confidence interval (Kruskal-Wallis non-parametric 1-way ANOVA). Hence, we might pose the question: what processes can account for the observed increase in DIC? The total change in DIC (dC/dt) can be evaluated by considering changes owing to: 1) uptake or release of CO2 through gas exchange (Cgasex); 2) physical mixing processes with a water mass with different DIC (Cmix), and; 3) biological processes (Cbio) such as primary production and respiration consuming and generating CO2, respectively, as well as calcification and CaCO3 dissolution i.e., dC=dt ¼ C gasex þC mix þ C bio 3.4.2. Gas exchange The surface seawater pCO2 of the newly formed EDW observed in February and March 2007 varied little (335.472.0 matm). Hence, the magnitude exchange of CO2 between the EDW and the atmosphere was mainly controlled by fluctuations in wind speed and a smaller effect from variations in atmospheric pCO2 owing to variations in air pressure and CO2 concentration of the air mass (  74 mmol mol  1) over this time period. Based on the average EDW DpCO2 and the average wind speed squared (from QuickSCAT data) between February 23 and March 11, the average sea–air CO2 flux equaled 28.5 710 mmol CO2 m  2 d  1 or 23% of the observed increase in the total EDW DIC inventory of 124782 mmol m  2 d  1 (Fig. 11). The average calculated flux for the EDW formation time period and location was thus significantly higher than the average magnitude of the fluxes calculated along the cruise tracks in January and March (1876 and 1479 mmol CO2 m  2 d  1, respectively). However, maximum fluxes along the cruise track were at times calculated to

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Fig. 11. Sea–air CO2 flux (black line) and associated error (gray dashed lines) calculated for outcropping EDW between February 23 and March 11. Wind speeds used in the flux calculation were based on average daily QuickSCAT measurements.

be as high as 104 mmol CO2 m  2 d  1 and 46 mmol CO2 m  2 d  1 in January and March, respectively, but none of these fluxes were sustained for synoptic periods longer than for several hours to about a day.

3.4.3. Mixing During the formation process of EDW, mixing is mainly vertical owing to gravitational instabilities and/or lateral along sloping isopycnal surfaces as a result of shear instabilities, i.e., slantwise convection (Joyce et al., 2009). According to Joyce et al., this issue, large cross-frontal fluxes across the GS north wall are required to account for the regional salinity structure in the EDW formation region, which implies that large lateral flux divergence of DIC is also likely and contributing to the DIC inventory of the EDW (Fig. 8). Water masses either from the cold side of the Gulf Stream or from below the permanent thermocline have higher DIC content than the EDW. These water masses also are less saline and have lower TA than the EDW. Thus, both vertical and lateral mixing would tend to lower salinity and TA while increasing the DIC of the EDW. Indeed, the EDW observed on March 11 showed statistically significant (Kruskal-Wallis non-parametric 1-way ANOVA) lower salinity and higher DIC than observed on February 23 (Table 2). No statistical difference was observed with respect to TA, the error associated with this measurement exceeding the observed change. To assess the potential influence of vertical or lateral mixing on the DIC of the EDW during the formation period, we used the salinity difference between the EDW sampled on Feb. 23 and March 11 as a constraint for the amount of mixing. The fraction of external water mixed into EDW was evaluated using a simple two-end-member mixing model based on conservation of mass (Mn) and assuming that salinity (Sn) behaved conservatively, i.e., M1  S1 þM2  S2 ¼M3  S3. In reality, salinity was not conserved because the EDW was exposed to the atmosphere and changed owing to evaporation and precipitation. Nevertheless, because the evaporation and precipitation balance was positive for the region, thus increasing salinity, and was far too small to significantly affect the salinity structure of the EDW (3.4 mm d  1 in Feb/Mar 2007) (Joyce et al., this issue; L. Thomas, personal communication) the assumption of conservative behavior was probably valid for the purpose of this evaluation. In this hypothetical mixing scenario, we first assumed vertical mixing with a second endmember below the thermocline with properties similar to those observed at 500 m (the maximum depth of EDW after formation) at station 25 (Fig. 8; T¼17.2 1C, S¼ 36.383, DIC¼2126.4 mmol kg  1, TA¼ 2385.3 mmol kg  1). In a second scenario, we assumed lateral mixing along sloping isopycnals in the northern end of the GS with an end-member defined by a density criteria of 26.55o sy o26.65 and its average properties (Fig. 8; T¼16.6 1C,

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S¼36.300, DIC¼2119.3 mmol kg  1, TA¼2380.2 mmol kg  1). The first scenario yielded an increase in DIC of 4.6 mmol kg  1 and a decrease in TA of 1.5 mmol kg  1 while the second scenario yielded an increase in DIC of 2.4 mmol kg  1 and a decrease in TA of 1.5 mmol kg  1. Hence, both of these hypothetical mixing scenarios illustrate that both vertical and lateral mixing with the two hypothetical end-members could contribute significantly to the EDW DIC budget. Although caution is advised in the interpretation of the details of this analysis, which is based on several assumptions that are hard to verify, it can be concluded that mixing is probably far more important than gas exchange in setting the initial DIC reservoir during the formation process and outcropping of EDW. However, this analysis does not take into account gas exchange occurring during the pre-conditioning of the water masses involved in the formation of EDW. 3.4.4. Biological processes The potential contribution from biological processes (i.e., primary production, respiration, CaCO3 formation and dissolution) to changes in DIC in the EDW during formation was difficult to assess because of low nutrient concentrations and small relative changes in concentration between the sampling dates as well as insufficient precision of total alkalinity measurements. The observed changes in nutrients were smaller or similar to the observed variability between measurements so little can be concluded based on these data. Using nitrate as an example, average nitrate ( 71 std) in the EDW was observed to increase from 2.870.4 mmol kg  1 to 3.0 70.3 mmol kg  1 between February 23 and March 11. Based on the mass balance approach previously applied and mixing with the hypothetical end member at 500 m depth at station 25, average nitrate was expected to increase to 3.3 mmol kg  1. If in fact nutrients were injected into the EDW from mixing with this end member, the fact that the observed average nitrate was lower than expected might indicate net nutrient uptake owing to primary production, but this would have a small effect on the ocean uptake of atmospheric CO2 since the injection of nutrients would also be accompanied by an elevated burden of remineralized DIC. In addition, any uptake owing to primary production would not be independent from the CO2 uptake accounted for by gas exchange and because CO2 is fixed into organic matter it would rather result in a decrease than an increase in seawater DIC. The biological mechanisms that could contribute to increasing DIC in the EDW are respiration and decomposition of organic material or dissolution of CaCO3 particles, but as previously stated the precisions of the measurements of nutrient concentration and TA were not sufficiently high to appropriately evaluate the effect of these processes on the present timescale.

4. Discussion and summary Prior to the present investigation, sea–air CO2 flux had never been measured or calculated from direct measurements of surface seawater CO2 parameters in the region of EDW formation during the time of formation. The reason for this can be attributed to the challenging working conditions experienced in this region during wintertime with more or less constant gales and high seas. The absence of these data was a major limitation in a previous study investigating the DIC budget of EDW (Bates et al. 2002), which served as the catalyst for the present investigation. The CO2 fluxes observed in the winter of 2006 and 2007 confirmed the previous understanding that the EDW formation region acts as a significant sink of CO2, but the magnitude of the flux may be greater than recent seawater pCO2 climatology fields suggest. Previous estimates of average sea–air CO2 flux in the EDW formation region

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during wintertime ranged from  5 to  10 mmol CO2 m  2 d  1 (Bates et al., 2001, 2002; Takahashi et al., 2002, 2009). The present study showed that the average flux calculated along the cruise tracks in this region in the wintertime of 2006 and 2007 was 70% and 40% higher than the previous upper estimate of  10 mmol CO2 m  2 d  1, respectively (Table 1). Mean sea–air CO2 flux calculated specifically for newly formed EDW in the time period February 23 and March 11 of 2007 was 185% higher than the previous upper estimate. However, the temporal duration for this flux estimate was relatively short, on the order of a few weeks. Furthermore, the formation process is brief, and the exact duration and the areal extent of outcropping EDW are variable. Fluxes may vary from year to year owing to different frequency and intensity of storms associated with variations in global climate modes such as the North Atlantic Oscillation (NAO) and El Nino Southern Oscillation (ENSO) (e.g., Bates et al., 2001; Gruber et al., 2002). In addition, fluxes may vary as a result of variability associated with the DIC and TA of the active modes of mixing, which ultimately will determine the DpCO2 of outcropping EDW and the atmosphere. For example, the average DIC of old EDW observed in February of 2007 equaled 2105.6711.0 mmol kg  1 compared to 2113.7714.8 mmol kg  1 and 2118.8717.0 mmol kg  1 along WOCE lines A20 and A22 in 1997, respectively, and 2096.67 16.1 mmol kg  1 and 2102.4720.2 mmol kg  1 along the same lines in 2003 (Andersson et al., in preparation). Similarly, variations in the DIC and TA of the EDW precursor and other modes involved in the formation process, e.g., water from the cold side of the GS, are important in controlling the surface seawater pCO2 of new outcropping EDW, and hence, the sea–air CO2 flux before, during, and after formation. Irrespective of these considerations, the calculated CO2 fluxes in 2006 and 2007 were higher than previous estimates. The GS region is a highly dynamic environment in wintertime. Thus, sea–air CO2 fluxes can be highly variable in time and space. In the winter of 2006 and 2007, substantial sea–air CO2 fluxes on the order of  40 to 50 mmol CO2 m  2 d  1, and sometimes higher, were regularly observed in association with pulses of cold air and high winds (Fig. 3). The high variability illustrates the difficulty involved in assigning a mean flux number to this region. Nevertheless, irrespective of the time of the year, the majority of the large CO2 flux events were exclusively associated with high wind events, although the maximum flux event observed in January 2006 was linked to a combination of low seawater pCO2 and intermediate wind speed. The BATS dataset does not indicate as large fluxes as those recorded during the winter of 2006 and 2007 in the EDW formation region. This probably does not mean that fluxes of this magnitude do not occur at BATS, but can rather be explained by the fact that the BATS data presented here represent bottle samples collected from hydrocasts and that it is not possible to conduct oceanographic work during the most intense wind events. Thus, the BATS dataset analyzed here may be underestimating the influx of CO2 in the wintertime, although the observed seasonal flux trend is accurate. The average sea–air CO2 fluxes for January and March for the past 10 years at BATS were 75% and 69% of the average fluxes recorded in January and March of 2006 and 2007 in the EDW formation region, respectively (Table 1). The observation of newly formed EDW during February and March of 2007 presented a unique opportunity to evaluate the initial properties of this water mass. If the newly formed EDWs were formed independently, their closely matching properties suggest that the formation process is controlled by a similar suite of mechanisms and active modes of mixing. Any observed variability may be explained by different contribution or minor differences in the processes and/or water masses involved in the formation process. However, if in fact the changes observed on March 11 compared to February 23 represented a time dependent

change along a Lagrangian pathway, sea–air CO2 exchange could account for about a quarter of the observed change in DIC between these two dates, and the remainder of the increase would be due to vertical and/or lateral mixing processes. This finding agrees with the conclusion of Joyce et al. (2009) based on shipboard SeaSOAR and ADCP data. Contribution from biological processes that would result in a net increase in DIC on the present time scale is most likely small and probably remains close to zero. In support of this, Bates et al. (2002) estimated a long-term increase of DIC in EDW during the 1990 s owing to biological processes at a rate of 0.55 mmol  1 kg  1 y  1. Notably, the EDW observed in June of 2007 in the vicinity of the region where EDW was formed during winter had very similar properties to the EDW observed in February and March. This suggests that very small changes occurred in the first several months after formation of EDW in the region of formation. The largest weakness of the present analysis was the fact that the EDW observations were not explicitly tracked from a Lagrangian perspective. Hence, the next obvious step to gain further insight to the mechanisms setting the initial DIC balance of EDW would be to actually track outcropping EDW and the evolution of DIC during the formation process using either drifters or gliders. One of the critical questions of interest to the scientific community is how much atmospheric and anthropogenic CO2 is taken up by the formation of EDW every year? Based on the present results of average CO2 flux in the newly formed EDW (  28.5710 mmol CO2 m  2 d  1) and assuming that the cumulative volume of EDW formed every year equaled 9.3 Sv y (¼ 9.3  365  86400  106 m3; Marshall et al., 2009), the average mixed layer depth during formation was 400 m, and the time the EDW remained in contact with the atmosphere ranged from 30 to 60 days, the total uptake of atmospheric CO2 by North Atlantic EDW formation would range from 0.007 to 0.015 Pg C y  1. However, this estimate does not distinguish between the natural background component and the anthropogenic CO2. Based on CO2 flux data collected from volunteer observing ships (VOS) in 2005, Watson et al. (2009) estimated that the North Atlantic Ocean bound between 101N and 651N served as a net sink of 0.2570.05 Pg C y  1. Adopting this estimate, sea–air CO2 flux into newly formed EDW alone would constitute 3–6% of the North Atlantic Ocean atmospheric CO2 sink. In addition, advection and eddy-diffusion transport light surface waters into EDW formation where they are transformed into denser mode waters and subducted into the subtropical thermocline. These surface waters carry an elevated burden of anthropogenic CO2 acquired via sea– air exchange over a much larger spatial and temporal domain, and EDW formation can thus act as a channel for sequestering this CO2 into the ocean interior (Levine et al., 2011). It is apparent that the formation of EDW (or STMW in general) is associated with a significant exchange of CO2 with the atmosphere. If the current estimates observed in the North Atlantic are representative for the formation of global subtropical mode water in the 5 subtropical gyres (i.e., North and South Atlantic and Pacific Oceans, Indian Ocean), they may account for several percent of the total oceanic CO2 uptake. A critical question that needs to be addressed though is how much of this CO2 is of anthropogenic origin? In addition to CO2 uptake during the formation process of STMW through gas exchange, additional CO2 of recent atmospheric origin may accumulate as the STMW ages and sinking organic material is remineralized within the STMW. However, unless the export of carbon or the rates of remineralization have changed over time this does not contribute to the net uptake of anthropogenic CO2 in the EDW. The effective removal of atmospheric CO2 and the time it remains stored depend on the rate and fate of STMW destruction and renewal, which are linked to climate modes such as NAO and

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ENSO (Bates et al., 2002; Gruber et al., 2002; Levine et al., 2011). One of the pillars of climate and carbon research is to identify and improve the understanding of processes and regions acting as CO2 sources or sinks. The formation and role of STMW need to be considered on the larger framework of the global carbon cycle, but our current understanding is still relatively scarce. The present investigation provides preliminary data for determining the quantitative role and importance of STMW in terms of atmospheric CO2 uptake and seawater CO2 dynamics. It is apparent that there are many important questions related to the formation, evolution, and destruction of this water mass, and consequently the fate of its CO2, that remain uncertain or unknown. For example, what are the physical and chemical properties of the precursor(s) of STMW, on what time scale does the STMW form, what is the areal extent of STMW formation, what is the duration the STMW remains in contact with the atmosphere, how is STMW destroyed or dissipated and what is the fate of its DIC inventory, what is the extent and variability of STMW formation on a global scale? It is our hope that the wealth of data collected during the 2006/2007 field campaign in the North Atlantic Ocean will continue to contribute to provide answers to some of these questions as they are critical in order to accurately evaluate the role of STMW as a sink of atmospheric and anthropogenic CO2.

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