What Drives the Latitudinal Gradient in Open-Ocean Surface Dissolved Inorganic Carbon Concentration?

Biogeosciences, 16, 2661–2681, 2019 https://doi.org/10.5194/bg-16-2661-2019 © Author(s) 2019. This work is distributed under the Creative Commons Attribution 4.0 License. What drives the latitudinal gradient in open-ocean surface dissolved inorganic carbon concentration? Yingxu Wu1, Mathis P. Hain2, Matthew P. Humphreys1,3, Sue Hartman4, and Toby Tyrrell1 1University of Southampton, National Oceanography Centre Southampton, Southampton, UK 2Earth and Planetary Sciences, University of California, Santa Cruz, USA 3Centre for Ocean and Atmospheric Sciences, School of Environmental Sciences, University of East Anglia, Norwich, UK 4National Oceanography Centre, Southampton, UK Correspondence: Yingxu Wu ([email protected]) Received: 12 August 2018 – Discussion started: 15 August 2018 Revised: 14 June 2019 – Accepted: 20 June 2019 – Published: 11 July 2019 Abstract. Previous work has not led to a clear under- 1 Introduction standing of the causes of spatial pattern in global surface ocean dissolved inorganic carbon (DIC), which generally in- The ocean absorbs about one-quarter of the anthropogenic creases polewards. Here, we revisit this question by inves- CO2 emitted every year (Le Quéré et al., 2018). It is tigating the drivers of observed latitudinal gradients in sur- the largest non-geological carbon reservoir (∼ 38000 Gt C; face salinity-normalized DIC (nDIC) using the Global Ocean Falkowski et al., 2000), containing 50 times as much carbon Data Analysis Project version 2 (GLODAPv2) database. We as the pre-industrial atmosphere, and thereby modulates the used the database to test three different hypotheses for the Earth’s climate system. Approximately 97 % of the carbon in driver producing the observed increase in surface nDIC from the ocean is in the form of dissolved inorganic carbon (DIC), low to high latitudes. These are (1) sea surface tempera- which is the total concentration of aqueous CO2 and bicar- ture, through its effect on the CO2 system equilibrium con- bonate and carbonate ions: stants, (2) salinity-related total alkalinity (TA), and (3) high- h i D ∗ C − C 2− latitude upwelling of DIC- and TA-rich deep waters. We find DIC CO2 HCO3 CO3 ; (1) that temperature and upwelling are the two major drivers. ∗ TA effects generally oppose the observed gradient, except where [CO2] refers to the sum of aqueous CO2 and undisso- where higher values are introduced in upwelled waters. ciated carbonic acid (H2CO3), with the latter being negligible Temperature-driven effects explain the majority of the sur- (Zeebe and Wolf-Gladrow, 2001). face nDIC latitudinal gradient (182 of the 223 µmol kg−1 in- Understanding what controls the distribution of oceanic crease from the tropics to the high-latitude Southern Ocean). DIC is essential for quantifying anthropogenic CO2 inva- Upwelling, which has not previously been considered as a sion (e.g., Gruber, 1998; Humphreys et al., 2016; Lee et al., major driver, additionally drives a substantial latitudinal gra- 2003; Sabine et al., 1999, 2002; Vázquez-Rodríguez et al., dient. Its immediate impact, prior to any induced air–sea CO2 2009) and consequent ocean acidification (e.g., Doney et al., exchange, is to raise Southern Ocean nDIC by 220 µmol kg−1 2009; Orr et al., 2005). As most marine organisms live in above the average low-latitude value. However, this immedi- the sunlit surface ocean, where CO2 exchange with the at- ate effect is transitory. The long-term impact of upwelling mosphere happens, the controls on surface ocean DIC in par- (brought about by increasing TA), which would persist even ticular merit investigation. if gas exchange were to return the surface ocean to the Many previous studies focused on the vertical, rather than same CO2 as without upwelling, is to increase nDIC by latitudinal, distribution of DIC. They investigated the contri- 74 µmol kg−1 above the low-latitude average. butions of the different “carbon pumps” – solubility pump, soft tissue pump, and carbonate pump (Cameron et al., 2005; Published by Copernicus Publications on behalf of the European Geosciences Union. 2662 Y. Wu et al.: Surface DIC latitudinal gradient Gruber and Sarmiento, 2002; Toggweiler et al., 2003a, b) lution or concentration due to precipitation or evaporation – to the pattern of DIC with depth. The solubility pump is (Friis et al., 2003); (5) warming and cooling, which alter based on the assumption that, at high latitudes where deep CO2 solubility and induce air–sea gas exchange that acts waters form, DIC is high because the low water temperature to reduce air–sea CO2 disequilibrium (Bozec et al., 2006; increases CO2 solubility. Lee et al. (2000) used this princi- Toggweiler et al., 2003a; Williams and Follows, 2011); and ple to predict salinity-normalized DIC (nDIC) from empiri- (6) the above processes (1–4) through their impact on TA – if cal functions of sea surface temperature and nitrate that var- high/low TA values are not matched by high/low DIC values ied seasonally and geographically. then the resulting low/high seawater pCO2 stimulates CO2 Key et al. (2004) depicted the global distribution of surface ingassing/outgassing until DIC matches TA (Humphreys et DIC using an earlier version (GLODAPv1) of the dataset al., 2018). The effects of equilibrium processes (the effects than we used in this study, noting that the surface DIC pattern through temperature and upwelled TA) change the surface is more similar to nutrients (including in the Southern Ocean, ocean DIC at which air–sea CO2 equilibrium occurs, so where both DIC and nutrients are enriched) than to salinity these effects can persist beyond the air–sea CO2 equilibrium – unlike total alkalinity (TA), whose pattern more closely re- timescale (months to a year; Jones et al., 2014). The effects sembles that of salinity (Fry et al., 2015). Using the data from of disequilibrium processes, such as direct DIC supply from the new GLODAPv2 database (Key et al., 2015; Olsen et al., upwelling, and biological uptake of DIC in response to up- 2016), surface DIC is confirmed here to have its highest val- welled nutrients (principally iron; Moore et al., 2016) can ues at high latitudes, like nutrients, and to reach its lowest persist no longer than the CO2 equilibrium timescale. values at low latitudes in each basin (Fig. 1a; more details Our study builds on previous work in several ways. First, in Sect. 3.1). Earlier studies (Lee et al., 2000; Toggweiler whereas many previous studies looked to understand the ver- et al., 2003a; Williams and Follows, 2011, Sect. 6.3 “What tical DIC distribution, our target is to understand the latitu- controls DIC in the surface ocean?”) suggested that tempera- dinal surface DIC distribution. Second, we identify the most ture is of primary importance in regulating surface DIC (e.g., important processes, not just the variables, driving the sur- Lee et al., 2000; Toggweiler et al., 2003a; Williams and Fol- face DIC distribution (Fig. 2). Third, we use a much larger lows, 2011; Humphreys, 2017). Under this assumption, sur- observational global dataset – GLODAPv2. face waters in cool regions at high latitudes should hold more We evaluate three main hypotheses as to which processes DIC than surface waters in the warm regions at low latitudes. cause the increase in surface DIC and nDIC from low to high Williams and Follows (2011) argued that another variable latitude (Fig. 1): also exerts control on the surface DIC distribution: TA sets the equilibrium capacity for seawater to hold DIC in solution 1. latitudinal variation in solar heating via its effect on sea (Omta et al., 2011; Humphreys et al., 2018), so higher surface surface temperature, and hence CO2 solubility; TA values may lead to higher DIC. Takahashi et al. (2014) 2. evaporation and precipitation, through their effects on explored the seasonal distribution of climatological surface TA; and DIC using seawater pCO2 from the Lamont Doherty Earth Observatory (LDEO) database and TA estimated from salin- 3. upwelling and winter entrainment through the introduc- ity, qualitatively attributing seasonal differences (on a re- tion of DIC- and TA-rich deep waters to the (sub)polar gional scale) to the greater upward mixing of high-CO2 deep surface oceans, when coupled with iron limitation of bi- waters in winter and summer biological carbon drawdown. ological uptake of DIC. They noted the potential for upwelling to alter surface DIC, but focused on DIC seasonality rather than its spatial vari- It is easier to constrain the dynamics of upwelling and ability. In recent years, the global surface DIC database has quantify their impact on surface DIC in the Southern Ocean greatly expanded (e.g., Bates et al., 2006; Sasse et al., 2013), (where upwelling has been more comprehensively studied; culminating in GLODAPv2 (Key et al., 2015; Olsen et al., e.g., Marshall and Speer, 2012; Morrison et al., 2015) than in 2016; Lauvset et al., 2016), but the drivers of the global sur- the subarctic North Atlantic and North Pacific oceans (where face DIC distribution have not yet been reassessed. upward transport occurs via deep mixing in the winter, com- The processes that influence the distribution of surface bined with upwelling in the North Pacific). The Southern DIC at the local scale can be divided into those which change Ocean also plays a crucial role in the global overturning DIC by direct addition or removal, and those which af- circulation (e.g., Marshall and Speer, 2012), and the global fect DIC indirectly. The direct processes include (1) biolog- carbon cycle (Landschützer et al., 2015; Mikaloff-Fletcher, ical carbon assimilation during primary production and re- 2015). Therefore, we focused on the Southern Ocean for the lease during remineralization (Bozec et al., 2006; Clargo et evaluation of the third hypothesis. A novel conclusion of this al., 2015; Toggweiler et al., 2003b; Yasunaka et al., 2013); study is that upwelling, whose global significance has pre- (2) transport of DIC-rich deep waters into the surface layer viously been overlooked, is very important in shaping the (Jiang et al., 2013; Lee et al., 2000); and (3) production and spatial distribution of surface ocean DIC, in part because up- export of CaCO3.

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