Carbon is present in various inorganic and organic forms in coastal waters (see Figure 16.2). Dissolved inorganic species include aqueous CO₂ (a combination of dissolved CO₂ and carbonic acid), bicarbonate and carbonate ions, and methane (CH₄); the first three carbon species are collectively referred to as dissolved inorganic carbon or DIC. The major particulate inorganic species is calcium carbonate (CaCO₃), also referred to as particulate inorganic carbon (PIC). Carbon is also present in various dissolved and particulate organic forms (DOC and POC). In shelf waters, the reduced carbon pool or total organic carbon pool (TOC) represents roughly 2% to 5% of the total carbon stock (Liu et al., 2010), and DOC constitutes more than 90% to 95% of this TOC (Vlahos et al., 2002).

Carbon is constantly transferred among these different pools and exchanged across the interfaces that demarcate coastal waters: the land-ocean interface, the air-sea interface, and the interface between coastal and open-ocean waters (see Figure 16.2). The internal carbon transformations within coastal regions include photosynthetic primary production, respiration, transfers between lower and higher trophic levels of the food web, exchanges between sediment and overlying water, biogeochemical processes in the sediment, and the formation and dissolution of CaCO₃. Major internal transformations are the conversion of DIC into organic carbon (POC and DOC), through primary production, and respiration throughout the water column, returning most of the organic carbon into inorganic forms (primarily DIC). Some POC settles out of the water column and becomes incorporated into the sediments where most of this material is respired through a range of different redox processes that produce DIC and, under certain circumstances, CH₄ (i.e., in the relative absence of electron acceptors other than CO₂). Both DIC and CH₄ are released back into the overlying water. POC that is not respired (referred to as refractory POC) can be buried in sediments and stored for a very long time. Some organisms form internal or external body structures of CaCO₃, which either dissolve or become incorporated into the sediments and are buried. This discussion will refer to long-term storage of buried POC and PIC in coastal sediments as permanent burial.

A major carbon exchange process along the ocean margin is the flux of CO₂ across the air-sea interface. The annual cycle of this flux is driven by 1) seawater warming and cooling, which affects CO₂ solubility; 2) the under- or oversaturation of CO₂ resulting from primary production, respiration, and CaCO₃ precipitation and dissolution; 3) the transport of DIC to and from the ocean surface (e.g., upwelling and convection); and 4) factors that influence the resistance to gas exchange across the air-sea interface (e.g., winds, sea ice extent, and surface films). The annual cycles of primary production, respiration, and air-sea CO₂ flux tend to be of larger magnitude and more variable in coastal waters than in the open ocean (Bauer et al., 2013; Liu et al., 2010; Muller-Karger et al., 2005; Thunell et al., 2007; Xue et al., 2016) and more pronounced in high latitudes. Other important exchange fluxes are organic and inorganic carbon inputs from land via rivers and estuaries (see Ch. 15: Tidal Wetlands and Estuaries), from tidal wetlands, and exchanges between the coastal and open oceans across the continental shelf break or the operationally defined open-ocean boundary of the coastal ocean. Net removal of carbon from direct interaction with the atmospheric reservoir can occur by export to the deep ocean or by permanent burial in coastal sediments.

Although continental shelves make up only 7% to 10% of the global ocean surface area, they are estimated to contribute up to 30% of primary production, 30% to 50% of inorganic carbon burial, and 80% of organic carbon burial (Dunne et al., 2007; Gattuso et al., 1998). As such, continental shelves have been argued to contribute disproportionately to the oceanic uptake of CO₂ (Cai 2011; Liu et al., 2010; Muller-Karger et al., 2005).

Carbon export, referring to the flux of organic and inorganic carbon from coastal waters to the deep ocean, can occur through the so-called “Continental Shelf Pump”—a term coined by Tsunogai et al. (1999) after they observed a large uptake of atmospheric CO₂ in the East China Sea. There are two distinct mechanisms underlying the Continental Shelf Pump (Fennel 2010). The first is physical in nature and thought to operate in mid- and ­high-latitude systems. In winter, shelf water is cooled more strongly than surface water in the adjacent open ocean because the former is not subject to deep convection. The colder shelf water is denser and experiences a larger influx of atmospheric CO₂; both density and the solubility of CO₂ increase with decreasing temperature. If this dense and carbon-rich water is transported off the shelf, it will sink due to its higher density, and the associated carbon will be exported to the deep ocean. The second mechanism relies on biological processes that concentrate carbon below the seasonal pycnocline (i.e., photosynthetic production of organic carbon and subsequent sinking). If the carbon-rich water below the seasonal pycnocline is moved off the shelf horizontally, carbon potentially could be exported if this water is transported or mixed below the seasonal thermocline. The depth to which the shelf-derived carbon can be exported will be different for POC, which will sink, and DOC and DIC, which primarily would be advected laterally. Both mechanisms for carbon export critically depend on physical transport of carbon-rich water off the shelf.

Carbon export flux from coastal waters to the deep ocean cannot be quantified easily or accurately through direct observation. Thus, the only available estimates of such export are indirect, using mass balances of POC and dissolved oxygen (Hales et al., 2006), mass balances of DOC (Barrón and Duarte 2015; Vlahos et al., 2002), mass balances of TOC and DIC (Najjar et al., 2018), and model estimates (Izett and Fennel 2018a, 2018b; Bourgeois et al., 2016; Fennel and Wilkin 2009; Fiechter et al., 2014; Mannino et al., 2016; Turi et al., 2014; Xue et al., 2013). If the total carbon inventory in a coastal system can be considered constant over a sufficiently long timescale (i.e., on the order of years), inferring carbon export is possible from using the sum of all other exchange fluxes across the system’s interfaces over that same period. Export to the open ocean must balance the influx of carbon from land and wetlands, its net exchange across the air-sea interface, lateral exchange caused by advection, and any removal through permanent sediment burial. The accuracy of the inferred export flux directly depends on the accuracy of the other flux estimates and of the assumption of a constant carbon inventory. Quantifying internal transformation processes (e.g., respiration and primary and secondary production) does not directly enter this budgeting approach but can elucidate the processes that drive fluxes across interfaces.

Current estimates of carbon fluxes across coastal interfaces come with significant uncertainties (Regnier et al., 2013; Birdsey et al., 2009). These uncertainties are caused by a combination of 1) small-scale temporal and spatial variability, which is undersampled by currently available means of direct observation, and 2) regional heterogeneity, which makes scaling up observations from one region to larger areas difficult. Contributing to variability in regional carbon budgets and export are geographical differences arising from variations in shelf width, the presence or absence of large rivers, seasonal ice cover, and latitude through its modulation of annual temperature and productivity cycles and of hydrography due to the rotation of the Earth (Sharples et al., 2017). Section 16.3 describes the regional characteristics of North American coastal waters and how these characteristics influence carbon dynamics. Available estimates of carbon fluxes are compiled in an attempt to estimate export.

The motivation for quantifying permanent burial of carbon and export of carbon from coastal waters to the deep ocean is that both processes remove CO₂ from the atmospheric reservoir. A more relevant but harder to obtain quantity in this context is the burial or export of anthropogenic carbon. The anthropogenic component of a given carbon flux is defined as the difference between its preindustrial and present-day fluxes. Thus, present-day carbon fluxes represent a superposition of the anthropogenic flux component and the natural background flux. Only total fluxes—the sum of anthropogenic and background fluxes—can be observed directly. Distinction between anthropogenic fluxes and the natural background is difficult to assess for coastal ocean fluxes and has to rely on process-based arguments and models (Regnier et al., 2013). Observation-based estimates of the global open ocean’s anthropogenic uptake have been made by Sabine et al. (2004), Sabine and Tanhua (2010), and Carter et al. (2017). Bourgeois et al. (2016) were the first to estimate coastal anthropogenic carbon uptake in their global model. Their estimates are presented in some detail in Section 16.3.5.

16.3.1 North American Atlantic Coast

The North American Atlantic Coast borders on a wide, geologically passive margin shelf that extends from the southern tip of Florida to the continental shelf of the Labrador Sea (see Figure 16.1). The shelf is several hundreds of kilometers wide in the north (Labrador shelf and Grand Banks) but narrows progressively toward the south in the Middle Atlantic Bight (MAB), which is between Cape Cod and Cape Hatteras, and the South Atlantic Bight (SAB), which is south of Cape Hatteras. The SAB shelf width measures only several tens of kilometers. Two major semi-enclosed bodies of water are the Gulf of Maine (GOM) and Gulf of St. Lawrence. Important rivers and estuaries north of Cape Hatteras include the St. Lawrence River and Estuary, the Hudson River, Long Island Sound, Delaware Bay, and Chesapeake Bay. South of Cape Hatteras, the coastline is characterized by small rivers and marshes. The SAB is influenced by the Gulf Stream, which flows northeastward along the shelf edge before detaching at Cape Hatteras and meandering eastward into the open North Atlantic Ocean. North of Cape Hatteras, shelf circulation is influenced by the confluence of the southwestward-flowing fresh and cold shelf-break current (a limb of the Labrador Current) and the warm and salty Gulf Stream (Loder et al., 1998). Because shelf waters north of Cape Hatteras are sourced from the Labrador Sea, they are relatively cold, fresh, and carbon rich, while slope waters (those located between the shelf break and the northern wall of the Gulf Stream) are a mix­ture of Labrador Current and Gulf Stream water. Exchange between the shelf and open ocean across the shelf break is impeded by the presence of the Gulf Stream south of Cape Hatteras and by shelf-break jets and fronts north of Cape Hatteras.

Air-sea fluxes of CO₂ exhibit a large-scale latitudinal gradient along the North American Atlantic Coast and significant seasonal and interannual variability. The net flux on the Scotian Shelf remains controversial. Shadwick et al. (2010), combining in situ and satellite observations, reported a large source of CO₂ to the atmosphere of 8.3 ± 6.6 grams of carbon (g C) per m² per year. In contrast, Signorini et al. (2013) estimated a relatively large sink of atmospheric CO₂, 14 ± 3.2 g C per m² per year, when using in situ data alone and a much smaller uptake, 5.0 ± 4.3 g C per m² per year, from a combination of in situ and satellite observations. The open GOM is a weak net source of 4.6 ± 3.1 g C per m² per year according to Vandemark et al. (2011) but with significant interannual variability, while Signorini et al. (2013) estimate the region to be neutral. The shallow, tidally mixed GOM regions (i.e., Georges Bank and Nantucket Shoals) are thought to be sinks, however (see Table 16.1; Signorini et al., 2013). The MAB and SAB are net sinks. Observation-based estimates for the MAB include sinks of 13 ± 8.3 g C per m² per year (DeGrandpre et al., 2002) and 13 ± 3.2 g C per m² per year (Signorini et al., 2013). Estimates for the SAB include sinks of 5.8 ± 2.5 g C per m² per year (Jiang et al., 2008) and 8.2 ± 2.9 g C per m² per year (Signorini et al., 2013). The change from neutral or occasional net source in the Scotian Shelf and GOM regions to net sink in the MAB arises because the properties of shelf water are modified during its southwestward flow by air-sea exchange, inflows of riverine and estuarine waters (Salisbury et al., 2008b, 2009), and exchange with the open North Atlantic across the shelf break (Cai et al., 2010a; Wang et al., 2013). Outgassing of CO₂ on the Scotian Shelf is driven primarily by warming of cold, carbon-rich shelf water, which still carries a pronounced signature of its Labrador Sea origin. The GOM, which is deeper than the Scotian Shelf and the MAB and connected to the open North Atlantic through a relatively deep channel, is characterized by a mixture of cold, carbon-rich shelf waters and warmer, saltier slope waters. Shelf water in the MAB is sourced from the GOM and thus is a mixture of shelf and slope water.

Shelf water in the SAB is distinct from that in the MAB and has almost no trace of Labrador Current water; instead, its characteristics are similar to those of the Gulf Stream, but its carbon signature is modified by significant organic and inorganic carbon and alkalinity inputs from coastal marshes (Cai et al., 2003; Jiang et al., 2013; Wang and Cai 2004; Wang et al., 2005). Herrmann et al. (2015) estimated that 59% of the 3.4 teragrams of carbon (Tg C) per year of organic carbon exported from U.S. East Coast estuaries is from the SAB. The subsequent respiration of this organic matter and direct outgassing of marsh-derived carbon make the nearshore regions a significant CO₂ source almost year-round. Despite the carbon inputs from marshes, uptake of CO₂ on the mid- and outer shelf during the winter months is large enough to balance CO₂ outgassing in the other seasons and on the inner shelf, making the SAB overall a weak net sink (Jiang et al., 2008).

North of Cape Hatteras, CO₂ dynamics are characterized by strong seasonality with solubility-driven uptake by cooling in winter and biologically driven uptake in spring followed by outgassing in summer and fall due to warming and respiration of organic matter (DeGrandpre et al., 2002; Shadwick et al., 2010, 2011; Signorini et al., 2013; Vandemark et al., 2011; Wang et al., 2013). Hydrography and CO₂ dynamics on the Scotian Shelf are influenced by the significant freshwater input from the St. Lawrence River. Riverine inputs of carbon and nutrients are relatively small in the GOM but can cause local phytoplankton blooms, CO₂ drawdown, and low-pH conditions (Salisbury et al., 2008a, 2009). Riverine and estuarine inputs become more important in the MAB with discharges from the Chesapeake Bay and the Delaware, Hudson, and Connecticut rivers (Wang et al., 2013). South of Cape Hatteras, seasonal phytoplankton blooms do not occur regularly and biologically driven CO₂ uptake is less pronounced than that further north (Wang et al., 2013), although sporadic phytoplankton blooms do occur because of intrusions of high-nutrient subsurface Gulf Stream water (Wang et al., 2005, 2013). The influence of riverine inputs is small and localized in the SAB (Cai and Wang 1998; Wang and Cai 2004; Wang et al., 2005).

Regional biogeochemical models reproduce the large-scale patterns of air-sea CO₂ flux with oceanic uptake increasing from the SAB to the GOM (Cahill et al., 2016; Fennel et al., 2008; Previdi et al., 2009). These model studies elucidate the magnitude and sources of interannual variability as well as long-term trends in air-sea CO₂ fluxes. Previdi et al. (2009) investigated opposite phases of the North Atlantic Oscillation (NAO) and found that the simulated air-sea flux in the MAB and GOM was 25% lower in a high-NAO year compared with that in a low-NAO year. In the MAB, the decrease resulted primarily from changes in wind forcing, while in the GOM, changes in surface temperature and new production were more important. Cahill et al. (2016) investigated the impact of future, climate-driven warming and trends in atmospheric forcing (primarily wind) on air-sea CO₂ flux (without considering the atmospheric increase in CO₂). Their results suggest that warming and changes in atmospheric forcing have modest impacts on air-sea CO₂ flux in the MAB and GOM compared with that in the SAB where surface warming turns the region from a net sink into a net source of CO₂ to the atmosphere. Model studies also illustrate the effects of interactions between biogeochemical transformations in the sediment and the overlying water column on carbon fluxes. For example, Fennel et al. (2008) showed that the effective alkalinity flux resulting from denitrification in sediments of the North American Atlantic Coast reduces the simulated ocean uptake of CO₂ by 6% compared to a simulation without sediment denitrification.

The passive-margin sediments along the Atlantic coast have not been considered an area of significant CH₄ release until recently (Brothers et al., 2013; Phrampus and Hornbach 2012; Skarke et al., 2014). Phrampus and Hornbach (2012) predicted that massive seepage of CH₄ from upper-slope sediments is occurring in response to warming of intermediate-depth Gulf Stream waters. Brothers et al. (2013) and Skarke et al. (2014) documented widespread CH₄ plumes in the water column and attributed them to gas hydrate degradation. Estimated CH₄ efflux from the sediment in this region ranges from 1.5 × 10^–5 to 1.8 × 10^–4 Tg CH₄ per year, where the uncertainty range reflects different assumptions underlying the conversion from CH₄ plume observations to seepage rates. The fraction of the released CH₄ that escapes to the atmosphere remains uncertain (Phrampus and Hornbach 2012).

16.3.2 North American Pacific Coast

The North American Pacific Coast extends from Panama to the Gulf of Alaska and is an active margin with varying shelf widths (see Figure 16.1). The continental shelf is narrow along the coasts of California, Oregon, and Washington, with a width on the order of 10 km but widening significantly in the Gulf of Alaska, where shelves extend up to 200 km offshore. In the Gulf of Alaska, freshwater and tidal influences strongly affect cross-shelf exchange, and the shelf is dominated by downwelling circulation. The region from Vancouver Island to Baja California is a classic eastern boundary current upwelling region influenced by the California Current System (Chavez et al., 2017). Winds drive a coastal upwelling circulation characterized by equatorward flow in the California Current and by coastal jets and their associated eddies and fronts that extend offshore, particularly off the coasts of Baja California, California, Washington, and Oregon. The northern California Current System experiences strong freshwater influences and seasonality in wind forcing that diminish in the southern part of the system. In addition to the Columbia River and the Fraser River, a variety of small mountainous rivers, with highly variable discharge, supply freshwater. The Central American Isthmus runs from Panama to the southern tip of Baja California and experiences intense and persistent wind events, large eddies, and high waves that combine to produce upwelling and strong nearshore mixing (Chapa-Balcorta et al., 2015; Franco et al., 2014). In addition to alongshore winds, strong seasonal wind jets that pass through the Central American cordillera create upwelling “hotspots” and drive production during boreal winter months in the gulfs of Tehuantepec, Papagayo, and Panama (Chapa-Balcorta et al., 2015; Chelton et al., 2000a, 2000b; ­Gaxiola-Castro and Muller-Karger 1998; Lluch-Cota et al., 1997). The California Current brings water from the North Pacific southward into the southern California and Central American Isthmus regions, while the California Undercurrent transports equatorial waters northward in the subsurface.

The net exchange of CO₂ with the atmosphere across the North American Pacific Coast is characterized by strong spatial and temporal variation and reflects complex interactions between biological uptake of nutrients and degassing of nutrient- and carbon-rich upwelled waters. A growing number of coastal air-sea flux studies have used extrapolation techniques to estimate fluxes across the coastal ocean on regional to continental scales. Observation-based studies of air-sea CO₂ flux suggest that estimates for the coastal ocean from Baja California to the Gulf of Alaska range from a weak to moderate sink of atmospheric CO₂ over this broad longitudinal range. Central California coastal waters have long been understood to have near-neutral air-sea CO₂ exchange because of their large and ­counter-balancing periods of efflux during upwelling conditions and influx during periods of relaxation and high primary productivity; this pattern is strongly modulated by El Niño–La Niña conditions (Friederich et al., 2002). Hales et al. (2005) used seasonal data to estimate an uptake of 88 g C per m² per year by Oregon coastal waters, which is about 15 times larger than the global mean of 6 g C per m² per year. Using data with greater temporal coverage, Evans et al. (2011) showed how large flux events can significantly alter the estimation of net exchanges for the Oregon shelf. After capturing a large and short-lived efflux event, their annual estimate was outgassing of 3.1 ± 82 g C per m² per year for this same region. The disparity illustrates the importance of basing regional flux estimates on observations that are well resolved in time and space. Capitalizing on the increased and more uniform spatiotemporal coverage of satellite data, Hales et al. (2012) estimated an annual mean uptake of 7.9 g C per m² per year between 22^o and 50^oN within 370 km offshore. The most northern estimates for the North American Pacific Coast by Evans et al. (2012) and Evans and Mathis (2013) determined influxes of 26 g C per m² per year for British Columbian coastal waters shoreward of the 500-m isobath and 18 g C per m² per year for Gulf of Alaska coastal waters shoreward of the 1500-m isobath.

Models for the upwelling region (Fiechter et al., 2014; Turi et al., 2014) reproduce the pattern of CO₂ outgassing nearshore and CO₂ uptake further offshore. They also illustrate the intense eddy-driven variability nearshore. Turi et al. (2014) simulate a weak source of 0.6 ± 2.4 g C per m² per year for the region from 30^ο to 46^οN, extending 800 km of shore, an amount which is inconsistent with the observations of Hales et al. (2012) that describe the same region as a sink of 7.9 g C per m² per year. Fiechter et al. (2014) simulate a source of atmospheric CO₂ of 0.6 Tg C per year for the region from 35^ο to 45^οN within 600 km of shore, an estimate which is in contrast to the observation-based estimate of a 14 Tg C sink published by Hales et al. (2012). Both models simulate strong outgassing within the first 100 km of shore, driven by intense upwelling of nutrient- and carbon-rich water, compensated by biologically driven CO₂ uptake from the atmosphere as upwelled nutrients are consumed by photosynthesis during subsequent offshore advection within several hundreds of kilometers of the coast. The disagreement in mean simulated fluxes may result partly from different choices of averaging region and period and differences in model forcing, such as the climatological forcing in Turi et al. (2014) versus realistic variability in Fiechter et al. (2014). Notable, however, is that observations for the Oregon shelf by Evans et al. (2015a) showed intense summer upwelling that led to strong outgassing with pronounced variability in air-sea fluxes but found only weak stimulation of primary production. The research team hypothesized that nutrient-rich waters might be subducted offshore at convergent surface temperature fronts before nutrients are fully consumed by primary producers.

Less is known about the air-sea flux of CH₄ along the North American Pacific Coast margin. Recent studies inventoried sedimentary sources of CH₄ hydrates, derived from terrestrial and coastal primary production, and suggested that extensive deposits along the Cascadia margin are beginning to destabilize because of warming (Hautala et al., 2014; Johnson et al., 2015).

Cross-shelf exchange of carbon occurs in the California Current System mostly in response to wind-driven circulation and eddies, but river plumes and tides also have been shown to increase offshore transport in the northern part of the system (Barth et al., 2002; Hales et al., 2006). Uncertainties in published estimates are high, ranging from very small (Ianson and Allen 2002; Pennington et al., 2010) to very high fractions of primary production (Hales et al., 2005; Turi et al., 2014), again as a result of the region’s large spatial and temporal variability.

16.3.3 Gulf of Mexico

The Gulf of Mexico (GMx) is a semi-enclosed marginal sea at the southern coast of the conterminous United States. The passive margin shelves of its northern portion are relatively wide (up to 250 km west of Florida), but, in contrast to shelf waters of the North American Atlantic Coast, those of the GMx are not separated from open-ocean waters by shelf-break fronts or currents. Ocean water enters the Gulf mainly through the Yucatan Channel, where it forms the northeastward meandering Loop Current (LC), which sheds anticyclonic eddies and exits the Gulf through the Florida Straits (Muller-Karger et al., 2015; Rivas et al., 2005). While shelf circulation is influenced primarily by local wind and buoyancy forcing, outer-shelf regions are at times influenced by LC eddies that impinge on and interact with the shelf (Lohrenz and Verity 2004). Riverine input is substantial in the northern GMx, where the ­Mississippi-Atchafalaya River System delivers large loads of freshwater, nutrients, and sediments.

Observational estimates indicate that the GMx, as a whole, is a weak net sink of atmospheric CO₂ with an annual average of 2.3 ± 0.96 g C per m² per year (Robbins et al., 2014). Robbins et al. (2014) also provide flux estimates, as follows, for smaller shelf regions, namely, the West Florida Shelf, the northern Gulf shelf, the western Gulf shelf, and the Mexico shelf. The West Florida Shelf and western Gulf shelf act as sources to the atmosphere, with estimated annual average fluxes of 4.4 ± 1.3 and 2.2 ± 0.6 g C per m² per year, respectively. The northern Gulf acts as a sink, with an estimated flux of 5.3 ± 4.4 g C per m² per year, and the Mexican shelf is almost neutral, with an estimated flux of 1.1 ± 0.6 g C per m² per year. Huang et al. (2015) estimated a larger uptake on the northern Gulf shelf of 11 ± 44 g C per m² per year (i.e., about twice the estimate of Robbins et al., 2014) and reported a much larger uncertainty. In an analysis that combines satellite and in situ observations, Lohrenz et al. (2018) estimated a similar uptake for the northern GMx of 13 ± 3.6 g C per m² per year. The overall carbon exchanges in the Gulf vary significantly from year to year because of interannual variability in wind, temperature, and precipitation (Muller-Karger et al., 2015).

Model-simulated air-sea CO₂ fluxes by Xue et al. (2016) agree relatively well with the estimates of Robbins et al. (2014), reproducing the same spatial pattern though their simulated Gulf-wide uptake of 8.5 ± 6.5 g C per m² per year is larger. This discrepancy results largely from a greater simulated sink in the open Gulf. Also, the uncertainty estimates of the model-simulated fluxes by Xue et al. (2016) are much larger than those of Robbins et al. (2014); the latter might be too optimistic in reporting uncertainties of the flux estimates.

Overall, the various observation- and model-derived estimates for Gulf regions agree in terms of their broad patterns, but existing discrepancies and, at times, large uncertainties indicate that current estimates need further refinement.

Quantitative understanding of CH₄ dynamics in GMx coastal and oceanic environments is limited. Solomon et al. (2009) speculated that deep CH₄ hydrate seeps in the Gulf potentially are a significant CH₄ source to the atmosphere. They estimated ocean-atmosphere fluxes from seep plumes of 1,150 ± 790 to 38,000 ± 21,000 g CH₄ per m² per day compared with 2.2 ± 2.0 to 41 ± 8.2 g CH₄ per m² per day for background sites. Subsequent acoustic analyses of bubble plume characteristics question the finding that CH₄ bubbles make their way to the surface (Weber et al., 2014), and the fate of CH₄ emissions from seeps and their overall contribution to atmospheric CH₄ remain uncertain.

16.3.4 North American Arctic

The North American Arctic coastal ocean comprises broad (~300 km) shallow shelves in the Bering and Chukchi seas, the narrower (<100-km) Beaufort Sea shelf, the Hudson Bay, and the extensive Canadian Arctic Archipelago (CAA). Shelf water enters these regions from the North Pacific and follows a large-scale pathway from its entrance into the North American Arctic through the Bering Strait via the Chukchi and Beaufort seas into the CAA and, ultimately, the North Atlantic (Carmack et al., 2006, 2015). Hudson Bay receives significant inputs of freshwater (Dery et al., 2005). Except for the southernmost Bering Sea, most of the coastal region is covered with sea ice from about October to June. Areas of persistent multiyear sea ice at the northernmost extent of the CAA are rapidly declining (Stroeve et al., 2012). Reoccurring polynyas (i.e., holes in the ice) are found in all three of its major regions (Smith and Barber 2007). The North American Arctic is sparsely populated with communities heavily reliant on subsistence fishing and hunting; the rapid regional changes associated with global warming are affecting these communities. Globally, the pace of increasing air temperatures is the highest in the North American Arctic and adjacent Arctic regions, resulting in significant reductions in both summer and winter sea ice cover that profoundly affect the marine ecosystems across the northern extent of the continent (Moore and Stabeno 2015; Steiner et al., 2015).

Coastal waters in the North American Arctic have been described consistently as a net sink for atmospheric CO₂ (Bates et al., 2006, 2011; Chen et al., 2013; Cross et al., 2014a; Dai et al., 2013; Else et al., 2008; Evans et al., 2015b; Laruelle et al., 2014; Mucci et al., 2010; Shadwick et al., 2011). This general trait is caused by low surface water pCO₂, the partial pressure of CO₂, relative to the atmosphere during ice-free months. These levels are set by the combination of low water temperatures and seasonally high rates of both ice-associated and open-water primary production (Cai et al., 2010b, 2014; Steiner et al., 2014), as well as by limited gas exchange through sea ice relative to open water (Butterworth and Miller 2016; Rutgers van der Loeff et al., 2014) during winter months when under-ice pCO₂ is higher. Suppressed gas exchange through sea ice has been a source of debate within the Arctic CO₂ flux community, likely a result of inconsistencies between methodologies and the challenge of data collection in such a harsh environment, particularly during winter. The typical approach of calculating air-sea CO₂ flux (from measured air-sea pCO₂ differences and gas transfer rates parameterized using wind speed relationships) can differ markedly from flux estimations determined by eddy correlations. The latter suggest high rates of CO₂ exchange relative to open-water fluxes (Else et al., 2011). Three arguments indicate that the high, initial eddy correlation–based fluxes may be overestimates: 1) the potential for unaccounted CO₂ and water vapor cross-correlation possibly affecting the measurement (Landwehr et al., 2014); 2) independent analysis of the ²²²Radon isotope showing near-zero gas exchange in areas covered by sea ice (Rutgers van der Loeff et al., 2014); and 3) recent demonstration of dampened gas-transfer velocities via concurrent, properly corrected eddy covariance–based fluxes and air-sea pCO₂ difference measurements in the Antarctic marginal ice zone supporting linear scaling methods that calculate fluxes using percent sea ice cover (Butterworth and Miller 2016).

However, despite the dampening effect of sea ice, its permeability is a known function of temperature (Golden et al., 2007). Therefore, as Arctic winter temperatures continue to rise, the role of wintertime air-ice CO₂ exchange may become increasingly important because rising temperatures may allow some degree of exchange to take place. To date, measurements of wintertime exchange have been limited to very few studies (Else et al., 2011, 2013; Miller et al., 2015). In recent years, the role of sea ice growth and decay has been shown to significantly affect the air-sea CO₂ flux (Rysgaard et al., 2007, 2009). During sea ice formation, brine rejection forms dense high-saline water that is exported from the surface layer. This process alters the ratio of total alkalinity to sea ice DIC and the underlying seawater, because DIC is a component of the brine whereas total alkalinity precipitates in the brine channels as a form of CaCO₃ known as ikaite (Dieckmann et al., 2008; Rysgaard et al., 2013). During sea ice decay, ikaite dissolves, leading to excess total alkalinity relative to DIC and undersaturation of CO₂ in meltwater.

Estimates of air-sea CO₂ flux in the Chukchi and Beaufort seas, Hudson Bay, and the western CAA all indicate atmospheric CO₂ uptake (Bates et al., 2006; Else et al., 2008, 2013; Gao et al., 2012; Mucci et al., 2010; Semiletov et al., 2007; Shadwick et al., 2011; see Table 16.1) with significantly higher uptake over the broad and productive Chukchi shelf. A recent synthesis of a decade of coastal ocean data collected within 400 km of land determined an annual mean uptake of 8.8 g C per m² per year (Evans et al., 2015b). Variability in wind patterns and sea ice cover affects the water column structure and connectivity between the surface ocean and overlaying atmosphere, thus influencing the magnitude of air-sea CO₂ exchange.

With regard to Arctic CH₄ fluxes, much more is known about the emission potential, distribution, and functioning of terrestrial sources (McGuire et al., 2009); knowledge of marine CH₄ sources is developing slowly due to sparse observations and the logistical challenges of Arctic marine research. The largest marine CH₄ source in the Arctic is dissociation of gas hydrates stored in continental margin sediments (Parmentier et al., 2013, 2015). As sea ice continues to retreat and ocean waters warm, CH₄ hydrate stability is expected to decrease with potentially large and long-term implications. An additional potential marine CH₄ source, unique to polar settings, is release from subsea permafrost layers, with fluxes from thawed sediments reported to be orders of magnitude higher than fluxes from adjacent frozen sediments (Shakhova et al., 2015).

16.3.5 Summary Estimates for CO₂ Uptake by North American Coastal Waters

Despite the variability in regional estimates discussed above and summarized in Table 16.1, North American coastal waters clearly act as a net sink of atmospheric carbon. Because of discrepancies among studies, these various regional estimates would be difficult to combine into one number with any confidence. Instead, this chapter considers estimates of net air-sea CO₂ exchange in North American coastal waters from two global data syntheses (Chen et al., 2013; Laruelle et al., 2014) and a ­process-based global model (Bourgeois et al., 2016; see Table 16.2). The data syntheses use a global segmentation of the coastal zone and associated watersheds known as MARCATS (MARgins and CATchments Segmentation; Laruelle et al., 2013), which, at a resolution of 0.5º, delineates a total of 45 coastal segments, eight of which surround North America. The data synthesis of Chen et al. (2013) is a summary of individual studies, whereas Laruelle et al. (2014) analyze the Surface Ocean CO₂ Atlas 2.0 database (Bakker et al., 2014) to derive regional estimates. The data syntheses of Chen et al. (2013) and Laruelle et al. (2014) estimate the North American coastal uptake to be 94.4 and 44.5 Tg C per year, respectively, and the process-based model of Bourgeois et al. (2016) estimates an uptake of 48.8 Tg C per year (see Table 16.2). Although there are significant regional discrepancies between the latter two estimates for the eastern tropical Pacific Ocean (i.e., the Central American Isthmus), the GMx, the Florida Upwelling region (actually covering the eastern United States), the Labrador Sea, and the CAA, the overall flux estimates for North America are in close agreement. This, and the fact that Laruelle et al. (2014) used a consistent methodology to estimate air-sea CO₂ flux, builds some confidence in these numbers.

The net CO₂ flux and its anthropogenic component from the process-based global model of Bourgeois et al. (2016) are also reported for a regional decomposition of the EEZs of the United States, Canada, and Mexico (see Table 16.3) in Table 16.4. The model simulates a net uptake of CO₂ in North American EEZ coastal waters (excluding the EEZ of the Hawaiian and other islands) of 160 Tg C per year with an anthropogenic flux contribution of 59 Tg C per year. This chapter adopts 160 Tg C per year as the net uptake by coastal waters of North America, excluding tidal wetlands and estuaries. Unfortunately, there are no formal error estimates for this uptake. Instead, estimates adopted here project an error by first noting that the Bourgeois et al. (2016) model is in good agreement with the more recent of the two observation-based estimates for the MARCATS regions of North America. Furthermore, the error estimate for the uptake by continental shelves globally is about 25%, with the North American MARCATS regions having mainly “fair” data quality (Laruelle et al., 2014). Hence, assuming an error of ±50% for the uptake by North American EEZ waters seems reasonable.

16.3.6 Summary Carbon Budget for North American Coastal Waters

Combining the atmospheric CO₂ uptake estimate with estimates of carbon transport from land and carbon burial in ocean sediments enables a first attempt at constructing a carbon budget for the North American EEZ (see Table 16.5). Carbon delivery to the coastal ocean from land via rivers and from tidal wetlands after estuarine processing (i.e., CO₂ outgassing and carbon burial in estuaries) is estimated to be 106 ± 30 Tg C per year (see Ch. 15: Tidal Wetlands and Estuaries). Estimates of carbon burial, based on the method of Dunne et al. (2007) for the regional decomposition of the North American EEZ, are reported in Table 16.4 with a total flux of 120 Tg C per year. Here these fluxes are considered to be an upper bound because they are substantially larger than other estimates. The Dunne et al. (2007) global estimates of organic carbon burial in waters shallower than 200 m are 19 ± 9 g C per m² per year, much larger than the estimates of 6 and 1 g C per m² per year by Chen (2004) and Muller-Karger et al. (2005), respectively, although areas are slightly different in the three studies. The organic carbon burial estimates of Dunne et al. (2007) for the GOM, MAB, and SAB (see Table 16.4) are larger by factors of 8, 17, and 3, respectively, than the best estimates of the empirical model of Najjar et al. (2018). However, due to different definitions of the boundary between coastal waters and the open ocean, the combined area of the GOM, MAB, and SAB in Najjar et al. (2018) is about a third of that in Dunne et al. (2007). Finally, Dunne et al. (2007) estimated the organic carbon burial in Hudson Bay to be 19 g C per m² per year, compared to a mean estimate of 1.5 ± 0.7 g C per m² per year of burial from sediment cores (Kuzyk et al., 2009). Given these results, SOCCR2 considers the estimates of Dunne et al. (2007) to be an upper bound and assumes that a reasonable lower bound is about an order of magnitude smaller, thus placing the North American organic carbon burial estimate at 65 ± 55 Tg C per year.

If these estimates of net air-sea flux, carbon burial, and carbon input from land are accurate, then the residual must be balanced by an increase in carbon inventory in coastal waters and a net transfer of carbon from coastal to open-ocean waters. In their global compilation, Regnier et al. (2013) report an increase in the coastal carbon inventory of 50 Tg C per year, which is a quarter of their estimated anthropogenic carbon uptake by air-sea exchange in the coastal waters of 200 Tg C per year. The latter estimate is uncertain. In their global modeling study, which did not account for anthropogenic changes in carbon delivery from land, Bourgeois et al. (2016) estimated an accumulation of carbon in the coastal ocean of 30 Tg C per year. This amount is a third of their estimated uptake of anthropogenic carbon from air-sea gas exchange in the coastal ocean of 100 Tg C per year and approximately half of their estimated cross-shelf export of anthropogenic carbon of 70 Tg C per year. The rate of carbon accumulation in the North American EEZ from the model of Bourgeois et al. (2016) is 50 Tg C per year (see Table 16.5). Here again, this chapter assumes an uncertainty of ±50%. The residual of 151 ± 105 Tg C per year is the inferred export of carbon to the open ocean (see Table 16.5). The fact that the error in this residual is large in absolute and relative terms emphasizes the need for more accurate carbon budgets for coastal waters of North America. The challenge, however, is that many of these terms are small compared to internal carbon cycling in coastal waters, which is dominated by primary production and respiration. Two separate estimates of primary production (see Table 16.4) are in broad agreement and reveal that terms in the Table 16.5 budget are just a few percent of primary production. This also emphasizes that small changes in carbon cycling in coastal waters can result in large changes in atmospheric uptake and transport to the open ocean.

16.4.1 Trends in Coastal Carbon Fluxes

Important questions with respect to coastal carbon fluxes include:

• What is the anthropogenic component of the CO₂ sink?

• How will the coastal ocean change as a CO₂ sink?

• How will changing climate and other forcings affect the total and anthropogenic flux proportions?

As stated in Section 16.2, when considering the ocean’s role in sequestering anthropogenic carbon, the relevant component is anthropogenic flux, not the total uptake flux. Neither quantifying the anthropogenic carbon flux component nor predicting its future trend is straightforward. Here the likely trends in total carbon fluxes are described; by definition, changes in total carbon fluxes imply changes in anthropogenic fluxes as well.

A direct effect of increasing atmospheric CO₂ will be an increase in net uptake by the coastal ocean. In addition to rising atmospheric CO₂ levels, changes in climate forcings (i.e., surface heat fluxes, winds, and freshwater input) may affect carbon fluxes in North American coastal waters. Ocean warming reduces the solubility of gases and thus directly affects gas concentrations near the surface; this likely will decrease the net air-sea flux of CO₂ by reducing the undersaturation of CO₂ (see Cahill et al., 2016, for the North American Atlantic Coast). Surface temperature increases also strengthen vertical stratification and thus impede vertical mixing, effects which will affect upward diffusion of nutrients and DIC. Enhanced stratification, therefore, could lead to decreases in both biologically driven carbon uptake and CO₂ outgassing. However, model projections for the northern GMx show that the direct effect of increasing atmospheric CO₂ overwhelms the other more secondary effects (Laurent et al., 2018). Furthermore, temperature trends in coastal waters around North America show complex patterns with some regions having cooled from 1982 to 1997 followed by warming from 1997 to 2013 (e.g., the MAB), some regions having warmed from 1982 to 1997 followed by cooling from 1997 to 2013 (e.g., the SAB and Gulf of Alaska), and other regions showing no consistent warming from 1982 to 2013 (e.g., the NAA; Liao et al., 2015). Temperature anomalies from a time series in the central California Current System show warm surface waters for the decade prior to 1997 followed by a prolonged cooler period until the strong surface warming associated with a marine heatwave and the 2015 to 2016 El Niño interrupted the cool anomalies (Chavez et al., 2017). However, deeper waters in the California Undercurrent have shown a multidecadal trend (1980 to 2012) toward warmer, saltier, ­lower-oxygen, and higher-CO₂ waters at a depth associated with increased northward transport of Pacific equatorial waters (Meinvielle and Johnson 2013).

Some studies suggest that trends in the air-sea pCO₂ gradient (ΔpCO₂) are indicative of a strengthening or weakening of the net CO₂ uptake by shelf systems, where an increasing ΔpCO₂, implying that ocean pCO₂ rises more slowly than atmospheric pCO₂, corresponds to increased net uptake and cross-shelf export (Laruelle et al., 2018). In their observation-based analysis of decadal trends in shelf pCO₂, Laruelle et al. (2018) found that coastal waters lag compared to the rise in atmospheric CO₂ in most regions. For North American coastal waters, they found that the MAB has an increase in ΔpCO₂ of 1.9 ± 3.1 microatmospheres (μatm) per year, a finding which means that in this region surface ocean pCO₂ does not increase or else increases at a rate that is substantially slower than in the atmosphere. For the shelves of the Labrador Sea, the Vancouver Shelf, and the SAB, they found rates of 0.68 ± 0.61 μatm per year, 0.83 ± 1.7 μatm per year, and 0.51 ± 0.74 μatm per year, respectively, implying that surface ocean pCO₂ does not increase or increases at a slower rate than atmospheric CO₂. The only North American coastal region that exhibits a negative trend is the Bering Sea, with –1.1 ± 0.74 μatm per year, meaning that surface ocean pCO₂ increases at a faster rate than in the atmosphere. Laruelle et al. (2018) concluded that the lag in coastal ocean pCO₂ increase compared to that in the atmosphere in most regions indicates an enhancement in the coastal uptake and export of atmospheric CO₂, although they did not investigate alternative explanations.

Trends in coastal ocean uptake of pCO₂ are highly variable regionally and result from a complex interplay of factors. In coastal upwelling systems, surface warming will increase the horizontal gradient between cold, freshly upwelled source waters and warm, offshore surface water, leading to a greater tendency for the subduction of upwelled water at offshore surface temperature fronts during periods of persistent and strong upwelling-favorable winds. The cumulative effect of these processes for the North American Pacific Coast may be greater and more persistent CO₂ outgassing nearshore and lower productivity offshore as upwelled nitrate is exported before it can be used by the phytoplankton community (Evans et al., 2015a). Rates of warming clearly are faster in higher latitudes, but predicting the net effect of these warming-induced changes in the North American Arctic is not easy. Furthermore, warming in the Arctic leads to reductions in ice cover and longer ice-free periods, both of which directly affect air-sea gas exchange (Bates and Mathis 2009). Another profound effect of Arctic warming is the melting of permafrost, which leads to the release of large quantities of CH₄ to the atmosphere, from both the land surface and the coastal ocean (Crabeck et al., 2014; Parmentier et al., 2013).

Changes in wind stress also directly affect air-sea gas fluxes because stronger winds intensify gas exchange. For example, for the North American Atlantic Coast, changes in wind stress were shown to significantly modify air-sea fluxes (Cahill et al., 2016; Previdi et al., 2009). Large-scale changes in wind patterns also affect ocean circulation with a range of implications (Bakun 1990). Upwelling-favorable winds along the North American Pacific Coast have intensified in recent years, especially in the northern parts of the upwelling regimes (García-Reyes et al., 2015; Rykaczewski and Checkley 2008; Rykaczewski et al., 2015; Sydeman et al., 2014), a change which has led to 1) shoaling of subsurface nutrient-rich waters (Aksnesa and Ohman 2009; Bograd et al., 2015), 2) increased productivity (Chavez et al., 2011, 2017; Jacox et al., 2015; Kahru et al., 2015), 3) higher DIC delivery to the surface (Turi et al., 2016), and 4) declining oxygen levels (Crawford and Peña 2016; Peterson et al., 2013; Bograd et al., 2015). In the North American Arctic, late-season air-sea CO₂ fluxes may become increasingly more directed toward the atmosphere as Arctic ­low-pressure systems with storm-force winds occur more often over open water, thus ventilating CO₂ respired from the high organic carbon loading of the shallow shelf (Evans et al., 2015b; Hauri et al., 2013; Steiner et al., 2013) and affecting net annual exchanges. The intense warming observed across the North American Arctic also influences mid-latitude weather patterns (Kim et al., 2014), with probable cascading effects on CO₂ exchanges through adjustments in the wind field.

16.4.2 Acidification Trends in North America’s Coastal Ocean

Increasing atmospheric CO₂ emissions lead to rising atmospheric CO₂ levels (see Figure 16.3) and a net ocean uptake of CO₂. Since about 1750, the ocean has absorbed 27% of anthropogenic CO₂ emissions to the atmosphere from fossil fuel burning, cement production, and land-use changes (Canadell et al., 2007; Le Quéré et al., 2015; Sabine and Tanhua 2010). As a result of this uptake, the surface ocean pCO₂ has increased (see Figure 16.3) and oceanic pH, carbonate ion concentration, and carbonate saturation state have decreased (Caldeira and Wickett 2003; Feely et al., 2004, 2009; Orr et al., 2005). Commonly called ocean acidification, this suite of chemical changes is defined more precisely as “any reduction in the pH of the ocean over an extended period, typically decades or longer, that is caused primarily by uptake of CO₂ from the atmosphere but also can be caused by other chemical additions or subtractions from the ocean” (IPCC 2011, p. 37). In addition to uptake of CO₂ from the atmosphere, variations in DIC concentrations and thus pH can be caused by biological production and respiration. Ocean acidification can significantly affect growth, metabolism, and life cycles of marine organisms (Fabry et al., 2008; Gattuso and Hansson 2011; Somero et al., 2016) and most directly affects marine calcifiers, organisms that precipitate CaCO₃ to form internal or external body structures. When the carbonate saturation state decreases below the equilibrium point for carbonate precipitation or dissolution, conditions are said to be corrosive, or damaging, to marine calcifiers. These conditions make it more difficult for calcifying organisms to form shells or skeletons, perform metabolic functions, and survive.

Acidification trends in open-ocean surface waters tend to occur at a rate that is commensurate with the rate of the increase in atmospheric CO₂ (see, for example, trends of atmospheric CO₂ in comparison to surface ocean pCO₂ at the Hawai‘i Ocean Time-series in Figure 16.3). Acidification in coastal waters is more variable because of a combination of changes in circulation and upwelling, larger-amplitude seasonal signals in production and respiration than in the open ocean, and atmospheric CO₂ uptake (see Figure 16.4; Feely et al., 2008, 2016, 2018; Chavez et al., 2017). In many coastal regions, pCO₂ rises more slowly than in the open ocean (see Section 16.4.1; Laruelle et al., 2018). Along the North American Pacific Coast, ­climate-driven changes in upwelling circulation result in coastal acidification events. As mentioned in Section 16.4.1, upwelling-favorable winds along this coast have intensified over recent years, especially in the northern parts of the upwelling regimes (García-Reyes et al., 2015; McClatchie et al., 2016; Rykaczewski and Checkley 2008; Rykaczewski et al., 2015; Sydeman et al., 2014). Intensified upwelling supplies deep water to the shelf that is rich in DIC and nutrients but poor in oxygen. Ocean acidification and hypoxia thus are strongly linked ecosystem stressors because low-oxygen, high-CO₂ conditions derive from the microbial respiration of organic matter (Chan et al., 2016; Feely et al., 2008, 2016, 2018). In the northern California Current System, pCO₂, pH, and aragonite saturation reach levels known to be harmful to ecologically and economically important species during the summer upwelling season (see Ch. 17: Biogeochemical Effects of Rising Atmospheric Carbon Dioxide; Barton et al., 2012, 2015; Bednaršek et al., 2014, 2016, 2017; Feely et al., 2008, 2016; Harris et al., 2013). In the Gulf of Alaska, aragonite saturation drops to near saturation values during the winter months when deep mixing occurs and surface ocean pCO₂ exceeds atmospheric pCO₂ (Evans and Mathis 2013). Along the Pacific Coast, 50% of shelf waters are projected to experience year-long undersaturation by 2050 (Gruber et al., 2012; Hauri et al., 2013; Turi et al., 2016).

Polar regions are naturally prone to acidification because of their low temperatures (Orr et al., 2005; Steinacher et al., 2009). In many Arctic coastal regions, pH and carbonate saturation state are naturally low relative to lower-latitude coastal settings. These low levels result from higher CO₂ solubility, the influence of multiple sources of freshwater (e.g., riverine, glacial melt, and sea ice melt) with varying CO₂ chemistries, and the high respiratory DIC content in bottom waters. The Beaufort and Chukchi Sea continental shelves experience inflows of naturally corrosive Pacific seawater with pH as low as 7.6 (Mathis et al., 2011). The main contributing factor to the relatively high rates of acidification in polar waters is retreating sea ice, which adds meltwater from multiyear ice and increases the surface area of open water, thereby enhancing the uptake of atmospheric CO₂ (Cai et al., 2010b; Steiner et al., 2013). These factors, in combination with increasing atmospheric CO₂ levels, have set a faster pace of ocean acidification in the Arctic than projected trends in other coastal regions (Feely et al., 2009; Mathis et al., 2015a). Models predict annual average aragonite undersaturation (i.e., favoring dissolution) for the Bering Sea and the Chukchi Sea by 2070 and 2030, respectively (Mathis et al., 2015a). The Beaufort Sea upper halocline and deep waters now regularly show aragonite undersaturation (Mathis et al., 2015a; Miller et al., 2014). These chemical seawater signatures are propagated via M’Clure Strait and Amundsen Gulf into the CAA and beyond (Azetsu-Scott et al., 2010; Turk et al., 2016; Yamamoto-Kawai et al., 2013). Model projections based on the IPCC high-CO₂ emissions scenario, Representative Concentration Pathway 8.5 (RCP8.5), suggest the Beaufort Sea surface water will become undersaturated with respect to aragonite around 2025 (Steinacher et al., 2009; Steiner et al., 2014). As these conditions intensify, negative impacts on calcifying marine organisms are expected to become a critical issue, reshaping ecosystems and fisheries across the North American Arctic domain (Mathis et al., 2015b; Moore and Stabeno 2015).

In the northern GMx, surface aragonite saturation states typically range from 3.6 to 4.5 and are thus well above the dissolution threshold (Wang et al., 2013; Wanninkhof et al., 2015). Here excessive nutrient inputs from the Mississippi River result in hypoxia and eutrophication-induced acidification of near-bottom waters (Cai et al., 2011; Laurent et al., 2017). Similar to the California Current System, low-oxygen and high-CO₂ conditions coincide and derive from microbial respiration of organic matter (Cai et al., 2011; Laurent et al., 2017; Feely et al., 2018). Currently, aragonite saturation states are around 2 in hypoxic bottom waters and thus well above the saturation threshold. Projections suggest that aragonite saturation states of these near-bottom waters will drop below the saturation threshold near the end of this century (Cai et al., 2011; Laurent et al., 2018).

Recent studies indicate that the northern regions of the North American Atlantic Coast (the MAB and GOM) are more prone to acidification than the SAB (Wang et al., 2013; Wanninkhof et al., 2015). Coastal waters in this region have, on average, lower pH and lower aragonite saturation states than more southern coastal regions. These properties are driven primarily by a decrease in mean total alkalinity of shelf water from the SAB northward to the GOM. Seasonal undersaturation of aragonite in subsurface water is occurring in the GOM with photosynthesis and respiration playing a major role in controlling the seasonal variability of aragonite saturation states; dissolution of aragonite might already occur in fall and winter (Wang et al., 2017). With a significant shellfish industry, the GOM displays the lowest pH and aragonite saturation levels along the East Coast in summer (Wang et al., 2013).

The research community has made tremendous progress in improving understanding and constraining rates of carbon cycling in coastal waters since SOCCR1 (CCSP 2007), primarily because of a greatly expanded suite of observations, process studies, and models. However, quantification of many coastal carbon fluxes remains a significant challenge. Carbon is constantly exchanged across the air-sea interface as well as the interfaces between land and coastal ocean, coastal and open-ocean waters, and water and sediment. Net exchange fluxes and trends are relatively small signals masked by a large and fluctuating background. At present, most of these fluxes are not quantified well enough to derive well-constrained carbon budgets for North American coastal waters or to project how those fluxes will change in the future due to various drivers.

This chapter focused primarily on the role of ocean margins in sequestering atmospheric CO₂ and coastal ocean acidification. In the coastal ocean, a net removal of carbon from direct interaction with the atmospheric reservoir can occur by export of dissolved or particulate carbon to the deep ocean or by permanent burial in sediments. Neither of these is easily observed or well quantified. The best-observed flux is gas exchange across the air-sea interface, although extracting the small net flux and its trend from a variable background remains a challenge. Ultimately, the removal of anthropogenic carbon is the relevant quantity for assessing the contribution of ocean margins to the uptake of anthropogenic carbon; however, the separation of anthropogenic fluxes from the natural background is thus far elusive for coastal waters.

Estimates of air-sea CO₂ fluxes currently provide the best evidence for the contribution of coastal waters to overall carbon uptake by the ocean. In the broad shelf system of the North American Atlantic Coast, shelf water is separated from the adjacent open ocean by persistent shelf break currents and density fronts. Available estimates suggest that the overall North American Atlantic Coast is a weak sink, with some subregions acting as sources (e.g., nearshore regions of the SAB), while others are either neutral (Scotian Shelf and GOM) or act as weak sinks (MAB and outer SAB). Large sections of the narrow shelf of the North American Pacific Coast are dominated by upwelling circulation, which leads to strong CO₂ outgassing near the coast. However, compensating for this outgassing is biologically driven uptake from upwelled nutrients further offshore. Recent estimates are consistent in suggesting that the region is a weak to moderate sink of atmospheric CO₂. The relatively wide shelves in the GMx are considered a weak net sink, with the West Florida Shelf and the western Gulf shelf acting as sources; the Mexico shelf being neutral; and only the northern shelf a clear sink that is driven largely by anthropogenic nutrient inputs from the Mississippi River. The wide, seasonally ice-covered shelves in the North American Arctic consistently are acting as a sink for atmospheric CO₂. The low ­surface-water pCO₂ in this region primarily results from low water temperatures and the decreased uptake of atmospheric CO₂ during a significant fraction of the year because of seasonal ice cover. Overall, North American coastal waters act as a sink, but regional variations and uncertainties are large.

Several drivers influence secular trends in coastal carbon fluxes and will continue to do so in the future. These drivers include rising atmospheric CO₂ levels, changes in atmosphere-ocean interactions (e.g., wind forcing and heat fluxes), changes in the hydrological cycle, and anthropogenic perturbations of global nutrient cycling (particularly, the nitrogen cycle). Coastal surface pCO₂ clearly does not closely track atmospheric pCO₂. Although there are a number of plausible mechanisms for potential future changes in coastal carbon uptake, the total effect cannot be predicted with any confidence. Regional model studies are beginning to address these challenges.

A major concern is coastal acidification, which can affect the growth, metabolism, and life cycles of many marine organisms, specifically calcifiers, and can trigger cascading ecosystem-scale effects. Most vulnerable are those organisms that precipitate aragonite, one of the more soluble forms of biogenic CaCO₃ in the ocean. Aragonite saturation states are routinely below saturation (i.e., favoring dissolution) in North American Arctic coastal waters. In the North American Pacific Coast region, atmospheric CO₂ uptake in combination with intensified upwelling that brings low-pH, low-oxygen water onto the shelves leads to aragonite levels below the saturation threshold in large portions of the subsurface waters. In the northern GMx, aragonite saturation states are well above the dissolution threshold. Although eutrophication-induced acidification occurs in bottom waters influenced by Mississippi River inputs of nutrients and freshwater, saturation levels remain well above the dissolution threshold.

Given the importance of coastal margins, both in contributing to carbon budgets and in the societal benefits they provide, further efforts to improve assessments of the carbon cycle in these regions are paramount. Critical needs are maintaining and expanding existing coastal observing programs, continuing national and international coordination and integration of observations, increasing development of modeling capabilities, and addressing stakeholder needs.