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Evidence for upwelling of corrosive "acidified" water onto the Continental Shelf

Richard A. Feely1*, Christopher L. Sabine1, J. Martin Hernandez-Ayon2, Debby Ianson3, Burke Hales4

1Pacific Marine Environmental Laboratory, National Oceanic and Atmospheric Administration, Seattle, WA 98115–6349, USA
2Instituto de Investigaciones Oceanologicas, Universidad Autonoma de Baja California, Km. 103 Carr. Tijuana-Ensenada, Ensenada, Baja Calivornia, MEXICO
3Fisheries and Oceans Canada, Institute of Ocean Science, Post Office Box 6000, Sidney, BC V8L 4B2, CANADA
4College of Oceanic and Atmospheric Sciences, Oregon State University, 104 Ocean Administration Building, Corvallis, OR 97331-5503, USA

Science, 320(5882), 10.1126/science.1155676, 1490–1492 (2008)
Copyright ©2008 by The American Association for the Advancement of Science. Further electronic distribution is not allowed.

Abstract

The absorption of atmospheric carbon dioxide (CO2) into the ocean lowers the pH of the waters. This so-called ocean acidification could have important consequences for marine ecosystems. To better understand the extent of this ocean acidification in coastal waters, we conducted hydrographic surveys along the continental shelf of western North America from central Canada to northern Mexico. We observed seawater that is undersaturated with respect to aragonite upwelling onto large portions of the continental shelf, reaching depths of ~40 to 120 meters along most transect lines and all the way to the surface on one transect off northern California. Although seasonal upwelling of the undersaturated waters onto the shelf is a natural phenomenon in this region, the ocean uptake of anthropogenic CO2 has increased the areal extent of the affected area.


Over the past 250 years, the release of carbon dioxide (CO2) from industrial and agricultural activities has resulted in atmospheric CO2 concentrations that have increased by about 100 parts per million (ppm). The atmospheric concentration of CO2 is now higher than it has been for at least the past 650,000 years, and is expected to continue to rise at an increasing rate, leading to pronounced changes in our climate by the end of this century (Siegenthaler et al., 2005). Since the beginning of the industrial era, the oceans have absorbed ~127 ± 18 billion metric tons of carbon as CO2 from the atmosphere, or about one-third of the anthropogenic carbon emissions released (Sabine and Feely, 2007). This process of absorption of anthropogenic CO2 has benefited humankind by substantially reducing the greenhouse gas concentrations in the atmosphere and minimizing some of the impacts of global warming. However, the ocean’s daily uptake of 22 million metric tons of CO2 has a sizable impact on its chemistry and biology. Recent hydrographic surveys and modeling studies have confirmed that the uptake of anthropogenic CO2 by the oceans has resulted in a lowering of seawater pH by about 0.1 since the beginning of the industrial revolution (Feely et al., 2004; Sabine et al., 2004; Orr et al., 2005; Caldeira and Wickett, 2005; Feely et al., 2008). In the coming decades, this phenomenon, called “ocean acidification,” could affect some of the most fundamental biological and geochemical processes of the sea and seriously alter the fundamental structure of pelagic and benthic ecosystems (Kleypas et al., 2006).

Estimates of future atmospheric and oceanic CO2 concentrations, based on the Intergovernmental Panel on Climate Change (IPCC) CO2 emission scenarios and general circulation models, indicate that atmospheric CO2 concentrations could exceed 500 ppm by the middle of this century, and 800 ppm near the end of the century. This increasewould result in a decrease in surface-water pH of ~0.4 by the end of the century, and a corresponding 50% decrease in carbonate ion concentration (Orr et al., 2005; Solomon et al., 2007). Such rapid changes are likely to negatively affect marine ecosystems, seriously jeopardizing the multifaceted economies that currently depend on them (Raven et al., 2005).

The reaction of CO2 with seawater reduces the availability of carbonate ions that are necessary for calcium carbonate (CaCO3) skeleton and shell formation for marine organisms such as corals, marine plankton, and shellfish. The extent to which the organisms are affected depends largely on the CaCO3 saturation state (W), which is the product of the concentrations of Ca2+ and CO32 divided by the apparent stoichiometric solubility product for either aragonite or calcite:

Ωarag = [Ca2+][CO32]/K'sp_arag (1)

Ωcal = [Ca2+][CO32–]/K'sp_cal (2)

where the calcium concentration is estimated from the salinity, and the carbonate ion concentration is calculated from the dissolved inorganic carbon (DIC) and total alkalinity (TA) measurements (11). In regions where Ωarag or Ωcal is > 1.0, the formation of shells and skeletons is favored. Below a value of 1.0, the water is corrosive and dissolution of pure aragonite and unprotected aragonite shells will begin to occur (Feely et al., 1988). Recent studies have shown that in many regions of the ocean, the aragonite saturation horizon shoaled as much as 40 to 200 m as a direct consequence of the uptake of anthropogenic CO2 (Feely et al., 2004; Orr et al., 2005; Caldeira and Wickett, 2005). It is shallowest in the northeastern Pacific Ocean, only 100 to 300 m from the ocean surface, allowing for the transport of undersaturated waters onto the continental shelf during periods of upwelling.

In May and June 2007, we conducted the North American Carbon Program (NACP) West Coast Cruise on the Research Ship Wecoma along the continental shelf of western North America, completing a series of 13 cross-shelf transects from Queen Charlotte Sound, Canada, to San Gregorio Baja California Sur, Mexico (Fig. 1). Full water column conductivity-temperature-depth rosette stations were occupied at specified locations along each transect (Fig. 1). Water samples were collected in modified Niskin-type bottles and analyzed for DIC, TA, oxygen, nutrients, and dissolved and particulate organic carbon. Aragonite and calcite saturation, seawater pH (pHSW), and partial pressure of CO2 (pCO2) were calculated from the DIC and TA data (11).

fig01

Fig. 1. Distribution of the depths of the undersaturated water (aragonite saturation < 1.0; pH < 7.75) on the continental shelf of western North America from Queen Charlotte Sound, Canada, to San Gregorio Baja California Sur, Mexico. On transect line 5, the corrosive water reaches all the way to the surface in the inshore waters near the coast. The black dots represent station locations.

The central and southern coastal region off western North America is strongly influenced by seasonal upwelling, which typically begins in early spring when the Aleutian low-pressure system moves to the northwest and the Pacific Highmoves northward, resulting in a strengthening of the northwesterly winds (Hickey, 1998; Pennington and Chavez, 2000). These winds drive net surface-water Ekman transport offshore, which induces the upwelling of CO2-rich, intermediatedepth (100 to 200 m) offshore waters onto the continental shelf. The upwelling lasts until late summer or fall, when winter storms return.

During the cruise, various stages and strengths of upwelling were observed from line 2 off central Vancouver Island to line 11 off Baja California, Mexico.We observed recent upwelling on lines 5 and 6 near the Oregon-California border. Coincident with the upwelled waters, we found evidence for undersaturated, low-pH seawater in the bottom waters as depicted by Ωarag values < 1.0 and pH values < 7.75. The corrosive waters reached mid-shelf depths of ~40 to 120 m along lines 2 to 4 and lines 7 to 13 (Fig. 1). In the region of the strongest upwelling (line 5), the isolines of Ωarag = 1.0, DIC = 2190, and pH = 7.75 closely followed the 26.2 potential density surface (Fig. 2). This density surface shoaled from a depth of ~150 m in the offshore waters and breached the surface over the shelf near the 100-m bottom contour, ~40 km from the coast. This shoaling of the density surfaces and CO2-rich waters as one approaches land is typical of strong coastal upwelling conditions (van Geen et al., 2000; Friederich et al., 2002; Ianson et al., 2003; Hales et al., 2005). The surface-water pCO2 on the 26.2 potential density surface was about 850 μatm near the shelfbreak and higher inshore (Fig. 2), possibly enhanced by respiration processes on the shelf (Ianson et al., 2003). These results indicate that the upwelling process caused the entire water column shoreward of the 50-m bottom contour to become undersaturated with respect to aragonite, a condition that was not predicted to occur in openocean surface waters until 2050 (Orr et al., 2005). On line 6, the next transect south, the undersaturated water was close to the surface at ~22 km from the coast. The lowest Ωarag values (<0.60) observed in the nearbottom waters of the continental shelf corresponded with pH values close to 7.6. Because the calcite saturation horizon is located between 225 and 400 m in this part of the northeastern Pacific (Feely et al., 2002), it is still too deep to shoal onto the continental shelf. Nevertheless, the calcite saturations values drop in the core of the upwelled water (Ωcal < 1.3).

fig02

Fig. 2. Vertical sections of (A) temperature, (B) aragonite saturation, (C) pH, (D) DIC, and (E) pCO2 on transect line 5 off Pt. St. George, California. The potential density surfaces are superimposed on the temperature section. The 26.2 potential density surface delineates the location of the first instance in which the undersaturated water is upwelled from depths of 150 to 200 m onto the shelf and outcropping at the surface near the coast. The red dots represent sample locations.

As noted, the North Pacific aragonite saturation horizons are among the shallowest in the global ocean (Feely et al., 2004). The uptake of anthropogenic CO2 has caused these horizons to shoal by 50 to 100 m since preindustrial times so that they are within the density layers that are currently being upwelled along the west coast of North America. Although much of the corrosive character of these waters is the natural result of respiration processes at intermediate depths below the euphotic zone, this region continues to accumulate more anthropogenic CO2 and, therefore, the upwelling processes will expose coastal organisms living in the water column or at the sea floor to less saturated waters, exacerbating the biological impacts of ocean acidification.

On the basis of our observed O2 values and estimated O2 consumption rates on the same density surfaces (Hales et al., 2005; Feely et al., 2002; Feely et al., 2004), the upwelled water off northern California (line 5) was last at the surface about 50 years ago, when atmospheric CO2 was about 65 ppm lower than it is today. The open-ocean anthropogenic CO2 distributions in the Pacific have been estimated previously (Sabine et al., 2004; Feely et al., 2002; Sabine et al., 2002). By determining the density dependence of anthropogenic CO2 distributions in the eastern-most North Pacific stations of the Sabine et al. (2002) data set, we estimate that these upwelledwaters contain ~31 ± 4 μmol kg–1 anthropogenic CO2 (fig. S2). Removing this signal from the DIC increases the aragonite saturation state of the waters by about 0.2 units. Thus, without the anthropogenic signal, the equilibrium aragonite saturation level (Ωarag = 1) would be deeper by about 50 m across the shelf, and no undersaturated waters would reach the surface. Water already in transit to upwelling centers carries increasing anthropogenic CO2 and more corrosive conditions to the coastal oceans of the future. Thus, the undersaturated waters, which were mostly a problem for benthic communities in the deeper waters near the shelf break in the preindustrial era, have shoaled closer to the surface and near the coast because of the additional inputs of anthropogenic CO2.

These observations clearly show that seasonal upwelling processes enhance the advancement of the corrosive deep water into broad regions of the North American western continental shelf. Because the region experiences seasonal periods of enhanced aragonite undersaturation, it is important to understand how the indigenous organisms deal with this exposure and whether future increases in the range and intensity of the corrosiveness will affect their survivorship. Presently, little is known about how this intermittent exposure to corrosive water might affect the development of larval, juvenile, and adult stages of aragonitic calcifying organisms or finfish that populate the neritic and benthic environments in this region and fuel a thriving economy. Laboratory andmesocosm experiments show that these changes in saturation state may cause substantial changes in overall calcification rates formany species of marine calcifiers including corals, coccolithophores, foraminifera, and pteropods, which are a major food source for local juvenile salmon (Kleypas et al., 2006; Green et al., 2004; Guinotte et al., 2003; Langdon and Atkinson, 2005; Spero et al., 1997; Riebesell et al., 2000; Zondervan et al., 2001; Seibel and Fabry, 2003; Delille et al., 2005; Engel et al., 2005). Similar decreases in calcification rates would be expected for ediblemussels, clams, and oysters (Green et al., 2004; Gazeau et al., 2007). Other research indicates that many species of juvenile fish and shellfish of economic importance to coastal regions are highly sensitive to higher-than-normal CO2 concentrations such that high rates of mortality are directly correlated with the higher CO2 concentrations (Gazeau et al., 2007; Ishimatsu et al., 2004). Although comprehensive field studies of organisms and their response to sporadic increases in CO2 along thewestern NorthAmerican coast are lacking, current studies suggest that further research under field conditions is warranted. Our results show that a large section of the North American continental shelf is affected by ocean acidification. Other continental shelf regions may also be affected where anthropogenic CO2-enriched water is being upwelled onto the shelf.

Acknowledgments

We thank Captain Richard Verlini and the crew of the R/V Wecoma for logistics support. We also thank D. Greeley, D. Wisegarver, P. Covert, and S. Barry for the DIC and TA measurements. Financial support for this work was provided by the National Oceanic and Atmospheric Administration’s Global Carbon Cycle Program and the National Aeronautical and Space Administration Ocean Biology and Biogeochemistry Program.

References and Notes

1. U. Siegenthaler et al., Science 310, 1313 (2005).

2. C. L. Sabine, R. A. Feely, in Greenhouse Gas Sinks, D. Reay, N. Hewitt, J. Grace, K. Smith, Eds. (CABI,
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3. R. A. Feely et al., Science 305, 362 (2004).

4. C. L. Sabine et al., Science 305, 367 (2004).

5. J. C. Orr et al., Nature 437, 681 (2005).

6. K. Caldeira, M. E. Wickett, J. Geophys. Res. Oceans 110, 365 (2005).

7. R. A. Feely et al., PICES Press 16, 22 (2008).

8. J. A. Kleypas et al., “Impacts of Increasing Ocean Acidification on Coral Reefs and Other Marine Calcifiers: A Guide for Future Research,” report of a workshop held 18 to 20 April 2005, St. Petersburg, FL, sponsored by NSF, NOAA, and the U.S. Geological Survey (2006).

9. S. Solomon et al., Eds, in Contribution of Working Group I to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change (Cambridge Univ. Press, Cambridge and New York, 2007).

10. J. Raven et al., “Ocean acidification due to increasing atmospheric carbon dioxide,” policy document 12/05 (The Royal Society, London, 2005).

11. The details of the analytical methods and calculations for the carbonate system and anthropogenic CO2 are given in the supporting online material.

12. R. A. Feely et al., Mar. Chem. 25, 227 (1988).

13. B. Hickey, in The Sea, A. R. Robinson, K. H. Brink, Eds. (Wiley, New York, 1998), vol. 2.

14. J. Timothy Pennington, F. P. Chavez, Deep Sea Res. Part II Top. Stud. Oceanogr. 47, 947 (2000).

15. A. van Geen et al., Deep Sea Res. Part II Top. Stud. Oceanogr. 47, 975 (2000).

16. G. E. Friederich, P. M. Walz, M. G. Burczynski, F. P. Chavez, Prog. Oceanogr. 54, 185 (2002).

17. D. Ianson et al., Deep Sea Res. Part I Oceanogr. Res. Pap. 50, 1023 (2003).

18. B. Hales et al., Global Biogeochem. Cycles 19, 10.1029/2004GB002295 (2005).

19. R. A. Feely et al., Global Biogeochem. Cycles 16, 1144 (2002).

20. R. A. Feely et al., J. Oceanogr. 60, 45 (2004).

21. C. L. Sabine et al., Global Biogeochem. Cycles 16, 1083 (2002).

22. M. A. Green, M. E. Jones, C. L. Boudreau, R. L. Moore, B. A. Westman, Limnol. Oceanogr. 49, 727 (2004).

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25. H. J. Spero et al., Nature 390, 497 (1997).

26. U. Riebesell et al., Nature 407, 364 (2000).

27. I. Zondervan et al., Global Biogeochem. Cycles 15, 507 (2001).

28. B. A. Seibel, V. J. Fabry, Adv. Appl. Biodivers. Sci. 4, 59 (2003).

29. B. Delille et al., Global Biogeochem. Cycles 19, GB2023 (2005).

30. A. Engel et al., Limnol. Oceanogr. 50, 493 (2005).

31. F. Gazeau et al., Geophys. Res. Lett. 34, L07603 (2007).

32. A. Ishimatsu et al., J. Oceanogr. 60, 731 (2004).

33. We thank Captain Richard Verlini and the crew of the R/V Wecoma for logistics support. We also thank D. Greeley, D. Wisegarver, P. Covert, and S. Barry for the DIC and TA measurements. Financial support for this work was provided by the National Oceanic and Atmospheric Administration’s Global Carbon Cycle Program and the National Aeronautical and Space Administration Ocean Biology and Biogeochemistry Program.

Supporting Online Material

www.sciencemag.org/cgi/content/full/1155676/DC1
Materials and Methods
Figs. S1 and S2
References

25 January 2008; accepted 13 May 2008
Published online 22 May 2008;
10.1126/science.1155676
Include this information when citing this paper.


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