Encyclopedia of Marine Geosciences

Living Edition
| Editors: Jan Harff, Martin Meschede, Sven Petersen, Jörn Thiede

Anoxic Oceans

  • Christian MärzEmail author
  • Hans-Jürgen Brumsack
Living reference work entry

Latest version View entry history

DOI: https://doi.org/10.1007/978-94-007-6644-0_216-2

Keywords

Black Shale Hydrogen Sulfide Benthic Foraminifera Organic Matter Degradation Shallow Marine Environment 
These keywords were added by machine and not by the authors. This process is experimental and the keywords may be updated as the learning algorithm improves.

Definition

A large body of open marine water free of dissolved molecular oxygen and potentially containing dissolved hydrogen sulfide.

Introduction

Today’s world oceans are dominated by oxic water masses, except for some marginal basins like the Black Sea, fjords, upwelling areas, and so-called coastal “anoxic dead zones.” However, since the formation of oceans on Earth, anoxic (oxygen-free) conditions have occurred frequently in parts of the global ocean, leading to the deposition of organic-rich sediments that subsequently generated oil or gas, the main energy source for our modern civilization. Here we will focus on deep ocean anoxia, their environmental controls, and their geological expressions.

The Proterozoic

Throughout large parts of the Precambrian, the oceans were anoxic. Oxygen-producing photosynthetic organisms like algae were only beginning to evolve, and atmospheric oxygen levels were orders of magnitude lower than at present. Due to this oxygen-poor atmosphere, sulfate was not produced in significant quantities by chemical weathering of magmatic rocks on land, and sulfate concentrations in the early oceans were low. In contrast, dissolved Fe concentrations were very high due to hydrothermal input related to ocean crust formation, reductive weathering on land, and possibly diagenetic Fe release from sediments. These oxygen-free, sulfate-poor but iron-rich (ferruginous) oceans persisted over much of the Proterozoic until the Great Oxidation Event (Canfield, 1998), which initiated the global deposition of banded iron formations (BIFs). When atmospheric oxygen concentrations reached a threshold, the oxidative weathering of sulfide minerals led to the establishment of the marine sulfate pool. Consequently, organic matter degradation by microbial sulfate reduction started to release hydrogen sulfide into the water column. As the dissolved Fe in the water column reacted with hydrogen sulfide to form pyrite, ferruginous conditions were gradually replaced by sulfidic (euxinic) conditions in the marine environment. With increasing atmospheric oxygen, the spatial extent of marine anoxia/euxinia was reduced, only persisting in certain ocean regions (Poulton et al., 2010). While the expansion and duration of Proterozoic deep ocean anoxia were never reached again, there were still periods in the Phanerozoic with widely distributed deep ocean anoxia.

The Phanerozoic

Several intervals of deep ocean anoxia occurred throughout the Phanerozoic (see review by Meyer and Kump, 2008). Their most prominent geological expressions include organic-rich deposits (“black shales”) from the Cambrian of China, the Ordovician of North America, the Silurian of North Africa, and the Devonian-Lower Carboniferous of the USA and Europe. Most of these Paleozoic deposits represent shallow marine environments, and synchronous records from the deep oceans are largely missing. In contrast, there is evidence for widespread deepwater anoxia around the Permian-Triassic boundary, coinciding with the most severe extinction event in Earth history (Meyer and Kump, 2008). In the Mesozoic, a number of rather short-termed but severe and widely recognized intervals of ocean anoxia occurred, the so-called oceanic anoxic events (OAEs; Schlanger and Jenkyns, 1976). Although their full global expansion is debated (e.g., Trabucho Alexandre et al., 2010), there is evidence that at least the North Atlantic and Tethys were periodically affected by intense deepwater anoxia to euxinia, whereas comparable records from the Pacific and Indian Oceans are scarce. The Toarcian Oceanic Anoxic Event (T-OAE) and the Cenomanian-Turonian Boundary Event (CTBE or OAE2) are considered to be the most widespread occurrences of Mesozoic marine anoxia (Jenkyns, 2010). The Cenozoic sediment record is comparably poor in black shale-type deposits. Organic-rich sediments related to the Paleocene-Eocene Thermal Maximum (PETM) are mostly limited to shallow marine environments (Cohen et al., 2007). However, the deep Arctic Ocean record shows clear signs of Paleocene and Eocene deepwater anoxia, most likely related to its restricted circulation and stable water column stratification. The most recent witnesses of short-termed deepwater anoxia are Pliocene-Pleistocene sapropels of the Eastern Mediterranean (Emeis et al., 2000; De Lange et al., 2008).

“Modern Analogues”

In the modern ocean, there are no true analogues for open marine black shale formation. However, a number of “near-analogue” environments can serve as natural laboratories to study some of the physical, chemical, and biological processes that interacted in past anoxic oceans (Demaison and Moore, 1980). The most prominent example is the Black Sea, Earth’s largest permanently euxinic marine basin (Degens and Ross, 1974). It is characterized by moderate primary productivity but strong salinity stratification and a very restricted connection to the Mediterranean. A similarly well-studied but less restricted and less euxinic marine environment is the Cariaco Basin, and also a number of fjords and bays around the world (e.g., Framvaren Fjord, Saanich Inlet, Kau Bay) exhibit multi-annual deepwater anoxia to euxinia (Richards, 1965; Tissot and Welte, 1984). Also coastal upwelling areas (e.g., off Namibia, Chile, Peru, Arabian Sea) share common features with past anoxic oceans as they display high seasonal primary productivity, leading to deepwater oxygen consumption and occasional development of sulfidic water masses (Copenhagen, 1953). Over the past decades, the study of geochemically similar but spatially restricted analogues has advanced our understanding of past anoxic oceans. A wide range of geochemical proxies is now at hand to unravel the paleoenvironmental setting from sedimentary archives.

Environmental Controls: Organic Matter Export, Water Column Stratification, Plate Tectonics, and Astronomical Forcing

Two fundamental environmental parameters are traditionally suggested as drivers of oceanic anoxia: high primary productivity inducing enhanced organic matter export to the seafloor and enhanced organic matter preservation by limited deepwater ventilation, i.e., low oxygen (Demaison and Moore, 1980). Modern end-members of this “productivity versus preservation” debate are the Black Sea, on the one hand (moderate productivity, severe basin restriction, salinity stratification, absence of molecular oxygen below redoxcline), and upwelling areas, on the other hand (full connection to the open ocean, seasonally very high productivity). Recent studies suggest that most black shale deposits do not allow a strict differentiation between organic matter productivity and preservation as dominant cause for organic carbon accumulation and anoxia. Past examples of restricted anoxic basins include the mid-Cretaceous proto-North Atlantic and the Paleogene Arctic Ocean. Past high-productivity regions with anoxic water masses developed in the mid-Cretaceous equatorial Atlantic as nutrients were “trapped” due to specific circulation patterns (Trabucho Alexandre et al., 2010). To create and maintain marine anoxia, a range of important biogeochemical feedback mechanisms needs to be considered. One well-known “anoxia-productivity feedback” is the efficient recycling of phosphate from marine sediments if bottom-water oxygen is strongly depleted (Ingall et al., 1993; Slomp et al., 2004; März et al., 2008). Recycled phosphate can then be reintroduced into the photic zone and sustain high-productivity and organic matter export.

Although the onsets and terminations of widespread ocean anoxia are still debated, they appear to be related to plate tectonics and volcanism (Meyer and Kump, 2008). Astronomical forcing can regulate the intensity of anoxia once favorable conditions have been created (De Lange et al., 2008). Plate tectonic constellations are responsible for the restriction of ocean basins (e.g., the mid-Cretaceous North and South Atlantic), reducing the renewal and oxygenation of deepwater masses (Erbacher et al., 2001). Volcanic activity introduces large amounts of the greenhouse gas CO2 into the atmosphere, leading to both global warming and enhanced chemical weathering (Turgeon and Creaser, 2008). Warming reduces the oxygen solubility in the oceans and the vertical mixing of its water masses, and chemical weathering enhances terrestrial nutrient input into the oceans (Adams et al., 2010). Global warming leads to eustatic sea-level rise, concomitant widespread flooding, and an expansion of shelf sea areas. Besides the increased potential for shelf anoxia, these shallow marine areas also trap a large fraction of siliciclastic detritus before it reaches the deep basins. Therefore, organic-rich deepwater deposits are often characterized by reduced terrestrial dilution (Arthur et al., 1988).

Recognizing Anoxia in Sedimentary Archives

As most multicellular organisms (such as benthic foraminifera) require oxygen to survive, anoxic environments are usually strongly deprived in macroscopic life. The sedimentological expressions of this lack of burrowing macrobenthos are finely laminated deposits without any bioturbation.

Both organic and inorganic geochemistry provides a number of proxies to detect anoxia/euxinia in different parts of a water column. For instance, a biomarker to detect euxinic conditions in the photic zone is isorenieratene, an organic pigment exclusively produced by phototrophic green and purple sulfur bacteria (Sinninghe Damsté et al., 1993).

A first-order geochemical approach to reconstruct marine anoxia was the organic carbon to total sulfur (C/S) ratio of Black Sea deposits (Berner, 1984). While normal, oxic marine sediments exhibited C/S ratios around 3, the production of H2S by organic matter degradation in the Black Sea, and its subsequent precipitation as metal sulfides (mostly pyrite, FeS2), resulted in C/S ratios <3. A better representation of the C-S-Fe interrelationship is a ternary diagram with the end-members organic C, total S, and reactive Fe, as it allows a distinction between anoxic systems where pyrite formation was limited by H2S or reactive Fe (Dean and Arthur, 1989).

Under anoxic ferruginous conditions, the sulfate depletion of the water column prevents any enrichment of sedimentary S and hence the recognition of water column anoxia using bulk sedimentary C-S-Fe proxies (Poulton and Canfield, 2011). To distinguish non-sulfidic from sulfidic conditions, a sequential Fe extraction scheme distinguishes between the relative amounts of the Fe pool that can potentially react, or already have reacted, with H2S (Poulton and Canfield, 2005). In addition to C-S-Fe systematic, the selective removal of sedimentary P relative to organic C under anoxic conditions results in increasing sedimentary C/P ratios under increasingly oxygen-depleted bottom-water conditions (Algeo and Ingall, 2007).

Among the most powerful redox proxies are trace metal enrichments/depletions relative to their lithogenic background contents (Brumsack, 2006). Redox-sensitive metals change their redox states under oxic, suboxic, and anoxic conditions (e.g., Mn, U, V); sulfide-forming metals precipitate sulfide phases/coprecipitate with pyrite if H2S is available (e.g., Mo, Cd, Zn). While questions remain in our understanding of natural trace metal systematics, several metals have emerged over the years as reliable redox indicators. While Mn tends to be depleted from sediments already under suboxic conditions, Re becomes enriched within sulfidic microenvironments in the sediment. Uranium is enriched if full anoxia is reached in the sediments, while Cd and Zn precipitate as sulfides under weakly sulfidic conditions. Molybdenum requires stronger euxinia in the water column to be transferred to particle-reactive thiomolybdates. The interpretation of trace metal records is complicated by the fact that some of these elements serve as micronutrients and are pre-concentrated via plankton (Böning et al., 2004) before burial. They may as well respond to specific redox conditions in the surface sediment, while others respond to redox changes in the water column. Early diagenetic processes pose another challenge, as trace metals may be redistributed within the sediment, biasing the primary redox record.

Beyond element contents and speciation in marine sediments, the stable isotopic signatures of metals like Mo and Fe are increasingly used as redox indicators (Anbar and Rouxel, 2007). The specific value of Mo isotopes is based on the quantitative removal of Mo from a water column under fully euxinic conditions without any isotope fractionation. Through isotope and mass balance calculations, this ultimately gives the percentage of the global oceanic water mass that was fully euxinic at a given point in time. The stable Fe isotope composition of marine sediments, in conjunction with total Fe enrichment patterns, is controlled by the input of isotopically light Fe from suboxic shelf sediments and the shuttling and distribution of this excess Fe within a suboxic redoxcline. Therefore, the depth of the redoxcline in a water column and the extent of underlying euxinic water masses can be reconstructed (Eckert et al., 2013). However, the full application of these isotope systems as paleo-redox proxies will require a better understanding of these trace metals and their isotopes in the marine environment under a variety of boundary conditions.

Conclusions

Our understanding of anoxic conditions in the oceans has greatly advanced over the past decades. The classic “productivity versus preservation” discussion sparked many of these advances and has motivated scientists to take a closer look at the underlying mechanisms causing, sustaining, and terminating ocean anoxia. This was achieved by combining studies of modern and ancient anoxic marine environments and by working across scientific disciplines. Nowadays, geoscientists can take advantage of a large set of analytical tools and paleoenvironmental proxies that led to a better understanding of, e.g., the causes for the development of anoxic conditions, their effects on organisms, global element cycles and the climate system, duration and spatial extent of anoxia and euxinia, the degree of anoxia or euxinia, and the reasons for their termination. Still, many questions remain about marine anoxia. Facing a warming ocean and intensifying oxygen depletion in coastal and open oceans, these questions are certainly relevant for the development and understanding of our future ocean.

Cross-References

Bibliography

  1. Adams, D. D., Hurtgen, M. T., and Sageman, B. B., 2010. Volcanic triggering of a biogeochemical cascade during oceanic anoxic event 2. Nature Geosciences, 3, 201–204.CrossRefGoogle Scholar
  2. Algeo, T. J., and Ingall, E., 2007. Sedimentary Corg:P ratios, paleocean ventilation, and Phanerozoic atmospheric pO2. Palaeogeography, Palaeoclimatology, Palaeoecology, 256, 130–155.CrossRefGoogle Scholar
  3. Anbar, A. D., and Rouxel, O., 2007. Metal stable isotopes in paleoceanography. Annual Reviews of Earth and Planetary Sciences, 35, 717–746.CrossRefGoogle Scholar
  4. Arthur, M. A., Brumsack, H.-J., Jenkyns, H. C., and Schlanger, S. O., 1988. Stratigraphy, geochemistry, and paleoceanography of organic carbon-rich Cretaceous sequences. In Ginsburg, R. N., and Beaudoin, B. (eds.), Cretaceous Resources, Events and Rhythms: Background and Plans for Research. Proceedings of ARW. Digne: Kluwer Academic Publishers.Google Scholar
  5. Berner, R. A., 1984. Sedimentary pyrite formation: an update. Geochimica et Cosmochimica Acta, 48, 605–615.CrossRefGoogle Scholar
  6. Böning, P., Brumsack, H.-J., Böttcher, M. E., Schnetger, B., Kriete, C., Kallmeyer, J., and Borchers, S. L., 2004. Geochemistry of Peruvian near-surface sediments. Geochimica et Cosmochimica Acta, 68, 4429–4451.CrossRefGoogle Scholar
  7. Brumsack, H.-J., 2006. The trace metal content of recent organic-rich sediments: implications for Cretaceous black shale formation. Palaeogeography, Palaeoclimatology, Palaeoecology, 232, 344–361.CrossRefGoogle Scholar
  8. Calvert, S. E., and Price, N. B., 1970. Minor metal contents of recent organic-rich sediment off South West Africa. Nature, 227, 593–595.CrossRefGoogle Scholar
  9. Canfield, D. E., 1998. A new model for Proterozoic ocean chemistry. Nature, 396, 450–453.CrossRefGoogle Scholar
  10. Cohen, A. S., Coe, A. L., and Kemp, D. B., 2007. The Late Palaeocene-Early Eocene and Toarcian (Early Jurassic) carbon isotope excursions: a comparison of their time scales, associated environmental changes, causes and consequences. Journal of the Geological Society of London, 164, 1093–1108.CrossRefGoogle Scholar
  11. Copenhagen, W. J., 1953. The Periodic Mortality of Fish in the Walvis Region. Investigational Report 14, Division of Fisheries Union of South Africa.Google Scholar
  12. De Lange, G. J., Thomson, J., Reitz, A., Slomp, C. P., Speranza Principato, M., Erba, E., and Corselli, C., 2008. Synchronous basin-wide formation and redox-controlled preservation of a Mediterranean sapropel. Nature Geosciences, 1, 606–610.CrossRefGoogle Scholar
  13. Dean, W. E., and Arthur, M. A., 1989. Iron-sulfur-carbon relationships in organic-carbon-rich sequences I: cretaceous western interior seaway. American Journal of Science, 289, 708–743.CrossRefGoogle Scholar
  14. Degens, E. T., and Ross, D. A., 1974. The Black Sea – Geology, Chemistry, and Biology. American Association of Petroleum Geologists Memoir 20, Tulsa, Oklahoma.Google Scholar
  15. Demaison, G. J., and Moore, G. T., 1980. Anoxic environments and oil source bed genesis. Organic Geochemistry, 2, 9–31.CrossRefGoogle Scholar
  16. Eckert, S., Brumsack, H.-J., Severmann, S., Schnetger, B., März, C., and Fröllje, H., 2013. Establishment of euxinic conditions in the Holocene Black Sea. Geology, 41, 431–434.CrossRefGoogle Scholar
  17. Emeis, K. C., Sakamoto, T., Wehausen, R., and Brumsack, H.-J., 2000. The sapropel record of the eastern Mediterranean Sea – results of Ocean Drilling Program Leg 160. Palaeogeography, Palaeoclimatology, Palaeoecology, 158, 371–395.CrossRefGoogle Scholar
  18. Erbacher, J., Huber, B. T., Norris, R. D., and Markey, M., 2001. Increased thermohaline stratification as a possible cause for an ocean anoxic event in the Cretaceous period. Nature, 409, 325–327.CrossRefGoogle Scholar
  19. Ingall, E. D., Bustin, R. M., and Van Cappellen, P., 1993. Influence of water column anoxia on the burial and preservation of carbon and phosphorus in marine shales. Geochimica et Cosmochimica Acta, 57, 303–316.CrossRefGoogle Scholar
  20. Jenkyns, H. C., 2010. Geochemistry of oceanic anoxic events. Geochemistry, Geophysics, Geosystems, 11, Q03004, doi:10.1029/2009GC002788.CrossRefGoogle Scholar
  21. Lyons, T. W., Anbar, A. D., Severmann, S., Scott, C., and Gill, B. G., 2009. Tracking euxinia in the ancient ocean: a multiproxy perspective and Proterozoic case study. Annual Reviews in Earth and Planetary Sciences, 37, 507–534.CrossRefGoogle Scholar
  22. März, C., Poulton, S. W., Beckmann, B., Küster, K., Wagner, T., and Kasten, S., 2008. Redox sensitivity of P cycling during marine black shale formation: dynamics of sulfidic and anoxic, non-sulfidic bottom waters. Geochimica et Cosmochimica Acta, 72, 3703–3717.CrossRefGoogle Scholar
  23. Meyer, K. M., and Kump, L., 2008. Oceanic euxinia in Earth history: causes and consequences. Annual Reviews in Earth and Planetary Sciences, 36, 251–288.CrossRefGoogle Scholar
  24. Poulton, S. W., and Canfield, D. E., 2005. Development of a sequential extraction procedure for iron: implications for iron partitioning in continentally derived particulates. Chemical Geology, 214, 209–221.CrossRefGoogle Scholar
  25. Poulton, S. W., and Canfield, D. E., 2011. Ferruginous conditions: a dominant feature of the ocean through Earth’s history. Elements, 7, 107–112.CrossRefGoogle Scholar
  26. Poulton, S. W., Fralick, P. W., and Canfield, D. E., 2010. Spatial variability in oceanic redox structure 1.8 billion years ago. Nature Geosciences, 3, 486–490.CrossRefGoogle Scholar
  27. Richards, F. A., 1965. Anoxic basins and fjords. In Riley, J. P., and Skirrow, G. (eds.), Chemical Oceanography. London: Academic, pp. 611–643.Google Scholar
  28. Schlanger, S. O., and Jenkyns, H. C., 1976. Cretaceous oceanic anoxic events: causes and consequences. Geologie en Mijnbouw, 55, 179–184.Google Scholar
  29. Sinninghe Damsté, J. S., Wakeham, S. G., Kohnen, M. E. L., Hayes, J. M., and De Leeuw, J. W., 1993. A 6,000-year sedimentary molecular record of chemocline excursions in the Black Sea. Nature, 362, 827–829.CrossRefGoogle Scholar
  30. Slomp, C. P., Thomson, J., and De Lange, G. J., 2004. Controls on phosphorus regeneration and burial during formation of eastern Mediterranean sapropels. Marine Geology, 203, 141–159.CrossRefGoogle Scholar
  31. Tissot, B. P., and Welte, D. H., 1984. Petroleum Formation and Occurrence. Berlin: Springer.CrossRefGoogle Scholar
  32. Trabucho Alexandre, J., Tuenter, E., Henstra, G. A., Van der Zwan, K., Van de Wal, R. S. W., Dijkstra, H. A., and De Boer, P. L., 2010. The mid-Cretaceous North Atlantic nutrient trap: black shales and OAEs. Paleoceanography, 25, PA4201, doi:10.1029/2010PA001925.CrossRefGoogle Scholar
  33. Turgeon, S. C., and Creaser, R. A., 2008. Cretaceous oceanic anoxic event 2 triggered by a massive magmatic episode. Nature, 454, 323–326.CrossRefGoogle Scholar

Copyright information

© Springer Science+Business Media Dordrecht 2015

Authors and Affiliations

  1. 1.School of Civil Engineering and Geosciences (CEGS)Newcastle UniversityNewcastle upon TyneUK
  2. 2.Institute for Chemistry and Biology of the Marine EnvironmentUniversity of OldenburgOldenburgGermany