Posted by: chrismaser | March 13, 2012


We take air for granted without realizing that the gases, which comprise Earth’s atmosphere, are leaking into space. This leakage explains many of the solar system’s mysteries—why Mars is red, for instance. The planet is red because its water vapor has been separated into hydrogen and oxygen; then, as the hydrogen drifted away, the surplus oxygen oxidized, which means it essentially rusted the rocks.1


To take something “for granted” is to be certain of the status quo, and the status quo is impossible in a world of infinite novelty. For example, if you could travel back to the Archaean Era, you would not recognize Earth as the same planet you inhabit. The sun, for example, was not as bright as it is now. By that time, however, the Earth’s crust had cooled enough for rocks and continental plates to begin forming. The atmosphere was likely composed of methane, ammonia, and other gases that would be toxic to most life today. Yet, it was early in the Archaean Era that life first appeared on Earth, as attested by oldest fossils of cyanobacteria bacteria (often thought of as “blue-green algae”) that date back to roughly 3.5 billion years ago, and are still among the oldest fossils known. Concentrations of atmospheric oxygen rose from negligible levels to about 21 percent of that present today and can be attributed to the cyanobacteria, which have also been tremendously important in shaping the course of evolution and ecological change throughout Earth’s history.

This increase in oxygen is thought to have occurred in six steps, measured in billions of years: 2.4, 2.45, 1.8, 0.6, 0.3, and 0.04 billion years ago, with a possible seventh step 1.2 billion years ago. The first step appeared to have been a decrease in the amount of dissolved nickel in the seawater, which could have stifled the methane-producing bacteria and set the stage for oxidation of the Earth’s atmosphere because the methane would have reacted with any oxygen and created carbon dioxide and water. The initial change in the Earth’s atmosphere took place 2.4 billions years ago, in what scientists call the Great Oxidation Event.

The timing of these steps coincides with the amalgamation of Earth’s landmasses into supercontinents. Whereas hydrothermal activity can buffer the inventory of oceanic-dissolved iron against shorter-term fluctuations in continental deposition because hydrothermal input is relatively constant over millennial, the collisions of continents required to form supercontinents produced huge mountains, which eroded quickly and thereby released large amounts of nutrients, such as iron and phosphorus, into the oceans. These nutrient pulses led to explosions of algae and cyanobacteria, which in turn caused marked increases in photosynthesis and thus the production of oxygen. Enhanced sedimentation during these periods buried large amounts of organic carbon and pyrite that not only prevented their reaction with free oxygen but also led to sustained increases in atmospheric oxygen.

In fact, much of the oxygen in the atmosphere we depend on was generated through the photosynthesis of cyanobacteria during the Archaean and Proterozoic Eras, the latter of which occurred 2.5 billion to 543 million years ago. (Proterozoic comes from the Greek roots proteros, “earlier” and zōilos, “of animals, from zōion, “living being.”) Moreover, the beginning of the Middle Proterozoic (16 million years ago) saw substantial evidence of oxygen accumulating in the atmosphere.2 The Proterozoic Era as a whole began 2.5 billion years ago and ended 543 million years ago. There is circumstantial evidence, however, that oxygen levels dropped significantly around 1.9 billion years ago and remained low for several million years before rising again, although the reason(s) remain obscure.3 And time passed.


Estimates based on molecular analysis suggest that diatoms arose in the Triassic period, perhaps as early as 250 million years ago, whereas the earliest well-preserved diatom fossils are from the Early Jurassic, 190 million years ago. Diatoms are single-celled algae that can live in colonies various shapes, from circles, to zigzags, to ribbons, and so on. “Diatom” is from the Greek diatomos, meaning “cut in half,” which refers to their distinctive two-part cell walls of silica, which is converted in intricately designed cell walls of glass.

Whereas the elaborate species-specific patterns of nano-scale to micro-scale pores, ridges, and tubular structures are genetically controlled, such external factors as salinity influence the density and pore size of the precipitated silica. The cell wall is produced in an acidic, silica-deposition vesicle, which is encased in an organic matrix rich in proteins and sugars that prevent the silica from dissolving in seawater. Consumption of this organic-rich matrix by bacteria accelerates the recycling of silicon within surface waters.

The Triassic ocean differed dramatically from modern oceans because the concentrations of atmospheric carbon dioxide were almost eight times higher than they are today. Consequently, the average global temperature was significantly higher. In addition, Africa and Europe were beginning to separate, which precipitated extensive flooding of continental shelves. This probably increased the amount of continental weathering, which released large amounts of nutrients and thus boosted phytoplankton activity. (“Phytoplankton” is from the Greek phyton “plant” and planktos, which means “wanderer” or “drifter.”)

Moreover, the absence of polar ice caps, as well as the smaller pole-to-equator temperature gradient, not only reduced ocean circulation but also increased stratification of the water column. Together, these factors decreased the oxygenation of the oceans and contributed to ocean anoxic events. (Anoxic waters are those in which dissolved oxygen is depleted.)

The emergence of diatoms and two other groups of larger phytoplankton resulted in a major shift in global cycling of organic carbon, which initiated an era of declining concentrations of atmospheric carbon dioxide and increasing concentrations of atmospheric oxygen. Although most of the plant organisms that compose phytoplankton are too small to be visible individually to the naked eye, when in large enough numbers, they appear as a green discoloration in the water due to the chlorophyll in their cells. These organisms are the basis of the aquatic food chain in both fresh and salt water.

Diatoms assumed their dominant role in the carbon cycle during the Cretaceous period, about 100 million years ago, a time when levels of atmospheric carbon dioxide were still roughly five times higher than they are today and oxygen levels were increasing. Concurrently, stratification of the oceans was decreasing as the supply of nutrients to surface waters was concentrating. The proliferation of diatoms and other photosynthetic organisms during this period increased the oxygenation of surface waters, with a concomitant decrease in the availability of iron, which coincided with the divergence of a second major lineage of diatoms. (Photosynthesis is the process by which chlorophyll-containing cells in green plants convert incident light to chemical energy and synthesize organic compounds from inorganic compounds, especially carbohydrates from carbon dioxide and water, with the simultaneous release of oxygen.)

The mass extinction at the end of the Cretaceous, 65 million years ago (thought to have been cause either by massive asteroid impacts or by increased volcanic activity), led to loss of about 85 percent of all marine species. Diatoms, however, survived relatively unscathed and began to colonize offshore areas, including the open ocean. By 50 million years ago, atmospheric oxygen had stabilized around today’s levels, which further reduced the amount of available pelagic iron. (“Pelagic” means “in the open ocean”) Coincidentally, the amount of atmospheric carbon dioxide continued to decline to near today’s levels. Diatom diversity peaked at the Eocene/Oligocene boundary, some 30 million years ago.4


The photosynthetic process is affected by various environmental factors. In the sea, temperature, the quality and quantity of light, and the availability of nutrients are the main variables that determine the ability of organisms to transform solar energy into chemical energy and fixed carbon. Nutrients and trace elements, especially iron, are present in ever-lower concentrations throughout much of the open ocean.5

These qualifications are true in a general sense with sea ice. Sea ice contains large communities of such life forms as bacteria, algae, worms, and crustaceans, which live in the briny channels that form in the ice. The algae alter the sea ice in a number of ways, such as the production of chemicals that can lower the freezing temperature of the ice or they can darken the ice so that it absorbs more sunlight, which lowers the albedo effect—the ability of the ice to reflect solar energy back into space.

These algae not only obtain their energy from photosynthesis but also form a major portion of the arctic food chain. They overwinter in nooks and crannies within the ice, and expand rapidly throughout the lower layers of ice, as the temperature rise in the spring. When the ice melts completely, the algae are released into the ocean, where they form massive blooms of biological activity that support the food chain, beginning with the plankton.6 “Plankton” is a mass of plant and animal organisms, generally microscopic, that float or drift in great numbers in fresh or salt water, usually near the surface, where they are eaten by fish and other aquatic animals.

Primary productivity in 30 to 40 percent of the world’s oceans is today limited by the availability of iron, particularly in pelagic regions of the Southern Ocean, equatorial Pacific, and North Pacific. (Primary production is the production of organic compounds from atmospheric or aquatic carbon dioxide, predominantly through the process of photosynthesis in green plants. Iron, in turn, is a critical component of the catalytic cycle that affects the amount of photosynthesis taking place in various parts of the ocean.) Exceedingly low levels of iron and high concentrations of such other essential nutrients as nitrate, phosphate, and silicic acid characterize these high-nutrient, low-chlorophyll regions. Pelagic diatoms reduce their iron requirements under iron-limiting conditions. Some species seem not only to have permanently modified their photosynthetic apparatus to require less iron but also to have replaced iron-requiring electron-transport proteins with equivalent ones that need copper.7

To determine whether iron might indeed limit the growth of phytoplankton—and thus photosynthesis—in large regions of the ocean, an area of 40 square miles in the open equatorial Pacific Ocean was enriched with iron. This enrichment doubled the plant biomass, which cause a threefold increase in chlorophyll and a fourfold increase in plant production. Similar increases were found in a chlorophyll-rich plume down-stream of the Galapagos Islands, which was naturally enriched in iron. These findings indicate that iron can act as a limiting factor in the productivity and biomass of phytoplankton in the ocean, which in turn affects the production of the oxygen required by aerobic (oxygen-breathing) life.8


There is today a teeming mass of organisms living beneath the seafloor that, by some estimates, account for as much one-third of all our planet’s biomass—which is the sheer weight of all its living organisms. Many of these microorganisms exist on such foods as plankton that once thrived in the sunlight of the ocean surface only to drift down in the water column to the seafloor. Others, such as bacteria, survive in nutrient-poor, suffocating sediments in the middle of the South Pacific Gyre, where they inhabit the seafloor in a circulating vortex the size of North America. The gyre is so far removed from any landmass wherefrom nutrients could wash to sea and stimulate productivity that it acts like a gigantic oceanic desert. In some places, where the microbes eke out an existence, the seafloor mud takes a million years to reach a depth of three inches. Therefore, a bacterium living six inches down in the seafloor sediments would be inhabiting two million years worth of accumulation. Such regions of low productivity in the middle of the world’s oceans are far more common than the nutrient rich coastal zones.

Although oxygen exists, at most, only in the upper one inch of nutrient-rich areas, where detritus reaches the seafloor, oxygen penetrates throughout the seafloor in the South Pacific Gyre to a depth of at least 263 feet. Here, it is thought bacteria either consume oxygen very slowly or they rely on naturally occurring radioactivity as their source of energy.

At the Juan de Fuca Ridge off the coast of Washington State, water flows in one volcano, through the oceanic crust, and out another volcano 31 miles away. The volcanoes are situated in a north-south orientation, as are most of the fractures in the oceanic crust, thus acting like a microbial “super-highway,” allowing them to be distributed easily by the flowing water. In addition, microbes creep 920 feet down into the seafloor, where they are flushed by water circulating through the ocean’s crust—the largest belowground reservoir on Earth, wherein water obeys thermodynamic laws in a perpetual balancing act, cycling the equivalent of the ocean’s entire volume every 500,000 years or so.

The North Pond, in contrast, is an area of bottom sediments about three miles below the surface adjacent to an underwater mountain marking the middle of the Atlantic Ocean. This locale is characterized by violent geologic activity as the ocean crust is born. Here, water is pushed quickly through the rocks and bottom sediments into the open ocean. Whereas water at the Juan de Fuca Ridge is 140 to 158 degrees Fahrenheit, water at the North Pond is only about 50 degrees Fahrenheit, but flows much faster. Here, organisms living beneath the seafloor are unlike any life forms known.

As it turns out, seafloor-dwelling microbes are far more diverse than scientists thought even a decade ago, and there are more archaea compared to bacteria is some locales. Archaea are a group of single-celled microorganisms that lack a nucleus or any other membrane-bound organelles with their cells—as opposed to bacteria. An “organelle” is specialized part of a cell, such as a nucleus, which has its own particular function.9


Series on Biodiversity:

• Earth Before Oxygen

• The Long, Slow Path To Life As We Know It

• From Whence Comes Today’s Biodiversity?

• What—Exactly—Is Biodiversity?

• Biodiversity—Our Social-Environmental Insurance Policy

• Endangering Our Environmental Insurance Policy

Related Posts:

• Biodiversity–The Variety Of Life

1. Composition, Structure, And Function

2. Disturbance Regimes

3. Cumulative Effects, Lag Periods, And Thresholds

4. Biological Diversity

5. Genetic Diversity

6. Functional Diversity

7. Nature’s Services–Ecological Wealth Across Generations

• Principle 2: All relationships are inclusive and productive

• Principle 6: All relationships are self-reinforcing feedback loops

• Principle 7: All relationships have one or more tradeoffs

• Oceans in Crisis—Meeting The Ocean

• The Link Between Nature’s Commons And Our Cultural Commons


1. David C. Catling and Kevin J. Zahnle. The Planetary Air Leak. Scientific American 300 (2009):36-43.

2. The preceding discussion of the origin of oxygen on Earth is based on: (1) Ian H. Campbell and Charlotte M. Allen. Formation of Supercontinents Linked To Increases In Atmospheric Oxygen. Nature Geoscience, 1 (2008):554-558; (2) C. Scott, T. W. Lyons, A. Bekker, and others. Tracing The Stepwise Oxygenation Of The Proterozoic Ocean. Nature, 452 (2008):456-459; (3) Introduction to the Archaean. (accessed on April 12, 2011); (4) James F. Kasting. Earth Sciences: Ups and Downs of Ancient Oxygen. Nature, 443 (2006):643-645; (5) Bacteria: Fossil Record. (accessed on April 12, 2009); (6) A.R. Palmer and J. Geissman
. Introduction to the Proterozoic Era. 2002.
proterozoic.html (accessed on January 13, 2010); (7) James F. Kasting and Janet L. Siefert. Life And The Evolution Of Earth’s Atmosphere. Science, 296 (2002):1066-1068; (8) James F. Kasting. Warming early Earth and Mars. Science, 276 (1997):1213-1215; (9) Pennsylvania State University. Deep Sea Rocks Point To Early Oxygen On Earth. (accessed on April 12, 2011); (10) Ernesto Pecoits, Stefan V. Lalonde, Dominic Papineau, and others. Oceanic Nickel Depletion And A Methanogen Famine Before The Great Oxidation Event. Nature, 458 (2009):750-753; (11) Colin Goldblatt, Timothy M. Lenton, and Andrew J. Watson. Bistability Of Atmospheric Oxygen And The Great Oxidation. Nature, 2006 Oct 12;443(7112):683-686; (12) Paul G. Falkowski, Miriam E. Katz, Andrew H. Knoll, and others. The Evolution Of Modern Eukaryotic Phytoplankton. Science, 305 (2004):354–360; and (13) Alessandro Tagliabue, Laurent Bopp, Jean-Claude Dutay, and others. Hydrothermal Contribution To the Oceanic Dissolved Iron Inventory. Nature Geoscience, 3 (2010) 252–256.

3. Nick Lane. First Breath: Earth’s Billion-Year Struggle For Oxygen. NewScientist, 2746 (2010):36-39.

4. The preceding discussion of diatoms is based on: (1) E. Virginia Armbrust. The Life Of Diatoms In The World’s Oceans. Nature, 459 (2009):185-192; (2) Ulf Sorhannus. A Nuclear-Encoded Small-Subunit Ribosomal RNA Timescale For Diatom Evolution. Marine Micropaleontology, 65 (2007):1–12; (3) Patricia A. Sims, David G. Mann, and Linda K. Medlin. Evolution Of The Diatoms: Insights From Fossil, Biological And Molecular Data. Phycologia, 45 (2006):361–402 (2006); (4). Engel G. Vrieling, Qianyao Sun, Mingwen Tian, and others. Salinity-Dependent Diatom Biosilicification Implies An Important Role Of External Ionic Strength. Proceedings of the National Academy of Science USA, 104 (2007):10441–10446 (2007); (5) Kay D. Bidle, Maura Manganelli, and Farooq Azam. Regulation Of Oceanic Silicon And Carbon Preservation By Temperature Control On Bacteria. Science, 298 (2002)1980–1984; (6) Michael W. Guidry, Rolf. S. Arvidson, and Fred T. MacKenzie. Pp 377–403. In: Evolution of Primary Producers in the Sea. (Paul G. Falkowski and Andrew H. Knoll, eds.) Elsevier, Amsterdam, Netherlands. 2007; (7) M.Katz, J.Wright, K.Miller, and others. Biological Overprint Of The Geological Carbon Cycle. Marine Geology, 217 (2005):323-338; (8) Paul G. Falkowski and Matthew J. Oliver. Mix And Max: How Climate Selects Phytoplankton. Nature Reviews Microbiology, 5 (2007):813–819; (9) Paul G. Falkowski, Miriam E. Katz, Allen J. Milligan, and others. The Rise Of Oxygen Over The Past 205 Million Years And The Evolution Of Large Placental Mammals. Science, 309 (2005):2202–2204; (10) Daniel L. Rabosky and Ulf Sorhannus. Diversity Dynamics Of Marine Planktonic Diatoms Across The Cenozoic. Nature, 457 (2009):183–187; and (11) Cretaceous–Tertiary extinction event.–Tertiary_extinction_event (accessed on August 1, 2010).

5. Jonathan P. Zehr and Raphael M. Kudela. Photosynthesis in the Open Ocean. Science, 326 (2009):945–946.

6. Alexandra Witze. Melting At the Microscale. Science News, 177 (2010):22-25.

7. The foregoing discussion of primary productivity is based on: (1) J. Keith Moore, Scott C Doney, David M Glover and Inez Y Fung. Iron Cycling And Nutrient-Limitation Patterns In Surface Waters Of The World Ocean. Deep Sea Research Part II: Topical Studies in Oceanography, 49 (2002):463–507; (2) Adrian Marchetti, Maria T. Maldonado, Erin S. Lane, and Paul J. Harrison, Iron Requirements of the innate Diatom Pseudo-nitzschia: Comparison of Oceanic (HNLC) and Coastal Species. Limnology and Oceanograph, 51 (2006):2092–2101; (3) William G. Sunda Dorothy G. Swift & Susan A. Huntsman. Low Iron Requirement For Growth In Oceanic Phytoplankton. Nature, 351 (1991):55–57; (4) Robert F. Strzepek and Paul J. Harrison. Photosynthetic Architecture Differs In Coastal And Oceanic Diatoms. Nature, 431 (2004)689–692; (5) Grasham Peers and Neil M. Price. Copper-Containing Plastocyanin Used For Electron Transport By An Oceanic Diatom. Nature, 441 (2006):341–344; (6) E. Virginia Armbrust. The Life Of Diatoms In The World’s Oceans. Nature, 459 (2009):185-192; (7) Hein J. W. De Baar, Jeroen T. M. De Jong, Dorothée C. E. Bakker, and others. Importance of Iron For Plankton Blooms and Carbon Dioxide Drawdown In the Southern Ocean. Nature, 373 (1995):412–415; and (8) Richard J. Geider and Julie Roche. The Role of Iron in Phytoplankton Photosynthesis, and the Potential For Iron-Limitation of Primary Productivity in the Sea. Photosynthesis Research, 39 (1994):275–301.

8. J.H. Martin, K.H. Coale, K.S. Johnson, and others. Testing The Iron Hypothesis In Ecosystems Of The Equatorial Pacific Ocean. Nature, 371 (2002):123–129.

9. The preceding discussion of seafloor life is based on: (1) Alexandra Witze. Deep Life. Science News 181 (2012):18-21; (2) Beth N Orcutt, Wolfgang Bach, Keir Becker, and others. Colonization of subsurface microbial observatories deployed in young ocean crust. International Society for Microbial Ecology 5 (2011):692–703; and (3) Archaea. (accessed February 15, 2012).

Text © by Chris Maser 2012. All rights reserved.

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