10.3 Open Ocean Biogeochemistry

Open Ocean Biogeochemical Processes

The open ocean covers roughly 70% of Earth's surface and accounts for about half of all global primary production. Understanding its biogeochemistry is central to understanding global carbon, nitrogen, and phosphorus cycles, and to predicting how climate change will reshape these systems.

Biogeochemical processes in open oceans

Primary production in the open ocean is driven by phytoplankton, single-celled photosynthetic organisms that convert dissolved inorganic carbon ($CO_2$) into organic matter. Two key groups are diatoms (silica-shelled, dominant in nutrient-rich waters) and coccolithophores (calcium carbonate-shelled, important in oligotrophic gyres). Together, marine phytoplankton fix roughly as much carbon as all terrestrial plants combined.

A few definitions to keep straight:

• Gross primary production (GPP): total carbon fixed by photosynthesis
• Net primary production (NPP): carbon fixed minus what phytoplankton lose to their own respiration. This is the organic matter actually available to the rest of the food web.

Nutrient limitation determines where and how much production occurs. Liebig's Law of the Minimum says that growth is controlled by whichever essential resource is in shortest supply, not by the total amount of all resources.

• The three major limiting nutrients are nitrogen (N), phosphorus (P), and iron (Fe).
• The Redfield ratio ($C:N:P = 106:16:1$) describes the average elemental composition of marine organic matter. Deviations from this ratio can indicate which nutrient is limiting.
• In high-nutrient, low-chlorophyll (HNLC) regions like the Southern Ocean, the subarctic Pacific, and the equatorial Pacific, macronutrients (N and P) are abundant but iron is scarce. Iron fertilization experiments in these regions have shown dramatic phytoplankton blooms, confirming iron as the limiting factor.

The biological pump is the set of processes that transport organic carbon from the sunlit surface ocean to depth. Here's how it works:

1. Phytoplankton fix $CO_2$ into organic matter via photosynthesis in the euphotic zone.
2. Dead cells, fecal pellets, and aggregates ("marine snow") sink through the water column.
3. Most of this sinking material is remineralized (decomposed back to $CO_2$ and dissolved nutrients) by bacteria and zooplankton before it reaches the deep ocean.
4. The fraction that escapes the surface layer is called export production.
5. A small portion reaches the seafloor and is buried in sediments, sequestering carbon on geological timescales.

The efficiency of the biological pump matters enormously for climate. The more carbon that reaches the deep ocean, the more $CO_2$ is kept out of the atmosphere.

Ocean circulation and nutrient distribution

Physical circulation controls where nutrients end up, which in turn controls where biology thrives.

Upwelling brings cold, nutrient-rich deep water to the surface, fueling high productivity:

• Coastal upwelling occurs when winds blow parallel to a coastline and Ekman transport pushes surface water offshore, drawing deep water up to replace it. The Peruvian (Humboldt) coast and the Benguela Current off southwest Africa are classic examples.
• Equatorial upwelling happens because the trade winds and the Coriolis effect drive surface water away from the equator on both sides, pulling nutrient-rich water upward along the equatorial band.

These upwelling zones, though small in area, are disproportionately productive and support major fisheries.

Deep water formation drives the global thermohaline circulation (sometimes called the "ocean conveyor belt"). Dense water forms when surface water becomes cold and salty enough to sink:

• North Atlantic Deep Water (NADW) forms in the Nordic Seas and Labrador Sea.
• Antarctic Bottom Water (AABW) forms in the Weddell and Ross Seas and is the densest water mass in the ocean.

This sinking ventilates the ocean interior, supplying oxygen to depth and redistributing nutrients globally.

Vertical nutrient profiles in the open ocean follow a characteristic pattern: surface concentrations are low (phytoplankton strip nutrients out), and concentrations increase with depth (remineralization releases nutrients back into solution). The concept of nutrient spiraling describes how nutrients cycle through repeated uptake and remineralization as water masses move along circulation pathways.

Dissolved gases also follow circulation-driven patterns:

• Oxygen minimum zones (OMZs) develop at intermediate depths (roughly 200–1000 m) where microbial respiration consumes oxygen faster than circulation can replenish it. Major OMZs exist in the eastern tropical Pacific, the Arabian Sea, and the Bay of Bengal.
• The solubility pump for $CO_2$ works because cold water dissolves more $CO_2$ than warm water. When cold, dense surface water sinks at high latitudes, it carries dissolved $CO_2$ into the deep ocean.
• Methane hydrates are ice-like deposits of methane trapped in water molecules, found in cold, high-pressure settings along continental slopes. They represent a large carbon reservoir, and their stability is sensitive to temperature changes.

Climate Change Impacts and Ocean-Climate Interactions

Climate change impacts on ocean biogeochemistry

Ocean acidification results from the ocean absorbing excess atmospheric $CO_2$. The chemistry proceeds through a series of equilibria:

$CO_2 + H_2O \leftrightarrow H_2CO_3 \leftrightarrow HCO_3^- + H^+ \leftrightarrow CO_3^{2-} + 2H^+$

As more $CO_2$ dissolves, more $H^+$ ions are produced, lowering pH. Since the Industrial Revolution, ocean surface pH has dropped by about 0.1 units (from ~8.2 to ~8.1), which corresponds to a roughly 26% increase in hydrogen ion concentration. The critical consequence is a reduction in carbonate ion ($CO_3^{2-}$) concentration, which lowers the carbonate saturation state. Calcifying organisms like corals, pteropods, and foraminifera depend on adequate carbonate saturation to build and maintain their shells and skeletons.

Ocean deoxygenation is driven by two reinforcing mechanisms:

• Warmer water has a lower capacity to hold dissolved oxygen (simple gas solubility).
• Warming strengthens ocean stratification (the density difference between surface and deep layers), which reduces vertical mixing and cuts off the oxygen supply to deeper waters.

The result is expanding OMZs, which compress the habitable depth range for aerobic organisms and alter biogeochemical cycling. Deoxygenation particularly affects the nitrogen cycle by promoting denitrification and anammox (anaerobic ammonium oxidation), which convert bioavailable nitrogen to $N_2$ gas, further limiting productivity.

Changes in primary production are expected to vary by region:

• Polar regions may see increases as sea ice retreats and growing seasons lengthen.
• Tropical and subtropical oceans may see decreases as stronger stratification reduces nutrient supply to the surface.
• Phytoplankton community structure is shifting toward smaller species (e.g., picoplankton over diatoms), which are better adapted to nutrient-poor conditions but produce less export flux. This reduces the efficiency of the biological pump.

Ocean's role as global carbon sink

The ocean currently absorbs roughly 25% of anthropogenic $CO_2$ emissions each year, making it the largest active carbon sink after the atmosphere itself. Two mechanisms drive this uptake:

• The solubility pump: $CO_2$ dissolves preferentially in cold, high-latitude surface waters, which then sink and carry dissolved carbon to depth.
• The biological pump: phytoplankton fix $CO_2$ into organic matter that sinks and is stored at depth.

Carbon cycle feedbacks could either strengthen or weaken this sink:

• Slowing thermohaline circulation would reduce the rate at which surface water (and its dissolved $CO_2$) is transported to depth.
• Shifts in marine ecosystems (e.g., toward smaller phytoplankton) could reduce biological carbon export.
• Warming of ocean sediments could destabilize methane hydrates, releasing methane (a potent greenhouse gas) and creating a positive feedback loop that amplifies warming.

Ocean-atmosphere interactions govern the rate of $CO_2$ exchange:

• Air-sea gas exchange is driven by the difference in $CO_2$ partial pressure ($pCO_2$) between the atmosphere and the ocean surface. Where ocean $pCO_2$ is lower than atmospheric, the ocean absorbs $CO_2$, and vice versa.
• Climate modes like ENSO affect this exchange. During El Niño events, reduced equatorial upwelling means less $CO_2$-rich deep water reaches the surface, temporarily increasing net ocean uptake in that region.
• Rising ocean heat content reduces $CO_2$ solubility over time, weakening the solubility pump.

Future projections carry significant uncertainty, but several trends are concerning:

• The ocean carbon sink may saturate as surface waters warm and acidify, reducing their capacity to absorb additional $CO_2$.
• Potential tipping points include a shutdown or major weakening of the Atlantic thermohaline circulation, which would reorganize global nutrient and carbon distributions.
• Climate models still struggle to represent the full complexity of biogeochemical feedbacks, particularly interactions between biology, chemistry, and circulation at regional scales.