10.5 Ocean geochemistry

Composition of seawater

Seawater's chemical makeup is the foundation of ocean geochemistry. The dissolved ions, trace elements, and gases in seawater control everything from marine biological productivity to global climate regulation. Getting a handle on these components is essential before diving into the larger cycles and processes covered in this unit.

Major ions in seawater

Sodium ($Na^+$) and chloride ($Cl^-$) dominate seawater composition, accounting for roughly 85% of all dissolved ions. The remaining major ions include sulfate ($SO_4^{2-}$), magnesium ($Mg^{2+}$), calcium ($Ca^{2+}$), and potassium ($K^+$).

A key concept here is the principle of constant proportions (Marcet's principle): the ratios of major ions to each other stay remarkably constant throughout the open ocean, even though total salinity can vary. This constancy results from the long residence times of these ions (millions of years for most), meaning they accumulate and mix far faster than they're added or removed.

• Total dissolved solids average about 35 g/kg, expressed as salinity in practical salinity units (PSU)
• Salinity varies locally due to evaporation, precipitation, river input, and ice formation, but the relative proportions of major ions stay the same

Trace elements in oceans

Trace elements are present at concentrations below 1 ppm, yet they're critical for biological processes. Iron ($Fe$), zinc ($Zn$), copper ($Cu$), and manganese ($Mn$) are among the most geochemically important.

Unlike major ions, trace element distributions are not uniform. Their concentrations vary dramatically with depth and location because of:

• Biological uptake in surface waters (depleting concentrations there)
• Scavenging onto sinking particles, which removes metals from the water column
• External inputs from rivers, atmospheric dust deposition, and hydrothermal vents

Some trace elements also serve as paleoproxies. Cadmium ($Cd$), for example, tracks phosphate distributions closely, so $Cd/Ca$ ratios in fossil foraminifera can reconstruct past ocean nutrient levels and productivity.

Dissolved gases in seawater

• Oxygen ($O_2$) is essential for aerobic marine life. Concentrations are highest at the surface (gas exchange with the atmosphere and photosynthesis) and drop to a minimum at intermediate depths where respiration consumes it, then increase slightly in deep water due to cold, oxygen-rich water masses sinking at high latitudes.
• Carbon dioxide ($CO_2$) plays a central role in both the ocean carbon cycle and ocean acidification. Its solubility increases in colder water, which is why high-latitude oceans are major $CO_2$ sinks.
• Nitrogen ($N_2$) is relatively inert in seawater but matters for nitrogen-fixing cyanobacteria that convert it to biologically available forms.
• Noble gases ($He$, $Ne$, $Ar$) are chemically unreactive, which makes them excellent tracers for physical processes like ocean circulation, air-sea gas exchange rates, and mixing.

Ocean circulation patterns

Ocean circulation distributes heat, nutrients, and dissolved gases across the globe. For geochemists, circulation patterns explain why chemical distributions look the way they do, and how quickly the ocean can respond to perturbations like increased $CO_2$ input.

Thermohaline circulation

This is the global-scale, density-driven circulation often called the global conveyor belt. Temperature and salinity together determine water density, and density differences drive the deep circulation.

The process works like this:

1. In the North Atlantic (and around Antarctica), surface water becomes cold and salty enough to become very dense.
2. This dense water sinks to form North Atlantic Deep Water (NADW) and Antarctic Bottom Water (AABW).
3. These deep water masses spread through the ocean basins along the seafloor.
4. Gradual upwelling and mixing eventually return this water to the surface, completing the loop.

The full circulation cycle takes roughly 1,000 years. This timescale matters because it controls how long carbon and other dissolved substances can be sequestered in the deep ocean before returning to contact with the atmosphere.

Surface currents vs deep currents

• Surface currents are driven primarily by wind. Wind stress on the ocean surface, combined with the Coriolis effect, produces Ekman transport, which moves water at an angle to the wind direction. Major surface currents include the Gulf Stream, Kuroshio Current, and the Antarctic Circumpolar Current.
• Deep currents are driven by density differences (thermohaline forcing) and steered by seafloor topography (bathymetry). Abyssal circulation is critical for transporting oxygen to the deep ocean and redistributing nutrients globally.

The distinction matters geochemically: surface currents move water (and its dissolved load) quickly over short timescales, while deep currents operate over centuries to millennia.

Upwelling and downwelling processes

Upwelling brings cold, nutrient-rich deep water to the surface, fueling high biological productivity. It occurs in two main settings:

• Coastal upwelling along the eastern boundaries of ocean basins (e.g., the Peru Current off South America, the Benguela Current off southwestern Africa), where winds push surface water offshore via Ekman transport
• Equatorial upwelling, driven by trade wind divergence at the equator

Downwelling pushes surface water downward, typically in subtropical gyres where wind-driven convergence piles up water. These regions become nutrient-poor "oceanic deserts" with low productivity because nutrients are pushed away from the sunlit zone.

Biogeochemical cycles in oceans

The oceans are central players in global biogeochemical cycles, regulating atmospheric composition and climate over timescales from years to millions of years. Marine biogeochemical cycles involve tightly coupled biological, chemical, and physical processes.

Carbon cycle in oceans

The ocean holds approximately 50 times more carbon than the atmosphere, making it the largest active carbon reservoir on Earth. Dissolved inorganic carbon (DIC) exists in three forms that are in chemical equilibrium:

$CO_2 \rightleftharpoons HCO_3^- \rightleftharpoons CO_3^{2-}$

At seawater's typical pH (~8.1), bicarbonate ($HCO_3^-$) dominates, making up about 90% of DIC.

Two major "pumps" transfer carbon from the surface to the deep ocean:

• Biological pump: Phytoplankton fix $CO_2$ through photosynthesis, and when they die or are consumed, sinking organic particles carry that carbon to depth. Most is remineralized (converted back to $CO_2$) in the water column, but some reaches the sediments.
• Carbonate pump: Organisms like foraminifera and coccolithophores build $CaCO_3$ shells. When these sink and dissolve at depth (below the lysocline), they release carbon back to the water. Shells that reach the seafloor above the carbonate compensation depth get buried, removing carbon on geological timescales.

Air-sea gas exchange of $CO_2$ is regulated by the difference in partial pressure ($pCO_2$) between the atmosphere and surface ocean, along with wind speed (which controls the gas transfer velocity).

Nitrogen cycle in marine environments

Nitrogen is often the limiting nutrient for primary production in much of the ocean. It cycles through several oxidation states, and the transformations between them are almost entirely mediated by microorganisms.

• Nitrogen fixation: Cyanobacteria (e.g., Trichodesmium) convert dissolved $N_2$ gas into bioavailable ammonium ($NH_4^+$), introducing "new" nitrogen to the ocean
• Nitrification: Aerobic bacteria oxidize $NH_4^+$ to nitrite ($NO_2^-$) and then to nitrate ($NO_3^-$), the most abundant form of bioavailable nitrogen in the ocean
• Denitrification: In low-oxygen environments, bacteria reduce $NO_3^-$ back to $N_2$ gas, removing bioavailable nitrogen from the system
• Anammox (anaerobic ammonium oxidation): Bacteria in oxygen minimum zones convert $NH_4^+$ and $NO_2^-$ directly to $N_2$, representing another major pathway of nitrogen loss

The balance between nitrogen fixation (input) and denitrification/anammox (removal) controls the ocean's total inventory of bioavailable nitrogen over time.

Phosphorus cycle in seawater

Phosphorus is critical for life because it's a structural component of DNA, RNA, and ATP. In seawater, it exists primarily as dissolved inorganic phosphate ($PO_4^{3-}$).

Unlike carbon and nitrogen, the phosphorus cycle has no significant atmospheric component. This makes it fundamentally different:

• Inputs come almost entirely from continental weathering delivered by rivers
• Removal occurs through burial in marine sediments, including formation of phosphorite deposits

Because of its slow input and removal, phosphorus is thought to be the ultimate limiting nutrient on geological timescales (millions of years), even though nitrogen limits productivity on shorter timescales in most of the modern ocean.

Ocean acidification

Ocean acidification is one of the most consequential changes happening to ocean chemistry right now. It directly threatens marine organisms that build carbonate shells and has cascading effects on marine ecosystems and global geochemical cycles.

Causes of ocean acidification

The primary driver is rising atmospheric $CO_2$ from fossil fuel combustion and land-use change. When $CO_2$ dissolves in seawater, it reacts with water to form carbonic acid:

$CO_2 + H_2O \rightarrow H_2CO_3$

Carbonic acid then dissociates:

$H_2CO_3 \rightarrow H^+ + HCO_3^-$

Those extra hydrogen ions ($H^+$) lower the pH. But there's a second, equally important effect: the added $H^+$ reacts with carbonate ions:

$H^+ + CO_3^{2-} \rightarrow HCO_3^-$

This reduces the concentration of carbonate ions, which is what organisms need to build $CaCO_3$ shells and skeletons.

• The ocean has absorbed roughly 30% of anthropogenic $CO_2$ emissions since the industrial revolution
• Surface ocean pH has already dropped by about 0.1 units (from ~8.2 to ~8.1), which represents a ~26% increase in $H^+$ concentration (remember, pH is a log scale)
• The current rate of acidification is approximately 100 times faster than anything in the last 300 million years of the geological record

Effects on marine ecosystems

• Reduced calcification rates in corals, mollusks, pteropods, and coccolithophores as carbonate ion concentrations decline
• Altered behavior and physiology in fish and invertebrates, including impaired sensory function in some species
• Disruption of food webs, particularly at the base where calcifying plankton are affected
• Increased dissolution of existing $CaCO_3$ sediments on the seafloor
• Changes in trace metal speciation and nutrient availability, with knock-on effects for biological communities

Future projections and impacts

• Under high-emission scenarios, models project a further pH decrease of 0.3 to 0.4 units by 2100
• This would push surface ocean conditions below the aragonite saturation horizon in large areas, particularly in polar regions where cold water already holds more $CO_2$
• Major shifts in marine biodiversity and ecosystem services are expected
• Feedbacks on the global carbon cycle could either amplify or dampen climate change
• Socioeconomic consequences include impacts on fisheries, aquaculture, and coastal communities that depend on healthy marine ecosystems

Hydrothermal vents

Hydrothermal vents are windows into the chemical exchange between the ocean and the solid Earth. They're found along mid-ocean ridges and back-arc basins, where tectonic activity drives fluid circulation through the oceanic crust. These systems significantly influence ocean chemistry and support unique ecosystems independent of sunlight.

Formation and characteristics

The process of hydrothermal circulation follows a clear sequence:

1. Cold seawater percolates downward through fractures in the oceanic crust.
2. Near the magma chamber (at depths of a few kilometers), the water heats to extreme temperatures (up to 400°C or more).
3. Hot water reacts with the surrounding basaltic rock, leaching metals and other elements while losing magnesium and sulfate.
4. The superheated, chemically altered fluid rises buoyantly back to the seafloor and vents into the cold ocean.
5. When the hot fluid meets near-freezing seawater (~2°C), dissolved minerals precipitate rapidly, forming chimney structures.

Two main vent types are distinguished by temperature:

• Black smokers (>300°C): The dark "smoke" is actually fine-grained metal sulfide particles ($FeS$, $CuS$, $ZnS$) precipitating on contact with cold water
• White smokers (<300°C): Produce lighter-colored precipitates, typically barium sulfate, silica, and calcium sulfate

Chemical composition of vent fluids

Vent fluids are dramatically different from normal seawater:

• Enriched in dissolved metals: $Fe$, $Mn$, $Cu$, $Zn$
• High concentrations of reduced compounds: hydrogen sulfide ($H_2S$), methane ($CH_4$), and hydrogen ($H_2$)
• Depleted in $Mg^{2+}$ and $SO_4^{2-}$ relative to seawater (magnesium is quantitatively removed during water-rock reactions, which is actually used as a diagnostic indicator)
• pH ranges from highly acidic (2-3 in black smokers) to alkaline (9-11 in some ultramafic-hosted systems like Lost City)
• Fluid composition reflects the subsurface rock type (basalt vs. ultramafic peridotite) and the temperature/pressure of reactions

Microbial communities in vents

Vent ecosystems run on chemosynthesis rather than photosynthesis. Microorganisms at the base of these food webs derive energy from chemical reactions rather than sunlight.

• Sulfur-oxidizing bacteria dominate many vent ecosystems, oxidizing $H_2S$ for energy. These often live as symbionts inside tubeworms and giant clams.
• Methanogens produce $CH_4$ from $CO_2$ and $H_2$, while methanotrophs consume $CH_4$ for energy
• Thermophilic and hyperthermophilic archaea thrive at temperatures up to 121°C, pushing the known limits of life
• Microbial activity directly influences mineral precipitation and element cycling at vents, creating feedback loops between biology and geochemistry

Marine sediments

Marine sediments are the ocean's long-term memory. They archive past ocean conditions, climate states, and biological productivity over millions of years. They also play active roles in element cycling through diagenetic reactions at and below the seafloor.

Types of marine sediments

• Terrigenous: Derived from continental weathering and erosion, transported by rivers, wind, and ice. Dominated by clay minerals, quartz, and feldspar.
• Biogenic: Composed of skeletal remains of marine organisms. Calcareous sediments come from foraminifera and coccolithophores ($CaCO_3$); siliceous sediments come from diatoms and radiolarians (opal/$SiO_2$).
• Authigenic: Formed in situ through chemical precipitation from seawater. Manganese nodules on the abyssal seafloor are a classic example.
• Volcanogenic: From submarine volcanic eruptions and ash falls.
• Cosmogenic: Extraterrestrial material such as micrometeorites. Volumetrically minor but useful as tracers.

Sediment composition and sources

The distribution of sediment types across the ocean floor follows predictable patterns:

• Continental margins are dominated by terrigenous input because of proximity to land sources
• Deep ocean basins are dominated by biogenic sediments (calcareous ooze above the CCD, siliceous ooze in high-productivity regions, and red clay in the most remote, low-productivity areas)

Lithogenic components include quartz, feldspars, and clay minerals. Biogenic components consist of $CaCO_3$, opal, and organic matter. Authigenic minerals like pyrite ($FeS_2$) and glauconite form through diagenetic reactions within the sediment.

Sediment distribution is controlled by distance from shore, water depth, ocean productivity, and the carbonate compensation depth (CCD), below which $CaCO_3$ dissolves faster than it accumulates.

Diagenesis in marine sediments

Diagenesis refers to the physical and chemical changes that occur in sediments after deposition. Early diagenesis is driven largely by microbial degradation of organic matter, which proceeds through a well-defined sequence of electron acceptors as each is exhausted:

1. Oxic zone: Aerobic respiration using $O_2$
2. Suboxic zone: Reduction of $NO_3^-$, then $Mn^{4+}$ oxides, then $Fe^{3+}$ oxides
3. Anoxic zone: Sulfate reduction ($SO_4^{2-}$), then methanogenesis ($CO_2$ reduction to $CH_4$)

This redox zonation develops with depth in the sediment and controls which authigenic minerals form. For example, pyrite ($FeS_2$) forms in the sulfate reduction zone where $H_2S$ reacts with reactive iron.

Other important diagenetic processes include:

• Carbonate dissolution below the lysocline and CCD
• Silica diagenesis: dissolution and reprecipitation of biogenic opal, eventually forming chert over geological time
• Phosphorite formation in areas of high organic matter flux and low oxygen

Isotope geochemistry in oceans

Isotope ratios are among the most powerful tools in ocean geochemistry. They serve as tracers for processes that are otherwise invisible, and as proxies for conditions that existed millions of years ago. Both stable and radioactive isotopes provide complementary information.

Stable isotopes in oceanography

Stable isotope ratios are reported in delta notation ($\delta$), expressed in per mil (‰) relative to a standard. Each isotope system tracks different processes:

• $\delta^{18}O$: Measured in foraminifera shells, this proxy reflects both water temperature and global ice volume. During ice ages, preferential storage of light $^{16}O$ in ice sheets leaves the ocean enriched in $^{18}O$.
• $\delta^{13}C$: Tracks carbon sources and biological productivity. Photosynthesis preferentially takes up $^{12}C$, so surface waters become enriched in $^{13}C$ during high productivity.
• $\delta^{15}N$: Indicates nutrient utilization and nitrogen cycling. Higher values suggest more complete nitrate consumption.
• $\delta^{30}Si$: Reflects silicic acid utilization by diatoms.
• $\delta^{11}B$: Serves as a proxy for paleo-pH because the boron isotope fractionation between boric acid and borate ion is pH-dependent.

Radioisotopes in marine systems

Radioactive isotopes decay at known rates, making them natural clocks and tracers:

• $^{14}C$ (half-life ~5,730 years): Used for dating marine carbonates and organic matter up to ~50,000 years. Also used to determine the "ventilation age" of deep water masses.
• $^{230}Th$ and $^{231}Pa$: Produced by uranium decay in seawater, these particle-reactive isotopes are used in paleoproductivity and circulation studies. The $^{231}Pa/^{230}Th$ ratio is a proxy for past changes in Atlantic overturning circulation.
• Radium isotopes ($^{226}Ra$, $^{228}Ra$): Trace submarine groundwater discharge and coastal mixing.
• Tritium ($^{3}H$) and $^{3}He$: Anthropogenic tritium from nuclear testing serves as a transient tracer for ocean circulation and ventilation rates on decadal timescales.
• $^{10}Be$: Records changes in cosmic ray flux and geomagnetic field strength.

Applications in paleoceanography

Combining multiple isotope systems allows reconstruction of past ocean conditions:

• Past ocean temperatures from $\delta^{18}O$ in foraminifera (with corrections for ice volume)
• Past ocean circulation patterns from $Nd$ isotopes ($\epsilon_{Nd}$) in ferromanganese crusts, which fingerprint different water masses
• Past atmospheric $CO_2$ from $\delta^{11}B$ in planktonic foraminifera or coral skeletons
• Ocean overturning rates from radiocarbon age differences between surface and deep waters
• Past productivity patterns from barite ($BaSO_4$) accumulation rates in sediments

Nutrient dynamics in oceans

Nutrient availability controls where and how much life the ocean can support. The distribution, cycling, and limitation patterns of nutrients are tightly linked to ocean circulation, biological activity, and human impacts.

Nutrient distribution patterns

The classic nutrient profile in the open ocean shows surface depletion and deep enrichment. Phytoplankton consume nutrients in the sunlit surface layer (photic zone), and as organisms die and sink, their decomposition releases nutrients back at depth. This creates a characteristic vertical profile with low concentrations at the surface and a maximum at intermediate depths (~1,000 m).

• Horizontal gradients also exist: coastal waters receive nutrient inputs from rivers and upwelling, while open-ocean gyres tend to be nutrient-poor
• High-latitude regions generally have higher surface nutrient concentrations because deep mixing and reduced light limit complete biological drawdown
• The Redfield ratio ($C:N:P = 106:16:1$) describes the average elemental composition of marine organic matter and provides a benchmark for understanding nutrient stoichiometry. Deviations from Redfield ratios indicate which nutrient is limiting or which processes (like nitrogen fixation or denitrification) are at work.

Limiting nutrients in oceans

Which nutrient limits productivity depends on the region and timescale:

• Nitrogen limits productivity across much of the subtropical and tropical ocean
• Phosphorus may be the ultimate limiting nutrient on geological timescales because its input depends solely on weathering (no atmospheric reservoir)
• Iron limits productivity in high-nutrient, low-chlorophyll (HNLC) regions like the Southern Ocean, equatorial Pacific, and subarctic Pacific. These areas have plenty of nitrogen and phosphorus but lack the iron needed for phytoplankton enzymes. Iron fertilization experiments have confirmed this.
• Silicon can limit diatom growth specifically, since diatoms require dissolved silicic acid ($Si(OH)_4$) to build their opal frustules
• Co-limitation by multiple nutrients simultaneously occurs in some ecosystems

Eutrophication in coastal waters

Eutrophication happens when excess nutrients, primarily nitrogen and phosphorus from agricultural runoff, sewage, and industrial discharge, enter coastal waters. The sequence of impacts follows a predictable pattern:

1. Nutrient loading stimulates excessive phytoplankton growth (algal blooms).
2. Some blooms produce toxins (harmful algal blooms, or HABs).
3. When the bloom dies, microbial decomposition of the massive organic matter load consumes dissolved oxygen.
4. Oxygen depletion (hypoxia, defined as $O_2$ < 2 mg/L) or complete oxygen loss (anoxia) develops in bottom waters.
5. Hypoxic conditions kill or displace fish and benthic organisms, creating "dead zones."

The Gulf of Mexico dead zone, fed by Mississippi River nutrient loads, is one of the largest and best-studied examples. Management strategies focus on reducing nutrient inputs at the source and improving wastewater treatment.

Trace metal cycling

Despite their extremely low concentrations (nanomolar to picomolar), trace metals are essential for marine life and exert significant control over ocean productivity. Their distributions reflect a complex interplay of sources, biological uptake, and removal processes.

Sources of trace metals

• Atmospheric deposition: Mineral dust (especially Saharan dust, a major source of iron to the Atlantic) and anthropogenic aerosols
• Riverine input: Dissolved and particulate metals from continental weathering, though much is removed in estuaries through flocculation
• Hydrothermal vents: Release dissolved $Fe$, $Mn$, $Cu$, and $Zn$ to the deep ocean. Recent research shows hydrothermal iron can be transported thousands of kilometers from vent sites when stabilized by organic ligands.
• Sediment resuspension and benthic flux: Particularly important in coastal and continental margin settings
• Anthropogenic sources: Industrial effluents, mining runoff, and atmospheric pollution

Scavenging processes in oceans

Scavenging is the primary removal mechanism for many trace metals. It involves:

• Adsorption onto sinking particles (both organic and inorganic), which carries metals to the seafloor
• Biological uptake and incorporation into cells and organic matter
• Co-precipitation with iron and manganese oxyhydroxides, which are efficient scavengers
• Complexation with organic ligands: This actually increases metal solubility and can protect metals from scavenging. For example, >99% of dissolved iron in the ocean is bound to organic ligands.

Residence times vary enormously among trace metals. Aluminum, which is strongly scavenged, has a residence time of only ~100-200 years, while molybdenum, which is conservative, has a residence time of ~800,000 years.

Biological importance of trace metals

Trace metals function as essential cofactors in key enzymes:

• Iron: Required for photosynthesis (photosystem I), nitrogen fixation (nitrogenase), and nitrate reduction. Its scarcity in HNLC regions directly limits ocean productivity.
• Zinc: Cofactor in carbonic anhydrase ($CO_2$ fixation) and alkaline phosphatase (phosphorus acquisition)
• Copper: Used in electron transport chains and oxidative enzymes like plastocyanin
• Cobalt: Necessary for vitamin $B_{12}$ synthesis, which many phytoplankton require but cannot produce themselves
• Manganese: Part of the oxygen-evolving complex in photosystem II, where water is split to produce $O_2$

Organic matter in oceans

Marine organic matter is the currency of the biological carbon pump. Its production, transformation, and fate determine how much carbon the ocean can sequester and how energy flows through marine food webs.

Sources of marine organic matter

The dominant source is phytoplankton primary production in the sunlit surface ocean, which fixes roughly 50 Gt of carbon per year. Additional sources include:

• Terrestrial inputs from rivers and coastal runoff (dissolved and particulate organic carbon)
• Atmospheric deposition of organic aerosols
• Chemosynthetic production at hydrothermal vents and cold seeps
• Release of dissolved organic matter (DOM) through viral lysis of cells, zooplankton sloppy feeding, and excretion

Degradation of organic compounds

Most marine organic matter is recycled rather than buried. Degradation proceeds through several pathways:

• Microbial remineralization: Heterotrophic bacteria break down organic matter in the water column and sediments, converting it back to $CO_2$, $NO_3^-$, $PO_4^{3-}$, etc.
• Photochemical degradation: UV light breaks down chromophoric dissolved organic matter (CDOM) in surface waters
• Enzymatic hydrolysis: Extracellular enzymes produced by bacteria break down large particulate organic molecules into smaller dissolved compounds that can be taken up

Organic compounds degrade at very different rates. Labile compounds like simple sugars and amino acids are consumed within hours to days. Semi-labile material persists for months to years. Refractory dissolved organic matter (RDOM) resists degradation and can persist in the deep ocean for 4,000-6,000 years, representing a significant long-term carbon reservoir.

Role in carbon sequestration

The ocean sequesters carbon through organic matter via several mechanisms:

• Biological pump: Sinking particulate organic carbon (POC) transfers carbon from the surface to the deep ocean. Only about 1-5% of surface production reaches the deep seafloor.
• Sediment burial: Organic matter that reaches the seafloor and gets buried removes carbon from the active cycle on geological timescales. Burial efficiency is highest in low-oxygen continental margin sediments.
• Dissolved organic carbon (DOC): The ocean holds ~660 Gt C as DOC, comparable to atmospheric $CO_2$. This is a massive but slow-cycling reservoir.
• Microbial carbon pump: Microbial processing transforms labile organic matter into refractory DOC, effectively sequestering carbon in dissolved form for millennia.
• Organo-mineral interactions: Organic matter that binds to mineral surfaces (especially clays and iron oxides) is protected from microbial attack, enhancing preservation in sediments.