8.1 Carbon Reservoirs and Fluxes

Earth's Carbon Reservoirs

Major Carbon Reservoirs and Their Sizes

The global carbon cycle moves carbon between five main reservoirs. Each one stores carbon in different chemical forms and on very different timescales, which matters for understanding how fast the climate can change.

• Atmosphere stores approximately 870 gigatons of carbon (GtC), primarily as carbon dioxide ($CO_2$) and methane ($CH_4$). This is the smallest of the major reservoirs, which is exactly why human emissions can shift its composition so quickly.
• Terrestrial biosphere contains about 2,000–3,000 GtC. This includes living plant biomass (~450–650 GtC) and soil organic matter (~1,500–2,400 GtC). Soils actually hold far more carbon than all living vegetation combined.
• Hydrosphere holds around 38,000 GtC, the vast majority in the oceans as dissolved inorganic carbon (DIC): $CO_2$, bicarbonate ($HCO_3^-$), and carbonate ($CO_3^{2-}$) ions.
• Lithosphere is the largest reservoir by far, storing over 75,000,000 GtC in sedimentary rocks (especially limestone and organic-rich shales) and the Earth's crust. Carbon here cycles on timescales of millions of years.
• Fossil fuel deposits contain approximately 5,000–10,000 GtC as coal, oil, and natural gas. These formed from ancient organic matter buried and transformed over geological time.

Importance of Reservoir Sizes

The relative size of a reservoir determines how sensitive it is to disturbance. The atmosphere is small enough that adding even a few hundred gigatons of carbon noticeably raises $CO_2$ concentrations. The lithosphere, by contrast, is so massive that changes to it are imperceptible on human timescales.

• Larger reservoirs exert a stronger influence on long-term global carbon balance.
• Smaller reservoirs (atmosphere, biosphere) can experience rapid compositional changes from human activities.
• Knowing reservoir sizes is essential for building accurate climate models and designing carbon mitigation strategies.

Carbon Transfers Between Reservoirs

Biological Processes

• Photosynthesis transfers carbon from the atmosphere to the biosphere. Plants, algae, and cyanobacteria use light energy to convert $CO_2$ and water into glucose ($C_6H_{12}O_6$) and oxygen. Global terrestrial photosynthesis fixes roughly 120 GtC per year.
• Respiration moves carbon back from the biosphere to the atmosphere. Organisms break down organic compounds to produce energy, releasing $CO_2$ as a byproduct. This includes both autotrophic respiration (by plants themselves) and heterotrophic respiration (by animals and microbes).
• Decomposition returns carbon from dead organisms to the atmosphere and soil. Bacteria and fungi break down organic matter through microbial respiration, releasing $CO_2$. In waterlogged, oxygen-poor environments, decomposition instead produces $CH_4$ (methanogenesis).

Geochemical Processes

• Oceanic $CO_2$ uptake occurs when $CO_2$ dissolves into surface waters. This is driven by the difference in $CO_2$ partial pressure between the atmosphere and the ocean surface. Once dissolved, $CO_2$ reacts with water to form carbonic acid and then bicarbonate ions.
• Carbonate formation transfers carbon from the hydrosphere to the lithosphere. Marine organisms like corals, foraminifera, and mollusks build calcium carbonate ($CaCO_3$) shells and skeletons. When these organisms die, their remains accumulate on the seafloor and eventually lithify into carbonate rocks such as limestone.
• Weathering releases carbon from the lithosphere. Chemical weathering of carbonate rocks (e.g., limestone) dissolves $CaCO_3$ and releases bicarbonate ions into rivers that flow to the ocean. Silicate weathering is particularly important because it consumes atmospheric $CO_2$ in the process:

$CaSiO_3 + 2CO_2 + H_2O \rightarrow Ca^{2+} + 2HCO_3^- + SiO_2$

• Volcanic eruptions transfer carbon from Earth's interior to the atmosphere by releasing $CO_2$ and other volcanic gases. This outgassing is a slow but steady carbon source that helps maintain the long-term carbon cycle over geological time.

Anthropogenic Processes

• Fossil fuel combustion rapidly transfers ancient carbon to the atmosphere. Burning coal, oil, and natural gas currently adds roughly 9–10 GtC per year, releasing $CO_2$ that had been locked away for millions of years.
• Land-use changes alter carbon storage in the biosphere. Deforestation releases carbon stored in vegetation and soils (contributing ~1–1.5 GtC per year to emissions), while reforestation and certain agricultural practices can sequester carbon back into biomass and soils.

Global Carbon Cycle Balance

Natural Carbon Cycle Equilibrium

Over geological timescales (millions of years), the natural carbon cycle maintains a dynamic equilibrium where carbon uptake and release roughly offset each other. Oceans and terrestrial ecosystems act as significant carbon sinks, while volcanic activity and respiration serve as natural carbon sources.

Different reservoirs respond to disturbances on very different timescales:

• The atmosphere adjusts within days to weeks as $CO_2$ mixes globally.
• Oceans may take centuries to millennia to fully equilibrate, because deep ocean mixing is slow.
• The lithosphere changes over millions of years through weathering and sedimentation.

This mismatch in response times is why a rapid pulse of carbon (like fossil fuel burning) can overwhelm the system's ability to rebalance.

Anthropogenic Disruptions and Feedback Mechanisms

Human activities have pushed the carbon cycle out of its natural equilibrium. Atmospheric $CO_2$ has risen from ~280 ppm (pre-industrial) to over 420 ppm, primarily from fossil fuel combustion and land-use changes.

Positive feedback mechanisms amplify the imbalance:

• Permafrost thawing releases stored $CH_4$ and $CO_2$, which accelerates warming, which thaws more permafrost.
• Forest dieback from drought or fire reduces carbon uptake while simultaneously releasing stored carbon.

Negative feedback mechanisms partially counteract the imbalance:

• $CO_2$ fertilization effect: higher atmospheric $CO_2$ can stimulate plant growth, increasing carbon uptake (though this effect has limits related to nutrient and water availability).
• Enhanced weathering: higher temperatures and $CO_2$ levels speed up chemical weathering of silicate rocks, which consumes $CO_2$.

Carbon Flux Quantification and Climate Predictions

Accurately quantifying carbon fluxes is essential for predicting future climate. Researchers use several approaches:

• Direct measurements of atmospheric $CO_2$ concentrations (e.g., the Mauna Loa record)
• Carbon isotope ratios ($^{13}C/^{12}C$) to distinguish fossil fuel carbon from natural sources
• Satellite observations of vegetation cover and ocean productivity

Significant uncertainties remain, especially around soil carbon dynamics and ocean-atmosphere gas exchange rates. Carbon cycle models integrate known fluxes, feedback mechanisms, and projected human activities to generate future climate scenarios and inform policy decisions.

Ocean's Role in the Carbon Cycle

Ocean Carbon Uptake and Storage

Oceans absorb approximately 25–30% of anthropogenic $CO_2$ emissions each year, making them the single largest active carbon sink. This uptake is governed by three interacting mechanisms:

• Solubility pump: $CO_2$ dissolves more readily in cold water, so high-latitude surface waters absorb more $CO_2$. As these cold, carbon-rich waters sink during deep water formation, they carry dissolved carbon to the deep ocean.
• Biological pump: Phytoplankton fix $CO_2$ through photosynthesis near the surface. When they die or are consumed, organic carbon sinks as particles ("marine snow") to the deep ocean, effectively exporting carbon from the surface.
• Thermohaline circulation (the global ocean conveyor belt) transports carbon-rich deep waters around the globe. Upwelling in certain regions brings nutrient-rich water back to the surface, fueling productivity.

Ocean Chemistry and Buffering Capacity

The ocean's carbonate system regulates seawater pH and its capacity to absorb additional $CO_2$. When $CO_2$ dissolves in seawater, it undergoes a series of equilibrium reactions:

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

The carbonate-bicarbonate system acts as a chemical buffer, neutralizing added acid or base to keep pH relatively stable. However, this buffering capacity has limits.

Ocean acidification occurs as excess atmospheric $CO_2$ dissolves in seawater. The added $H^+$ ions lower pH (ocean surface pH has already dropped from ~8.2 to ~8.1 since pre-industrial times, a ~26% increase in hydrogen ion concentration). Simultaneously, carbonate ion ($CO_3^{2-}$) concentrations decrease, making it harder for calcifying organisms like corals, pteropods, and shellfish to build and maintain their $CaCO_3$ structures.

Climate Change Impacts on Ocean Carbon Cycling

Several climate-driven changes threaten the ocean's role as a carbon sink:

• Temperature: Warmer water holds less dissolved $CO_2$, reducing the efficiency of the solubility pump.
• Salinity and stratification: Freshwater input from melting ice increases surface stratification, which limits mixing and reduces the transport of carbon to depth.
• Biological productivity: Changes in nutrient upwelling and ocean temperatures alter phytoplankton communities and the strength of the biological pump.
• Circulation changes: A potential weakening of thermohaline circulation under warming scenarios could significantly reduce deep ocean carbon storage.

These factors create feedback loops that complicate climate predictions. Warmer oceans absorb less $CO_2$ (and may even release stored $CO_2$), which amplifies atmospheric warming, which further warms the oceans. Accurately modeling these feedbacks remains one of the major challenges in climate science.