Ocean Zones & Physical Chemistry

0.1 Depth Zonation of the Ocean

The ocean is divided into five major depth zones, each with distinct physical, chemical, and biological characteristics. These zones are defined primarily by light penetration, temperature, pressure, and nutrient availability.

Zone Boundaries as Ecological Transitions

Zone boundaries are not sharp lines but gradients. The euphotic zone boundary (base of epipelagic) is defined by the 1% light level — the depth where photosynthetically active radiation (PAR) falls to 1% of its surface value. This varies with water clarity: ~20 m in turbid coastal waters to ~200 m in the clearest subtropical gyres. The compensation depth, where gross photosynthesis exactly balances algal respiration, typically lies at 1–5% of surface PAR.

Understanding these zones is foundational for ocean biogeochemistry and connects directly toclimate-biodiversity interactions andclimate science, since the ocean's vertical structure controls the biological carbon pump that sequesters atmospheric CO\(_2\).

0.2 Light Attenuation: Beer-Lambert Law

Light intensity decreases exponentially with depth due to absorption and scattering by water molecules, dissolved organic matter (CDOM), phytoplankton pigments, and suspended particles. The governing law is the Beer-Lambert law of exponential attenuation:

where \(I_0\) is the surface irradiance (typically 400–2000 \(\mu\)mol photons/m\(^2\)/s for PAR), and \(K_d\) is the diffuse attenuation coefficient (m\(^{-1}\)).

Derivation: Wavelength Dependence of \(K_d\)

The attenuation coefficient is the sum of absorption and scattering contributions:

Pure water absorption \(a_w(\lambda)\) dominates the wavelength structure. Red light (\(\lambda \approx 700\) nm) is strongly absorbed: \(a_w(700) \approx 0.65\;\text{m}^{-1}\), meaning 50% absorption in just 1.1 m. Blue light (\(\lambda \approx 475\) nm) penetrates deepest: \(a_w(475) \approx 0.015\;\text{m}^{-1}\), reaching 1% at ~300 m in clear ocean water. This is why the deep ocean appears blue.

Euphotic Zone and Compensation Depth

The euphotic zone depth \(z_{eu}\) is where irradiance falls to 1% of the surface value. Setting \(I(z_{eu}) = 0.01 \cdot I_0\):

For clear ocean water (\(K_d = 0.03\)): \(z_{eu} \approx 154\) m. For coastal water (\(K_d = 0.20\)): \(z_{eu} \approx 23\) m.

The compensation depth \(z_c\) is where gross photosynthesis\(P_g\) equals phytoplankton respiration \(R\). Below \(z_c\), cells consume more carbon than they fix. For a P-I curve \(P_g = P_{\max} \tanh(\alpha I / P_{\max})\), the compensation irradiance \(I_c\) satisfies \(P_{\max} \tanh(\alpha I_c / P_{\max}) = R\), typically at 1–5 \(\mu\)mol photons/m\(^2\)/s.

0.3 Hydrostatic Pressure

Pressure increases linearly with depth in the ocean, governed by hydrostatic equilibrium. The water column exerts a weight proportional to its height and density:

Derivation: Pressure at Key Depths

One atmosphere of additional pressure is added for every 10.33 m of seawater:

At the bottom of the Mariana Trench (Challenger Deep, \(z = 10{,}994\) m):

Biochemical Effects of Pressure

High pressure profoundly affects macromolecular structure. By Le Chatelier's principle, pressure favours states with smaller molar volume. The volume change of a reaction determines its pressure sensitivity:

Key effects on biomolecules:

• Protein unfolding: The hydrophobic core of proteins has a larger molar volume than the unfolded state (water fills cavities). At >200 atm, many mesophilic enzymes lose activity. \(\Delta V_{\text{unfold}} \approx -50\;\text{to}\;-100\;\text{mL/mol}\)
• Membrane compression: Lipid bilayers become more ordered and rigid. Deep-sea organisms incorporate more unsaturated fatty acids to maintain fluidity (homeoviscous adaptation).
• TMAO accumulation: Trimethylamine N-oxide increases linearly with depth in fish (approximately 0.04 mol/atm). It stabilises proteins by strengthening the hydration shell around hydrophobic groups.

0.4 Temperature Structure: Mixed Layer, Thermocline, Deep Ocean

The ocean's vertical temperature structure consists of three layers:

• Mixed layer (0–100 m): Warm, well-mixed by wind, waves, and convection. Temperature ~20–28 °C in tropics, ~5–15 °C at mid-latitudes. Nearly isothermal due to turbulent mixing.
• Thermocline (100–1000 m): Rapid temperature decrease, acting as a barrier to vertical mixing. The permanent thermocline exists year-round; a seasonal thermocline forms and erodes above it.
• Deep ocean (below ~1000 m): Cold (1–4 °C) and nearly uniform. Filled by thermohaline circulation — cold, dense water sinking at high latitudes (North Atlantic Deep Water, Antarctic Bottom Water).

Heat Diffusion in the Thermocline

The thermocline can be modelled as a balance between downward diffusion of heat from the warm surface and upwelling of cold deep water. At steady state, the advection-diffusion equation gives:

The solution is an exponential profile:

With \(\kappa \approx 10^{-4}\;\text{m}^2/\text{s}\) (diapycnal diffusivity) and\(w \approx 10^{-7}\;\text{m/s}\) (\(\sim 3\) m/yr upwelling), the thermocline scale height is \(h = 10^{-4}/10^{-7} = 1000\) m, consistent with observations. This connects directly to ocean circulation discussed in Climate Science Module 2.

0.5 Dissolved Oxygen Profile & the Oxygen Minimum Zone

Dissolved oxygen shows a distinctive vertical profile with three regimes:

• Surface (0–100 m): Near saturation (~6–8 mL/L) due to air-sea gas exchange and photosynthetic production. Saturation follows Henry's law, decreasing with temperature (warmer water holds less O\(_2\)).
• OMZ (200–1000 m): O\(_2\) drops to <0.5 mL/L in the most intense OMZs (eastern tropical Pacific, Arabian Sea, Bay of Bengal). Biological oxygen demand from bacterial decomposition of sinking organic matter exceeds the rate of O\(_2\) supply by lateral advection and mixing.
• Deep ocean (>1000 m): O\(_2\) gradually increases to 4–6 mL/L, replenished by thermohaline circulation carrying cold, O\(_2\)-saturated water from the polar surface to the deep ocean interior.

O\(_2\) Consumption Rate Derivation

At steady state, the vertical O\(_2\) distribution balances advection, diffusion, and biological consumption:

The bacterial decomposition follows sinking organic matter flux, which decays as a power law (Martin curve, see Module 2):

Where \(R_0 \approx 5\text{--}20\;\mu\text{mol O}_2/\text{kg/yr}\) at the base of the euphotic zone (\(z_0 = 100\) m). The stoichiometry of organic matter remineralisation follows the Redfield ratio:\( (\text{CH}_2\text{O})_{106}(\text{NH}_3)_{16}(\text{H}_3\text{PO}_4) + 138\text{O}_2 \to 106\text{CO}_2 + 16\text{HNO}_3 + \text{H}_3\text{PO}_4 + 122\text{H}_2\text{O} \). Thus 138 moles of O\(_2\) are consumed per 106 moles of organic carbon remineralised.

0.6 Salinity, pH, and Nutrient Profiles

Pycnocline and Density Stratification

The pycnocline is the zone of rapid density increase with depth, controlled by both temperature (thermocline) and salinity (halocline). Seawater density is computed from the equation of state:

Ocean pH and the Carbonate System

Surface ocean pH is ~8.1, decreasing with depth to ~7.6–7.8 due to CO\(_2\) released by remineralisation. The carbonate chemistry equilibria:

As CO\(_2\) increases from remineralisation, pH drops and carbonate ion concentration decreases. Below the carbonate compensation depth (CCD, ~4500 m), the water is sufficiently undersaturated that CaCO\(_3\) dissolves faster than it accumulates — no carbonate sediment is preserved. This has implications forcarbon cycle feedbacks.

Redfield Ratio: C:N:P = 106:16:1

Alfred Redfield (1934) discovered that the elemental composition of marine plankton and the dissolved nutrients in deep water follow a remarkably consistent ratio:

This ratio emerges because plankton assimilate nutrients in fixed proportions during growth, and deep-water nutrient concentrations reflect cumulative remineralisation. The deviation from Redfield reveals nutrient limitation:

• If \(\text{N:P} < 16\) in surface water: nitrogen-limited (most of the ocean)
• If \(\text{N:P} > 16\): phosphorus-limited (Mediterranean, some coastal)
• Iron limitation in HNLC regions (Southern Ocean, subarctic Pacific) overrides N and P

The nutricline — the depth of rapid nutrient increase — typically coincides with the thermocline (~100–500 m), since nutrients accumulate where remineralisation occurs and vertical mixing is suppressed by stratification.

0.7 Seawater Equation of State & Stability

The full UNESCO equation of state (EOS-80) relates seawater density to temperature, salinity, and pressure through a polynomial with 41 coefficients. For most oceanographic applications, the simplified linear approximation suffices for small perturbations from a reference state:

Potential density \(\sigma_\theta\) removes the effect of adiabatic compression by referencing all parcels to surface pressure. This is essential for assessing vertical stability: a water column is stably stratified when \(\partial \sigma_\theta / \partial z > 0\)(density increases with depth).

Brunt-Väisälä Frequency

The buoyancy frequency \(N\) quantifies the strength of stratification and the frequency of internal gravity waves:

When \(N^2 > 0\), the water column is stable and resists vertical mixing. A displaced parcel oscillates about its equilibrium depth with period \(T = 2\pi/N\). In the pycnocline,\(T \approx 10\) minutes; in the deep ocean, \(T \approx 1\) hour.

Strong stratification (\(N^2\) large) in the tropics inhibits vertical nutrient supply to the surface, explaining why tropical gyres are oligotrophic despite high solar radiation. Conversely, weak stratification at high latitudes allows deep winter convection and nutrient replenishment, fueling productive spring blooms (see Module 1).

Thermohaline Circulation and Ventilation

The deep ocean is ventilated by the thermohaline (meridional overturning) circulation. Cold, saline surface water at high latitudes (principally the North Atlantic and around Antarctica) becomes dense enough to sink to the abyss:

• North Atlantic Deep Water (NADW): forms in the Greenland-Iceland-Norwegian seas, sinks to ~2000–4000 m, flows southward. Temperature ~2–4 °C, salinity ~34.9 psu.
• Antarctic Bottom Water (AABW): forms on Antarctic continental shelf (brine rejection from sea ice formation), sinks to the very bottom (>4000 m). Temperature <0 °C, salinity ~34.7 psu. The densest water mass in the ocean.
• Antarctic Intermediate Water (AAIW): subducts at the Antarctic Convergence (~50°S), spreads northward at ~700–1200 m. Low salinity (~34.3 psu), high O\(_2\).

The global overturning circulation takes approximately 1000–1500 years for a full cycle. This sets the timescale for deep-ocean ventilation and is crucial for understanding oceanic carbon sequestration and climate dynamics. Water parcels that sink in the North Atlantic today will not return to the surface for a millennium, carrying with them any dissolved CO\(_2\) and nutrients they contain.

0.8 Sound Propagation & the SOFAR Channel

Sound speed in seawater depends on temperature, salinity, and pressure through an empirical relationship:

Temperature dominates near the surface (warmer = faster), while pressure dominates at depth (deeper = faster). This creates a sound speed minimum at ~700–1200 m depth (the SOFAR channel), where sound is trapped by refraction and can propagate for thousands of kilometres with minimal attenuation.

Marine mammals exploit this: blue whale calls at the SOFAR channel axis can be detected across entire ocean basins. The minimum sound speed \(c_{\min} \approx 1480\;\text{m/s}\) occurs where\(\partial c / \partial z = 0\):

By Snell's law, sound rays bend toward regions of lower sound speed, trapping acoustic energy in the channel. This phenomenon is exploited in ocean acoustic tomography to measure large-scale temperature changes, directly relevant to climate monitoring.

The SOFAR channel also has biological significance: marine mammals use low-frequency sound for long-range communication. Blue whales produce calls at 10–20 Hz that can travel over 1000 km through the SOFAR channel. Anthropogenic noise from shipping, sonar, and seismic surveys is increasingly disrupting these communication channels, with documented effects on whale behaviour, stress hormones, and stranding events. The sound intensity from large container ships reaches 190 dB re 1 \(\mu\)Pa at 1 m, significantly raising the ambient noise floor in the 20–200 Hz frequency band throughout the world's oceans.

Ocean Depth Profile: Multi-Variable Cross-Section

The following diagram shows the vertical distribution of temperature, dissolved oxygen, light intensity, pH, and pressure across the five ocean depth zones:

0.9 Dissolved Gas Solubility & Henry's Law

The solubility of gases in seawater is governed by Henry's law, which states that at equilibrium, the dissolved gas concentration is proportional to its partial pressure in the overlying atmosphere:

For O\(_2\), solubility decreases approximately 2% per °C of warming: at 0 °C, saturated seawater holds ~8 mL/L; at 30 °C, only ~4.5 mL/L. This temperature dependence is a key driver of ocean deoxygenation under climate change, as discussed inClimate Science Module 6.

For CO\(_2\), the solubility is enhanced by chemical reactions with water (hydration to carbonic acid), making the ocean the largest active carbon reservoir (~38,000 Gt C dissolved inorganic carbon). The temperature dependence of CO\(_2\) solubility follows:

The Revelle factor \(R_f \approx 10\) describes the buffer capacity: a 1% increase in atmospheric CO\(_2\) produces only a ~0.1% increase in dissolved inorganic carbon, because most added CO\(_2\) is neutralised by reaction with carbonate ions. As ocean CO\(_2\) uptake continues, \(R_f\)increases (buffer capacity decreases), reducing the ocean's ability to absorb future emissions.

Simulation: Ocean Depth Profiles

Quantitative depth profiles for temperature, dissolved oxygen, light intensity (Beer-Lambert), pressure, and pH, with compensation depth calculation and thermocline modelling:

Simulation: Sound Speed Profile & SOFAR Channel

The sound speed profile showing the SOFAR channel minimum, with temperature and pressure contributions decomposed:

Simulation: Thermocline Advection-Diffusion Model

Testing the Munk (1966) model: the balance between upwelling \(w\) and turbulent diffusion\(\kappa\) sets the thermocline structure. We also explore how changes in diffusivity and upwelling rate affect the temperature profile:

References

1. Munk, W. H. (1966). Abyssal recipes. Deep-Sea Research, 13(4), 707–730.
2. Redfield, A. C. (1934). On the proportions of organic derivatives in sea water and their relation to the composition of plankton. James Johnstone Memorial Volume, 176–192.
3. Kirk, J. T. O. (2011). Light and Photosynthesis in Aquatic Ecosystems (3rd ed.). Cambridge University Press.
4. Sarmiento, J. L. & Gruber, N. (2006). Ocean Biogeochemical Dynamics. Princeton University Press.
5. Yancey, P. H., Gerringer, M. E., Drazen, J. C., Rowden, A. A., & Jamieson, A. (2014). Marine fish may be biochemically constrained from inhabiting the deepest ocean depths. PNAS, 111(12), 4461–4465.
6. Jamieson, A. J. (2015). The Hadal Zone: Life in the Deepest Oceans. Cambridge University Press.
7. Emerson, S. & Hedges, J. (2008). Chemical Oceanography and the Marine Carbon Cycle. Cambridge University Press.