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Physical oceanography is the branch of oceanography and geophysical fluid dynamics that studies the physical state, properties, and motions of the ocean, together with its exchanges of energy, momentum, and matter with the atmosphere, cryosphere, and lithosphere. The global ocean is a stratified, turbulent, rotating layer of fluid that covers about 71% of Earth's surface (3.61 × 10¹⁴ m²), has a volume of about 1.335 × 10¹⁸ m³ (a mass of about 1.4 × 10²¹ kg), and has a mean depth of about 3,688 m.

The specific heat capacity of seawater (c_p ≈ 3,990 J kg⁻¹ K⁻¹) is about four times that of dry air, and the mass of the ocean is about 270 times that of the atmosphere. As a result, the total heat capacity of the ocean is roughly 1,000 times that of the atmosphere. The ocean is therefore the main heat reservoir of the Earth system: it has taken up about 90% of the excess energy accumulated by the planet as a result of anthropogenic radiative forcing (see Earth's energy imbalance), and it carries up to about 1.5 to 2 petawatts (PW; 1 PW = 10¹⁵ W) of heat poleward from the tropics, shaping regional climate and weather patterns.

The dynamics of the ocean are governed largely by the distribution of seawater density (ρ), which controls buoyancy forces, horizontal pressure gradients, and vertical stability.

Unlike an ideal gas, seawater has a strongly non-linear equation of state. Under the international standard TEOS-10 (Thermodynamic Equation of Seawater – 2010), all thermodynamic properties of seawater are derived from a single Gibbs function of salinity, temperature, and pressure. Three state variables are used in practice:

1. Absolute Salinity (S_A, in g/kg), the mass fraction of dissolved material in seawater. The global mean is about 34.9 g/kg (equivalent to a practical salinity of about 34.7). The major ions occur in nearly constant proportions (the principle of constant proportions): by mass, chloride makes up 55.0%, sodium 30.6%, sulfate 7.7%, magnesium 3.7%, calcium 1.2%, potassium 1.1%, and bicarbonate 0.4%.
2. Conservative Temperature (Θ, in °C), proportional to potential enthalpy h⁰ (Θ = h⁰/_c__p⁰, where c_p⁰ = 3,991.87 J kg⁻¹ K⁻¹). Like potential temperature, it removes the warming caused by adiabatic compression (about 0.1 to 0.2 °C per kilometre of depth). Unlike potential temperature, it is very nearly conserved when water parcels mix, which makes it the preferred measure of ocean heat content.
3. Sea pressure (p, in decibars), the absolute pressure minus atmospheric pressure. One decibar (10⁴ Pa) is close to the pressure exerted by 1 m of seawater, from the hydrostatic relation ∂_p_/∂_z_ = −_ρg_ (with z positive upward).

For small perturbations about a reference density ρ₀, the change in density can be written as:

where α = −(1/_ρ_)(∂_ρ_/∂Θ) is the thermal expansion coefficient, β_S = (1/_ρ_)(∂_ρ_/∂_S__A) ≈ 7.8 × 10⁻⁴ kg/g is the haline contraction coefficient, and κ is the compressibility. Pure water has its maximum density at 3.98 °C. Dissolved salt lowers both the temperature of maximum density and the freezing point, the former more rapidly; above a practical salinity of about 24.7, the temperature of maximum density lies below the freezing point. Typical seawater (S_A ≈ 35 g/kg) therefore becomes steadily denser as it cools, right down to its freezing point of about −1.9 °C.

Because α varies strongly with temperature and pressure, two non-linear effects are important, particularly at high latitudes:

• Cabbeling: the thermal expansion coefficient increases with temperature, from about 0.5 × 10⁻⁴ K⁻¹ at 0 °C to about 3.3 × 10⁻⁴ K⁻¹ at 30 °C at the surface. Lines of constant density on a temperature–salinity diagram are therefore curved, so that when two water masses of equal density but different temperature and salinity mix, the mixture is denser than either parent water mass and tends to sink.
• Thermobaricity: the thermal expansion coefficient also increases with pressure; equivalently, cold seawater is more compressible than warm seawater. A cold parcel pushed downward through warmer surroundings compresses more than its surroundings and becomes relatively denser, which can enhance deep convection in regions such as the Weddell and Greenland seas.

Solar heating at the surface and the sinking of dense water at high latitudes divide the ocean vertically into three broad layers:

1. The surface mixed layer, typically 10 to 200 m thick but much deeper in winter at high latitudes, is kept nearly uniform in temperature, salinity, and density by wind stirring, breaking waves, Langmuir circulation, and convective cooling.
2. The pycnocline, from about 100 to 1,000 m, is a strongly stratified layer in which density increases rapidly with depth. At low and middle latitudes this is mainly due to the decrease in temperature across the permanent thermocline; at high latitudes it is mainly due to the increase in salinity across the halocline.
3. The deep and abyssal ocean, below about 1,000 m, contains about 75% of the ocean's volume. It is filled with cold (about −1 to 4 °C), weakly stratified water masses formed in polar regions.

The stability of the water column is measured by the buoyancy frequency N, the natural frequency at which a parcel displaced vertically will oscillate:

where ρ_θ is potential density. In the pycnocline, N ≈ 10⁻² s⁻¹, corresponding to an oscillation period of about 10 minutes. Strong stratification suppresses vertical mixing: the diapycnal (cross-density) diffusivity in the thermocline is only about 10⁻⁵ m²/s. Large-scale currents therefore flow almost along surfaces of constant density (isopycnals), where mesoscale eddies mix properties with a much larger diffusivity of about 10³ m²/s. Stratification also supports internal waves, with frequencies between the inertial frequency f and N.

The circulation of the upper ocean, from the surface down through the thermocline to about 1,000 m, is driven mainly by the momentum transferred from the prevailing winds. The wind stress τ (in N/m²) is commonly estimated with a bulk formula:

• *_τ_** = ρ_a C_D |U₁₀| U₁₀

where ρ_a is the density of air (about 1.2 kg/m³), C_D is a dimensionless drag coefficient (about 1 to 2.5 × 10⁻³, increasing with wind speed), and U₁₀ is the wind velocity 10 m above the sea surface.

In 1905, Vagn Walfrid Ekman, prompted by Fridtjof Nansen's observation during the Fram expedition that Arctic sea ice drifted 20° to 40° to the right of the wind, solved the equations for the wind-driven frictional boundary layer at the top of the ocean. In this Ekman layer, typically a few tens of metres to about 100 m thick, the vertical divergence of turbulent stress is balanced by the Coriolis force:

where f = 2Ω sin φ is the Coriolis parameter at latitude φ and Ω = 7.2921 × 10⁻⁵ rad/s is Earth's rotation rate. In Ekman's idealized solution, which assumes a constant eddy viscosity, the current decreases exponentially with depth while turning progressively to the right of the wind in the Northern Hemisphere (to the left in the Southern Hemisphere), forming the Ekman spiral; the surface current flows at 45° to the wind. Observed deflections are generally smaller.

Integrating these equations from the base of the layer, where the stress vanishes, to the surface, where it equals the wind stress, gives the Ekman transport M_E (in kg m⁻¹ s⁻¹), a result that does not depend on how the eddy viscosity varies with depth:

or, in vector form, M_E = (τ × k)/_f_. The net Ekman transport is therefore directed 90° to the right of the wind stress in the Northern Hemisphere (f > 0) and 90° to the left in the Southern Hemisphere (f < 0).

Because the winds vary with latitude, with easterly trade winds between about 0° and 30° and westerlies between about 30° and 60°, Ekman transports converge in some regions and diverge in others. Integrating the continuity equation over the Ekman layer shows that this convergence or divergence forces a vertical velocity w_E at the base of the layer:

• Convergence and downwelling (w_E < 0) in the subtropical gyres: between about 15° and 45° latitude, the trade winds on the equatorward side and the westerlies on the poleward side both drive Ekman transport toward the centre of the subtropical basin near 30°. The converging water raises sea level by about 1 m relative to the surroundings, depresses the thermocline by hundreds of metres, and is pumped downward at about 30 to 50 m per year (Ekman pumping).
• Divergence and upwelling (w_E > 0) in the subpolar gyres and at the equator: in subpolar basins (about 45° to 65°), between the westerlies and the polar easterlies, and along the equator, where the change in sign of f causes the trade winds to drive water poleward in both hemispheres, Ekman transport diverges. The thermocline domes upward and cold, nutrient-rich water is drawn toward the surface (Ekman suction).

Below the Ekman layer, friction is small (a small Ekman number), and for large-scale flows the Rossby number Ro = _U_/(fL) is also small (about 10⁻³ for gyre-scale flow), so inertia is negligible as well. The interior ocean is therefore in geostrophic balance between the horizontal pressure gradient, set by the slope of the sea surface and of the isopycnals, and the Coriolis force:

In a layer of fluid of thickness H in which the relative vorticity ζ = ∂_v_/∂_x_ − ∂_u_/∂_y_ is small compared with the planetary vorticity f, the large-scale potential vorticity _f_/_H_ is conserved. This leads to the linear vorticity balance:

where β = ∂_f_/∂_y_ = 2Ω cos _φ_/_a_ is the northward gradient of the Coriolis parameter, a is Earth's radius, and β ≈ 1.6 × 10⁻¹¹ m⁻¹ s⁻¹ at 45° latitude.

In a subtropical gyre, Ekman pumping pushes the top of the interior water column downward (w_E < 0), while the vertical velocity is near zero at the base of the thermocline. The water column is therefore squashed (∂_w_/∂_z_ < 0). To conserve _f_/_H_ as H decreases, the column must move toward smaller f, that is, toward the equator. Integrating the vorticity balance over depth and including the Ekman transport gives Harald Sverdrup's 1947 Sverdrup balance for the depth-integrated meridional transport V:

Across most of an ocean basin, the north–south transport is therefore set locally by the curl of the wind stress. Ocean volume transports are measured in sverdrups (Sv), where 1 Sv = 10⁶ m³/s, roughly the combined discharge of all the world's rivers.

In the North Atlantic subtropical gyre, the Sverdrup balance requires a broad, slow equatorward flow across the basin interior, totalling roughly 30 Sv at speeds of a few centimetres per second or less. Conservation of mass requires an equal poleward return flow. This return flow does not occur across the interior or along the eastern boundary; instead, it is concentrated in a narrow (about 50 to 100 km wide), deep, fast (about 1 to 2.5 m/s) current along the western side of each basin. This western intensification was first explained by Henry Stommel in 1948 using bottom friction, and extended by Walter Munk in 1950 using lateral friction. In steady state, the vorticity balance becomes:

Over a subtropical gyre, the wind continuously supplies negative (anticyclonic) vorticity. In the interior this is accommodated by equatorward motion toward smaller f. Water returning poleward, however, must gain planetary vorticity, and in a steady gyre this requires a source of positive (cyclonic) vorticity, which only friction against a coastline can provide in a narrow current. A poleward current along a western coast, with zero velocity at the coast and a maximum just offshore, has cyclonic shear, so friction there supplies vorticity of the required sign; a poleward current along an eastern coast would have anticyclonic shear and could not close the balance. Mathematically, only a western boundary layer gives a solution that decays away from the coast, with a width of δ_S = _r_/_β_ in Stommel's model or δ_M = (A_H/_β_)^1/3 in Munk's, where r is a bottom-friction coefficient and A_H a lateral eddy viscosity.

As a result, ocean basins show a marked east–west asymmetry between their boundary currents:

Comparison of western and eastern boundary currents

The Leeuwin Current off Western Australia is an exception: a poleward-flowing eastern boundary current.

In the North Atlantic, the Gulf Stream leaves the coast near Cape Hatteras and continues northeastward as the North Atlantic Current. Western Europe has much milder winters than eastern North America at the same latitude, a contrast long attributed to heat carried by the Gulf Stream. Modelling studies have shown that atmospheric processes contribute substantially: heat carried by the prevailing southwesterly winds from the ocean, and stationary atmospheric Rossby waves forced partly by the Rocky Mountains, which tend to place a cold trough over eastern North America and a warm ridge over western Europe. Some studies attribute most of the east–west contrast to these atmospheric mechanisms, while others emphasize the role of ocean heat transport by the subpolar gyre and the Atlantic meridional overturning circulation (AMOC) in warming Europe relative to the zonal average and in keeping the Nordic Seas largely free of winter sea ice.

Beneath the wind-driven gyres, the deep ocean is ventilated by the global meridional overturning circulation (MOC), also called the thermohaline circulation and popularized as the "great ocean conveyor" by the geochemist Wallace Smith Broecker.

At the sea surface, the buoyancy flux B₀ depends on the net heat flux into the ocean (Q_net, in W/m²) and on the net freshwater flux from evaporation minus precipitation, river runoff, and ice melt (E − P − R, in m/s):

where positive B₀ represents a gain of buoyancy by the ocean. Only in a few high-latitude regions is buoyancy loss strong enough to erode the stratification (N² ≤ 0) and trigger open-ocean deep convection or dense overflows down continental slopes that fill the deep basins:

• North Atlantic Deep Water (NADW): salty subtropical water, made saline by net evaporation over the tropical and subtropical Atlantic, partly because water vapour is exported across Central America to the Pacific, is carried northward into the Nordic Seas and the Labrador and Irminger seas. There, intense winter heat loss to cold Arctic air makes it dense enough to sink. Dense water formed in the Nordic Seas spills over the Greenland–Scotland Ridge through the Denmark Strait and the Faroe Bank Channel, entraining surrounding water as it descends, and combines with Labrador Sea Water formed by convection to about 2,000 m. The resulting NADW (about 2 to 4 °C, S_A ≈ 35.0 to 35.2 g/kg) occupies depths of about 1,500 to 4,000 m and flows southward along the western side of the Atlantic in the Deep Western Boundary Current at about 15 to 18 Sv.
• Antarctic bottom water (AABW): forms in coastal polynyas on the continental shelves of Antarctica, mainly in the Weddell Sea, the Ross Sea, and along the Adélie Coast. Katabatic winds blowing off the ice sheet push newly formed sea ice offshore, keeping the polynyas open so that ice continues to form rapidly. Because salt is largely excluded from the ice crystal lattice, freezing releases dense brine into the shelf water (brine rejection), producing High-Salinity Shelf Water at the surface freezing point (about −1.9 °C). This water mixes with Circumpolar Deep Water at the shelf break and descends the continental slope to the abyssal floor at depths greater than 4,000 m. AABW, with temperatures generally below 0 °C, is the densest widespread water mass in the ocean and spreads northward beneath NADW into the Atlantic, Indian, and Pacific basins.

It is often assumed that sinking at high latitudes drives the overturning circulation by density differences alone. However, as argued by Sandström in 1908 and revisited by Harold Jeffreys, Munk, and Carl Wunsch, a fluid that is heated and cooled at the same level, in this case the sea surface, supports only a weak, shallow overturning. Without a supply of mechanical energy to mix heat downward or to lift dense water back up through the stratification, the cold water formed at high latitudes would gradually fill the ocean up to a thin warm surface layer.

Surface buoyancy fluxes therefore determine where deep water forms and set the ocean's density structure, while the strength of the overturning, roughly 20 to 30 Sv in each of its main cells, depends on two processes that return deep water to the surface:

1. Diapycnal mixing. For abyssal water entering the deep Indian and Pacific oceans to rise across stable density surfaces, heat must be mixed downward by small-scale turbulence. In the simplest model, upwelling at speed w^* balances downward diffusion: w^* ∂_ρ_/∂_z_ ≈ ∂/∂_z_ (κ_ρ ∂_ρ_/∂_z_). Maintaining the required diffusivity (about 10⁻⁴ m²/s on average, much higher over rough topography) needs about 2 terawatts (2 × 10¹² W) of mechanical power, supplied by two sources:

• Tides (about 1 TW in the deep ocean): tidal currents flowing over rough topography such as mid-ocean ridges and seamounts generate internal tides (internal waves at tidal frequencies) that break and mix the deep water column.
• Wind (about 1 TW, highly uncertain): storms generate near-inertial internal waves in the mixed layer, some of which propagate downward and break in the interior.

1. Wind-driven upwelling in the Southern Ocean. In the latitude band of the Drake Passage (about 56° to 63°S), the Southern Ocean forms an unbroken channel around the globe. Above the depth of the major submarine ridges (about 2,000 m), there are no continental barriers to support a net east–west pressure gradient (∮ ∂_p_/∂_x_ d_x_ = 0). The strong westerlies drive a northward Ekman transport of about 30 Sv across the Antarctic Circumpolar Current (ACC), the largest current on Earth, which carries about 130 to 175 Sv eastward. This transport is replaced by Circumpolar Deep Water, which rises from 2,000 to 3,000 m along steeply sloping isopycnals to the surface with little need for diapycnal mixing; this upwelling is partly offset by the transport of mesoscale eddies. At the surface, the upwelled water divides into two branches. In the lower cell, cooling and brine rejection near Antarctica convert it into AABW. In the upper cell, freshening and warming convert it into Antarctic Intermediate Water and Subantarctic Mode Water, which flow northward and close the upper limb of the Atlantic overturning.

Together, Southern Ocean upwelling and diapycnal mixing link the ocean basins into a single global overturning circulation, often idealized as two interconnected cells centred on the Antarctic Circumpolar Current. Outside the deep western boundary currents, deep flow speeds are only about 0.1 to 1 cm/s. Radiocarbon and chlorofluorocarbon tracers indicate that deep water has been out of contact with the atmosphere for a few hundred years in the deep Atlantic and for about 1,000 to 1,500 years in the deep North Pacific.

The AMOC is unusual among ocean circulations. In the other basins, the ocean carries heat poleward in both hemispheres, but in the South Atlantic heat is carried northward toward the equator. Combined with heat gained in the tropics, this gives a northward heat transport of about 1.25 ± 0.35 PW at 26.5°N, measured continuously since 2004 by the RAPID-MOCHA mooring array. The AMOC also transfers anthropogenic heat and dissolved inorganic carbon into the deep ocean.

The atmosphere responds within hours to days to changes in sea surface temperature (SST), while the upper ocean adjusts to changes in the wind over months to years through the propagation of waves. In the tropics this makes the ocean and atmosphere a tightly coupled system. The dominant mode of year-to-year climate variability on Earth, recurring irregularly every 2 to 7 years, is the El Niño–Southern Oscillation (ENSO) in the equatorial Pacific.

At the equator, the Coriolis parameter is zero and changes sign, while its gradient is at a maximum (β = 2Ω/_a_ ≈ 2.3 × 10⁻¹¹ m⁻¹ s⁻¹). This forms an equatorial waveguide whose width is set by the equatorial Rossby radius of deformation, L_eq = √(_c_/_β_) ≈ 250 to 350 km. Here c = √(_g_′_H__e) ≈ 2 to 2.8 m/s is the phase speed of the first baroclinic gravity wave, where _g_′ = g Δ_ρ_/_ρ_₀ is the reduced gravity across the thermocline and H_e is an equivalent depth. Changes in the wind excite two main types of wave on the thermocline:

• Equatorial Kelvin waves are non-dispersive waves trapped at the equator that travel only eastward at about 2.5 m/s, crossing the Pacific (about 15,000 km) in about two to two and a half months. Westerly wind anomalies generate downwelling Kelvin waves, which deepen the thermocline and warm the surface; easterly anomalies generate upwelling Kelvin waves, which raise the thermocline and cool the surface.
• Equatorial Rossby waves travel westward, with maximum amplitude a few degrees off the equator. Long Rossby waves travel at c_R = −_c_/(2_n_ + 1), where n = 1, 2, ... is the meridional mode number, so the gravest mode travels at one third of the Kelvin wave speed and takes roughly 6 to 9 months to cross the basin. At the western boundary, Rossby waves reflect as equatorial Kelvin waves.

In 1969, Jacob Bjerknes linked the oceanic El Niño, an anomalous warming of the eastern and central equatorial Pacific, with the atmospheric Southern Oscillation described by Gilbert Walker, a seesaw in surface pressure between Darwin, Australia, and Tahiti that is measured by the Southern Oscillation Index (SOI). He proposed a positive feedback, now called the Bjerknes feedback:

1. A warm SST anomaly in the eastern or central equatorial Pacific reduces the east–west SST gradient across the basin.
2. Because the east–west pressure gradient and the atmospheric Walker circulation are driven by this SST gradient, the pressure difference between Tahiti and Darwin decreases and the easterly trade winds weaken.
3. Weaker trade winds reduce westward surface currents and equatorial upwelling in the east, and excite downwelling Kelvin waves that deepen the eastern thermocline. These changes further warm the eastern Pacific, reinforcing the initial anomaly.

In the climatological mean state, the easterly trade winds blow westward along the equator from South America toward Indonesia.

• Sea-level and thermocline slope: the westward wind stress piles warm water into the Indo-Pacific Warm Pool, raising sea level in the west by about 40 to 60 cm relative to the east. The thermocline lies at about 150 to 200 m in the west, where SSTs reach 28 to 30 °C, but only about 30 to 50 m in the east.
• Eastern Pacific cold tongue: just north of the equator, the Coriolis effect deflects the wind-driven flow northward, and just south of the equator it is deflected southward, so the trade winds cause equatorial Ekman divergence. The diverging surface water is replaced by water upwelled from 50 to 100 m depth. Because the eastern thermocline is shallow, this water is cold and rich in nutrients, producing the eastern Pacific cold tongue (SSTs of about 20 to 23 °C) and supporting high biological productivity, including the Peruvian anchoveta fishery.
• Walker circulation: warm, moist air rises in deep convection over the western Pacific warm pool, bringing heavy rainfall to Indonesia and northern Australia. It flows eastward near the tropopause (about 200 hPa), sinks over the cool eastern Pacific, and returns westward at the surface as the trade winds.

An El Niño event typically begins when a build-up of warm water in the equatorial Pacific coincides with bursts of westerly wind over the western Pacific, often associated with the Madden–Julian oscillation.

• Thermocline flattening and reduced upwelling: as the trade winds weaken or reverse, downwelling Kelvin waves carry warm water eastward. The east–west sea-level slope flattens and the eastern thermocline deepens by tens of metres, by as much as about 100 m in strong events. Even where winds still drive some upwelling, the thermocline is now so deep that the upwelled water comes from the warm surface layer rather than the cold, nutrient-rich layer below. SSTs in the central and eastern Pacific rise by about 1.5 to 4 °C above normal, and coastal productivity and anchoveta populations collapse.
• Monitoring index: operational centres such as NOAA's Climate Prediction Center use the Oceanic Niño Index (ONI), the three-month running mean SST anomaly in the Niño 3.4 region (5°N to 5°S, 170°W to 120°W). An El Niño episode is declared when the ONI is at or above +0.5 °C for at least five consecutive overlapping three-month periods. Events are often informally described as weak (+0.5 to +0.9 °C), moderate (+1.0 to +1.4 °C), strong (+1.5 to +1.9 °C), or very strong (+2.0 °C or more). Eastern Pacific (canonical) events, with the largest warming in the Niño 3 and Niño 1+2 regions, are distinguished from Central Pacific ("Modoki") events, with the largest warming in the Niño 4 region.
• Global teleconnections: as the warm water shifts east, the rising branch of the Walker circulation shifts with it and the SOI becomes negative. Heating from deep convection over the central and eastern Pacific generates stationary atmospheric Rossby wave trains, such as the Pacific–North American teleconnection pattern, and strengthens the subtropical jet stream. Typical impacts include drought and wildfires in Indonesia, Australia, India, and the Amazon; heavy rain and flooding in coastal Peru and Ecuador and along the southern tier of the United States; mild winters in northern North America; and suppressed Atlantic tropical cyclone activity because of increased vertical wind shear over the Caribbean.

The end of an El Niño and the transition to La Niña (ONI at or below −0.5 °C) are explained by two related mechanisms. In the delayed oscillator, upwelling Rossby waves generated off the equator during El Niño travel west, reflect at the western boundary as upwelling Kelvin waves, and return to raise the eastern thermocline several months later. In the recharge–discharge oscillator proposed by Fei-Fei Jin, the anomalous wind-stress curl during El Niño drives a poleward Sverdrup transport, V = k · (∇ × τ)/_β_, that drains warm water from the equatorial band and leaves the equatorial thermocline unusually shallow.

Once the equatorial heat content has been discharged, the system can overshoot into La Niña:

• Stronger trade winds and an expanded cold tongue: the trade winds strengthen (positive SOI), piling warm water into the western Pacific and raising the eastern thermocline to within about 10 to 20 m of the surface. Enhanced upwelling spreads cold water (about 1 to 2 °C below normal) rich in nutrients and dissolved inorganic carbon across the central and eastern Pacific.
• Opposite teleconnections: convection strengthens over the Maritime Continent, bringing heavy rain and flooding to Australia, Southeast Asia, and northern South America, and drought to the southwestern United States, the Horn of Africa, and parts of southeastern South America. Global mean surface temperature is temporarily lowered, and reduced wind shear over the tropical Atlantic favours active Atlantic hurricane seasons.

Typical conditions during the phases of ENSO

• Ocean current
• Ocean heat content
• Indian Ocean Dipole
• Pacific decadal oscillation
• Sea level rise