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Meteorology is the branch of the atmospheric sciences and geophysical fluid dynamics that investigates the physical, dynamical, thermodynamic, and radiative processes governing the behavior of a planetary atmosphere. While operationally associated with synoptic and mesoscale weather forecasting, modern meteorology integrates compressible fluid mechanics, atmospheric thermodynamics, radiative transfer, cloud physics, and boundary layer meteorology to model how thermal energy, momentum, water vapor, and trace chemical species circulate through the Earth system.

In the hierarchy of Earth system science, the atmosphere functions as a compressible, stratified fluid shell on a rotating sphere heated differentially by solar radiation. Understanding the governing equations of atmospheric motion is essential for predicting both short-term tropospheric weather transients and the long-term statistical reorganization of global climate.

A foundational distinction in atmospheric physics separates weather from climate based on dynamical timescales, phase-space geometry, and the mathematical classification of numerical weather prediction versus climate models as formulated by meteorologist Edward Lorenz.

Weather is the instantaneous, short-term thermodynamic and kinematic state of the atmosphere (principally within the troposphere and lower stratosphere) at a specific spatial coordinate and time. It is completely specified by the local atmospheric state vector X(x, t):

X(x, t) = ( u, v, w, T, P, ρ, q_v, q_c, q_i )

comprising the three-dimensional wind velocity components (zonal u, meridional v, and vertical w), absolute temperature (T), barometric pressure (P), air density (ρ), specific humidity (q_v), and liquid/ice cloud condensate mixing ratios (q_c, q_i) evolving over timescales of minutes, hours, days, and synoptic weeks.

Mathematically, forecasting the weather is a Cauchy initial value problem (termed Lorenz Predictability of the First Kind): integrating the non-linear partial differential equations of fluid motion forward in time from an observed initial atmospheric state X(x, t₀). As discovered by Edward Lorenz in 1963, because the advective terms (V · ∇V) and moist diabatic phase transitions in the atmospheric equations are strongly non-linear, the atmosphere behaves as a deterministic chaotic system exhibiting sensitive dependence on initial conditions.

Any infinitesimal observational uncertainty or unresolved sub-grid turbulence in the initial state (_δ_X₀) grows exponentially with time t at a rate governed by the system's leading positive Lyapunov exponent (λ₁ > 0):

|_δX(t)| ≈ |_δX₀| · e^λ₁t

Furthermore, because atmospheric kinetic energy at mesoscales (<500 km, where _E_(k) &propto; k^-5/3) cascades upscale into synoptic planetary Rossby-wave scales (>1,000 km, where geostrophic turbulence follows a k^-3 enstrophy cascade), errors at the scale of individual convective thunderstorms double every few hours and contaminate synoptic-scale wave phases within days. Consequently, even with dense satellite data assimilation and ensemble forecasting, the intrinsic theoretical limit of deterministic day-to-day weather predictability is strictly bounded at approximately 10 to 14 days.

Climate, by contrast, is the long-term statistical ensemble distribution of the atmospheric state vector X(x, t), conventionally evaluated over a standard 30-year reference period (Climatological Standard Normals) established by the World Meteorological Organization (WMO). In dynamical systems theory, while an individual weather forecast traces a single trajectory along the leaves of the Lorenz strange attractor, climate represents the invariant probability density distribution (including the mean, seasonal cycle, interannual variance, spatial covariance modes, and extreme value theory tail probabilities) across the entire geometry of the attractor.

Predicting climate (and its response to external perturbations) is therefore a boundary value problem (Lorenz Predictability of the Second Kind). Rather than depending on the exact positions of individual cyclones on day t₀, the long-term statistical geometry of the climate attractor is governed by external radiative forcing and slowly evolving Earth system boundary conditions: solar irradiance, atmospheric concentrations of greenhouse gases (CO₂, CH₄, N₂O) and aerosols, sea surface temperatures and deep ocean heat content, cryosphere ice-sheet topography, and continental albedo and vegetative roughness.

The physical behavior of the atmosphere is governed by a coupled set of six fundamental conservation equations known as the primitive equations, which adapt the compressible Navier-Stokes equations, conservation of mass, and first law of thermodynamics to a thin fluid shell on a sphere of radius a (~6,371 km) rotating at constant angular velocity Ω = 7.2921 × 10^-5 rad/s:

1. Zonal and Meridional Momentum Equations (in a Rotating Non-Inertial Frame):

1. The Hydrostatic Approximation (Vertical Momentum): Because the vertical depth of the troposphere (H &sim; 10 km) is two to three orders of magnitude smaller than the horizontal wavelength of synoptic systems (L &sim; 1,000 to 10,000 km, yielding a small aspect ratio _H_/_L_ &ll; 1), vertical accelerations (_Dw_/_Dt_) are negligible compared to buoyancy and gravity (except inside kilometer-scale non-hydrostatic convective thunderstorm updrafts). The vertical momentum equation reduces to exact hydrostatic balance between the upward vertical pressure-gradient force and downward gravity (g = 9.81 m/s²):

1. Mass Continuity Equation: Expresses the conservation of total air mass in three-dimensional compressible flow:

1. Equation of State for Moist Air (Ideal gas law and Virtual temperature): Because water vapor (M_v = 18.015 g/mol) has a lower molar mass than dry air (M_d = 28.964 g/mol, ratio ε = M_v/_M__d = R_d/_R__v ≈ 0.622), moist air is less dense and more buoyant than dry air at the same temperature and pressure. To retain the dry gas constant (R_d = 287.05 J·kg^-1·K^-1) in the state equation, meteorologists define virtual temperature (T_v), the temperature that dry air would need to attain to equal the density of moist air with specific humidity q_v:

1. Thermodynamic Energy Equation and Potential temperature (&theta;): Derived from the first law of thermodynamics (c_pdT - αdP = J, where c_p = 1,004 J·kg^-1·K^-1 is isobaric specific heat, α = 1/_ρ_, and J is the diabatic heating rate per unit mass from radiation and latent heat release). Defining potential temperature (&theta;) via Poisson's equation as the temperature an air parcel would have if compressed or expanded adiabatically from pressure P to a standard reference pressure P₀ = 1,000 hPa:

1. Water Vapor Continuity Equation: Dq_v / Dt = S_q (where S_q represents evaporative sources minus condensation/precipitation sinks, governed by the Clausius-Clapeyron relation).

The ultimate thermodynamic driver of all atmospheric motion is the latitudinal imbalance of radiative heating across the spherical surface of the Earth. At a mean orbital distance d̄ = 1 AU, the top of the atmosphere intercepts a parallel beam of solar irradiance with flux density S₀ ≈ 1,361 W/m².

At any given latitude &phi;, seasonal solar declination angle δ (ranging from -23.44° at the December solstice to +23.44° at the June solstice due to Earth's axial obliquity), and diurnal hour angle h, the instantaneous top-of-atmosphere solar irradiance I_TOA incident upon a horizontal unit surface area is proportional to the cosine of the solar zenith angle (&theta;_z), governed by the spherical law of cosines:

I_TOA = S₀ · ( d̄ / d )² · cos(&theta;_z) = S₀ · ( d̄ / d )² · ( sin_&phi;_ · sin_δ_ + cos_&phi;_ · cos_δ_ · cos_h_ )

Integrating this relation from sunrise (hour angle -h₀, where cos_h_₀ = -tan_&phi;_·tan_δ_) to sunset (+h₀) yields the daily average insolation Q̄_day(&phi;, δ):

Q̄_day = ( S₀ / &pi; ) · ( d̄ / d )² · ( h₀ · sin_&phi;_ · sin_δ_ + cos_&phi;_ · cos_δ_ · sin_h_₀ )

Because &theta;_z is small in the tropics (beam hits nearly perpendicularly, concentrating radiative energy over a minimal surface footprint and traversing a short optical path length through the atmosphere) and large near the poles (beam strikes at an oblique glancing angle, spreading the same photon flux over a vast surface area A = A₀ / cos_&theta;__z), annual mean incident insolation at the poles (~175 W/m²) is less than half that at the equator (~415 W/m²). Furthermore, this geometric insolation contrast is amplified by the latitudinal gradient of planetary albedo (α_p): dark tropical oceans absorb ~85% to 90% of incident sunlight, whereas the high-albedo polar cryosphere (snow, sea ice, and continental ice sheets, where α ≈ 0.60 to 0.85) reflects the majority of incident shortwave radiation back into space.

While absorbed shortwave solar radiation (ASR = _Q̄_(1 - α_p)) varies steeply from equator to pole, outgoing longwave radiation (OLR = σT_e⁴) emitted to space varies much more weakly with latitude. Consequently, plotting zonal-mean ASR minus OLR as a function of latitude &phi; reveals two distinct planetary radiative regimes separated near 38°N and 38°S latitude:

• Low-Latitude Radiative Surplus (38°S to 38°N): Annual absorbed solar radiation exceeds outgoing longwave emission (ASR > OLR), continuously accumulating thermal energy in the tropics and subtropics.
• High-Latitude Radiative Deficit (38° to 90° N and S): Outgoing longwave thermal emission to space exceeds absorbed solar radiation (OLR > ASR), continuously cooling the middle and polar latitudes.

To prevent the tropics from heating indefinitely and the poles from cooling toward absolute zero, the coupled ocean-atmosphere fluid envelope operates as a global heat engine that transports approximately 5.5 to 6.0 petawatts (10¹⁵ W) of thermal energy poleward across the ~35° to 40° latitude circles in each hemisphere. While wind-driven and thermohaline ocean currents dominate poleward heat transport in the deep tropics (peaking at ~2 PW near 15° to 20° latitude), the atmosphere accounts for 75% to 80% (~4 to 5 PW) of the total poleward energy transport across the mid-latitudes (35° to 60°).

Atmospherically, this meridional energy flux is carried as moist static energy (MSE), the sum of dry enthalpy (sensible heat), gravitational potential energy, and evaporative latent heat of water vapor:

MSE = c_pT + g · z + L_v · q_v

In 1735, English meteorologist George Hadley proposed that the equator-to-pole heating gradient drives a single, hemisphere-spanning thermally direct convective loop in each hemisphere, with warm air rising at the equator, flowing all the way to the poles aloft, sinking at the poles, and returning equatorward along the surface.

On a rotating planet, however, a single equator-to-pole cell is dynamically impossible due to the conservation of angular momentum. The absolute axial angular momentum per unit mass (M) of an air ring located at latitude &phi; (where distance to the Earth's rotational axis is r = _a_·cos_&phi;_) moving with zonal (eastward) velocity u is the sum of planetary angular momentum and relative atmospheric angular momentum:

M = ( Ω · a · cos_&phi;_ + u ) · a · cos_&phi;_ = Ωa²cos²&phi; + _ua_cos_&phi;_

In the frictionless upper troposphere, an air ring ascending at the equator with zero initial zonal wind (&phi; = 0, u = 0, so M_eq = Ωa²) and moving poleward to latitude &phi; while conserving its axial angular momentum (_M_(&phi;) = M_eq) must accelerate eastward to a zonal velocity u_M(&phi;) as its radius of rotation _a_·cos_&phi;_ shrinks:

u_M(&phi;) = Ω · a · ( (1 - cos²&phi;) / cos_&phi;_ ) = Ω · a · ( sin²&phi; / cos_&phi;_ )

Evaluating this angular-momentum-conserving wind speed demonstrates why the Hadley cell cannot reach the poles: by latitude &phi; = 30° (where Ωa ≈ 465 m/s at the equator), u_M(30°) reaches 134 m/s (~483 km/h); by &phi; = 60°, u_M(60°) would reach 698 m/s (~2,510 km/h, Mach 2.3); and as &phi; → 90°, u_M → ∞. Long before reaching mid-latitudes, this extreme vertical and horizontal wind shear triggers vigorous baroclinic instability, breaking the zonal flow into turbulent Rossby waves and terminating the axisymmetric overturning cell near 25° to 30° latitude. As derived analytically in the Held-Hou model (1980), combining angular momentum conservation with thermal wind balance and energy conservation restricts the poleward boundary of the Hadley cell (&phi;_H) to scale with the square root of the Rossby radius / thermal forcing:

&phi;_H ≈ ( ( 5 · g · H_t · Δ_H ) / ( 3 · Ω² · a² · T₀ ) )^1/2 ≈ 25° to 30°

where H_t is the tropopause height (~15 to 17 km) and Δ_H / T₀ is the fractional equator-to-pole radiative-equilibrium potential temperature difference. Consequently, Earth's rotation fractures the hemispheric circulation into three distinct latitudinal cells in each hemisphere.

The Hadley cell is a powerful, thermally direct meridional overturning circulation (where warm, buoyant air rises and cold, dense air sinks, converting thermal potential energy directly into kinetic energy) that dominates the tropics and subtropics:

• Ascending Branch (The Intertropical Convergence Zone, ITCZ): Driven by intense tropical solar heating and high sea surface temperatures, moisture-laden surface trade winds from the Northern and Southern Hemispheres converge along the thermal equator inside a planetary trough of low surface barometric pressure known as the ITCZ (or equatorial doldrums). Forced upward inside deep cumulonimbus convective towers, the ascending saturated air cools at the moist adiabatic lapse rate (~4 to 5 K/km), condensing massive quantities of water vapor. This condensation produces the planet's heaviest equatorial precipitation belts (sustaining tropical rainforest biomes) while releasing ~2.5 × 10⁶ J/kg of latent heat into the middle and upper troposphere, providing the buoyant enthalpy that drives the entire cell upward to the tropical tropopause (~16 to 18 km altitude).
• Poleward Upper-Tropospheric Branch and Subtropical Jet: Reaching the stable tropopause barrier, the air diverges poleward in both hemispheres. As this upper-tropospheric air moves toward ~30° latitude, the Coriolis force deflects it strongly eastward, concentrating axial angular momentum into the subtropical jet stream (~200 hPa).
• Descending Branch (The Subtropical ridge / Horse Latitudes): As the poleward-moving air radiates longwave infrared energy to space and cools while encountering convergence aloft near 25° to 35° latitude, it becomes negatively buoyant and sinks back toward the surface (large-scale atmospheric subsidence). As this dry upper-tropospheric air descends into higher barometric pressures, it undergoes adiabatic compression warming at the dry adiabatic lapse rate (~9.8 K/km). This warming suppresses cloud formation, lowers relative humidity, and spikes the atmospheric vapor pressure deficit (VPD), creating semi-permanent belts of high surface barometric pressure (the subtropical highs or horse latitudes) that dictate the geographic location of the world's major hot arid deserts (including the Sahara, Arabian, Kalahari, Australian, Sonoran, and Atacama deserts).

Situated in the mid-latitudes between ~30° and ~60° latitude, the Ferrel cell (first identified by American meteorologist William Ferrel in 1856) is a thermally indirect Eulerian-mean circulation: when wind velocities are averaged at fixed latitudes and pressure levels over time, cool air appears to rise near 60° latitude and warm air sinks near 30° latitude, with poleward flow near the surface and equatorward flow aloft.

Because warm air sinks and cold air rises in the Eulerian mean, the Ferrel cell consumes rather than generates kinetic energy. Crucially, in modern dynamical meteorology, the Ferrel cell is not driven like a mechanical gear by the rising and sinking branches of the adjacent Hadley and Polar cells. Instead, the mid-latitudes are dominated by a steep meridional temperature gradient (the baroclinic zone) that constantly spawns transient and stationary baroclinic eddies (synoptic extratropical cyclones, anticyclones, and meandering Rossby waves).

Applying Reynolds decomposition to partition any atmospheric variable A into its zonal-time mean (Ā) and fluctuating eddy departure (_A, such that A = Ā + _A), mid-latitude cyclones systematically transport warm air poleward and cold air equatorward (generating a strong positive poleward eddy heat flux, _v'&theta;̄ > 0) while simultaneously transporting eastward momentum poleward into the mid-latitudes (generating a strong eddy momentum flux convergence, -_∂_(_u'v̄)/_∂y_ > 0). In the zonal-mean momentum equation, this convergence of westerly eddy momentum aloft at ~45° to 55° latitude must be balanced by an equatorward Coriolis torque (f · v̄ < 0, requiring equatorward mean flow v̄ < 0 aloft) and a poleward frictional return flow in the surface boundary layer (v̄ > 0), statistically inducing the indirect Ferrel cell.

When mid-latitude circulation is transformed from fixed Eulerian coordinates into isentropic coordinates (averaging along surfaces of constant potential temperature &theta;) or analyzed via the Transformed Eulerian Mean (TEM) residual circulation (v̄^*, w̄^*, which incorporates the eddy Stokes drift):

v̄^* = v̄ - (1 / ρ₀) · ( _∂_( ρ₀ · _v'&theta;_'̄ / (_∂&theta;̄_/_∂z_) ) / ∂z )

the indirect Ferrel cell disappears entirely, revealing that net Lagrangian mass and parcel transport actually flows in a single continuous, thermally direct loop from the tropics to the poles in each hemisphere.

Operating poleward of ~60° latitude to the North and South Poles (90°), the polar cell is a weak, shallow thermally direct overturning circulation restricted to a lower tropopause height (~8 to 9 km):

• Polar Subsidence: Over the frigid, high-albedo polar ice sheets and sea-ice packs, persistent radiative heat loss to space cools the lower atmosphere, increasing air density and driving slow atmospheric subsidence over the poles. This subsidence creates domes of high surface pressure known as the polar highs (producing extremely cold, dry polar desert conditions over interior Antarctica and the high Arctic).
• Subpolar Ascent at the Polar Front: Cold, dense air flows equatorward out of the polar highs along the surface until it reaches ~60° latitude, where it collides with warmer, moister subtropical air flowing poleward within the surface branch of the Ferrel cell. The sharp density and temperature boundary separating these two air masses is the polar front. Forced to ascend along the sloping frontal surface in the subpolar low-pressure belt (such as the Icelandic Low, Aleutian Low, and the circum-Antarctic low-pressure trough), the rising air precipitates and diverges poleward near the polar tropopause to complete the loop.

Because meteorological observations are referenced to the solid rotating surface of the Earth (a non-inertial reference frame), Newton's second law of motion requires the inclusion of two apparent inertial forces: the centrifugal force (-Ω × (Ω × r), which points perpendicular to the Earth's rotational axis and is combined vectorially with true Newtonian gravitation g^* to define effective local gravity g, shaping the Earth into an oblate spheroid) and the Coriolis force (per unit mass):

F_Coriolis / m = -2Ω × V

For quasi-horizontal atmospheric motion (where vertical velocity w &ll; u, v), only the local vertical component of the planetary vorticity vector (f = 2_Ω_·sin_&phi;_, known as the Coriolis parameter or planetary vorticity) deflects horizontal winds, producing a horizontal Coriolis acceleration of magnitude f_ · V_H directed strictly perpendicular to the velocity vector (90° to the right of motion in the Northern Hemisphere where f > 0, and 90° to the left of motion in the Southern Hemisphere where f < 0; vanishing to zero directly at the equator where &phi; = 0° and reaching a maximum of f = ±1.458 × 10^-4 s^-1 at the poles). Because the Coriolis force acts at 90° to the velocity vector, it can only change the direction of the wind and can never perform mechanical work or alter kinetic wind speed (V · (-2Ω × V_) = 0).

Physically, the horizontal Coriolis force acts equally on both north-south (meridional) and east-west (zonal) air parcel motions through two complementary mechanisms:

1. Meridional (North-South) Motion (Conservation of Axial Angular Momentum): Because the Earth is a rotating sphere of radius a, the tangential eastward linear velocity of the solid surface (V_Earth(&phi;) = Ω · a · cos_&phi;_) is maximum at the equator (~465 m/s or ~1,674 km/h) and decreases with the cosine of latitude to 0 m/s at the poles. An air parcel moving poleward from low latitudes carries high initial eastward angular momentum into a latitude circle whose underlying ground is spinning more slowly, causing the parcel to outrun the ground toward the east (deflecting right in the NH, left in the SH). Conversely, a parcel moving equatorward from high latitudes enters a faster-spinning latitude belt and lags behind toward the west.
2. Zonal (East-West) Motion (Anomalous Centrifugal Deflection): An air parcel blowing toward the east (u > 0) rotates around the Earth's axis faster than the solid ground beneath it (Ω_total = Ω + _u_/(_a_·cos_&phi;_)). Consequently, the parcel experiences an excess outward centrifugal force (2_Ωu_) perpendicular to the Earth's axis that is not balanced by the oblate spheroid's gravitational slope; projecting this outward centrifugal excess onto the local horizontal tangent plane pulls the eastward-moving parcel equatorward (to the right/south in the NH; to the left/north in the SH). Conversely, a westward-blowing parcel (u < 0) rotates slower than the ground, experiencing a centrifugal deficit that pulls it poleward.

The relative importance of inertial advection versus planetary rotation in any atmospheric flow of characteristic velocity U and horizontal length scale L is quantified by a dimensionless ratio known as the Rossby number (Ro):

Ro = (Inertial Acceleration) / (Coriolis Acceleration) = ( U² / L ) / ( f · U ) = U / ( f · L )

For large-scale synoptic weather systems in the extratropical free troposphere (above the ~1 km frictional boundary layer), U &sim; 10 m/s, L &sim; 1,000 km (10⁶ m), and f &sim; 10^-4 s^-1, yielding Ro ≈ 0.1 &ll; 1. Because Ro is small and friction is negligible aloft, the horizontal momentum equations reduce to geostrophic balance, an equilibrium between the horizontal pressure-gradient force and the opposing Coriolis force:

f · v_g = (1 / ρ) · (∂P / ∂x)   and   f · u_g = -(1 / ρ) · (∂P / ∂y)   ↔   V_g = (1 / (ρf)) · (k × ∇_HP) = (g / f) · (k × ∇_PZ)

where Z is geopotential height on an isobaric surface. Rather than blowing directly from high pressure to low pressure, geostrophic winds blow strictly parallel to isobars (lines of constant pressure or geopotential height), circulating clockwise around high-pressure anticyclones and counterclockwise around low-pressure cyclones in the Northern Hemisphere (reversed in the Southern Hemisphere, as summarized by Buys Ballot's law).

Combining horizontal geostrophic balance with vertical hydrostatic equilibrium (_∂P_/_∂z_ = -_ρg_) and the ideal gas law yields the thermal wind relation, which governs the vertical shear of the geostrophic wind (_∂_V_g/_∂_ln_P_) as a function of the horizontal temperature gradient on an isobaric surface (∇_PT_v):

∂u_g / _∂_(ln P) = ( R_d / f ) · ( ∂T_v / ∂y )_P   and   ∂v_g / _∂_(ln P) = -( R_d / f ) · ( ∂T_v / ∂x )_P

Because temperature in the troposphere decreases poleward (∂T_v/_∂y_ < 0 in the NH) and pressure P decreases upward (_d_ln_P_ < 0), the thermal wind relation dictates that westerly (west-to-east) wind speed must increase with altitude throughout the troposphere, reaching a maximum right at the tropopause (~200 to 300 hPa, or 9 to 12 km altitude) above the sharpest horizontal temperature gradients (fronts) to form the planetary jet streams (specifically the subtropical jet stream near 30° latitude and the eddy-driven polar front jet stream near 50° to 60° latitude, where core velocities exceed 50 to 100 m/s).

Within the lowest 0.5 to 2.0 kilometers of the troposphere, known as the planetary boundary layer (PBL), aerodynamic surface drag and turbulent eddy viscosity (F_r) decelerate the wind below its geostrophic speed. Because the Coriolis force (fV) is directly proportional to wind speed V, slowing the wind weakens the Coriolis deflection while the horizontal pressure-gradient force remains unchanged. This three-way balance among pressure gradient, Coriolis force, and turbulent friction (the Ekman spiral) deflects surface winds across the isobars at an angle of ~10° to 20° over smooth oceans and ~25° to 45° over rough continents toward low pressure.

Combining the meridional pressure belts established by the three circulation cells with this boundary-layer Coriolis deflection generates Earth's three primary belts of prevailing winds (which are named by meteorological convention according to the compass direction *from* which they blow):

• 1. The Trade winds (Tropical Easterlies, 0° to ~30° N and S): Along the floor of the Hadley cell, surface air flows equatorward from the subtropical high-pressure ridges (~30°) toward the equatorial low-pressure trough (ITCZ, ~0°). Deflected westward by the Coriolis force (to the right in the Northern Hemisphere and to the left in the Southern Hemisphere), these steady, persistent lower-tropospheric winds blow from the northeast in the Northern Hemisphere (Northeast Trades) and from the southeast in the Southern Hemisphere (Southeast Trades). The trade winds drive the westward equatorial ocean currents, maintain the Walker circulation, and steer tropical waves and tropical cyclones.
• 2. The Prevailing Westerlies (~30° to ~60° N and S): Along the surface branch of the mid-latitude Ferrel belt, air flows poleward from the subtropical highs (~30°) toward the subpolar low-pressure belt (~60°). Deflected eastward by the Coriolis force, these winds blow predominantly from the southwest in the Northern Hemisphere and from the northwest in the Southern Hemisphere, steering synoptic extratropical cyclones, warm and cold fronts, and atmospheric rivers from west to east across North America, Eurasia, and the Southern Ocean. Because the Southern Hemisphere between 40°S and 60°S is nearly devoid of continental landmass friction, the Southern Hemisphere westerlies achieve extreme sustained velocities (the Roaring Forties, Furious Fifties, and Screaming Sixties) that drive the Antarctic Circumpolar Current.
• 3. The Polar easterlies (~60° to 90° N and S): Within the surface branch of the Polar cell, cold, dense air flows equatorward out of the polar highs (90°) toward the subpolar lows (~60°). Strong polar Coriolis deflection turns this outflow westward, generating frigid northeasterly winds in the Arctic and southeasterly winds off Antarctica (reinforced over sloped Antarctic and Greenland ice-sheet topography by gravity-driven katabatic winds) that converge with the warmer mid-latitude westerlies along the baroclinic polar front.