The Atlantic branch of the ocean's thermohaline circulation transports heat northward and ventilates the deep ocean, and coupled-model experiments synthesized by Manabe and Stouffer in 1999 showed that it weakens or collapses when the northern North Atlantic is freshened or when atmospheric CO2 rises, that their model possesses both an active and a weak inactive overturning equilibrium, and that recovery from strong forcing can take a millennium or more. The review left open whether this behaviour is robust and applicable to the real climate system, identifying six coupled-model uncertainties: the use of modern rather than glacial or deglacial boundary conditions; the poorly constrained magnitude of effective vertical diapycnal mixing, including spurious numerical mixing; coarse ocean resolution and idealized bottom topography; unknown sensitivity of the response to the rate of CO2 increase; unverified mechanisms and timescales of multi-century recovery; and the possible roles of flux adjustments and initialization. A quarter century of later work has retired part of the model-fidelity framing - freshwater-hysteresis experiments have since been run without flux adjustment at eddy-permitting resolution, glacial and deglacial hosing experiments exist, and diapycnal mixing has been mapped from observations - but the central questions remain open and actively researched: whether the real Atlantic overturning has multiple stable equilibria and where its tipping thresholds lie; why comprehensive current models diverge irreconcilably between recovery and persistent weakening under identical freshwater forcing, for reasons not explained by resolution or parameterization; how the transient weakening depends on the rate and path of the forcing; and which mechanisms set the multi-century-to-millennial recovery timescale.

Background

The circulation and its diagnostic. The thermohaline circulation (THC) is the part of the ocean circulation maintained by density gradients - set jointly by temperature and salinity - rather than directly by wind stress. In the Atlantic its dominant expression is the Atlantic meridional overturning circulation (AMOC): warm, saline upper-ocean water flows northward, releases heat to the atmosphere, becomes dense, and returns southward at depth as North Atlantic Deep Water. This overturning carries of order 101510^{15} W of heat northward and ventilates the abyssal ocean, so its strength matters for European and North Atlantic climate, for ocean heat and carbon uptake, and for global sea level. Its strength is conventionally measured as the maximum of the Atlantic meridional overturning streamfunction (the northward volume transport, zonally integrated across the basin, per unit depth, as a function of latitude and depth), quoted in sverdrups (1 Sv=106 m3 s−11\ \mathrm{Sv}=10^{6}\ \mathrm{m^{3}\ s^{-1}}). Deep sinking occurs through wintertime convective overturn in the Labrador and Nordic seas; a stable low-salinity surface layer (a halocline) suppresses that convection and weakens the overturning.

Multiple equilibria. Box-model theory (Stommel 1961) showed that a two-basin loop forced by surface heating, cooling, and freshwater exchange can have two stable regimes: in one the flow is strong and salinity differences are large, in the other the flow is weak; the mechanism is the salt-advection feedback, whereby a stronger current transports more salt into the sinking region and reinforces itself, while temperature differences are damped by heat loss to the atmosphere. A central question since has been whether the coupled ocean-atmosphere system - and the real ocean - likewise admits several stable overturning equilibria, in particular a weak or reverse "inactive" mode alongside the observed active mode, and how far the present climate sits from the thresholds between them. Hysteresis - the failure to retrace the same circulation strength when the forcing is reversed - is the standard model diagnostic of such multiple equilibria.

The 1999 synthesis. By the 1990s, coupled ocean-atmosphere general circulation models (GCMs, which solve the three-dimensional fluid equations of atmosphere and ocean on a rotating sphere) could sustain multicentennial integrations. To control climate drift in such models at then-standard resolution, the Geophysical Fluid Dynamics Laboratory coupled model used by Manabe, Stouffer, Spelman and Bryan (1991) employed flux adjustments: fixed, geographically and seasonally varying corrections to the ocean's surface heat and freshwater fluxes. With this model Manabe and Stouffer (1994) ran transient experiments in which atmospheric CO2 increases at a compounded 1% per year until doubling or quadrupling and is then held fixed; Delworth, Manabe and Stouffer (1993) documented intrinsic multidecadal overturning variability in the unforced 1000-year control. Manabe and Stouffer (1999, Tellus B), the source of the present problem, synthesized the full experiment suite: (i) discharging 0.1 Sv0.1\ \mathrm{Sv} of freshwater ("hosing") uniformly over the 50∘50^\circN-70∘70^\circN North Atlantic for 500 years from the modern-like control state creates a halocline, suppresses convective sinking, weakens and shoals the overturning, and produces large regional cooling with a maximum south of Greenland, remote Southern Ocean cooling, and little global-mean change; (ii) the model has two stable overturning equilibria, and the persistence of the inactive mode depends on the effective vertical diapycnal diffusion coefficient (the mixing of heat and salt across density surfaces), including the spurious mixing contributed by the numerical scheme that advects tracers, compared against open-ocean tracer-release estimates - experiments in which a passive tracer is injected and its spread tracked - of order 0.10.1-0.2 cm2 s−10.2\ \mathrm{cm^{2}\ s^{-1}} (Ledwell, Watson and Law 1993); (iii) under the 1% per year CO2 ramps, enhanced poleward water-vapour transport freshens the high latitudes (an additional poleward freshwater supply of roughly 0.28 Sv0.28\ \mathrm{Sv} north of 50∘50^\circN over 500 years in the quadrupling run): in the doubling experiment the overturning weakens by more than a factor of two and recovers within about five centuries, whereas in the quadrupling experiment it nearly vanishes for centuries and regains its original intensity only after roughly 1600 years, the recovery being attributed to slow warming of the deep water; and (iv) the unforced control exhibits intrinsic multidecadal variability with coherent surface-temperature and salinity signals propagating along the subpolar gyre (the counter-clockwise surface circulation of the northern North Atlantic). That the collapse threshold could depend on the rate of CO2 increase had been anticipated in an intermediate-complexity model (a simplified global model with a coarse ocean) by Stocker and Schmittner (1997).

The unresolved agenda. The review closed by asking whether this simulated behaviour is robust and applicable to the real climate system, listing the open uncertainties: the freshwater experiments started from a simulated modern state rather than glacial or deglacial boundary conditions (ice-sheet geometry, sea-ice cover, sea level); the stability and recovery behaviour depends on the magnitude of effective vertical mixing, explicit and numerical, which was poorly constrained; coarse resolution and idealized bottom topography could distort the timing and magnitude of the response; the magnitude of weakening depends on the rate of CO2 increase, with the quantitative sensitivity undetermined; the slow-recovery mechanisms were postulated but not verified; and the results depended on flux adjustments and on a particular initialization, whose roles were unknown. These items, individually and jointly, define the problem stated below.

Problem Statement

Determine whether the thermohaline-circulation behaviour simulated by Manabe and Stouffer (1999) - freshwater- and CO2-forced weakening of the Atlantic overturning, possible transition to a weak inactive mode, and multi-century recovery - is robust and applicable to the real climate system. Concretely, determine:

  1. Multiple equilibria and thresholds. Whether the coupled ocean-atmosphere system, and the real Atlantic, possess more than one stable overturning equilibrium (an active mode with northern deep-water formation and a weak or reverse inactive mode); where the critical freshwater-input and CO2-concentration thresholds between these regimes lie; and which properties of the effective diapycnal mixing, model resolution and topography, advection numerics, initialization, and the presence or absence of flux adjustment control the existence and width of the bistable regime.

  2. Rate and path dependence. How the magnitude and timescale of overturning weakening depend on the rate at which atmospheric CO2 - and the associated intensification of the hydrological cycle and of high-latitude freshwater input - increases, across the range of plausible emission scenarios, including whether faster forcing produces more severe weakening and whether a forcing trajectory can cross thresholds that a slower approach to the same final forcing would not.

  3. Recovery. By which processes - slow warming of deep and bottom water, changes in the density stratification of the water column, high-latitude freshening, re-establishment of convection - and on what timescales the overturning recovers after strong transient forcing, including the conditions under which recovery is delayed or does not occur.

  4. Boundary-condition dependence. How glacial or deglacial background states (ice-sheet geometry, sea-ice cover, sea level, closed or open Bering Strait) modify the overturning's response to freshwater input and the spatial pattern of the associated surface cooling.

These components jointly support one objective - establishing the robustness, timescales, and real-world applicability of the simulated overturning responses - and are mutually dependent: the meaning of the transient weakening (2) and of recovery (3) is fixed by the stability landscape (1), whose determination must itself be shown robust to model fidelity and boundary conditions (1, 4). The four components are: the number and location of the stable overturning equilibria and the thresholds between them and the factors that control the existence and width of the bistable regime (1); how the magnitude and timescale of weakening depend on the rate and path of the forcing (2); the processes and timescales of recovery after strong transient forcing (3); and how glacial or deglacial background states modify the response (4).

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