Historical analogues of the recent extreme minima observed in the Atlantic meridional overturning circulation at 26°N

Abstract

Observations of the Atlantic meridional overturning circulation (AMOC) by the RAPID 26°N array show a pronounced minimum in the northward transport over the winter of 2009/10, substantially lower than any observed since the initial deployment in April 2004. It was followed by a second minimum in the winter of 2010/2011. We demonstrate that ocean models forced with observed surface fluxes reproduce the observed minima. Examining output from five ocean model simulations we identify several historical events which exhibit similar characteristics to those observed in the winter of 2009/10, including instances of individual events, and two clear examples of pairs of events which happened in consecutive years, one in 1969/70 and another in 1978/79. In all cases the absolute minimum, associated with a short, sharp reduction in the Ekman component, occurs in winter. AMOC anomalies are coherent between the Equator and 50°N and in some cases propagation attributable to the poleward movement of the anomaly in the wind field is observed. We also observe a low frequency (decadal) mode of variability in the anomalies, associated with the North Atlantic Oscillation (NAO). Where pairs of events have occurred in consecutive years we find that atmospheric conditions during the first winter correspond to a strongly negative Arctic Oscillation (AO) index. Atmospheric conditions during the second winter are indicative of a more regional negative NAO phase, and we suggest that this persistence is linked to re-emergence of sea surface temperature anomalies in the North Atlantic for the events of 1969/70 and 2009/10. The events of 1978/79 do not exhibit re-emergence, indicating that the atmospheric memory for this pair of events originates elsewhere. Observation of AO patterns associated with cold winters over northwest Europe may be indicative for the occurrence of a second extreme winter over northwest Europe.

Appendix: Decomposition of the AMOC

Ocean models typically output the northward velocity for each grid box, which allows us to exactly compute the AMOC ($\varPsi$) in the model, i.e.

$$\begin{aligned} \varPsi (y,z) = \int _{z}^{0} \int _{x_{w}}^{x_{e}} v(x,y,z') \quad dx dz'. \end{aligned}$$

(1)

In order to make comparisons with the observations taken at 26°N we may also compute components of the transport corresponding to those measured by the 26°N observing array, namely the Florida Straits transport, $\varPsi _{FST}$, the geostrophic (or thermal wind) transport, $\varPsi _{geo}$ and the Ekman transport, $\varPsi _{ekm}$.

$\varPsi _{FST}$ is computed by integrating the meridional velocity, $v$, through the Florida Straits (between Florida, $x_{w}$, and the Bahamas, $x_{Bh}$), and from the maximum depth of the Florida Straits $H_F$ to the surface,

$$\begin{aligned} \varPsi _{FST} = \int _{H_F}^{0} \int _{x_{w}}^{x_{Bh}} v(x,z') \quad dx dz'. \end{aligned}$$

(2)

$z'$ is a dummy integration variable. $\varPsi _{geo}$, the baroclinic geostrophic component arising from zonal density gradients across the Atlantic basin is

$$\begin{aligned} \varPsi _{geo}(z) = \int _{z}^{0} \int _{x_{Bh}}^{x_e} (v_{geo} - \bar{v}_{comp}) \quad dxdz', \end{aligned}$$

(3)

where $v_{geo}$ and $\bar{v}_{comp}$ are

$$\begin{aligned} v_{geo}(x,z) = -\frac{g}{\rho ^{*} f} \int _{-H(x)}^{z} \frac{\partial \rho }{\partial x} \quad dz' \end{aligned}$$

(4)

and

$$\begin{aligned} \bar{v}_{comp}(x,z) = \frac{1}{H(x)} \int _{-H(x)}^{0} v_{geo}(x,z') \quad dz' + \bar{v}_{FST} \end{aligned}$$

(5)

respectively, $x_{e}$ is the easternmost extent of the Atlantic (i.e. Africa), $H(x)$ is the maximum depth of the basin as a function of longitude, $g$ being the Earth’s gravitational acceleration, $\rho$ the in-situ density, $f$ the Coriolis parameter, and $\rho ^{*}$ a reference density. $\bar{v}_{FST}$ is

$$\begin{aligned} \bar{v}_{FST}= \frac{\varPsi _{FST}}{A} \end{aligned}$$

(6)

with $A$ being the cross-sectional area of the Atlantic basin east of the Bahamas.

We define $\varPsi _{ekm}$, the Ekman (wind driven) component, here as a function of latitude and depth compensated by a section mean return flow to ensure no net transport,

$$\begin{aligned} \varPsi _{ekm}(y,z) = \int _{-H_{max}(y)}^{z} \int _{x_w}^{x_e} (v_{ekm} - \bar{v}_{ekm}) \quad dxdz', \end{aligned}$$

(7)

where $v_{ekm}$ and $\bar{v}_{ekm}$ are

$$\begin{aligned} v_{ekm} = -\frac{1}{(\rho ^{*} f L \varDelta _{z})} \int _{x_w}^{x_e} \tau _{x} dx \end{aligned}$$

(8)

and

$$\begin{aligned} \bar{v}_{ekm} = -\frac{1}{(\rho ^{*} f A)} \int _{x_w}^{x_e} \tau _{x} dx \end{aligned}$$

(9)

respectively, $L$ being the basin width, $\varDelta _{z}$ the Ekman depth, and $H_{max}(y)$ the latitudinally dependent maximum depth of the basin. The Ekman depth, $\varDelta _{z}$, which defines the base of the Ekman layer in which the wind driven transport occurs we chose to be 100 m. The choice of $\varDelta _{z}$ does not strongly affect the resulting overturning profile. Note that the compensation term associated with the Ekman transport could equally be added to $\bar{v}_{comp}$, but that it will be small compared with the other terms.

Therefore at $26.5^{\circ}\hbox {N}$ the AMOC transport can be considered as the sum of these components plus a residual term, $\varPsi _{res}$,

$$\begin{aligned} \varPsi = \varPsi _{FST} + \varPsi _{geo} + \varPsi _{ekm} + \varPsi _{res}. \end{aligned}$$

(10)

where $\varPsi _{res}$ can be obtained by rearranging Eq. (10). If averaged over a time longer than a few cycles of the local inertial period the residual term is small (order 1 Sv), and can be ignored. It should be noted, however, that this term can dominate the AMOC variability at near-inertial time scales (Blaker et al. 2012; Sevellec et al. 2013).