OSOcean ScienceOSOcean Sci.1812-0792Copernicus PublicationsGöttingen, Germany10.5194/os-14-633-2018 On the role of the North Equatorial Counter Current during a strong El NiñoThe role of the NECC in a strong El NiñoWebbDavid Johndjw@noc.ac.ukhttps://orcid.org/0000-0001-7084-8566National Oceanography Centre, Southampton, SO14 3ZH, UKDavid John Webb (djw@noc.ac.uk)11July20181446336604December201715December20172June201822June2018This work is licensed under the Creative Commons Attribution 4.0 International License. To view a copy of this licence, visit https://creativecommons.org/licenses/by/4.0/This article is available from https://os.copernicus.org/articles/14/633/2018/os-14-633-2018.htmlThe full text article is available as a PDF file from https://os.copernicus.org/articles/14/633/2018/os-14-633-2018.pdf
An analysis of archived data from the NEMO 1/12th degree global ocean model
shows the importance of the North Equatorial Counter Current (NECC) in the
development of the strong 1982–1983 and 1997–1998 El Niños. The model
results indicate that in a normal year the core of warm water in the NECC is
diluted by the surface Ekman transport, by geostrophic inflow and by tropical
instability waves. During the development of the 1982–1983 and 1997–1998 El Niños,
these processes had reduced effect at the longitudes of warmest
equatorial temperatures and to the west. During the autumns of 1982 and 1997,
the speed of the NECC was also increased by a stronger-than-normal annual
Rossby wave. The increased transport of warm water by the NECC due to these
changes resulted in warm water reaching the far eastern Pacific and appears
to have been a major factor in moving the centre of deep atmospheric
convection eastwards across the Pacific.
The works published in this journal are distributed under
the Creative Commons Attribution 4.0 License. This licence does not affect
the Crown copyright work, which is reusable under the Open Government
Licence (OGL). The Creative Commons Attribution 4.0 License and the OGL are
interoperable and do not conflict with, reduce or limit each other.
Introduction
Studies of the tropical Pacific often focus on the equatorial waveguide and
the propagation of equatorial Kelvin waves generated by westerly wind events
i.e.. The
study reported here starts in a similar manner, focusing on the waveguide. It
uses data from a long run of the NEMO 1/12th degree computer model of the
global ocean and starts by calculating the average sea surface temperatures
in the equatorial band as a function of longitude and time.
During the strong El Niño events of 1982–1983 and 1997–1998, the results
show warm water propagating eastwards from the warm pool region of the west
Pacific across to the South American coastline. A different type of event,
the warm pool El Niños or oscillations , is seen
in other years but these are limited to the western and central Pacific.
The strong El Niño events propagate eastwards at a speed of about
0.6 m s-1. The equatorial Pacific is highly stratified, with the
warmest water concentrated in the top 200 m, so a speed of 0.6 m s-1
is comparable with the speed of a number of equatorial Kelvin
wave modes whose first zero occurs similarly at around 200 m. There is
therefore some justification in connecting the propagation of the warm
features with the propagation of equatorial Kelvin waves.
Except that this is highly unlikely.
Simple waves, like equatorial Kelvin waves, transport momentum and energy but
they cannot easily transport quantities like temperature and salinity,
qualities associated with individual particles in the medium. Such advection
can only occur if the waves are highly non-linear so that particle velocities
are comparable with the phase velocity. This occurs in breaking waves and, to
a lesser extent, in tidal bores but, as far as the author is aware, no one
has reported evidence of an equivalent major feature in the near-surface
layers of the equatorial Pacific.
In order to clarify the situation, model archived data
are used to calculate the flux of warm water across 180 and
240∘ E as a function of time during the period 1980 to 1985. This
period includes the strong 1982–1983 El Niño.
The study by showed that sea surface temperatures (SSTs)
greater than 28 ∘C are required for the onset of widespread
deep convection over the tropical ocean. They also showed that at times
temperatures of over 29.5 ∘C may be required. For this reason, the
study concentrates on water temperatures that exceed 28 ∘C.
The analysis shows that during the 1982–1983 El Niño, the main flux of
warm water in the model did not occur within the equatorial waveguide.
Instead, it occurs further north, at the latitude of the eastward-flowing
North Equatorial Counter Current (NECC).
was possibly the first to suggest that the
NECC had the potential to transport significant amounts of heat eastwards in
the tropical Pacific. The NECC continues to be important in his later papers
i.e., but in his theory of the El Niño
he also introduced the idea of equatorial Kelvin waves
triggering the El Niño. It is this aspect of his work that has been
developed most by later authors.
To return to the NECC, the model results studied here indicate that in most
years the Ekman transport, the geostrophic inflow and tropical instability
waves carry warm water away from the core of the NECC and replace it with
cooler water from the north and south. As a result, the core temperature of
the NECC is significantly reduced.
During periods when an El Niño is developing, the trade winds retreat
eastwards and they are replaced by a region of low or westerly winds. The
model shows that one result of the low winds is that at the longitudes
affected, the Ekman transport at the latitude of the NECC is reduced. The
strength of the geostrophic inflow is also reduced as is the strength of the
tropical instability waves. The latter is probably in part due to the
reduction and change in direction of the surface current at the Equator.
As a consequence, while the El Niño develops, the NECC transports much
warmer water than normal past the region of low winds. This transport of warm
water occurs near the latitudes of the subtropical convergence in the
atmosphere. Thus, although the present study does not include an atmospheric
model, it is likely that this is one, or possibly the main, factor moving the
region of deep atmospheric convection and low winds further east. The process
is then repeated, moving the convection region, the region of low winds and
warmer-than-normal water, steadily eastwards across the ocean.
Further support for this argument is obtained by tracing particles during a
strong El Niño and a non-El Niño year. The results from the model
show that during the non-El Niño year, water particles are rapidly mixed
out of the NECC but during the strong El Niño year they stay within the
NECC and are transported further east.
The analysis also shows that the NECC is affected by annual Rossby waves
which propagate westwards across the equatorial Pacific. These increase the
speed of the NECC at all longitudes, but, in particular, it is found that the
wave at 6∘ N, arrives in the western Pacific in the middle of the year, the time
that the classic strong oceanographic El Niños usually start. In 1982, the
amplitude of the wave in the western Pacific was greater than normal, and this
may have started the strong 1982–1983 El Niño.
The structure of the paper is as follows. Section 2 describes the underlying
numerical model, and Sect. 3 uses the results to plot the time series of
average sea surface temperatures in the equatorial band during the period 1980
to 2000. Such a time series clearly shows the strong El Niño events
of 1982–1983 and 1997–1998 as well as the weaker warm pool events in
intermediate years.
Section 4 focuses on the eastward advection of warm water to determine when
and where this occurs during the period 1980 to 1985. The NECC is found to be
primarily responsible but there is a large year-to-year variability. Section 5
therefore then starts by examining the effects of Ekman transport, the
geostrophic inflow and tropical instability waves on the transport of
warm water by NECC.
The section also investigates the varying strength of the NECC itself and
this is developed further in Sect. 6, which follows up Wyrtki's idea
that El Niños are connected to the difference in sea level across the NECC.
Up to this point, the analysis makes extensive use of Hovmöller diagrams, but
to give a more geographical overview of events, Sect. 7 relates the results
to plots of sea surface temperature, elevation, currents and the wind
stress vectors, in the northern spring, summer, autumn and winter of 1982.
Section 8 then investigates the mixing processes using particle tracks
started in the central Pacific in the autumns of 1981 and 1982.
The final analysis section briefly reports on a similar analysis of the
strong 1997–1998 El Niño. Although this El Niño starts differently,
the role of the NECC, the annual Rossby wave and mixing processes is found to
be similar to the 1982–1983 period. The paper closes with a review of the
main results of the study.
The NEMO 1/12∘ global ocean model
The ocean model discussed here is one of the family of NEMO models
, all with a similar code base but with different choices
of horizontal and vertical grid resolution and of the many options for
representing the underlying ocean physics. The present model uses a
non-uniform grid based on a longitude grid spacing of 1/12∘ along
the Equator. In the Southern Hemisphere and in the Indian and Pacific oceans,
the latitude spacing is chosen so that each of the grid boxes has the same
width and height. In the North Atlantic and Arctic, a more complex scheme is
used to prevent the convergence of the grid near the pole.
Average model sea surface temperature (∘C) during 1981,
showing the averaging region used for Figs. and .
The model has 75 layers in the vertical. Their nominal thicknesses range from
1 m at the surface to 204 m in the lowest layer, but as the
ocean surface moves up and down each of the layers expands or contracts
slightly to allow for this. The nominal thicknesses are based on an analytic
formula which ensures a smooth transition between the strongly stratified
surface layers, which need to be well resolved, and the weakly stratified
deep ocean for which less resolution is necessary. One consequence of this is
that 35 layers are used to resolve just the top 300 m. In the
equatorial Pacific, this covers most of the major current systems.
This high-resolution version of the NEMO ocean model was developed for
coupling with a similar high-resolution version of the UK Meteorological
Office atmospheric model, the aim being to create an improved coupled model
for both weather prediction and climate change research. However, before
coupling the two models, a number of test runs were carried out using just
the ocean model. The data analysed here come from run 6, the last and longest
of these tests.
The surface boundary conditions used for run 6 are those of
together with the Drakker DFS5.2 atmospheric fields
described by . The Drakkar datasets, like the ECMWF
reanalysis datasets on which they are based, start from 1958. This is also
the start date of run 6.
In a previous analysis of run 6, compared observed
temperatures in the equatorial Pacific Niño regions with those from the
model. The Niño regions are a series of standard ocean areas in the
equatorial Pacific often used in El Niño studies .
This analysis showed that there was good agreement between observations and
the model. It also showed that this was not due to the existence of a strong
feedback loop – the actual sea surface temperature somehow controlling the
model SST via its effect of the atmospheric layers
closest to the ocean surface.
On the basis of this analysis, and the additional confidence in the model
code which comes from its successful use over many years for oceanic and
climate research, it appears reasonable to make further use of the model
archive data in the present study of processes affecting the El Niño.
Data from the model are available in the form of averages over each 5-day
period during the model run. The analysis reported here concentrates on 1981,
as a typical non-El Niño year, and on the 1982–1983 and
1997–1998 periods, during which strong El Niños developed. The 5-day
averages miss the diurnal variations in sea surface properties which may help
to trigger convection. Observations from the west Pacific
Fig. 3 show temperature variations of 1–2∘
during calm conditions, but this resulted in only a few convective clouds.
During windier, cloudier conditions, the diurnal variation was usually less
than a degree.
As background, Fig. shows the model average SST during 1981. In
the west, the North Pacific warm pool shows a large region with average
temperatures above 28 ∘C. Similar temperatures are also found over a
large region of the South Pacific. The figure also shows the region of the
equatorial waveguide, extending from 5∘ S to 5∘ N, which is
analysed in the next section of this paper. In the west, this has average
temperatures above 28 ∘C, but in the east, where the Equatorial
Undercurrent outcrops, temperatures fall below 23 ∘C.
Figure shows the average values of temperature, salinity and
velocity near the surface in a section at 200∘ E (160∘ W).
The figure illustrates the strong stratification of the surface layers and
the fact that the primary currents of the region, the Equatorial Current, the
Equatorial Undercurrent and the North Equatorial Counter Current all lie
close to the ocean surface.
North–south sections at 200∘ E (160∘ W) of the
average values during 1981 of (a) temperature (∘C),
(b) salinity, and (c) east and (d) north
components of velocity (m s-1).
Eastward component of wind stress (Pa) at the Equator between
140 and 280∘ E (80∘ W).
The archive datasets also contain the ocean surface wind stress and
precipitation fields used to force the model. The wind stress field
(Fig. ) shows that in the central Pacific during the second half
of 1992, the normal trade winds, which extend to the Equator, retreated
eastwards. They were replaced by a region of low winds and, near the
dateline, periods of westerly winds. On inspection, the westerlies were often
associated with cyclonic flows, probably due to convection, over warm water
to the north and south of the Equator. The low wind stress region can also be
seen in Figs. to from 1982 and
Figs. to from 1997.
Figure also shows that, in the west Pacific, the winds are
often light or westerly and that the strongest westerly winds often occur on
either side of the new year. Plotted on a geographical grid, the wind stress
vectors show that this is often due to winds converging on the South Pacific
Convergence Zone. An example can be seen in Fig. .
The eastern limit of the low wind region along the Equator appears to be
associated with the eastern limit of the main deep convection region in the
atmosphere and the associated region of high precipitation rates. However,
the forcing data show that often there is limited precipitation near the
Equator even when the wind stress is low.
However, under these conditions, there are usually bands of strong wind stress
convergence to the north and south of the Equator close to areas with high
rates of precipitation. Figure allows for this wider
precipitation region by plotting the average precipitation rate between
12∘ S and 12∘ N, as a function of longitude and time. The
figure shows that during the second half of 1982 precipitation moved from the
western to the central Pacific. The central Pacific values then declined
early in 1983 after which there was a short period of increased precipitation
near the eastern boundary. The main precipitation then moved back to the
western Pacific.
In the South Pacific, precipitation may be enhanced by the South Pacific warm
pool, especially late in the year after the Sun has crossed the Equator. When
precipitation increases in the eastern Pacific, it often occurs along the line
of the Intertropical Convergence Zone (ITCZ).
Precipitation (kg m-2 day-1) averaged between
12∘ S and 12∘ N.
Time series of mean temperatures in the equatorial band
In the standard picture of an El Niño, warm water from the western
Pacific moves to the central equatorial Pacific and, in the more extreme
cases, all the way to the South American coastline. This warming of the
central and eastern Pacific moves the location of the dominant atmospheric
convection region eastwards and, because of the amount of heat energy
involved, results in large-scale changes in the atmospheric circulation.
Figure plots the sea surface temperature, averaged between
5∘ S and 5∘ N, as a function of longitude and for the
period 1980 to 2000. During this entire period, temperatures in the
western equatorial Pacific hover around 30 ∘C but in the eastern
Pacific the mean temperature can vary between 22 and 30 ∘C.
Average model sea surface temperature (∘C) in the equatorial
Pacific between 5∘ S and 5∘ N during the period 1980
through 1999.
The longest periods of warm temperatures in the eastern Pacific occur during
the strong El Niños of 1982–1983 and 1997–1998 when the area is affected by
features that have propagated in from the western Pacific. These El Niños
are also abrupt: the warm temperature fronts extend rapidly all across the
equatorial ocean and then equally rapidly retreat to the far west.
Weaker warm pool events are observed in 1987 and
in 1991–1992 but these mainly affect surface temperatures in the central and
western Pacific. They are also much more incremental, with the temperature front
only moving gradually east during the periods 1985 to 1987 and 1989 to 1992.
Following 1992, the front only retreats to near its 1991 position.
Expanded view of model sea surface temperature, averaged between
5∘ S and 5∘ N, for the equatorial Pacific between 120 and
280∘ E (80∘ W), and (a) between 1980 and 1985,
(b) between 1995 and 2000, and (c) when coupled to a high-resolution atmospheric model.
The period between 1980 and 1985 is expanded in Fig. , together
with a similar figure for the period including the 1997–1998 El Niño,
emphasising the similarities between the two strong El Niños. The third
figure comes from a similar period from a fully coupled run, where the same
ocean model was coupled to a high-resolution atmospheric model. In this case,
the ocean temperatures are slightly too low, but it shows that the detailed
structures seen in the figures are not the result of forcing by, possibly
anomalous, surface boundary conditions but arise as natural variabilities of
the coupled system.
As well as the main El Niño event, each of the figures shows a series of
fine-scale wave-like features, with an east–west wavelength of 5–10∘
and a period of about a month. The features grow in amplitude during
the autumn of each year and die out in the spring when their westward phase
velocity tends to be reduced or even reversed. Their wavelength and period
are consistent with them being due to the passage of tropical instability waves (TIWs).
As each TIW passes, it advects warm water into the equatorial band, but
after it has passed upwelling at the Equator will return the sea surface
temperature to its earlier value.
The second feature to note is the annual signal which primarily affects the
eastern and central Pacific. Near the South American coastline, temperatures
are at a maximum early each year, a comparison with the wind field
(i.e. Fig. ) indicating that this occurs after periods when the
eastward component of the wind stress has dropped to near zero. Thus, they are
probably partly due to a reduction in equatorial upwelling. However, this is
also the time when the area of warm water off Central America has moved
furthest south. An example is shown in Fig. .
Further west, the period of low winds occurs later, and this may explain why
the temperature maximum in the annual signal occurs later towards the central
Pacific. Alternatively, it may be due to an annual wave triggered by the
changes in the east.
Each of these features warrants further study, but in the rest of this paper
the focus is on the strong El Niño events when temperatures of
29 ∘C and more are found all across the ocean. In a normal year, as
illustrated in Figs. and , the average temperatures
in the equatorial waveguide, between 220 and 240∘ E,
lie between 24 and 27.5 ∘C. As shown in Fig. , warmer
temperatures are found at 8∘ N, but on average
these lie below 28 ∘C, and there is not enough heat available
locally to explain the warming of the whole of the eastern Pacific.
Vertically integrated flux of water (m2 s-1) with
temperature (a) greater than 28 ∘C crossing longitude
180∘ E, (b) greater than 29 ∘C crossing
180∘ E, (c) greater than 28 ∘C crossing
210∘ E and (d) greater than 28 ∘C crossing
240∘ E. The figure is blank where the flux is zero. Rectangles
enclose the regions of Table .
An increase in the local surface heat flux might produce the observed
increases in model SST. However, see Figs. 11 and 13
investigated these fluxes in two of the key Niño regions and found that
in each case the net heating was less during the development of the
1982–1983 El Niño than in the two previous years. Significant errors in
the model's vertical mixing scheme could also be responsible, but this is
unlikely. Instead, it is much more probable that the increases in both the
model SST and observations are due to advection of heat by the ocean.
Zonal advection of warm water near the Equator
In an attempt to clarify how heat was advected in the equatorial Pacific, the
integrated flux of water across a series of longitudes was calculated, with
the constraint that the temperature had to be greater than a given minimum
value. Figure shows the results plotted as a time series at
longitudes 180, 210 and 240∘ E, for a
minimum temperature of 28 ∘C and, at longitude 180∘ E,
for a minimum temperature of 29 ∘C. The figure also includes a
series of boxes, covering time periods and latitude ranges of interest, the
total flux in each period being summarised in Table .
The figure shows that most of the eastward flow of warm water occurs at the
latitudes of the North Equatorial Counter Current. The fluxes are largest in
the autumn of each year, the temperatures in the spring occasionally being
below the minimum temperature. The largest transport occurs during 1982, in
the period when the 1982–1983 El Niño is developing. Thus, Fig. 5a and b
show stronger-than-normal flows of > 28 ∘C water at
180∘ E over most of 1982 and of > 29 ∘C in late summer. At
210 and 240∘ E, the flow starts later in 1982 and continues until the year's end.
In the equatorial band, there are long periods when the water at all depths
is too cool to contribute to the flux calculations. Longitude 180∘ E
is an exception for a minimum temperature of 28 ∘C, but the flow is
predominantly westwards, as is expected for the latitudes of the Equatorial
Current. At 180 and 210∘ E, large eastward fluxes are observed
during 1982, a small event occurring in late summer and a major event just
before the end of the year. The major event is also seen at 210∘ E
but is missing at 240∘ E. However, it does show up when the minimum
temperature is reduced to 26 ∘C.
South of the Equator, an eastward transport of warm water is also seen at the
latitudes of the South Equatorial Counter Current. This is unexpected as,
east of 180∘ E, the westward-flowing South Equatorial Current is
usually thought to be contiguous with the westward-flowing Equatorial
Current. However, the flow is weak and reversing, and does not appear to be
connected with the 1982–1983 El Niño, so it is not considered further
here.
Volume transport of water for the longitudes, temperature classes,
latitude bands and time periods denoted in Fig. . The net flux is
given both in units of 1012 m3 and in terms of the number of
degrees longitude that would be covered by the same volume if it was in a
layer 100 m thick which extended from 5∘ S to 5∘ N.
Table shows that, at 210∘ E, the NECC transports a
total of 310 × 1012 m3 of water warmer than
28 ∘C between the spring and end of 1982. To give an idea of the potential impact
of this volume of water, the table also shows the corresponding span of
longitude that it would cover if it was contained in a surface layer 100 m
thick extending from 5∘ S to 5∘ N.
The table also includes the contribution of the equatorial band for the same
longitude and over roughly the same time period. Although the El Niño is
often described as resulting from increased eastward heat advection in the
equatorial waveguide, the contribution of the NECC is seen to be roughly
4 times larger than the contribution from currents close to the Equator. This
is also true for water warmer than 28 ∘C at 180∘ E.
In each of the three longitudes shown, the flux of water warmer than
28 ∘C is not enough for a layer 100 m thick to extend from
5∘ S to 5∘ N and all the way to South America (at
270∘ E), but it is sufficient to have a significant impact.
In summary, prior to the peak of the 1982–1983 El Niño, sufficient warm
water was advected by the North Equatorial Counter Current and the reversed
Equatorial Current to produce significant warming of the eastern equatorial
Pacific. Although it was not possible to provide a full heat
budget
This is because of problems in reproducing the model mixing
using only the data available in the 5-day average datasets.
, there appears
to be no reason to look for any other mechanism advecting warm nutrient-poor
water into the eastern equatorial Pacific prior to an El Niño event.
Whereas most discussions of the El Niño focus on the role of the
equatorial waveguide, these results show that, in the model, strong eastern
Pacific El Niño events, like the 1982–1983 El Niño, occur primarily as
a result of heat transported by the North Equatorial Counter Current. Given
the good agreement between the model and observations, discussed earlier,
this is also likely to be true for the real ocean.
Heat transport and the NECC
These results raise the question of why the NECC transports so much heat in
an El Niño year. Alternatively, because, as pointed
out, the NECC has its source in the west Pacific warm pool, the real question
is “Why does the NECC transport so little heat in a non-El Niño
year?”.
Here, it is argued that the transport is reduced by the combined effect of
the Ekman transport, the geostrophic inflow, both parts of the tropical cell,
and tropical instability waves. All of these processes have the ability to
remove warm water from the core of the NECC and to replace it with cooler
water from the north or south.
Unfortunately, our theoretical understanding of these three components and the
NECC itself is poor. As a result, the argument made here has to be based on a
mixture of theory and analysis of the model results.
The tropical cell
This is probably the best understood part of the problem. As discussed by
, an easterly wind acting along the Equator produces a
rise in sea level on the western boundary of the ocean. This results in a
pressure gradient along the Equator which, when a steady state has been
reached, exactly balances the surface wind stress.
Solution of Stommel's model of the tropical cell, for a surface
layer 150 m deep, of constant density and with a vertical kinematic
viscosity of 100 cm2 s-1. Velocity contours are at
intervals of 5 cm s-1. Stream function contours are at intervals of
2 m2 s-1. Positive values are in red.
If τ is the wind stress, p the pressure in the ocean, x the distance
east and z the depth, then this balance is given by
τ=∫dz∂p/∂x.
This pressure gradient also affects the upper layers of the ocean north and
south of the Equator, where it results in a geostrophic flow vg towards
the Equator:
ρf∫dzvg=∫dz∂p/∂x,=τ,
where ρ is density and f the Coriolis term.
As the east–west pressure gradient changes only slowly with latitude, close
to the Equator it can be considered as a constant, which means that the
integral of vg tends to plus or minus infinity as 1/f as the
Equator is approached.
Away from the Equator, the surface wind stress generates Ekman transport,
ve, such that
-ρf∫dzve=τ.
This integral also tends to plus or minus infinity as the Equator is
approached. However, the singularity in ve exactly balances that
in vg, so as shown in Fig. , overall, the solution is
well behaved. In fact, it is so well behaved, that along with the Ekman
transport away from the Equator and the geostrophic inflow towards the
Equator, the solution also includes the Equatorial Current and the Equatorial
Undercurrent.
Unfortunately, the theory has a major flaw. treated the
ocean surface layer as one of constant density and with a constant vertical
viscosity. He then found that if the vertical viscosity was reduced enough to
generate a realistic undercurrent speed, then the width of the undercurrent
was only a fraction of a degree, whereas in reality it is a few degrees wide.
If the viscosity is increased by a factor of 10, as it has been for
Fig. , then a reasonable width can be obtained, but the maximum
undercurrent velocity is only 25 cm s-1 instead of a more realistic 150 cm s-1.
The solution to the problem was eventually found by who
showed that it was necessary to introduce stratification. When this was
included, upwelling near the Equator was limited by the rate at which heat
could diffuse downwards. This has a number of effects. First, it increases the
range of latitudes over which upwelling occurs. Secondly, the Ekman transport
reduces sea level near the Equator and results in a compensating rise of
the density surfaces below. This in turn increases the temperature gradient
and aids the downward diffusion of heat.
The main point though is that near the Equator the poleward Ekman transport
due to the wind is balanced by a shallow geostrophic inflow. This can be seen
in Fig. , where at 6∘ N the Ekman transport has speeds
of order 5 cm s-1, and above 150 m the geostrophic inflow has
speeds of around 1 cm s-1. At the same longitude, the core of the
NECC lies in the top 150 m, with speeds of only 20–30 cm s-1. As a
result, given the size of the Pacific Ocean, even small secondary flows can
have a significant influence on the core water transported by the NECC.
Ekman transport
Figure shows the eastward component of wind stress at
6∘ N plotted as a function of longitude and time. It shows a
regular pattern each year, the wind stress in the central Pacific being
largest during the northern spring and weakest in summer. The year 1982 is
unusual as the stress drops to near zero near the dateline for a large part
of the summer and autumn. This is shown more clearly in Fig. .
At 6∘ N, the Coriolis term equals 1.52 × 10-5 s-1,
so if the water density is taken as 1024 kg m-3,
then from Eq. (), the northward transport due to a westward wind
stress of 1 Pa is 64.2 m2 s-1. From Figs.
and , the westward stress in the middle of the ocean lies around 0.1 Pa in
spring, dropping to half that value in autumn. A value of 0.05 Pa
will generate a northward Ekman transport of 3.2 m2 s-1, equivalent
to 0.36 Sv per degree of longitude. This appears small compared to the
NECC transport in a normal year (∼ 20 Sv; see Fig. ),
but a over longitude span of 20∘ or more it will become significant.
Eastward component of wind stress (Pa) at 6∘ N in the
equatorial Pacific between 140 and 280∘ E (80∘ W),
between 1980 and 1985.
Figure shows that in the autumn of 1982 the zonal wind stress
was small over much of the central Pacific. During this period, the Ekman
transport would have been much less effective in cooling the warm core of the
NECC. The figure also shows that the reduction in central Pacific winds at
this time is consistent with the arrival of warm water shown in Fig. .
Geostrophic inflow
The meridional component of geostrophic transport, Vg, is related
to the zonal gradient of P, the vertical integral of the pressure p, by
the equations
Vg=(1/(ρf))∂P/∂x,P=∫-300zsshdzp(z),
where zssh is the height of the sea surface, z is depth, and x is the
coordinate in the zonal direction. The lower limit of 300 m was
chosen because the horizontal gradient of pressure is small at greater depths
and the limit is below the normal depth of the NECC.
Eastward component of wind stress (Pa) at 180∘ E,
6∘ N.
In Figs. and , these variables are plotted as
functions of longitude and time at latitude 6∘ N. As expected, the
integrated pressure field is usually greater in the west than in the east, a
typical mean gradient between 260 and 170∘ E being
0.035 Pa m-1. This corresponds to a southward transport of
2.2 m2 s-1 or 0.25 Sv per degree of longitude.
A second feature that might have been expected near 6∘ N is the
annual Rossby wave which shows up in the integrated
pressure field. This has a minimum which starts at the eastern boundary each
northern winter and which reaches the western Pacific in the following late
summer and autumn. In most years, the wave tends to die out west of
200∘ E, but in 1982, during the development of the El Niño,
this does not happen.
In this year, there is also a lowering of sea level close to the western
boundary, similar to an event seen in 1980. Towards the end of the year, there
is a second rapid reduction in sea level which affects the western and
central Pacific. As discussed later, there are similar drops in sea level
which occur at the Equator at the same time.
The Rossby wave and the other sea level changes are significant in that they
are large enough to reduce and change the sign of the east–west pressure
gradient. Thus, especially in 1982, they can significantly affect the
flushing of the NECC by the geostrophic inflow.
Pressure integral (106 Pa m) defined in Eq. () at
6∘ N (after subtracting a constant equal to the same integral but
with a constant density of 1024 kg m-3 and zero surface elevation).
Vertical lines are due to shallow topography.
As in Fig. , Figs. and show
short-wave features which appear to be the result of tropical instability waves.
Figure shows that the meridional transport due to the features
can reach values of over 50 m2 s-1 or 5.6 Sv degree-1,
sufficient to have a significant effect on the core of the NECC. This aspect
is discussed further in the following section.
Tropical instability waves
Tropical instability waves are wave motions observed north and south of the
Equator in the Pacific and Atlantic oceans. They show up most clearly in the
surface temperature field as fronts between the cooler equatorial water and
warmer water to the north and south. In the Pacific, they are most noticeable
in the eastern Pacific in the late northern summer and autumn.
Northward component of geostrophic transport (m2 s-1), defined by
Eq. (), at 6∘ N.
Understanding of the waves has come primarily through numerical model
studies. used a two-layer model and showed that the
waves' growth was due primarily to a barotropic instability resulting from
the strong shear between the Equatorial Current and the North Equatorial
Counter Current. , using a multilayer model, confirmed the
importance of barotropic instability but also found that baroclinic
instability was involved when the amplitude became large.
Surface temperature (∘C) from the model in late September
(a) 1981 and (b) 1982. (Values below 20.5 ∘C
are combined.)
However, this picture was not supported by the study of
. They analysed observations made during the
Hawaii-to-Tahiti shuttle experiment and found that the main instability laid
just south of the Equator and was due to the shear between the Equatorial
Undercurrent and the South Equatorial Current. Also, unlike Philander and
Cox, they found an instability between the Equatorial Current and the NECC in
the northern winter and a baroclinic instability of the NECC during the
northern spring.
These inconsistencies have never been properly explained, but later studies,
both observational and numerical see the Menkes paper for more
references, support the earlier analysis of Philander and Cox.
Figure shows the model surface temperature fields for late
September in 1981 and 1982. The first shows a series of well-developed
tropical equatorial waves just north of the Equator starting near
250∘ E and extending west to beyond 210∘ E. The
corresponding velocity field shows a series of oval anticyclonic eddies with
an west–east width of about 10∘ with southern and northern limits at
approximately 1.5∘ N the 7.5∘ N. The eddies tend to be
confined to the top 300 m, the 28 ∘C isotherm which is
at a depth of ∼ 20 m at the Equator, dropping to around 200 m
in the centre of each eddy. Below 200 m, the eddy
signature drops off rapidly, so although there is some displacement of the
isotherms near 500 m, the isotherm displacements are very small below that depth.
The eddies are affected by the tropical cell. As a result, at a depth of
30 m, maximum northward velocities near 5∘ N are ∼ 1 ms-1
and maximum southward velocities ∼ 0.6 ms-1.
In contrast, at 108 m, maximum northward velocities are
∼ 0.85 ms-1 and maximum southward velocities ∼ 1 ms-1.
The rms northward transport variability Vrms (defined
in Eq. ) along latitude 6∘ N. Units are
m2 s-1.
The surface current at the Equator (m s-1), averaged between
1∘ S and 1∘ N. Negative values correspond to the normal
westward-flowing Equatorial Current.
TIW variability
Given the potential impact of tropical instability waves on the NECC, it is
useful to have a measure of how their ability to advect water north or south
changes with time. This may be achieved by estimating the rms variance of
the northward velocity about its mean value.
Let V300 be the northward transport in the top 300 m of the ocean:
V300=∫dzv.V‾, its value over a range of longitudes, and Vrms, the
rms variance, are then given by
V‾=HV300,Vrms=HV300-V‾,
where H() is a Hann smoothing filter with a width of 20∘ of longitude.
The result at 6∘ N is shown in Fig. . In most years,
the rms transport after smoothing has values around
30 m2 s-1, consistent with the peak values discussed previously.
However, what is very significant is the region of very low variability that
starts in the west, in mid-1982, and moves across the Pacific during the
latter part of the year. The variability then stays low for a large fraction of 1983.
As the generation of TIWs is partly associated with the Equatorial Current,
it is possible that the low variability results from the reduction in the
Equatorial Current as the El Niño develops and the low wind stress region
moves east.
Figure plots the strength of the surface Equatorial Current as
a function of longitude. The region of reduced activity of tropical
instability waves, seen in 1982, fits very closely with the region of reduced
and reversed currents at the Equator.
This region of low TIW variability can also be seen in Fig. . In
September 1981, tropical instability waves are mixing cooler equatorial
water into the NECC between 180 and 200∘ E. In
September 1982, these are not present and the warm core of the NECC is
advected much further east before such mixing events occur.
Eastward transport (m2 s-1) of the NECC across
180∘ E, plotted as a function of latitude and time. The transport
here is defined as the integral of the eastward component of velocity from
the surface to the first level where it is negative. It is zero if the
surface velocity is westward.
Total transport (Sv) between 3 and 8∘ N of the NECC across
180∘ E.
The North Equatorial Counter Current
Any attempt to define the strength of the eastward-flowing NECC is
complicated by the fact that near the Equator it is often connected to the
eastward-flowing Equatorial Undercurrent and that at times the wind-driven
current at the Equator may reverse direction. For that reason, it is
convenient to define the NECC as the region lying between 3 and 8∘ N
where the surface velocity is eastward and its transport as the integral down
to the depth where the velocity first changes sign.
Figure plots the transport defined in this way across
180∘ E but extended to cover the region from the Equator to
10∘ N. It shows that when defined in this way the transport is
highly variable. Examination of the velocity field when the eastward flux is
zero showed that it was primarily due to oceanic eddies. The large values
seen near the Equator arise from the reversal of the surface current at the
Equator together with a contribution from the Equatorial Undercurrent.
The year 1982 is seen to be unusual. First, the current is continuous,
consistent with the reduction in current variability discussed previously.
Secondly, the peak and average transport in the current appear to increase
with the peak value reaching 140 m2 s-1. Thirdly, the latitude of
the current core appears to move southward, lying near 5∘ N rather
than the 7∘ N that predominates in 1981 and 1982.
Figure plots the total transport between 3 and
8∘ N. This shows that the total transport of the NECC averages
between 15 and 20 Sv, but in 1982 it doubles to between 30 and 40 Sv.
Differences in El Niño years
The results presented so far show that the NECC can at times transport large
amounts of water with temperatures above 28 ∘C eastwards across the
Pacific. They also show that in most years this does not occur because
tropical instability waves, the Ekman transport and the geostrophic inflow
combine to dilute the warm core of the NECC with cooler water from the north and south.
However, in an El Niño year, once the region of low wind stress has
started moving eastwards, the strength of these processes is reduced in the
ocean to its north and west. As a result, the core of the NECC passing the
eastern boundary of the low wind region is much warmer than normal, and as it
continues eastwards it has the potential to trigger new episodes of deep
atmospheric convection. As a result, the region of deep atmospheric convection
may progress steadily eastwards.
This poses the question, “Why does an El Niño not occur every year?” or,
given that the processes that start an El Niño have not been discussed,
“Why is every El Niño not a strong El Niño like the one in 1982–1983?”.
One possibility, originally proposed by and supported by
the results of the last section, is that the year-to-year differences are, in
part, a result of changes in the strength of the NECC. For this reason, the
next sections consider the year-to-year differences in more detail.
Wyrtki's NECC estimate based on sea levels
estimated changes in the strength of the NECC from sea
level measurements made at Kiritimati (Christmas Island; 01∘52′ N,
157∘24′ W) on the equatorial ridge and Kwajalein Atoll
(8∘43′ N, 167∘44′ E) in the counter current trough. He
found that the height difference was largest and the NECC presumably
strongest during the El Niños of 1957–1958, 1963–1964 and 1967–1968.
The model data were analysed in the same way, and it was found that the SSH
difference between 3 and 9∘ N correlated with the average surface
currents between those latitudes. In particular, they both showed significant
increases during the same periods in the autumns of 1982 and 1997, when the
strong El Niños shown in Fig. were developing.
Wyrtki's analysis showed that the change in NECC strength was due primarily
to the lowering of SSH in the counter current trough. As shown in
Fig. a, at 168∘ E, the longitude of Kiritimati,
the model results agree with this.
They show a reduced sea level in the trough, at both 6 and 9∘ N,
during the second half of 1982 but a roughly constant sea level near the
Equator during the same time period. At other times, sea level at
3∘ N is usually slightly above that at the Equator, a result of the
equatorial trough and ridge that develop when the westward-flowing Equatorial
Current is present.
Sea surface height (m) at longitudes: (a) 168∘ E,
(b) 230∘ E, and latitudes: (black) the Equator, (red)
3∘ N, (green) 6∘ N and (blue)
9∘ N.
The model SSH (m) at the Equator as a function of time and
longitude.
Further east at 230∘ E (130∘ W)
(Fig. b), sea level differences between latitudes are
generally smaller and there is a strong annual signal, especially at
6 ∘N. In the second half of 1982, there is again a large sea level
difference between the Equator and 9∘ N. However, at this longitude,
the main slope lies further north between 6 and 9∘ N.
It also arises primarily from an increase in sea level near the Equator, the
sea level in the trough at 9∘ N remaining relatively constant.
Thus, the model agrees with Wyrtki's result for Kiritimati but it also
indicates that the full picture is much more complex. To understand more, it
is convenient to investigate the changes in sea level with both longitude and
time at each latitude.
The annual wave and other processes
Figures to show the sea level plotted as a
function of longitude and time at the Equator, at 6∘ N and at 9∘ N.
The model SSH (m) at 6∘ N as a function of time and
longitude.
Starting with the Equator, the figures show that, except during the
1982–1983 El Niño event, the east–west slope remains relatively constant. Eastward
travelling equatorial Kelvin waves occur at regular intervals which, east of
their generation region, produce increases in sea level which return to
normal after the wave has passed.
In this figure, the El Niño event starts in the middle of 1982 when the
sea level in the west drops and the region of maximum sea level moves to
approximately 190∘ E. The initial movement may be associated with a
Kelvin wave, but the maximum then remains fixed, despite further Kelvin
waves, until near the end of 1982 when sea level drops rapidly all along the
Equator. This collapse is certainly associated with a Kelvin wave.
The model SSH (m) at 9∘ N as a function of time and
longitude.
Following the collapse, sea level stays low throughout 1983, recovers
slightly in 1984 and only returns to normal at the start of 1985.
At 9∘ N, sea level again shows a mean east–west slope, but at large
scales it is highly variable, the east–west differences being largest in the
middle of 1981, 1983 and 1984, and smallest at the end of 1982. The latter
occurred around the period when sea level at 9∘ N dropped along the
Equator. However, unlike the Equator, sea level rapidly recovers in 1983 to a
value in the west even higher than in 1981.
Sea level at 9∘ N also shows short-period and short-wavelength
Rossby-wave-like features moving westward at all times. The features may be
partly due to tropical instability waves, but the region is also affected by
eddies along the edge of the North Equatorial Current.
The annual signal at 9∘ N is strong and to first order appears to
consist of two main components. The first is a change independent of
longitude which has its maximum in the middle of each year. The second is a
set of westward travelling waves, an example of which is the minimum in sea
level that starts at the eastern boundary in the autumn of 1981 and which
reaches 200∘ E at the end of 1982.
At 6∘ N, sea level also shows an annual variation, but here the
signal appears to be dominated by the westward travelling annual wave. Like
the wave at 9∘ N, this starts at the eastern boundary late in the
year. It reaches 260∘ E (100∘ W), the approximate
longitude of the Galapagos Islands, in the northern spring where it is
associated with a minimum in sea level. It then moves westward more rapidly,
the leading edge reaching the western boundary in the middle of the year and the trailing
edge arriving before the end the year.
The propagation of the 1982 minimum in sea level at 6∘ N appears to
be unusual in that at 230∘ E the minimum is similar to the value
in 1981, but in the region west of 180 ∘E the minimum is much lower.
As shown in Fig. , at 168∘ E, the passage of
the wave results in the sea level at 6∘ N being similar to that at
9∘ N at a time when sea level at the Equator remains high. Thus,
although the meridional pressure difference across the NECC remains roughly
constant, the current is squeezed into a path closer to the Equator, where the
Coriolis term is smaller. As the current is in geostrophic balance, its
transport per unit depth must increase.
(a) Surface temperature and wind stress vectors.
(b) Sea level (SSH) and velocity vectors are from the 29 March 1982 archive
dataset. Each archive dataset contains averages over the previous 5 days of
the model run.
Thus, a stronger-than-normal annual Rossby wave will move the core of the NECC
towards the Equator, increasing the speed of the current and the flux of warm
water to the east. The increased speed will help to reduce the effect of
tropical instability waves and the other mechanisms on the core temperature
of the NECC. If the core temperature is high enough and the flux large
enough, this may then trigger new episodes of deep atmospheric convection further
east in the Pacific.
Development of the El Niño during 1982
The discussion so far has concentrated on individual physical processes with
only limited discussion of the overall development of the El Niño. To
give more context, the following sections briefly discuss some of the other
events that occurred in the equatorial Pacific during 1982 and how these may
be connected to the processes discussed above.
29 March
Figure shows fields of sea level and surface
temperature together with the surface velocity and wind stress vectors for
29 March 1982. At this time, the minimum in the annual Rossby wave at both
6 and 9∘ N is still in the eastern Pacific, where it contributes to
the minimum in the counter current trough near 240∘ E. The vector
plot shows that the NECC is a weak feature, except between 140 and
160∘ E, where it runs along the northern flank of a region of
maximum sea level.
The figure also shows the Equatorial Current in the central Pacific with sea
level ridges to north and south on which maxima can be seen due to tropical
instability waves. The warmest temperatures at the Equator are found
north-east of New Guinea and this is also the region where sea level along
the Equator is highest.
The figure shows winds flowing in a south-east direction along the north cost
of New Guinea. Such winds often occur early in the year when there is strong
convection in the South Pacific Convergence Zone. In this case, convection
over warm water appears to have generated cyclones both north and south of
the Equator generating, for just one 5-day averaging period, an extended
region of westerly winds along the Equator.
In the ocean, the Equatorial Current was still present early in the month but
has now disappeared. It is not re-established, but instead, during May and
June, there are periods with a reversed Equatorial Current between 150 and
170∘ E.
29 June
By the end of June, the annual Rossby wave at 6∘ N has reached the
western Pacific and the wave at 9∘ N has reached 230∘ E. As
a result, the counter current trough is deeper and more uniform throughout
the central and eastern Pacific. The equatorial ridge in the eastern Pacific
is also more developed, and this, together with the changes in the trough,
results in a much stronger NECC all across the Pacific.
(a) Surface temperature and wind stress vectors.
(b) SSH and velocity vectors are from the 29 June 1982 archive dataset.
Colours and vector scales are as in Fig. .
In the west, sea level at the Equator has dropped slightly, but this is the
time that, in Fig. , the maximum sea level is in the process
of moving from around 150 to 190∘ E (170∘ W). The region of
low winds has started to expand, temperatures have risen, including along the
line of the NECC, and the current is more effective at transporting warm
water to the east, beyond the region of low winds.
27 September
This lies in the middle of the time period when the maximum SSH along the
Equator lies near 190∘ E. The temperatures within the region
increase with time. The region also spreads north and south on both sides of
the Equator. One consequence of this is the region of higher-than-normal sea
level at 200∘ E, 6∘ N seen in Fig. .
By this time, the NECC has grown in strength and its path shifts northwards as
it crosses the ocean, starting near 4∘ N and reaching 8∘ N near 240∘ E.
This is also a this period when westerly wind bursts develop. These can be
seen in Fig. and the resulting Kelvin waves in
Fig. . However, these occur in the region where the mean wind is now
westerly and there is no evidence that the resulting eastward surface current
along the Equator in this region is significantly different from that to be
expected from the average westerly wind.
31 December
In October and early November 1982, the central and western equatorial Pacific
was a region of light westerly winds interspersed by stronger westerly wind
bursts. In contrast, in the eastern Pacific, a strong trade wind continued to blow.
In late November, the pattern changed and strong westerly winds blew along the
Equator on either side of the dateline. These result in an eastward-flowing
surface current which continues until the end of the year (Fig. ),
advecting the warm water patch at the Equator towards the east.
By the end of the year, the trade winds are starting to be re-established
north of the Equator. In the west, the Equatorial Current is reforming, and
by 10 January, it is again established in the east Pacific. As a result, in
the following weeks, the patch of warm water at the Equator moves back
westwards.
However, as shown in Fig. , this is the time that
precipitation is most established in the central Pacific. Precipitation
remains high in the central Pacific during the remainder on the Southern
Hemisphere summer after which it first moves closer to South America before
the high precipitation region is re-established in the western Pacific.
Particle tracking
A useful alternative view of the processes can be obtained using the TRACMASS
particle tracking program . Figure
shows the results of seeding the NECC at 200∘ E in June 1981 and 1982.
This is the time when in 1982 the NECC was carrying water warmer than
29 ∘C into the western Pacific.
In 1982, each model grid box along the line shown, with a temperature greater
than 29 ∘C, was seeded with a single particle. In 1981, there was no
water along the line with this temperature so boxes were seeded where the
temperature was greater than 27.8 ∘C.
(a) Surface temperature and wind stress vectors.
(b) SSH and velocity vectors are from the 27 September 1982 archive dataset.
Colours and vector scales are as in Fig. .
(a) Surface temperature and wind stress vectors.
(b) SSH and velocity vectors are from the 31 December 1982 archive
dataset. Colours and vector scales are as in Fig. .
In 1981, the water was initially carried east but before reaching the far
eastern Pacific most particles moved north, where they were carried westward
by the North Pacific Subtropical Gyre. The remainder moved south and were
carried westward by the Equatorial Current.
By contrast, in 1982, the much warmer water was carried predominantly to the
east with a significant quantity reaching as far as the Galapagos Islands.
Particles were carried to the north and south, but fewer were lost in this
way than in 1981.
In two other runs (not shown), particles were seeded in the Niño 1 and 2
regions, which are to the south-east of the Galapagos Islands. The tracking
program was then run backwards in time, from early in 1982 and 1983, to
determine where the water came from.
In 1982, the water was found to have a local origin, some coming from
upwelling regions near the coast. In contrast, in 1983, a significant amount
came from just north of the Galapagos, apparently displaced by the water
entering the region shown in Fig. . Observations made in 1981
and late 1982 show a similar movement of warm, low-nutrient surface water
southwards across the Equator at this time .
Water particle positions plotted every 5 days, starting from 24 June
(a) 1981 and (b) 1982, and running to the end of the year.
The date is denoted by the colour of each dot. For the initial state, one
particle was placed at the centre of each model grid cell lying along the
black line, having a water temperature greater than
(a) 27.8 ∘C and (b) 29 ∘C, there being no
water with a temperature greater than 29 ∘C along the line
in 1981.
In a final test, water particles were tracked moving eastward along the
Equator. The model showed that in late October 1982, following the reduction
and reversal of the winds, the Equatorial Current at 200∘ E also
reversed direction. The region of water warmer than 29 ∘C, shown in
Figs. and , then started moving eastwards
along the Equator.
This ocean was seeded as before, and Fig. shows the initial
seeding line and the later particle positions. The particles are seen to move
eastward but they do not progress far. By the end of the year, none of the
particles have passed 240∘ E and many have turned back westward.
The 1997–1998 El Niño
As a check that the above results were not unique, the analysis was repeated
for the strong 1997–1998 El Niño. Key results are shown in Fig.
and in Figs. to .
The temperature contours in Fig. b indicate that the
1997–1998 El Niño started much earlier in the year than the 1982–1983 one. During the
northern spring months, the region with SST values above 28 ∘C
extended eastwards until the middle of the year when it had reached 230 ∘E, the
limit being similar to that of the warm pool events of 1987–1988 and 1991–1992 (see Fig. ).
In 1997, the temperature contours then briefly retreat westwards before being
overtaken by the main El Niño event which carries water with a
temperature greater than 28 ∘C to the eastern boundary.
Figure shows the flux of water with temperatures greater
than 28 ∘C across 180, 210 and 240∘ E. The first
corresponds to a longitude well within the warm pool event, the second to a
longitude near its the eastern limit and the third to a longitude beyond the
limit.
Water particle positions plotted every 5 days between 2 October 1982
and the end of the year, the date being denoted by the colour of each dot.
For the initial state, one particle was placed at the centre of each model
grid cell, lying along the black line, having a water temperature of greater
than 29 ∘C.
At 180∘ E, the transport of warm water by the NECC in 1995 and 1996
is seen to be greater than that in the year preceding the
1982–1983 El Niño. This may indicate that the western Pacific was warmer
during the later period. The NECC continues with a similar flux of warm water
during early 1997, but in early spring, at about the time the temperature
front of Fig. reaches 180∘ E, there is a significant
flux of warm water in the equatorial band and this continues until the end of
the year.
At 210∘ E, in the equatorial band, there is a single pulse of warm
water in the middle of the year, when the warm water of Fig. reaches this
longitude, but for the rest of the year the main transport is at the
latitudes of the NECC. At the end of 1997, as in 1982, there is again a short
pulse of warm water in the equatorial band.
At 240∘ E, during 1997, warm water is only advected by the NECC,
and as shown in both figures for 180 and 210∘ E,
this is associated with a movement of the NECC towards the Equator.
Overall, the results indicate that the strong 1997–1998 El Niño was
different in that it developed from a warm pool event whose maximum occurred
around the middle of the year. However, the NECC was again involved in the second half of
the year, transporting warm water eastwards well beyond the limit of the warm
pool event and eventually into the eastern Pacific.
Dilution processes during 1997–1998
Figure a shows the eastward wind stress at 6∘ N,
responsible for the Ekman transport contribution to the dilution of the NECC
at that latitude. During the early part of 1997, winds in the central Pacific
are weaker than in the corresponding period of 1982 and in the west the winds
are mainly either near zero or westerlies. As in 1982–1983, the region of low
winds lies well to the west of the warm water boundary seen in Fig. b.
During the early autumn, there is a second period of stronger westerlies in
the west, but these are weaker than in 1982. Winds remain low or westerly
across the whole of the Pacific, until late in the year when the normal
pattern of easterly winds starts to return.
The pressure integral term (Fig. b) shows the influence of
a stronger-than-normal annual Rossby wave in 1997 similar to the wave that
occurred in 1982. The integral also has values which are much lower than
normal along the western boundary, during the autumn of 1997 and all across
the ocean towards the end of the year. As in 1982, the changes in the west
and at the end of the year appear correlated with changes in sea level along
the Equator (Fig. ) occurring at the same time.
The variability of the pressure gradient (Fig. c) shows
that the tropical instability waves are reduced in intensity during the
development of the 1997–1998 El Niño, again as they were in 1982–1983.
The region of reduced variability starts in the west in early spring 1997 and
gradually extends eastwards until the end of the year. In the following year,
the region slowly retreats and is replaced, again as before, by a
stronger-than-normal set of waves that develops in the central and eastern Pacific.
The plot by Fig. 4 based on mooring data from the
Equator at 220∘ E (140∘ W) shows a similar reduction in
the strength of the tropical instability waves during the second half of 1997
and the early months of 1998.
Vertically integrated flux of water (m2 s-1) with
temperature (a) greater than 28 ∘C crossing longitude
180∘ E, (b) greater than 28 ∘C crossing
210∘ E and (c) greater than 28 ∘C crossing
240∘ E. The figure is blank where the flux is zero.
Sea level during 1997–1998
Figure shows the sea surface temperature field in late
September 1997. Although the El Niño started in a different way, at this
stage, the temperature field is very similar to that from the 1982–1983 event.
This will have been partly due to the reduction in the dilution processes
discussed in the last section. However, following the analysis of
the 1982–1983 El Niño, it may also be due to changes in the strength of the NECC resulting
from changes in sea level at the Equator and at latitudes close to the NECC.
Figure shows sea levels at the Equator, 6 and
9∘ N during the period 1995–2000. In most years, sea level at the
Equator shows the expected increase from east to west due to the trade winds.
Sometimes, as in early 1996, there is a small reversal in slope close to the
western boundary. When this does occur, it is usually near the turn of the
year when winds on there are often westerly (Figs.
and ). Equatorial Kelvin waves are also seen in the sea
level figure but, as before, they are not observed in the surface temperature
plot (Fig. b).
In early 1997, the maximum in sea level moved away from the western boundary.
The behaviour is similar to the 1982 event but this time the maximum dies
away and there is a period of reduced east–west slope, with lower-than-normal
sea levels in the west and higher-than-normal sea levels in the eastern Pacific.
In the autumn, a second region of high sea levels develops near
190∘ E and, as in 1982, it remains in approximately the same
position until late in the year. It then moves slightly eastwards before sea
level again falls rapidly all along the Equator.
Sea level at 6∘ N also has strong similarities with the earlier
period, the annual Rossby wave reaching the western Pacific in the second
half of 1997. There is also a reduction of equatorial sea level near the
western boundary at this time which spreads eastward to extend along much of
the Equator by the end of the year. This similar to the change in the
pressure integral discussed in the previous section. A plot of the difference
in sea level between 6∘ and the Equator shows that this is
approximately constant in the spreading region as the sea levels drop.
At 9∘ N, sea level also shows a reduction in east–west slope during
late 1997 (and early 1998). In the west, sea level is higher than at
6∘ N, implying a westward current probably due to a southward
extension of the North Pacific Gyre. In the central and eastern Pacific, sea
level is lower than at 6∘ N, this implying that the path of
eastward-flowing NECC has moved further north here.
Developments during 199716 March
Figure is included partly to illustrate the
cross-equatorial wind flows that often occur near the beginning and end of each
year in the western Pacific. The winds cross the Equator north of New Guinea
and are responsible for many of the positive values seen west of the dateline
in plots of the eastward wind stress along the Equator (Figs. and ).
After crossing the Equator, the winds continue towards the South Pacific
Convergence Zone, where deep atmospheric convection events are expected to be
strongest at this time of year. Westerly winds along the Equator may also be
produced by cyclones close to Indonesia and, once the warm water front has
moved further east, as a result of cyclones that develop north or south of
the Equator over the warm water.
In mid-March, the strong westerly wind event lasted for almost 20 days
(Fig. ) and as seen in Fig. this
resulted in a strong current along the Equator advecting a surface water mass
with temperatures of up to 30 ∘C.
Values at 6∘ N, during the period 1995 to 2000, of
(a) the eastward component of the wind stress (Pa), (b) the
pressure integral (106 Pa m) of Eq. () at 300 m and
(c) the rms northward transport variability Vrms
(m2 s-1).
Surface temperature (∘C) from the model in late
September 1997 (values below 20.5 ∘C are combined).
The figure corresponds to the period when the warm pool event was developing.
At this time, the NECC is relatively weak. It is advecting some water warmer
than 28 ∘C to the east, but tropical instability waves are well
developed near 180∘ E, and these are rapidly mixing away warm water
from the core of the current.
Sea surface height (m) during the period 1995 to 2000 at
(a) the Equator, (b) 6∘ N and
(c) 9∘ N.
29 June
Figure corresponds to the end of the western and
central Pacific warm pool event and the start of the eastern Pacific
El Niño. At the Equator, the winds are predominantly from the east, but
earlier in the month a strong reversed Equatorial Current developed in the
central Pacific which appeared to be closely linked with the NECC. The
current is still present here and is associated with a ridge of high sea
level along the Equator. Sea surface temperatures of greater than
29.5 ∘C are found in the central Pacific along the Equator and to
the north.
At 8∘ N, the counter current trough is well developed, resulting in
a well-developed NECC over most of the width of the Pacific. At the
longitudes where there is a ridge along the Equator, the two currents appear
to combine generating a single broad reversed Equatorial Current.
27 September
The overall picture (Fig. ) is similar to that of
September 1982. The counter current trough remains well developed, and there
is a strong NECC carrying warm water to the far western Pacific. The
eastward-flowing reversed Equatorial Current has disappeared and in the
eastern Pacific the westward-flowing Equatorial Current has returned. There
is some upwelling of cooler water near the Galapagos but, compared with the
same time in 1982, tropical instability waves are less developed.
At the Equator, the region of low winds extends to 210∘ E. Near the
dateline, where sea surface temperatures are high, the wind stress vectors
show convergence at the Equator. This may be connected with continuing deep
atmospheric convection in the region.
Discussion
This paper is the result of a preliminary analysis of archived data from an
early run of a high-resolution global ocean model. A previous comparison with
observations from the equatorial Pacific indicated that the model behaved
well and thus provides some measure of confidence in the present results.
The analysis shows that, during the development of the strong 1982–1983 El Niño,
the North Equatorial Counter Current dominated the transport of
water with temperatures greater than 28 ∘C. This was also true during
the development of the strong 1997–1998 El Niño at longitudes east of
130∘ E. The path of the NECC lies close to the latitude of the
Intertropical Convergence Zone so the atmosphere is likely to be very
sensitive to warm water carried eastwards by the NECC.
Eastward component of wind stress (Pa) at the Equator between
140 and 280∘ E (80∘ W).
The analysis also showed that the movement of warm water along the equatorial
band during the two strong El Niños was very different from the warm pool
events of 1987–1988 and 1992–1993, and the similar event of early 1997. In
the strong events, warm water spreads rapidly eastwards across the Pacific
after which it equally rapidly retreats. The warm pool events are more
incremental, with the westward extent usually extending slowly from one year to
the next to a maximum near 230∘ E, after which there is a
relatively small retreat.
During the growth of the strong El Niños, and also during the warm pool
event of early 1997, the core temperature of the NECC is higher than usual.
This is associated with a reduction at 6∘ N of the Ekman transport,
geostrophic inflow and tropical instability waves, all of which can remove
warm water from the core of the NECC and replace it with cooler water from
north and south.
The reduction in the Ekman transport is associated with reduced winds in
regions where deep atmospheric convection appears to have moved out over the
ocean. The reduction in the strength of the tropical instability waves,
potentially the most important process, is associated with a reduction in the
strength of the Equatorial Current. This can also result from reduced
easterly winds. The reduction in geostrophic inflow may partially result from
the reduction in Ekman transport but the model results also show that it is
connected with the passage of the annual Rossby wave.
During the growth of the strong El Niños, the NECC is observed to move
closer to the Equator and become stronger. The model results indicate that this
is also a result of the passage of a stronger-than-normal annual Rossby wave.
In the west, the wave deepens the counter current trough, thus increasing the
strength of the NECC. In the central Pacific, the wave moves the northern
boundary of the current southwards but produces little change in the
north–south pressure difference across the current. As the Coriolis term
drops to zero at the Equator, this will inevitably increase the speed of the current.
In both of the two strong El Niños and in the warm pool event of early 1997,
sea level at the Equator developed a maximum in the middle of the ocean. Once this
had formed, its position remained relatively fixed despite the continuing
eastward extension of the pool of warm water. Near the end of the event the
maximum moved slightly eastwards before sea level dropped all along the
Equator. The reason for the behaviour is not understood but the maximum was
usually associated with the warmest patch of water lying at the Equator.
The forcing fields show that periods with strong westerly winds occurred at
the Equator during the development of both the two strong El Niños
studied and the warm pool event. North of New Guinea, on either side of the new
year, this was often due to a cross-equatorial airflow towards the South
Pacific Convergence Zone.
At other times, strong westerlies and associated cyclones were only found
above regions where the water temperature was already above 28 ∘C
and deep atmospheric convection is likely to have occurred. The strong
westerlies at the Equator did drive a reversed Equatorial Current but this
was confined primarily to the region of warm ocean. The westerly winds also
generated equatorial Kelvin waves, but there is no evidence that these caused
a significant extension of the warm water region.
Ocean mechanisms
The results highlight two oceanic mechanisms that are important during the
development of a strong El Niño.
The first is the Rossby wave mechanism that increases the speed of the North
Equatorial Counter Current. In the west, the annual Rossby wave deepens the
counter current trough. In the central Pacific, it moves the NECC closer to
the Equator into a region where the Coriolis term is smaller.
(a) Surface temperature and wind stress vectors.
(b) Sea level (SSH) and velocity vectors are from the 16 March 1997 archive
dataset. Each archive dataset contains averages over the previous 5 days of
the model run.
(a) Surface temperature and wind stress vectors.
(b) SSH and velocity vectors are from the 29 June 1997 archive dataset.
Colours and vector scales are as in Fig. .
On the basis of the present model results, the timing of the strong El
Niños is almost certainly due to the annual Rossby waves, the arrival
of the wave at 6∘ N in the western Pacific in the middle of the year, lowering sea
level and triggering the increased transport by the NECC. As it crosses the
Pacific, the flow of warm water is aided by slower-moving Rossby waves until it
arrives in the far eastern Pacific around the new year.
The second mechanism involves the changes which result in less dilution of
the warm water core of the North Equatorial Counter Current. Once an
El Niño has started, the low winds around the Equator and the collapse of
the equatorial current mean that the diluting effects of the Ekman transport,
the geostrophic return flow and tropical instability waves are all reduced in intensity.
(a) Surface temperature and wind stress vectors.
(b) SSH and velocity vectors are from the 27 September 1997 archive dataset.
Colours and vector scales are as in Fig. .
When these two mechanisms are active, they both allow the NECC to carry warm
water much further east than normal, and it does so at a latitude where the
atmosphere may be particularly sensitive to extra surface warming. In the
cases studied, water with temperatures greater than 28 ∘C was
transported past the region of low winds and deep atmospheric convection to
longitudes where it could trigger new episodes of deep atmospheric
convection. This almost certainly had the result of extending the low wind
region after which the processes can be repeated.
A third potential oceanic mechanism, that is not fully understood, involves
the sea level maximum, and associated temperature maximum, that develops at
the Equator in the central Pacific. In the model, the maximum was generated in
both strong El Niños and independently in the warm pool event of early 1997.
It is of interest because, as discussed by , as well as
the pressure gradient along the Equator having the potential to generate
currents whenever the opposing wind stress drops, the drop in sea level north
and south of the Equator will result in eastward-flowing geostrophic
currents. These will transport warm water eastwards, independently of the
local winds, and will continue as long as there is a ridge in sea level along the Equator.
Of the three mechanisms, the Rossby wave mechanism is probably of greatest
importance. This is because the annual Rossby waves are generated in the
eastern Pacific well before any extra advection of warm water occurs in the
west. Thus, a better theoretical understanding of the growth and propagation
of the waves, plus measurements made as they first develop, should allow
useful predictions to be made early each year as to the probability of a
strong El Niño.
At the time of publication, the archived data are
freely available at
http://gws-access.ceda.ac.uk/public/nemo/runs/ORCA0083-N06/means/ (last
access: 1 July 2018). The NEMO ocean model code and its documentation are
available from http://forge.ipsl.jussieu.fr/nemo/wiki/Users (last
access: 1 July 2018).
The author is on the advisory board of Ocean Science.
Acknowledgements
I wish to acknowledge the support of the Marine Systems Modelling group and
the aid of Andrew Coward at the UK National Oceanography Centre, part of the
Natural Environment Research Council, where much of this research was carried
out – as part of the NERC projects ODYSEA (grant number: NE/M006107/1) and
SMURPHS (NE/N005686/1). I also wish to acknowledge the earlier role of the
Australian CSIRO Division of Fisheries and Oceanography. Without the financial
support and the professionalism and enthusiasm of staff at both centres this
work would not have been possible.
Finally thanks to the reviewers and topic editor for their conscientious reviews
and helpful comments.
Edited by: Matthew Hecht
Reviewed by: Lauren Kuntz and two anonymous referees
References
Barber, R. T. and Chavez, F. P.: Biological Consequences of El Niño,
Science, 222, 1203–1210, 1983.
Chen, S., Wu, R., Chen, W., Yu, B., and Cao, X.: Genesis of westerly wind
bursts over the equatorial western Pacific during the onset of the strong
2015–2016 El Niño, Atmos. Sci. Lett., 17, 384–391, 2016.
Cox, M.: Generation and Propagation of 30-day waves in a numerical model of
the Pacific, J. Phys. Oceanogr., 10, 1168–1186, 1980.
de Vries, P. and Döös, K.: Calculating Lagrangian trajectories using
time-dependent velocity fields, J. Atmos. Ocean. Tech., 18, 1092–1101, 2001.Dussin, R., Barnier, B., and Brodeau, L.: The making of Drakkar forcting set
DFS5, DRAKKAR/MyOcean Report 01-04-16, LGGE, Grenoble, France, https://www.drakkar-ocean.eu/publications/reports/report_DFS5v3_April2016.pdf
(last access: 1 July 2018), 2014.Evans, J. L. and Webster, C. C.: A variable sea surface temperature threshold
for tropical convection, Aust. Meteorol. Oceanogr. J., 64, S1–S8, 10.22499/2.6401.007, 2014.Hu, S. and Fedorov, A. V.: The extreme El Niño of 2015–2016: the role of
westerly and easterly wind bursts, and preconditioning by the failed 2014 event,
Clim. Dynam., https://doi.org/10.1007/s00382-017-3531-2, in press, 2017.King, B. A., Allison, M., Alderson, S. G., Read, J. F., Smithers, J., and Webb,
D. J.: SeaSoar data from the western equatorial Pacific Ocean, Report 291,
Institute of Oceanographic Sciences, Deacon Laboratory, http://nora.nerc.ac.uk/id/eprint/115310
(last access: 1 July 2018), 1991.Kug, J.-S., Jin, F., and An, S.-I.: Two types of El Niño Events: Cold Tongue
El Niño and Warm Pool El Niño, J. Climate, 22, 1499–1515, 10.1175/2008JCLI2624.1, 2009.Large, W. and Yeager, S.: Diurnal to Decadal Global Forcing for Ocean and
Sea-Ice Models: The Data Sets and Flux Climatologies, NCAR Technical Note TN-460+STR,
National Centre for Atmospheric Research, Boulder, Colorado, 2004.Levine, A. F. Z. and McPhaden, M. J.: How the July 2014 easterly wind burst
gave the 2015–2016 El Niño a head start, Geophys. Res. Lett., 43, 6503–6510,
10.1029/2005GC001115, 2016.
Luther, D. and Johnson, E.: Eddy energetics in the upper equatorial Pacific
during the Hawaii-to-Tahiti Shuttle Experiment, J. Phys. Oceanogr., 20, 913–944, 1990.
Madec, G.: NEMO ocean engine (Draft edition r6039), Note du Pole de modelisation 27,
Institut Pierre-Simon Laplace (IPSL), France, 2014.
McCreary, J.: A Linear Stratified Ocean Model of the Equatorial Undercurrent,
Philos. T. Roy. Soc. Lond. A, 298, 603–635, 1981.
Menkes, C., Vialard, J., Kennan, S., Boulanger, J.-P., and Madec, G.: A Modeling
Study of the Impact of Tropical Instability Waves on the Heat Budget of the
Eastern Equatorial Pacific, J. Phys. Oceanogr., 36, 847–865, 2006.Moum, J., Lien, R.-C., Perlin, A., Nash, J., Gregg, M., and Wiles, P.: Sea
surface cooling at the Equator by subsurface mixing in tropical instability
waves, Nat. Geosci., 2, 761–765, 10.1038/NGEO657, 2009.
Myers, G.: On the Annual Rossby Wave in the Tropical North Pacific Ocean, J.
Phys. Oceanogr., 9, 663–674, 1979.
Philander, S. G. H.: Instabilities of zonal equatorial currents, J. Geophys.
Res., 83, 3679–3682, 1978.Stommel, H.: Wind-drift near the Equator, Deep-Sea Res., 6, 298–302,
10.1016/0146-6313(59)90088-7, 1960.Trenberth, K. E.: The Definition of El Nino, B. Am. Meteorol. Soc., 78,
2771–2777, 10.1175/1520-0477(1997)078<2771:TDOENO>2.0.CO;2, 1997.Webb, D. J.: A comparison of sea surface temperatures in the Equatorial Pacific
Nino regions with results from two early runs of the NEMO 1/12∘ Ocean
Model, Research and Consultancy Report No. 55, National Oceanography Centre,
Southampton, UK, http://nora.nerc.ac.uk/id/eprint/513264 (last access: 1 July 2018), 2016.Webb, D. J.: A Note on Stommel's Theory of the Tropical Cell, Research and
Consultancy Report No. 63, National Oceanography Centre, Southampton, UK,
http://nora.nerc.ac.uk/id/eprint/520018 (last access: 1 July 2018), 2018.
Wyrtki, K.: Teleconnections in the Equatorial Pacific Ocean, Science, 180, 66–68, 1973.
Wyrtki, K.: Equatorial Currents in the Pacific 1950 to 1970 and Their Relation
to the Trade Winds, J. Phys. Oceanogr., 4, 372–380, 1974.
Wyrtki, K.: El Niño – The Dynamic Response of the Equatorial Pacific Ocean
to Atmospheric Forcing, J. Phys. Oceanogr., 5, 372–380, 1975.
Wyrtki, K.: Sea level During the 1972 El Niño, J. Phys. Oceanogr., 7, 779–787, 1977.
Wyrtki, K.: The Response of sea Surface Topography to the 1976 El Niño,
J. Phys. Oceanogr., 9, 1224–1231, 1979.