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Chapter 2. Ocean observations 2.1 Observational methods Before we introduce the observational instruments and methods, we will first introduce some definitions related to observations. Accuracy: The difference between a result obtained and the true value. Precision: Ability to measure consistently within a given data set (variance in the measurement itself due to instrument noise). Generally the precision of oceanographic measurements is better than the accuracy. 2.1.1 Measurements of depth. Each oceanographic variable, such as temperature (T), salinity (S), density , and current , is a function of space and time, and therefore a function of depth. In order to determine to which depth an instrument has been lowered, we need to measure ``depth''. Meter wheel. The wire is passed over a meter wheel, which is simply a pulley of known circumference with a counter attached to the pulley to count the number of turns, thus giving the depth the instrument is lowered. This method is accurate when the sea is calm with negligible currents. In reality, ship is moving and currents are strong, the wire is not straight. The real depth is shorter than the distance the wire paid out. Measure pressure. Derive depth from hydrostatic relation: where g=9.8m/s2 is acceleration of gravity and is depth. (i) Protected and unprotected reversing thermometer developed especially for oceanographic use. They are mercury-in-glass thermometers which are attached to a water sampling bottle. The pressure was measured using the pair of reversing thermometers - one protected from seawater pressure by a vacuum and the other open to the seawater pressure. They were sent in a pair down to whatever depth, then flipped over, which cuts off the mercury in an ingenious small glass loop in the thermometer. They were brought back aboard and the difference between the mercury column length in the protected and unprotected thermometers was used to calculate the pressure. Depth accuracy 0.5% or 5m, whichever is the greater. (ii) Electrical strain-gauge pressure transducer which uses the change of electrical resistance of metals with mechanical tension. A resistance wire is firmly connected to a flexible diaphragm, to one side of which the in situ hydrostatic pressure is applied. As the diaphragm flexes with change of pressure, the tension in the wire changes and so does its resistance, which is measured to provide a value for the pressure and therefore depth. Accuracy 0.1%.

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2.1.2 Measurements of temperature. (a) Bathythermograph. A liquid-in-metal thermometer causes a metal point to move in one direction over a smoked or gold plated glass slide which is itself moved at right angles to this direction by a pressure sensitive bellows. The instrument is lowered to its permitted limit in the water (60, 140 or 270m) and then brought back. Since pressure is directly related to depth, the line scratched on the slide forms a graph of temperature against depth. It is read against a calibration grid to an accuracy of 0.2k and 2m if well calibrated. Advantage: continuous T(z). Less accurate. This is an old method. (b) Expendable Bathythermograph (XBT). Widely used. Uses a thermistor as temperature-sensitive element. The thermistor is in a small streamlined weighted casing which is simply dropped over the ship's side. It is connected by a fine wire to a recorder on the ship which traces the temperature of the water in a graphical plot against depth. The latter is not sensed directly but is estimated from the time elapsed since release, using the known rate of sink of the freely falling thermistor casing. These XBTs are available for depth ranges from 200m to 1800m. Use aircraft: 300m--800m. This is an old method. Note: Recently, errors were found in the XBT falling rates (different companies have different falling rates) in temperature datasets (the data were used in IPCC AR4, 2007); now these errors were corrected.

(c) CTD--Conductivity, temperature, and depth (actually pressure). T is measured uses a thermistor mounted close to the conductivity sensor. This will be discussed a bit more in the next subsection. (d) Protected reversing mercury thermometer. These were invented by Negretti and Zamba in 1874. Since it is protected from the sea water pressure, the length of mercury is determined from temperature. As described above, it is attached to a water sampling bottle. When the bottle is closed to collect the sample the thermometer is inverted. Then the mercury is cut off in an ingenious small glass loop in the thermometer. Accuracy is 0.004C and precision is 0.002C. (e) Thermistors chains consisting of a cable with a number of thermistor elements at intervals are sometimes moored along with current meters to record the temperature at number of points in the water column. A ``data logger'' samples each thermistor sequentially at intervals and records temperatures as a function of time. Quality varies significantly. The best thermistors commonly used in oceanographic instruments have an accuracy of 0.002C and precision of 0.0005-0.001C. [Thermistor can also be instrumented on drifting buoys.] (f) Satellite. Direct observations have space and time limitations. Satellite observations can provide large spatial and temporal scale data. Advanced Very High Resolution Radiometer (AVHRR) on board of NOAA satellite, can measure SST with accuracy of 0.1-0.3k. [Multi-channel: 0.58-0.68 (visible), 0.725-1.10 (near-infra-red),

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thermal infra-red (3.65-3.93 , 10.3-11.3 ,11.5-12.5 ). Problem: Cloud vapor absorption. Inaccurate when there are clouds. Tropical Rainfall Measuring Mission (TRMM)--Microwave Imager (TMI), measure SST, 0.2C difference compare with buoy data. Spatial and temporal resolutions: 25x25 km and daily since 1997. TMI can penetrate clouds and thus are not contaminated by clouds; but the data quality can be affected by strong rainfall. [Polar orbiting: 500-800km height. Geostationary: 36,000km.] (g) Acoustic tomography. Acoustic tomography maps changes in ocean temperature using changes in sound speed along paths between acoustic sources and receivers. It was used in two somewhat different modes - in concentrated regional experiments where an attempt is made to reconstruct the full three-dimensional temperature field, and over global paths to monitor changes in the average temperature along very long paths. The first has been used to good effect in winter convection regions, where in situ ship observations have been very difficult to obtain due to the small size of convection features and the poor weather in the interesting part of the year. The second is being used for global climate monitoring. Acoustic Thermometry of Ocean Climate (ATOC) website. (h) Buoy: Some Drifters are also instrumented to measure T.

Pressure of the ocean increases greatly downward. A parcel of water moving from one pressure to another will be compressed or expanded. When a parcel of water is compressed adiabatically, that is, without exchange of heat, its temperature increases. (This is true of any fluid or gas.) When a parcel is expanded adiabatically, its temperature decreases. The change in temperature which occurs solely due to compression or expansion is not of interest to us - it does not represent a change in heat content of the fluid. Therefore if we wish to compare the temperature of water at one pressure with water at another pressure, we should remove this effect of adiabatic compression/expansion. Definition:``Potential temperature'' is the temperature which a water parcel has when moved adiabatically to another pressure. In the ocean, we commonly use the sea surface as our "reference" pressure for potential temperature - we compare the temperatures of parcels as if they have been moved, without mixing or diffusion, to the sea surface. Since pressure is lowest at the sea surface, potential temperature (computed at surface pressure) is ALWAYS lower than the actual temperature unless the water is lying at the sea surface.

2.1.3. Measurements of salinity. (a) Laboratory. Evaporate and weigh residual (oldest method). (b) Laboratory. Classical (Knudsen) method. Determine amount of chlorine, bromine and iodine to give "chlorinity", through titration with silver nitrate. Then relate salinity

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to chlorinity: S = 1.80655 Cl. Accuracy is 0.025. This method was used until the International Geophysical Year in 1957. Water sample. Not convenient on board ship. (c) Measure conductivity. Conductivity of seawater depends strongly on temperature, somewhat less strongly on salinity, and very weakly on pressure. If the temperature is measured, then conductivity can be used to determine the salinity. Salinity as computed through conductivity appears to be more closely related to the actual dissolved constituents than is chlorinity, and more independent of salt composition. Therefore temperature must be measured at the same time as conductivity, to remove the temperature effect and obtain salinity. Accuracy of salinity determined from conductivity: 0.001 to 0.004. Precision: 0.001. The accuracy depends on the accuracy of the seawater standard used to calibrate the conductivity-based measurement. How is conductivity for calculating salinity measured? (c.1) For a seawater sample in the laboratory, an ``autosalinometer'' is used, which gives the ratio of conductivity of the seawater sample to a standard solution. The standard seawater solutions are either seawater from a particular place, or a standard (Potassium Chlorine) KCl solution made in the laboratory. The latter provides greater accuracy and has recently become the standard. Because of the strong dependence of conductivity on temperature, the measurements must be carried out in carefully temperature-controlled conditions. (c.2) CTD. From an electronic instrument in the water, either inductive or capacitance cells are used, depending on the instrument manufacturer. Temperature must also be measured, from a thermistor mounted close to the conductivity sensor. In a CTD, a unit consisting of conductivity, temperature, and pressure sensors is lowered through the water on the end of an electrical conductor cable which transmits the information to indicating and recording units on board ship. The digital transmitting units have claimed accuracies of 0.005 (conductivity accuracy expressed as equivalent salinity accuracy), 0.005K and 0.15% of full-scale depth. Calibration procedures include matching the temperature and conductivity sensor responses. From conductivity, T, and depth, we obtain salinity with depth. (d) Satellite. NASA Aquarius mission: Measure sea surface salinity. Launched June 2011: http://aquarius.nasa.gov/. 2.1.4. Measurements of density. The standard laboratory method, using a weighing bottle, to determine density is not practical at sea because of the motion of the ship. Usually it is calculated from the equation of state of sea water. (T,S,P).

2.1.5. Measurements of currents. Goals for measuring large-scale circulation is to understand the circulation and variability in three dimensions. They are important for understanding climate variability, since they are directly related to heat and salt transports.

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Typical horizontal current speeds in the ocean range from about 200 cm/s in the swift western boundary currents (Gulf Stream, Kuroshio, Somali current), through 10-100 cm/s in the equatorial currents, to a fraction of 1 cm/s in much of the surface layer and in the deep waters. Vertical speeds are estimated to be very much less, of the order of 10-5 cm/s. Why? As we will see later in the Dynamics section: Vertically, pressure gradient force is basically balanced by gravitational force, vertical scale is much smaller than the horizontal scale. Measurements methods: Lagrangian methods: The path followed by each fluid particle is stated as a function of time. Measurement follows fluid parcels. Eulerian methods: The velocity (speed and direction) is stated at every point in the fluid. 2.1.5 a. Direct current measurements. Surface drifters (Lagrangian). (a) Ship drift. The earliest maps of ocean circulation came from ship drift calculations, based on speed through the water and heading. Drift bottles, drift cards - released in large quantities in early part of last century through WWII, combined with ship drift calculations (which are still used quite profitably especially given the current excellent state of navigation using GPS satellites). (b) Drift pole. a wooden pole a few meters long and weighted to float with only 0.5--1 meter emergent, is often used to determine surface currents near landmarks. (c) Drift buoy. Extending the drift-pole idea to the open ocean we have the freely drifting buoy with a radio transmitter so that its position can be determined by radio direction finding from the shore, or tracked by the satellite--satellite-tracked buoy. Part of the buoy is above the surface, can be affected by winds. To make sure the buoys drift with the water and to minimize the wind effect, they are frequently fitted with a subsurface drogue to provide additional water drag and more effective coupling with the water motions. The drifting buoys may also be instrumented to measure and transmit surface water properties, atmospheric pressure, etc. Surface drifters with drogues below the surface (``parachutes'') follow the current just below the surface with minimum windage problems. TOGA and WOCE drifters are drogued at 15 m, and use a drogue design which was chosen for its minimum slippage. A portion of drifters is also drogued at about 100-150 m, but it is not clear what they are measuring. (Drifter maps: http://www.aoml.noaa.gov/phod/dac/dacdata.html) Subsurface floats (Lagrangian). Subsurface floats are either tracked acoustically (SOFAR floats which are sound sources and which are tracked by moored receivers, or RAFOS floats which receive sound from moored sound sources) or are tracked periodically by satellite navigation when they pop to the surface. RAFOS are cheaper than SOFAR floats by removing the sound sources. ``Pop up'' floats are cheaper than RAFOS by removing the sound devices. They pop up regularly to communicate with the satellite about their positions, by inflating a bladder and then sink down to a desired level to float. Global

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deployments for WOCE are concentrating on the 800- 1000 meter level. Concentrated deployments of acoustically-tracked floats have been made over the years in the Gulf Stream region and in the North Atlantic Current. (http://wfdac.whoi.edu). Current meters (Eulerian). Current meters are deployed on fixed moorings. Most of them use a rotor and a vane and a compass. The number of turns per minute for the rotor is proportional to the current speed. The current direction is determined by the vane and the compass. Information on current meter moorings recently deployed, and for historical information can be obtained from: WOCE Current Meter Data Assembly Center at Oregon State University. It includes maps, data, and a pointer to a set of averages and statistics maintained in the U.K. (http://kepler.oce.orst.edu). Rotor current Meter (RCM)--Aanderaa RCM. Accuracy: 1--a few cm/s. Within 10% range. Acoustic Doppler Current profiling (ADCP). ADCP on board a ship measures currents relative to a moving ship. It sends out an acoustic pulse which is then reflected back to the ship by particles in the water (such as plankton). The Doppler shift of the returned signal makes it possible to compute the ship's speed relative to the water. There are generally several beams at angles to each other--usually 3-4 beams to determine both speed and direction. Using a 4-element sensor head, an ADCP is capable of resolving both speed and direction of the water movements relative to the sensor. ADCP is originated as doppler speed logs for ships - to measure the speed of the ship through the water. With very precise information from navigation about the ship's speed, heading, and motion, the ship's motion relative to the earth can be subtracted and the speed of the water measured. The range of an ADCP is about 300 meters, depending on the frequency and efficiency of scattering. ADCPs are used in ship mountings, on lowered instrument packages and on moorings as current meters. The acoustic doppler current profiler data assembly center at the U. Hawaii provides online information and data. ( http://ilikai.soest.hawaii.edu/sadcp/index.html.) By controlling the acoustic beams, ADCP can measure currents at different depths below the ship. Moorings: upward and downward looking ADCP. WOCE 150kHz, 75kHZ. Accuracy: 1-a few cm/s, within 10%. Now coastal: 1200kHZ. Accuracy 0.9 cm/s or larger.

How does it work? The ADCP measures water currents with sound, using a principle of sound waves called the Doppler effect. A sound wave has a higher frequency, or pitch, when it moves to you than when it moves away. You hear the Doppler effect in action when a car speeds past with a characteristic building of sound that fades when the car passes. The ADCP works by transmitting "pings" of sound at a constant frequency into the water. (The pings are so highly pitched that humans and even dolphins can't hear them.) As the

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sound waves travel, they ricochet off particles suspended in the moving water, and reflect back to the instrument. Due to the Doppler effect, sound waves bounced back from a particle moving away from the profiler have a slightly lowered frequency when they return. Particles moving toward the instrument send back higher frequency waves. The difference in frequency between the waves the profiler sends out and the waves it receive is called the Doppler shift. The instrument uses this shift to calculate how fast the particle and the water around it are moving.

Sound waves that hit particles far from the profiler take longer to come back than waves that strike close by. By measuring the time it takes for the waves to bounce back and the Doppler shift, the profiler can measure current speed at many different depths with each series of pings. What platforms are needed? ADCPs that are bottom-mounted need an anchor to keep them on the bottom, batteries, and an internal data logger. Vessel-mounted instruments need a vessel with power, a shipboard computer to receive the data, and a GPS navigation system (so the ship's own movements can be subtracted from the current data). ADCPs have no external read-out, so the data must be stored and manipulated on a computer. Software programs designed to work with ADCP data are available. Advantages and limitations? Advantages: (i) In the past, measuring the current depth profile required the use of long strings of

current meters. This is no longer needed. (ii) Measures small scale currents (iii) Unlike previous technology, ADCPs measure the absolute speed of the water, not

just how fast one water mass is moving in relation to another. (iv)Measures a water column up to 1000m long Disadvantages: High frequency pings yield more precise data, but low frequency pings travel farther in the water. So scientists must make a compromise between the distance that the profiler can measure and the precision of the measurements. ADCPs set to "ping" rapidly also run out of batteries rapidly If the water is very clear, as in the tropics, the pings may not��� hit enough particles to produce reliable data. Bubbles in turbulent water or schools of swimming marine life can cause the instrument

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to miscalculate the current. Users must take precautions to keep barnacles and algae from growing on the transducers. 2.1.5 b. Indirect current measurements - geostrophic method. Temperature and salinity are measured to provide density profiles, which can then be used to compute the vertical shear of geostrophic currents perpendicular to the line connecting a station pair. With the assumption of a level of no motion at a specific depth, or with measurements (or inferences) of the absolute velocity at least one level for that station pair, the velocity profile can be constructed for the station pair. Inferences come from mapping of various properties, along vertical cross-sections, or on maps (usually isopycnal surfaces). Tracers with independent sources and sinks are the most useful--these include various salinity and temperature themselves, nutrients, oxygen, chlorofluorocarbons, tritium, helium-3 (with deep hydrothermal sources as well as surface sources), carbon-14,and other tracers. These types of measurements are made from research ships. Temperature profiling is also done regularly from ships of opportunity (including many merchant vessels), using XBT's (see below), providing information on temporal variability. Direct velocity measurements could be those from a large enough set of subsurface floats, or suitably averaged acoustic doppler current profiling simultaneous with the geostrophic measurement.

Geostrophic balance. Coriolis force: Due to the earth's rotation, the Coriolis force acts on a moving body. In the northern hemisphere (NH), the Coriolis force directs toward the right of the motion. In the southern hemisphere (SH), it is to the left of the motion. Pressure gradient force: From high to low pressure. Geostrophic balance. Large-scale ocean circulation obeys geostrophic balance, which is the pressure gradient force balances the Coriolis force.

(1) where is current speed - geostrophic current component that is perpendicular to the forces, the earth's angular speed ( ), is latitude, and is the direction along the two station pairs. We refer as the Coriolis parameter. In the NH, the motion is in the direction with pressure ``high'' to its ``right'' and ``low'' to its left.

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Figure 2.1 is a schematic diagram showing geostrophic method for measuring the geostrophic current. Figure 2.1. Schematic diagram showing geostrophic method.

The above discussion suggests that: p3, which is flat, is a level of no motion. Current derived from the above equation is relative to p3. That is,

€

V (z) =V (r) +1fρ∂p∂x

=V (r) +1fα∂p∂x,

is specific volume, is a reference depth.

To illustrate the derivation process, Figure 2.1a shows a schematic diagram for measurements at stations A and B between two pressure surfaces: P1 and P(r). For a station-pair A and B,

€

V (z) =V (r) +1fαPB (z) − PA (z)

Δx.

€

V (z) =V (r) +1fΦB −ΦA

Δx.

A B x

h1

p1

p2

x PGF CF

Geostrophic current: into the paper. 1000m

h2 p3 V3=0

y

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Note that Boussinesq approximation (density change is neglected in horizontal pressure gradient terms) is used in these two equations. Introducing the concept of

geopotential

€

Φ = gdz0

z

∫ , and thus

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δΦ = gδz, we have:

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PA (z) =1α0

δΦA ,PB (z) =1α0

δΦB , and thus (see Fig. 2.1a):

€

αPB (z) − PA (z)

Δx= α0 ×

1α0

δΦB −δΦA

Δx=(ΦB( p1) −ΦB(r)) − (ΦA (p1) −ΦA (r ))

Δx.

Therefore,

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V (z) =V (r) +1f(ΦB (p1) −ΦB (r )) − (ΦA( p1) −ΦA(r))

Δx.

Based on hydrostatic balance:

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dp = −ρgdz, and

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αdp = −gdz, (note that

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α is not a constant here, based on Boussinesq approximation), we have,

€

αdpp(r )

p1∫ = − gdz

0

h∫ = −(Φp1 −Φp(r )), and

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(Φp1 −Φp(r)) = αdpp1

p(r)∫ .

For simplicity, we denote

€

Φp1 −Φp(r ) =Φ, and

€

ΦA = αAdp,p1

p(r)∫

ΦB = αBdp.p1

p(r)∫

€

(z) =V (r) +1fΦB −ΦA

Δx.

Note that P(z) is the pressure on a geopotential surface (relative to geoid), and is on a pressure surface. We call -``dynamic height'' or ``geopotential distance''. Unit: J/kg-work required for lifting a unit mass from sea level to height z against the force

of gravity. [ ] In oceanography, we often produce ``Dynamic topography map'' to infer the geostrophic currents. It can be obtained by calculating according to observed

profiles. Then we can get the geostrophic current relative to the reference depth. Usually, we assume the reference depth at 500db (~500m), 1000m, etc., as the level of ``no motion'', where .

Sometimes we use ``dynamic meter--dyn m'' to represent . Dynamic meter is D= /10 and 1 geopotential meter is referred to as ``D= /g= /9.8''. 1 dynamic m is

=10J/kg, which is very close to ``1m'', because: 10J/kg=10 N m/kg=10 kg × m/s2 × m/kg=10 m2/s2. /g=1 dynamic meter, then =10 because g=9.8~10. The geopotential meter is a measure of the work required to lift a unit mass from sea level to a height z against the force of gravity. Pressure level 1 dbar is close to 1 m. [ , 1 dbar=104Pascal,

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=1000kg/m3, g=10 m/s2, then z=1 m.]

Fig. 2.1a. Summary of the geostrophic method:

where

€

Φ = αdpp(z )

p(r)∫ ,, is specific volume.

Step 1: Use observed T, S, P derive density and thus at stations A and B; Step 2: Calculate from P(r) to P(z) by integrating

€

Φ = αdpp(z )

p(r)∫ ,

Step 3: Calculate geostrophic current (i) V(r)=0. Reference level is the level of no motion. (ii) V(r) ≠ 0. Then it can be inferred from observations by direct methods.

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We can also use surface currents (say geostrophic current derived from satellite altimetry) as currents at the reference level. Figures 2.2 and 2.3 show the currents obtained from the geostrophic method.

Concepts for baroclinic and barotropic motions.

Fig. 2.2. Dynamic topography of the Pacific Ocean, 0/1000db.

Fig. 2.3. Dynamic topography of the North Atlantic Ocean, 100/700db.

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Baroclinic motion: Depth dependent--vertical shear. Barotropic motion: Depth independent--vertical average of the entire water column. [NOTE: For answering a student’s question: “Near the western boundary, friction becomes important. Can we still use “geostrophic method there?” Note that near the western boundary, friction is important. The alongshore geostrophy, however, still applies, by adopting typical boundary layer assumptions (boundary current is narrow, alongshore component is much larger than the across-shore component). Equations of motion in the boundary layer, after the scale analysis and expressed in coordinates, are

at lowest order. An additional assumption, implicit in the above equations, is that effects of boundary curvature appear at higher order; this is valid, provided the boundary slope changes on a scale larger than the Rossby radius. So, you also see the WBC in the geostrophic current map.] Advantage for geostrophic method: Take advantage of the observed hydrographic data, because direct current measurements are often expensive. Disadvantage-drawback: (i) Assumption of ``level of no motion'', when direct current measurement is not available at the reference level. Although current generally has a small amplitude at the ``reference level'', say 1000dbar, it can affect the vertically integrated ``transport'' calculation by a large amount. (ii) Filter out the ``barotropic mode''--only measures the baroclinic current relative to the reference level. Surface topography derived from satellite altimetry, using geostrophic relation, contains both barotropic and baroclinic modes. Satellite (TOPEX/POSEIDON) altimetry provides a measure of the sea surface height relative to the earth's geoid. The sea surface height measurement is directly related to the pressure and hence to the geostrophic currents at the sea surface. It is sufficiently accurate to provide a measure of the variability of geostrophic currents, and may eventually provide a measure of the mean flow. More complete information can be obtained starting at the (http://diu.cms.udel.edu/woce/dacs.html/satellite). Search ``TOPEX/POSEIDON'' at the web. They only provide surface geostrophic current. Sea level from island and coastal stations. Sea level measurements, although scattered in space, are used as long time series to indicate overall change, and have some limited use to calibrate altimetry measurements. Provide surface geostrophic current.

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2.1.5. c. Integrated observing systems. (http://www.pmel.noaa.gov/tao) TAO/TRITON: In the past few years, a large amount of information has become more readily available online, and in near real-time. These systems are primarily satellite based, but also include the ambitious integrated observing system of the tropical Pacific (TAO array - Tropical Ocean and Atmosphere). The El Nino theme page listed here is a window to many of the NOAA, Navy and other governmental analyses. El Nino theme page: accessing distributed information related to El Nino. TRITON (TRIangle Trans Ocean buoy Network) array: TRITON buoys have been deployed since 1998. (http://www.jamstec.go.jp/jamstec/TRITON/real_time/top.html)

Fig. 2.4a TRITON array in the east Indian Ocean and western Pacific.

Fig. 2.4b. TAO/TRITON array in the Pacific.

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Research Moored Array for African-Asian-Australian Monsoon Analysis and Prediction (RAMA)

Fig. 2.4c Schematic of the IndOOS. Green squares indicate the locations of RAMA moorings. Tide gauges are indicated by blue dots. Argo floats and surface drifters are indicated by a single symbol, although many of each are spread throughout the basin (141 drifters and 296 Argo floats as of 31 Aug 2008). XBT and expendable conductivity/ temperature/depth (XCTD) sections sampled by ships of opportunity are shown as black lines. Most of these lines are sampled 12–18 times per year at along-track intervals of approximately 1° (though the Australia–Sumatra line and the Australia–Mauritius–South Africa are sampled more frequently to measure details of ocean circulation). Nationally sponsored regional observing systems (ROS) are shown in white ovals: IMOS, LOCO, Arabian Sea (ASEA), and Bay of Bengal (BOB). Process studies (PS) are shown in blue ovals: MISMO, VASCO–Cirene, and INSTANT. The satellite in the upper-right symbolizes the constellation of Earth-observing satellites for SST, surface winds, sea level, and other important oceanic and atmospheric parameters. From McPhaden et al. (2009; BAMS) ARGO floats: (http://www.argo.ucsd.edu/) Argo is a "Weather System for the Ocean" that provides realtime ocean temperature and salinity for use in climate and fisheries research, and more. Argo consists of a network of

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oceanic robotic probes covering the Earths oceans, with a total of 3000 probes planned. The probes measure salinity and temperature at depths down to 2 km, surfacing once every 10 days to transmit the collected data via satellite. As of May 2006, 2,453 of the 3000 probes are deployed and active. The program is a collaboration between 50 research and operational agencies from 18 countries, with the United States contributing over half the total funding (as of December 2004). The data collected from the network is made freely available. Argo floats are designed to drift at a fixed pressure (usually around 1000 meters depth) for 10 days. After this period, the floats move to a profiling pressure (usually between 1000 and 2000 meters deep) then rise, collecting profiles of pressure, temperature, and salinity data on their way to the surface. Once at the surface, the floats remain there for under a day, transmitting the data collected by satellite back to a ground station and allowing the satellite to determine their surface drift. They then sink again and repeat their mission. The floats have a nominal lifetime of five years, and will yield valuable information about large-scale ocean water property distributions and currents, including their variability over time scales from seasonal to the duration of the array. The program is named after the Greek mythical ship Argo which Jason and the Argonauts use on their quest for the Golden Fleece. The name was chosen to emphasize the complementary relationship of the project with the Jason-1 satellite altimeter mission.

Fig. 2.4d

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2.2 Description of observed ocean circulation. Based on the observations, now we have a good picture of ocean circulation patterns for the world oceans, especially near the oceanic surface. Next, we will describe the general oceanic circulation in the world oceans. The Pacific (Figure 2.2). Generally, the depth-integrated (or vertical mean) circulation has a similar pattern as the circulation shown in figures 2.2 and 2.3, the geostrophic currents relative to a depth of no motion. The most prominent feature is: a strong subtropical gyre in the northern hemisphere (NH). [Time mean pattern--There are seasonal variations in the strength of the currents.

Figure 2.5. Schematic diagram showing major current systems in the Pacific.

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Figure 2.6. A schematic diagram showing the subsurface equatorial currents. Ignore the “numbers” on the figure. The westward flowing SEC at the surface is not shown. The green and brown color shows eastward flow, with the brown color being stronger than the green. The EIC flows westward.

The strong subtropical gyre (STG) in the NH, consists of North Equatorial Current (NEC) with strongest flow near 15oN, the Philippines Current, the Kuroshio--the strong WBC, the North Pacific Current (NPC), and the California current (see schematic figure 2.5). In the SH, the circulation is weaker but the STG is clear. Associated currents: East Australian current--WBC, South Pacific Current adjacent to Antarctic Circumpolar Current (ACC), Peru/Chile Current, and South Equatorial Current (SEC). In the South Pacific, there is ACC. Near the equator, there are SEC flows westward, North Equatorial Counter Current (NECC) flows eastward near 5oN, and SECC flows eastward near 5oS-10oS. These ``counter'' currents flow against the weak westward wind in the ITCZ (intertropical convergence zone in the atmosphere--Doldrums). The subpolar gyre (SPG)--cyclonic in the NH. Associated currents: NPC, Alaskan Current, Alaskan Stream, southern part of East Kamchatka Current, and Oyashio.

Equatorial Under Current--EUC. Subsurface currents structures that do not show up in the vertically integrated flow in Figures 2.2 and 2.3. Field observations show (Figure 2.6): EUC--a swift eastward-flowing ribbon of water extending over a distance of more than 14,000km along the equator with a thickness of only 200m and a width of at most 400km, with a core at 200m depth. EIC--Equatorial intermediate current, NSCC, and SSCC (North and South subsurface counter current).

Eastern Pacific--the home of El Nino. Normal year--upwelling of cold subsurface water into the surface--Cold tongue due to easterly trades. El Nino year: upwelling is reduced--SST increases. As discussed earlier, large-scale east-west basin extent facilitates strong air/sea interaction.

Pacific-Indian Ocean connection: Indonesian Throughflow.

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Figure 2.7a. The Atlantic. The Atlantic.

There are STGs in both the NH and SH, and Subpolar gyre in the NH. (Figure 2.7) Currents associated with the NH STG: NEC centered at 15N, the Antilles Current to the west and the Caribbean Current through the American Mediterranean Sea, the Florida Current, the Gulf Stream--swift WBC filled with eddies, the North Atlantic Current--Azores Current, Portugal and Canary Currents. The SH STG: SEC which is centered in the SH but extends just across the equator through the western boundary, the Brazil Current, the South Atlantic Current, and the Benguela Current. The SPG: is modified by interaction with the Arctic Ocean, to the extent that it is hardly recognizable as a gyre. It involves North Atlantic Current, Irminger Current, the East and West Greenland Currents, and the Labrador Current, with substantial water exchange with the Arctic Mediterranean Sea through the North Atlantic Current (and its extension into the Norwegian Current) and part of it is separated into the East Greenland Current. Research results suggest that the warmer NAC is important in warming and thus melting the Arctic Sea Ice. Meanwhile, fresher water from the

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Arctic Ocean through EGC (and East Iceland Current) also affects deepwater formation and thus the thermohaline circulation in the North Atlantic Ocean.

Figure 2.7b. Equatorial and south Atlantic. In the Southern Ocean, ACC prevails. Equatorial current systems: SEC and Equatorial Counter Current (ECC) that flows eastward against the weak westward wind. Cross-equatorial currents at the western boundary: North Brazil Current, Guyana Current. EUC: eastward-flowing maximum speed 120 cm/s, core 100m depth, strongest in the west and weakens toward the east. Connection with the Indian Ocean: Agulhas Current. Agulhas Current and eddies transport heat and salt into Atlantic. There may be interaction between AO and IO via coastal Kelvin waves. Circulation has seasonal-interannual-decadal variabilities. Continents bound the E-W Atlantic, both air/sea interaction and continental monsoon effects are important for the variabilities. [Note: Currents derived from CTD data have the largest discrepancies from the depth-integrated currents (due to wind forcing--Sverdrup balance) in the Atlantic ocean (the discrepancy is the largest because of the recirculation of NADW, which destroyed the assumption of ``level of no motion''). Figure 2.3.] The Indian Ocean. The Indian Ocean circulation subjects to strong seasonal variations due to the monsoon wind forcing, which is southwest monsoon (SWM) and northeast monsoon (NEM). Therefore, its annual mean picture does not make much

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sense in the north Indian Ocean. In the south IO, however, circulation is less affected by the monsoon. The STG exists in the south IO throughout the year as in the other two oceans. Associate with it are SEC, East Madagascar Current and Mozambique Current to join the Agulhas Current (WBC), South Indian Ocean current (adjacent to ACC). Leeuwin current in the east, flows against the wind. Fig 2.8.

Figure 2.8. The Indian Ocean current systems. Left: Jan, July. Right: Apr-may (top), and Sep-Oct (bottom). Late Northeast monsoon (Mar-April): NEC, Somali Current, ECC, and South Java Current. There are anti-cyclonic circulations in the Arabian Sea and the Bay of Bengal. Late Southwest monsoon (Sep-Oct): SC (feeds by EACC--east African coastal current) reverses to flow northward (acting as WBC). Southwest monsoon current is in opposite direction as the NEC. SJC also reverses direction to flow northwest (rather than southeast ward). South of 10s circulation remains the same. Bay of Bengal: cyclonic. Equatorial surface jets--Wyrtki jets occur during spring and fall. Swift:100cm/s eastward jets, peak during May and November. Figure 2.9. The EUC exists only during winter--early spring, when surface winds are easterlies as in PO and AO. During other seasons, it disappears unless during the Indian Ocean zonal model years, the EUC reappears during summer-fall. Interacting with Pacific by the Indonesian Throughflow, and with Atlantic by the Agulhas Current.

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Figure 2.9. Indian Ocean currents. The Arctic. The Arctic Ocean belongs to a class of ocean basins know as Mediterranean seas. The Mediterranean Sea is defined as a part of the world ocean which has only limited communication with the major ocean basins (Pacific, Atlantic, and Indian) and where the circulation is dominated by thermohaline forcing. Fig 2.10: A schematic diagram showing the circulation in the Mediterranean seas. (a) E-P>0, salinity increases and thus density increases. Heavy surface water sinks to the bottom. Produce frequent deepwater renewal--oxygen available for marine life.

E-P >0 E-P<0

Pycnocline

O2

Marine life Salinity increases: concentration basin

Marine life Devoid below Pycnocline. Dilution basin.

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Since salinity increases as it passes through the Mediterranean Sea, it is also known as a concentration basin. Circulation: Surface water flows into the Mediterranean Sea from the adjacent oceans through the sill. At lower-layer the Mediterranean water flows out of the basin into the adjacent ocean, where it sinks to the density level it belongs. (b) (E-P<0), due to rainfall and land run-off, deep water renewal is inhabited. Inflow of oceanic water is usually the only renewal process of significance. If the sill is narrow or the deepwater volume is large, deepwater renewal is not always sufficient to prevent the depletion of oxygen in the deep basins. There is a strong pycnocline. Most life is devoid below the pycnocline. It is known as a dilution basin. Circulation: Surface fresher water is fresher due to precipitation. It flows out of the Mediterranean and enters the adjacent ocean. At lower layer, the ocean water flows into the Mediterranean Sea. The Arctic Ocean is a dilution basin. Observations in the Arctic ocean is not worse than the southern ocean where research vessels are difficult due to the lack of islands to fuel the ships. Arctic: collection of rainfall and snowfall data on the stable platforms of drifting ice stations is available. Precipitation is low in the region of the Polar area but significant in the subpolar regions which are dominated by the West Wind Drift and its associated high storm frequency. The major contribution of polar mass comes from precipitation over Siberia and the resulting river run-off, estimated in total as 0.2Sv (Sverdrup; 1 Sv=106 m3/s). [Others precipitate on ice and export out so that they do not contribute to the mass budget in the Arctic.] With evaporation over ice being comparatively low, the Arctic Mediterranean Sea is a dilution basin: its outflow is fresher on average than its inflow. See figure 2.11 for surface circulation pattern and 2.12 for sea ice coverage and export.

Figure 2.11. Arctic circulation.

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Figure 2.12. Sea ice extent.

Thermohaline circulation--concept. The circulation patterns we previously discussed in the PO, AO and IO are primarily wind-driven. How deep does the wind-driven circulation extend in the interior of the North Pacific's subtropical region? Using patterns of properties on isopycnals, it is possible to trace a subtropical gyre down to about 2000 meters, with poleward shrinkage throughout this depth. The western boundary current can go deeper than 2000m. As we just discussed, circulation in the Arctic Ocean is primarily thermohaline driven, as any other Mediterranean seas in the world oceans (such as the Mediterranean Sea in the North Atlantic; Red Sea and Persian Gulf in the IO).

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Thermohaline circulation is the movement of water when its density is changed by a change of temperature or of salinity. Examples: change of salinity--Arctic Ocean and the Mediterranean Sea. Change of temperature: North Atlantic. (i) The case in the Arctic Ocean: Fresher river water due to precipitation exceeds evaporation drives the outflow and thus exporting of sea ice into North Atlantic at the surface, and the North Atlantic water flows into the Arctic beneath the outflow. It is a dilution basin. (Salinity decrease causes density decrease). Surely, the Arctic sea ice motion is also significantly affected by the wind forcing. (ii) The Mediterranean Sea in the North Atlantic, the Red Sea and Persian Gulf in the IO are all concentration basins. Their E is much larger than P and thus salinity increases, causes density increase near the surface. The heavy near surface water sinks to the bottom, and flows out of the basin into the Adjacent oceans. Thus, thermohaline circulation can be much deeper than the wind driven circulation, depending on the density of deep water that is formed at the oceanic surface. At the surface, the ocean water flows into the Mediterranean Seas to compensate for the water loss due to evaporation. [Circulation within these basins are also significantly affected by the wind forcing.] (iii) North Atlantic. The two cases discussed above represent the density change due to ``salinity'' change: either decrease or increase. In the North Atlantic, the strong surface cooling can cools the SST considerably and thus increases the density. An example is in the Labrador Sea, where surface cooling increases the density and thus causes ``convection'' in the ocean, contributing to the formation of the North Atlantic Deep Water (NADW). The NADW flows southward out of the Atlantic, which is important for world ocean thermohaline circulation. 2.3. Water masses and T-S diagram. 2.3.1. Water masses Water mass is body of water with a common formation history, having its origin in a physical region of the ocean. Water masses are physical entities with a measurable volume and therefore occupy a finite volume in the ocean. In their formation region they have exclusive occupation of a particular part of the ocean. Elsewhere they share the ocean with other water masses with which they mix. The total volume of a water mass is given by the sum of all its elements regardless of their location (Tomczak 1999). Atlantic water masses. (i) Abyssal water masses.--Formed by thermohaline process. There are five sources of Abyssal Atlantic water masses, classified in a bit of an over simplification as: Antarctic Intermediate Water (AAIW), Antarctic Bottom Water (AABW), Mediterranean Overflow Water (MOW), Labrador Sea Water (LSW) and Nordic Sea Overflow Water (NSOW), and the last three contribute to the NADW,

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which fills the depth range between 1000m and 4000m. Two of these are the original southern sources, and the other three are the North Atlantic sources.

The deep thermohaline circulation of the Atlantic involves (i) flow of waters from the southern hemisphere into the North Atlantic, (ii) modification and convection of waters in the North Atlantic and its adjacent seas, and (iii) outflow in a thick deep layer--NADW. This deep layer affects the world ocean, where it can be tracked through its high salinity signature since the North Atlantic is the most saline of all the oceans. The deep layer flowing out of the North Atlantic is called NADW and is notable for its vertical salinity maximum, vertical oxygen maximum (actually two) and vertical nutrient minima. [Figure 2.13: y-z section at 25W for salinity in the Atlantic Ocean] Figure 2.13: Meridional section of CTD salinity at 25W of the Atlantic Ocean. Orange color: high salinity waters; blue: low salinity waters. It shows salinity maximum of MOW (30-40N at 1000m), salinity minimum of Labrador Sea Water (LSW; 40-60N at 1500-2000m). It also shows salinity minimum of AAIW (south of 20N at 500-1000m) and overall salinity maximum of NADW (south of 20N and 1500-3000m).

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Southern source: AAIW and AABW. Inflow to the North Atlantic occurs in two layers - a relatively warm one and a cold one, sandwiching the outflowing NADW. The upper layer is composed of thermocline and Antarctic Intermediate Water (evidenced by a salinity minimum in the vertical at about 800 m) AAIW: depth (500m--1000m), resulting from its source at the sea surface near the tip of South America in the Antarctic Polar Front zone, which is normally between 50S-60S. The formation of AAIW can be explained very simply through the Ekman transport process and the divergence and convergence of water masses. The winds over Antarctica are called the polar easterlies where winds blow from the east to the west. This creates a counter-clockwise surface current near the coast of Antarctica, called the Antarctic Coastal Current. Ekman transport causes the water to push towards the left of the surface motion in the Southern Hemisphere. Thus, this westward directed coastal current in Antarctica will push the water towards Antarctica. At the same time there is a strong current north of the Antarctic Coastal Current, called the Antarctic Circumpolar Current (ACC) created by the strong westerlies in this region which flows clockwise around Antarctica. Again, Ekman transport will push this water to the left of the surface motion, meaning away from Antarctica. Because water just offshore of Antarctica is being pushed away and into Antarctica, it leads to the Antarctic Divergence region. Here, upwelling of North Atlantic Deep Water (NADW) takes place. Some of this water diverges northward and is heated and loses salinity due to precipitation. When this water moves northward to encounter the substantially warmer and buoyant Subantarctic waters, it sinks to form AAIW. The AAIW can be tracked as a salinity minimum up to the subtropical gyre in the North Atlantic and North Pacific. The other southern source for Abyssal Atlantic Water is at the ocean bottom - the Antarctic Bottom Water (AABW). This is actually deep water (not bottom water) from the South Atlantic sector of the Antarctic. (The true bottom waters, formed in the Weddell Sea and Ross Sea, are thought not to escape very far northward, mainly because of topography that confines them. The Antarctic Bottom Water in the South and North Atlantic is essentially the same as what we call Circumpolar Deep Water in the Pacific and Indian Oceans.) Some research, however, suggests that some Ross sea deep water can be traced to the central south Pacific. The AABW is apparent as a cold, lower salinity bottom layer and extends northward in the North Atlantic up to the Gulf Stream latitude. Below 4000m depth, all Atlantic Ocean basins are occupied by AABW. It upwells into the southward-flowing NADW layer above it (and hence does not reach the sea surface). Northern Source for Abyssal Atlantic Water: MOW, LSW, NSOW--form NADW. In the North Atlantic, the components of NADW are formed at three sites, all involving intermediate depth convection: inside the Mediterranean Sea, in the Labrador Sea and in the Greenland Sea. The NADW occupies 1000m-4000m. The MOW. The Mediterranean Sea is connected to the North Atlantic through the narrow Strait of Gibraltar. Flow is into the Mediterranean at the sea surface in the Strait. Within the Mediterranean there is large evaporation and cooling and production of dense water. This flows out into the North Atlantic at the bottom of the Strait. The

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resulting Mediterranean Water (or Mediterranean Outflow or Overflow Water) in the North Atlantic is found at mid-depth and is marked by its salinity maximum both in the vertical and in the horizontal along isopycnals. The MOW forms the upper part of the North Atlantic Deep Water. (In the tropical Atlantic, Wust referred to the salinity maximum core of the NADW as Upper North Atlantic Deep Water - UNADW clearly originates as MOW.) MOW properties at Gibraltar are 13C, 38.45 psu and 29.07 sigma theta. This is actually denser than the bottom water of the North Atlantic. As the plume of MOW exits Gibraltar, it moves to the north (boundary to the right, in the Kelvin wave sense) and down the slope. Vigorous entrainment as it moves down the slope reduces its high salinity and density. It finally equilibrates after mixing at a depth of about 1000 m. LSW and NSOW. The two northern parts of NADW are Labrador Sea Water and Nordic Sea Overflow Water (also Arctic Bottom Water--ABW). (The latter is also referred to by its three separate components resulting from overflows at three separate sites into the North Atlantic, and also sometimes as Greenland-Iceland-Norwegian Sea overflow water or GIN-Sea overflow.) In undiluted form NSOW is restricted to the immediate vicinity of Greenland-Iceland-Scotland Ridge. Its major contribution is to the formation of the lower-part of the NADW. Labrador Sea Water is formed in the western Labrador Sea through convection to about 1500-2000 meters in late winter. This forms a relatively homogeneous water mass within the Labrador Sea. Much attention has been focused in recent years on both the formation of LSW and on its changing properties (temperature changing from 3.5C to about 2.9C on decadal time scales). Labrador Sea Water spreads out into the North Atlantic, filling both the subpolar gyre and entering the subtropical gyre. Within the subpolar gyre, it is marked by a salinity minimum in the vertical. Within both the subpolar and subtropical gyres it is marked by an oxygen maximum in the vertical. Within both gyres it is also marked by a thickness maximum resulting from its convective source. Figure 2.14.

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Figure 2.14. Zonal salinity section at 47N of the Atlantic Ocean.

LSW moves southward along the western boundary of the North Atlantic as the upper part of the Deep Western Boundary Current, above the denser water that originates in the Nordic Seas. The oxygen maximum of the LSW persists to the tropical Atlantic where Wust referred to it as ``Middle North Atlantic Deep Water''. The densest part of the NADW is formed through convection in the Greenland Sea offshore of the East Greenland Current and ice edge. This portion of NADW has many different names, but we will call it Nordic Sea Overflow Water. Flow into the Nordic Seas is in the Subpolar Mode Water layer, east of Iceland. The Greenland Sea convection is usually to intermediate depths, and is colder and denser than convection in the Labrador Sea, hence producing the denser part of the NADW. (Field experiments, involving acoustic tomography to track convection through the winter, have examined the Greenland Sea convection. Much attention has been given to climatic variations in convection in the Greenland Sea.) Outflow of the dense NSOW occurs in three locations along the Greenland-Faroe ridge: through Denmark Strait (between Greenland and Iceland), across the Iceland-Faroe ridge (between Iceland and the Faroe Islands), and through the Faroe-Shetland

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channel (between the Faroe and Shetland Islands). All three locations are relatively shallow sills. The dense water flowing southward over the sills plunges downward in plumes, entraining (mixing with) surrounding water in the process. This modifies the properties of the NSOW as it enters the North Atlantic. The deep part of the NADW thus produced as NSOW fills the western North Atlantic through the lower part of the Deep Western Boundary Current. The DWBC crosses under the Gulf Stream. The presence of the DWBC is not obvious in isopycnal slopes crossing the Gulf Stream. Now that we have mapped many different properties, the DWBC is apparent as a core layer of higher oxygen. The DWBC was predicted in a theory by Stommel and Arons, and then its presence was detected through the earliest use of deep tracked floats. The DWBC continues to the South Atlantic. One feature of the DWBC is that is exhibits recirculations to the north on its offshore side. These spread the DWBC waters into the ocean interior. Transports for the various components of the NADW overturn have been computed. The net transport involved in the overturn is 15 to 20 Sv. Part of the NADW flows out of the Atlantic, joins the ACC and is an important part of the global ocean circulation.

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Figure 2.15. Subpolar Mode water (SPMW). The two northern sources of NADW are formed from surface waters that flow northward in the Gulf Stream and North Atlantic Current and then eastward and northward in the subpolar region. These surface waters cool and freshen along this path towards the ultimate intermediate-depth convection regions. Cooling of the subpolar surface waters creates very thick mixed layers in the subpolar gyre. These thick layers are called Subpolar Mode Water, in analogy with the Subtropical Mode Water of the subtropical gyres, as will be discussed next. SPMW temperature range from about 14C near the North Atlantic Current, to 8C where SPMW enters the Norwegian Sea, to 4C where SPMW enters the Labrador Sea. Ventilation of the subtropical gyre--Water masses in the thermocline and surface layer: thermohaline and wind effect. Thermohaline effect. As sea surface temperature generally decreases poleward (Fig. 2.16), surface density increases poleward. Therefore, surface waters in the northern part of the subtropical gyre are denser than in the southern part. As the anticyclonic circulation advects the higher density waters southward, they must either change to lower density due to surface heating or slide below the less dense surface waters to the south. Since heat fluxes in the eastern parts of subtropical gyres, outside of the upwelling zone in the eastern boundary currents, are usually quite small, generally the surface waters slide down. This process has been called subduction. Consequently, thermocline waters in the tropics have the same density as the surface water in the subtropics. The subtropical gyre is thus ``ventilated''. We also call it thermocline ventilation.

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Figure 2.16. Annual mean global SST.

Subtropical Mode Water (``Eighteen Degree Water''). The isopycnal bowl immediately south of the Gulf Stream created by the Gulf Stream itself and its recirculation south of the Gulf Stream is a site of wintertime convection. Convection is favored here for two reasons - the isopycnal bowl has reduced vertical stratification compared with other regions simply because of the bowl, and heat losses to the atmosphere in this region are very large because of the conjunction of the warm Gulf Stream waters and cold, dry air blowing off the North American continent. The convected water mass is called ``Eighteen Degree Water'' because of its dominant temperature. It is the Subtropical Mode Water that is associated with the Gulf Stream. (There is an STMW for each of the subtropical gyres' western boundary currents.) The Eighteen Degree Water is spreaded by the circulation into the whole of the western subtropical gyre, even though the only location where it outcrops is close to the Gulf Stream. ``Mode'' means relatively large volume on a volumetric T-S diagram. Wind effect: Central Water and Subtropical Underwater. Subduction in the Ekman convergence region of the STG moves water from the sea surface southward into subsurface layers of the subtropical gyre. Central Water is the general name for the whole of the thermocline. It was suggested that the subtropical thermocline properties are set by this subduction process, which was further elaborated by Stommel and then received its name and formal theory from Luyten, Pedlosky and Stommel (1983). The surface waters of the central subtropical gyre are very salty as a result of evaporation under the atmospheric high-pressure region. As this salty water subducts southward beneath the water that is not quite as saline, it forms a salinity maximum in the vertical. This salinity maximum is typical of every subtropical gyre, and is sometimes called Subtropical Underwater. The Pacific. As in the Atlantic, the "water masses" which are created by subduction are the Central Water (water in the thermocline, spreads along a continuous temperature/salinity range), and Subtropical Underwater (salinity maximum arising

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from subduction of the very high central subtropical gyre waters beneath the fresher waters which lie to the south). Subtropical mode water. In the western subtropical North Pacific, the main thermocline (pycnocline) is interrupted by a ``thermostad'' (pycnostad), which is a region of lower vertical gradients of temperature, salinity and density, compared with the thermocline above and below. Such a thermostad is typical of the major subtropical western boundary current recirculation regions in each ocean. This pycnostad in the North Pacific is referred to as "Subtropical Mode Water". The STMW in the North Pacific is in the temperature range of 16-19C and is found just south of the Kuroshio Extension. Ventilation of the Subpopar gyre. Sea ice formation in the Okhotsk ventilates the upper portion of the intermediate density layer of the North Pacific. Vertical mixing can be shown to be locally very important for extending the ventilated waters downward in the water column, especially along the Kuril Islands. The ventilated waters of the western subpolar gyre enter the subtropical gyre mainly along the western boundary, probably as meanders and intrusions from the separating Oyashio. Because they are quite fresh, they appear as a salinity minimum in the subtropical gyre. This water or the full intermediate layer is often called North Pacific Intermediate Water (NPIW). Deep waters of the North Pacific. Below the intermediate, ventilated layer lies the nearly homogeneous deep water layer, between about 2000 and 4000 meters. Its origin is basically upwelling of the southern source bottom waters (sometimes known as Circumpolar Water). This is the oldest deep water in the world ocean, and is fairly well mixed. So Pacific does not produce its own DEEP water, unlike the Atlantic. Additionally, the Bottom Water produced in the Weddell and Ross Seas do not enter the north Pacific Ocean, due to the bottom topography--the Pacific-Antarctic Ridge. Some of the Ross sea water, however, is suggested to flow to the central South Pacific. The Indian Ocean. Abyssal water mass. AABW fills the Indian Ocean below about 3800m (many authors call it Circumpolar Water). In the Atlantic it is called AABW. So here we call it AABW. From 3800m to about 1500-2000m is occupied by Indian Deep Water (IDW). Based on water mass properties the transition from Bottom to Deep Water is gradual, and some authors refuse to use Bottom and IDW, just say lower and upper deep water. The Red Sea and Persian Gulf waters are high T and high S Mediterranean water in concentration basins. The former has a core around 400--700m and the latter around 200-300m in the Arabian Sea. Recent studies suggested that the RSW flows out of the IO along the western boundary of the IO, and the PGW seldom gets out of the Arabian Sea. Just mixes with the Arabian Sea water. Water masses of the thermocline and surface. Indian Central Water (ICW) as in the other two oceans. ICW is a subtropical water mass formed and subducted in the

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Subtropical Convergence. Surface water masses: Bay of Bengal water--fresh due to river runoff. 2.3.2. T-S diagram. Plots of salinity as a function of temperature are called T-S diagram. Here is why the plots are so useful: water properties, such as temperature and salinity, are formed only when the water is at the surface or in the mixed layer. Heating, cooling, rain, and evaporation all contribute. Once the water sinks below the mixed layer, temperature and salinity can change only by mixing (with adjacent water masses). Thus water from a particular region has a particular temperature associated with a particular salinity, and the relationship changes little as the water moves through the deep ocean. Water masses can therefore be identified by their temperature and salinity (T-S) combinations. All other properties of sea water such as oxygen, nutrients etc. are affected by biological and chemical processes and therefore non- conservative. Figure 2.17. T-S diagram for observed profile at 9S of the Atlantic Ocean.

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Figure 2.17