Chapter 2: Ocean Observations

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Chapter 2. Ocean observations

2.1 Observational methods

With the rapid advancement in technology, the instruments and methods for measuring oceanic circulation and properties have been quickly evolving. Nevertheless, it is useful to understand what types of instruments have been available at different points in oceanographic development and their resolution, precision, and accuracy. The majority of oceanographic measurements so far have been made from research vessels, with auxiliary measurements from merchant ships and coastal stations.

Fig. 2.1 Research vessel.

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 deployed, we need to measure ``depth''. Depth measurements are often made with the measurements of other properties, such as temperature, salinity and current.

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, research vessels are moving and currents might be strong, and thus 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

1 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. Pairs of reversing thermometers carried on Nansen bottles were the primary source of subsea measurements of temperature as a function of pressure from around 1900 to 1970. 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%. (iii) Quartz crystal: Very accurate pressure measurements can be made using a quartz crystal, whose frequency of oscillation depends on pressure. This technology is used in modern CTDs. Temperature must be accurately measured for the best pressure accuracy. In CTDs, a thermistor is part of the quartz pressure transducer. The accuracy is _0.01% and precision is _0.0001% of full-scale values. (For more details: See chapter S16 of Talley textbook.)

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 angle 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.

(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 seawater 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), 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 Thermometry of Ocean Climate (ATOC): Acoustic thermometry maps changes in ocean temperature using changes in sound speed along paths between acoustic sources and receivers. It can be used to monitor changes in the average temperature along very long paths at basin or global scale. It has been applied to measure the thermal field in the North Pacific basin from 1996-2006. See http://staff.washington.edu/dushaw/atoc.html: “ATOC is directed at using the travel time data obtained from a few acoustic sources and receivers located throughout the North Pacific basin to study the climatic variability of the thermal field at the largest scale.” They concluded that “…the experiment was a success, with transmissions occurring between 1996 and 2006…The marine mammal/biology problem was formally determined based on extensive scientific studies to be not significant for the acoustic sources employed by ATOC.”

Fig. 2.2 Acoustic sources & receivers in the North Pacific.

(h) Drifting buoys: 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 (i.e., without losing or gaining 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 solely due to compression or expansion is usually not of much interest to climate scientists, because 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 the effect of adiabatic compression/expansion.

Definition:``Potential temperature'' is the temperature that a water parcel has when it moves adiabatically to a reference 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 to the sea surface adiabatically without mixing or diffusion. Since pressure is the 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.

Note that when oceanic mixing occurs (without change of external heat flux, e.g. heat flux at air-sea interface), the temperature (or potential temperature) of the mixture of two water parcels with different T &S does not equal to the average temperature of the two original water parcels, while heat content remains the same. This is because heat capacity (specific heat) also varies with varying temperature and salinity values. For this reason, “conservative temperature” is defined in TEOS-10, which more precisely scales with heat content and insensitive to pressure (also see: https://www.nature.com/scitable/knowledge/library/key-physical-variables-in-the- ocean-temperature-102805293/). The difference between potential temperature and conservative temperature is usually well within ±0.05ºC for most ocean waters (although the difference can be large for warm fresh waters). Application of “conservative temperature” is not required in this course.

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 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.

TEOS-10 (Thermodynamical equation of seawater 2010) proposed new standard: Use Practical Salinity (SP) measured by conductivity to derive Absolute salinity (SA), which takes into account of the varying compositions of seawater (i.e., compositions of salinity vary in different regions of the world’s oceans) instead of constant compositions. SA has a unit of g/kg which measures the dissolved materials in seawater (salt content). See lecture 4 file for the relationship between SP and SA. In this course, details about SA & its calculation are not required. However, in your homework please spell out explicitly whether you are using practical

salinity (PSU) or absolute salinity (g/kg).

(d) Satellite. NASA Aquarius mission: Measure sea surface salinity. The Aquarius satellite was launched in June 2011 and ended in 2015. It provided weekly and monthly sea surface salinity (SSS) from September 2011-April 2015. The NASA’s soil moisture active passive (SMAP) mission also measures SSS, and its version 3.0 products provide 70km and 40km resolution SSS since April 2015. The ESA’s Soil Moisture and Ocean Salinity (SMOS) satellite provides 9day and monthly SSS since January 2010. Accuracy: Aquarius SSS is of the best quality: ~0.2psu; SMAP & SMOS ~ 0.21-0.23psu in open ocean between 40S-40N.

Bao, S., Wang, H., Zhang, R., Yan, H., & Chen, J. (2019). Comparison of satellite‐ derived sea surface salinity products from SMOS, Aquarius, and SMAP. Journal of Geophysical Research: Oceans, 124. https://doi.org/10.1029/ 2019JC014937

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.

The goals for measuring large-scale circulation is to understand the three dimensional circulation and variability. They are directly related to heat and salt transports and thus are important for understanding climate variability and climate change. 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, and the 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 currents. 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) Drifting 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) Drifting 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 are also drogued at about 100-150 m, but it is not clear what they are measuring. (For global drifters’ data, see: 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 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, 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 at the data center of the IPRC, the University of Hawaii (http://apdrc.soest.hawaii.edu/datadoc/woce_cm.php). Rotor current Meter (RCM) Aanderaa RCM. Accuracy: 1--a few cm/s. Within 10% range.

Acoustic Doppler Current profiling (ADCP). ADCP carried by 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

7 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.

For current measurements, 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 (https://uhslc.soest.hawaii.edu/sadcp/). By controlling the acoustic beams, ADCP can measure currents at different depths below the ship. For moorings: upward and downward looking ADCP are mounted to measure currents above and below the ADCP mounted depth. 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 currents with sound waves, using a principle of the Doppler effect. A sound wave has a higher frequency (or pitch) when it moves toward 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 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 receives 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,

8 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 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 geostrophic velocity profile can be constructed for the station pair.

Inferences of currents a specific depth 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 XBTs, 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 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: Direct from high to low pressure. Geostrophic balance. Large-scale ocean circulation obeys geostrophic balance, which is the pressure gradient force balances the Coriolis force. −�� + = 0, (1a) �� + = 0. (1b) In the above, (u,v) are zonal and meridional components of geostrophic current, is the Coriolis parameter, is the earth's angular speed ( ), is latitude, and (x,y) are zonal and meridional axises of cartesian coordinate system. In the NH (SH), the motion is in the direction with pressure ``high'' to its ``right'' (left) and ``low'' to its left (right). (Equations (1a) and (1b), together with the hydrostatic equation shown below, will be derived later in chapter 3.)

x h1 B A

x PGF CF p2

V3=0 p3 1000db

Geostrophic current: into the paper.

Figure 2.3a: schematic diagram showing geostrophic method.

Equations (1a)-(1b) are in z coordinate (units: m). In physical oceanography, pressure coordinate is often used instead of z vertical coordinate (note: pressure is also directly measured, e.g., CTD). Below, we derive equations (1a)-(1b) in pressure coordinate.

Fig. 2.3b. Schematic diagram illustrating pressure surfaces in z coordinate.

From Fig. 2.3b, we have:

() () [ ] = [ ] ( ) . For infinitely small ��, the above equation yields, ( ) = - ( ) ( ) . Using hydrostatic equation (i.e., hydrostatic balance showing upward pressure gradient force is balanced by downward gravitational force), =-�� (where g=9.8m/s2 is acceleration of gravity), and times 1/� on both sides of the above equation, we obtain

( ) = - ( ) ( ) =- (-�g) ( ) = ( ) = ( ) , and � = �� is “geopotential”. Thus, ( ) =( ) (1c) . Similarly, we can obtain ( ) =( ) (1�) .

Using ( ) to replace ( ) in equation (1a) yields, −�� + ( ) = 0. (2a) Similarly, we can obtain: �� + ( ) = 0. (2b) Equations (2a) and (2b) are geostrophic balance in pressure coordinate. By manipulating equations (2a)-(2b), we obtain zonal and meridional components of geostropic currents:

� = � (�� ) , (3a) � �� = - � (�� ) . (3b) � ��

If is known Φ, we will be able to obtain geostrophic currents from equations (3a)-(3b). How is Φ = gz related to hydrographic observations (T, S, p)? As mentioned earlier, since many hydrographic observations are available in the upper ocean, and geostrophic method derives current shear between one pressure level (denoted p1) and a reference pressure level p(r): v1 − v(r). Introducing geopotential anomaly �Φ = �� and use hydrostatic balance = −��, we have �� = −�� = −�d� = − d�, and thus d� = −��, (�) where � = 1/� is specific volume. From equation (4) we obtain geopotential anomaly between p1 and p(r) ∆Φ at station A & station B, () �(�) ∆�� = �� = −�� = ��. (��) () �(�) � �(�) �(�) ∆�� = �� = −�� = ��. (��) � �(�) �(�)

Since the geostrophic method calculates the geostrophic currents between station pairs, we use equation (6) below for our discussion. Geostrophic current shear between p1 and p(r) using hydrographic observations at Stations A and B is:

V1 – V(r)= � (�� ) = � ∆��∆�� . (�) � �� ∆�

Geostrophic currentIn equation shear (6) ,between ∆�� and p1∆� and� can p(r) be calculatedusing observations from equations at (5a) and (5b); f is the Coriolis parameter at the geostrophic current location (mid-point between A and B), and Stations A and∆ �B is is: the distance between A and B. Note that equation (6) only provides the ! #$ ! ∆( )∆( V1 – V(r)“magnitude= ( ” of) geostrophic= current! shear." Regarding(6) its direction, the geostrohic current is perpendicular" #% &to line" AB and∆ %flows in the direction with high Φ values to its right (left) in NH (SH).

p1 Fig. 2.3c. Schematic diagram A B showing isobaric surfaces p1 & ∆& P(r), z levels and distance between stations A and B.

P(r) Note: • f in equation (6) should be the Coriolis parameter at the location of current V1, which is in the middle of A & B; • Equation (6) only provides the “magnitude” of V shear. The direction is determined using the rules stated on Slide 16. 12

Usually, we assume the reference depth at 500db, 1000db, etc., as the level of ``no motion'', where . Fig. 2.3d below provides a schematic illustration. p1

A B

x CF PGF z

z y

0 x V(r) P(r) 1000db (~1000m)

Geostrophic current direction: flows into the paper.

Fig. 2.3d. Northern Hemisphere (NH).

In physical 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.

Sometimes we use ``dynamic meter-dyn m'' to represent . Dynamic meter is D= /10 and 1 dynamic meter is =10m2/s2=10J/kg. Geopotential height is �(�) = = . . Values of D and Z are close, and 1 Dynamic meter is very close to ``1m''. Pressure level 1 dbar is close to 1 m. Modern publications: D and Z are often used interchangeably, differing only by ~ 2%.

Summary of using geostrophic method to calculate geostrophic current with hydrographic observations at station A and station B:

Step 1: With observed values of T, S and p at station A and station B, calculate density � and � using equation of state (will be given if appears in exams or homework); then obtain � = and � = ;

Step 2: Obtain ∆�� and ∆�� by integrating � and � from a pressure surface P(z) to reference pressure level P(r):

�(�) ∆� = � ��, � ∫�(�) � �(�) ∆�� = ∫�(�) ��.

Step 3: Calculate geostrophic current shear using V1 – V(r)= � ∆∆. � ∆� V(r)=0 if assumes level of no motion, or V(r) is known from direct observations or inference.

Oceanographers often use specific volume anomaly � by removing a constant value, to achieve more accurate calculation of the geostrophic current. � = �(�, �, �) − �(35,0, �), and () ∆� = ��. ∫() In this class, using � is not required.

Figures 2.4 and 2.5 show the currents obtained by geostrophic method.

Fig. 2.4. Dynamic topography of the Pacific Ocean, 0/1000db based on 36,356 observational profiles. Units: dyn-cm

From Klaus Wyrtki, 1961: Fluctuations of dynamical topography in the Pacific Ocean. J. of Physical Oceanography.

Fig. 2.5. Dynamic topography of the North Atlantic Ocean, 100/700db. Adapted from Pickard and Emery.

Concepts for baroclinic and barotropic motions.

Baroclinic motion: Depth dependent--vertical shear: vertical integral from ocean bottom to surface equals 0. Barotropic motion: Depth independent--vertical average of the entire water column.

NOTE: Near the western boundary, friction becomes important. The alongshore geostrophy, however, still holds, 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. This will be further demonstrated in later chapters.

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/Jason) altimetry provides a measure of the sea surface height relative to the earth's geoid since October 1992. 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. They only provide surface geostrophic current. Satellite altimetry is viewed as the most successful satellite mission, which provides continuous global sea level data for about 3 decades (~27yrs up to 2019).

Sea level data from island and coastal tide gauge stations, although scattered in space, are used as long time series to indicate overall change, and have some limited use to calibrate altimetry measurements. They can be used to calculate surface geostrophic current along coasts.

2.1.6 The integrated observing systems In recent decades, a large amount of information has become more readily available online and in near real-time. These systems are primarily satellite based, such as satellite SST (section 2.1.2), sea surface salinity (section 2.1.3), satellite ocean surface topography mission that measures sea surface height (https://science.nasa.gov/earth- science/oceanography/physical-ocean/ocean-surface-topography) (e.g., Topex/Poseidon & Jason series) and others. However, integrated observations also include in-situ observing system, such as the Global Tropical Moored Buoy Array: (https://climatedataguide.ucar.edu/climate-data/tropical-moored-buoy-system-tao-triton- pirata-rama-toga).

The Pacific: TAO/TRITON (Tropical Atmosphere-Ocean array/Triangle Trans-Ocean Buoy Network) array. The El Niño theme page listed here is a window to many of the NOAA, Navy and other governmental analyses. El Niño theme page: accessing distributed information related to El Niño (https://www.pmel.noaa.gov/elnino/). The moorings observe temperature, salinity, and some also observe current Profiles (i.e., T, S and V as a function of depth z). Meanwhile, they also measure surface atmospheric variables (surface air temperature wind, relative humidity, shortwave and longwave radiation & sea level pressure). TAO array played an important role in monitoring El Niño since the 1980s.

TRITON (TRIangle Trans Ocean buoy Network) array: TRITON buoys have been deployed since 1998, most of them are in the western Pacific but two are in the east Indian Ocean (http://www.jamstec.go.jp/jamstec/TRITON/real_time/top.html).

Fig. 2.6. Tropical moored buoy array in the three world’s oceans.

Fig. 2.7. ATLAS (Autonomous Temperature Line Acquisition System) mooring for TAO array: measuring surface atmospheric fields and oceanic properties (e.g., T, S, V) at various depth.

T-Flex (Tropical-Flex) moorings: technology advancement: allow greater flexibility with respect to number and types of instruments that may be deployed, and utilize more commercial off-the-shelf instrumentation.

For details about ATLAS vs T-Flex moorings please see: https://www.pmel.noaa.gov/pubs/PDF /frei4747/frei4747.pdf

In the tropical Atlantic, PIRATA (Prediction and Research Moored Array in the Tropical Atlantic) array was first established in the mid-1990s: https://www.pmel.noaa.gov/gtmba/pmel-theme/atlantic-ocean-pirata

The Indian Ocean Research Moored Array for African-Asian-Australian Monsoon Analysis and Prediction (RAMA)array was the newest tropical buoy array, which was initiated in 2004 (https://www.pmel.noaa.gov/gtmba/pmel-theme/indian-ocean- rama).

Fig. 2. 8. RAMA array and integrated observing systems over the Indian Ocean. Schematic of the Indian Ocean Observing System (IndOOS). From McPhaden et al.(2009; BAMS).

ARGO floats: (http://www.argo.ucsd.edu): Argo is a “Weather System for the Ocean” that provides real time ocean temperature and salinity for use in climate and fisheries research and more. Argo consists of a network oceanic robotic probes covering Earth’s oceans. 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 August 30, 2019, there were 3854 floats in the ocean (Fig. 2.9).

Fig. 2.9. Argo floats in the world’s oceans on Sep 8, 2020.

The floats drift at a fixed pressure (usually around 1000 meters depth) for ~9 days. After this period, the floats move to a profiling pressure (usually between 1000 and 2000

18 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 the large-scale ocean water property distributions and currents, including their variability over time scales from seasonal to the duration of the array. Argo data have good global coverage since 2004.

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.

CPIESs and Underwater Gliders – newer technology (IndOOS decadal review: Chapter 24: http://www.clivar.org/sites/default/files/IndOOS%20Review%20White%20Paper_2nd%2 0draft%20combined_20180110_clear.pdf) Current and pressure-sensor-equipped inverted echo sounders (CPIESs) can be used to obtain long-duration, high temporal resolution time series of boundary current transport and velocity structure. CPIESs are deployed on the seafloor in a rigid anchor which can last for one to five years. Each CPIES measures the round-trip surface-to-bottom acoustic travel time, near-bottom pressure and temperature, and the horizontal currents 50 m above the seabed. These measurements can be used together with a region’s hydrography -- from shipboard casts, Argo floats, or gliders -- to obtain time series of absolute

19 geostrophic velocity profiles from the surface to the bottom and their corresponding transports. Underwater gliders are small (50-kg mass, 2-m length), buoyancy-driven vehicles optimized for extended endurance. Gliders propel themselves by changing buoyancy, sinking or rising through the water column and adjusting altitude to use lift from the hull and wings to project vertical motion into the horizontal. This allows the vehicles to navigate from point to point while they profile vertically through the water column. Glider frequently come to the sea surface to communicate with mission control by using the Iridium satellite phone system. Antennas can be housed in a tail fin (Slocum), in a tail string (Seaglider) or in a wing (Spray). Gliders carry flexible payloads that can include sensors for temperature, conductivity (for salinity), velocity profiles, dissolved oxygen, chlorophyll, shear microstructure, nitrate, zooplankton, marine mammals, etc.. Gliders move slowly (20 km/day) to provide finely-spaced profiles, with the capability to maintain continuous sampling for periods typically lasting up to half a year. They can thus resolve scales of a few kilometers and weeks over sections of hundreds of kilometers length. Geostrophic transports can be reliably calculated from density sections using the thermal wind relationship. Direct transports can be measured using glider-mounted ADCPs. Gliders have been successfully used for sustained, multi-year quantification of eastern (e.g., California Current) and western (e.g., Kuroshio Current and Gulf Stream) boundary currents. For more details about gliders, please read the review paper below. Meyer, D., 2016: Glider Technology for Ocean Observations: A Review. Ocean Sci. Discuss., doi:10.5194/os-2016-40, 2016

2.2 The observed ocean circulation

Based on the various observations, we have obtained a good picture of the general circulation patterns for the world oceans, especially near the ocean surface.

The Pacific (Figure 2.11). Generally, the depth-integrated (or vertical mean) circulation has a similar pattern as the circulation shown in figure 2.4 above, which is the geostrophic current relative to a depth of no motion. The most prominent features are the strong subtropical gyres (STG), which prevail the subtropical northern hemisphere (NH) and subtropical southern hemisphere (SH). These figures show their time mean patterns; there are seasonal, interannual and longer term variations in the strength of the STGs.

The strong subtropical gyre (STG) in the NH has clockwise circulation, consists of the North Equatorial Current (NEC) with strongest flow near 15oN, the Philippines Current, the Kuroshio--the strong western boundary current (WBC), the North Pacific Current (NPC), and the California current (see schematic figure 2.11). In the SH, the STG is weaker and exhibits counter-clockwise circulation. Its associated currents are: the 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, the 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 blowing wind in the ITCZ (intertropical convergence zone in the atmosphere--Doldrums).

The subpolar gyre (SPG)—( counter-clockwise) in the NH. Its associated currents are: the NPC, Alaskan Current, Alaskan Stream, southern part of East Kamchatka Current, and Oyashio.

Figure 2.11. Schematic diagram showing major current systems in the Pacific.

Fig. 2.12. Cross-sectional sketch of the equatorial current system in the central Pacific Ocean (170°W). Shown in crosshatch are the North and South Equatorial Countercurrent (NECC and SECC), subsurface Equatorial Intermediate Current (EIC), North and South Subsurface Countercurrents (NSCC and SSCC), and Equatorial Undercurrent (EUC ). Eastward flow is colored green or brown, and all westward flow is white, including the North Equatorial Current (NEC) north of 5°N and the South Equatorial current (SEC ) south of 5°N and outside the EIC. Black numbers in italics were observations from January 1984 to June 1986 (latitude 165°E), and bold red numbers were observations from April 1979 to March 1980 (latitude 155°W), with both representing transports in Sverdrups (Sv =106 m3/s). Figure from https://www.researchgate.net/figure/Cross-sectional-sketch-of-the-equatorial-current-system-in- the-central-Pacific-Ocean_fig14_227056371.

The Equatorial Under Current (EUC). Field observations also show prominent currents in subsurface equatorial oceans (Figure 2.12). The EUC is a swift eastward-flowing ribbon of water extending over a distance of more than 14,000km along the Pacific equator with a thickness of only 200m and a width of at most 400km, with a core located at ~200m depth. The Equatorial intermediate current (EIC), NSCC, and SSCC (North and South subsurface counter current) are also evident.

The eastern equatorial Pacific Ocean is the home of El Nino. During normal years (no El Nino or La Nina), easterly trade winds cause upwelling, bringing the colder subsurface water into the surface layer and causing eastern equatorial Pacific cold tongue. During El Nino years: upwelling is reduced and SST increases. As discussed earlier, the large-scale east-west basin extent of the Pacific facilitates strong air/sea interaction.

Pacific-Indian Ocean connection: the Indonesian Throughflow. The warm and fresher tropical Pacific water flows into the Indian Ocean via a few straits of the Indonesia (Fig. 2.13). Fig. 2.13. Passages of the Indonesian Throughflow, which carries the warm and fresher tropical Pacific water into the South Indian Ocean. Figure from http://ocp.ldeo.columbia.ed u/res/div/ocp/projects/insta nt/projinfo.html.

The Atlantic Ocean. Similar to the Pacific, the STGs prevail the upper-ocean circulation in the subtropical regions of the NH and SH, and the Subpolar gyre exists in the NH. (Figure 2.14). The major currents associated with the NH STG are: the NEC centered near 15N, the Antilles Current east of 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 consists of: the SEC, the Brazil Current, the South Atlantic Current, and the Benguela Current.

The SPG is modified by the interaction between the North Atlantic and the Arctic Ocean, to the extent that it is hardly recognizable as a gyre. It involves the North Atlantic Current (NAC), Irminger Current, the East and West Greenland Currents (EGC and WGC), 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. Existing studies suggest that the warmer NAC is important in warming and thus melting the Arctic Sea Ice. Meanwhile, fresher water and sea ice exported from the Arctic Ocean through EGC (and East Iceland Current) also affect deep water formation and thus the thermohaline circulation in the North Atlantic Ocean.

Compared to the currents derived from the CTD data using geostrophic method of Fig. 2.5, large discrepancy occurs in the SPG region. Why? This is because at 700db,

23 significant current exists, which destroyed the assumption of ``level of no motion''. Recall that the geostrophic method measures current “shear” relative to the reference level.

In the southern Atlantic Ocean, ACC prevails.

Figure 2.14a. The equatorial and North Atlantic.

The Equatorial current system includes: the SEC and Equatorial Counter Current (ECC) that flows eastward against the weak westward wind; near the western boundary, there are cross-equatorial currents: the North Brazil Current and Guyana Current; the EUC flows eastward, with maximum speed of ~120 cm/s, core depth at ~100m, strongest in the west and weaker toward the east.

Connection with the Indian Ocean: via the Agulhas Current. The Agulhas Current and eddies transport heat and salt into Atlantic. There may be interaction between AO and IO via coastal Kelvin waves. The circulations have seasonal-interannual-decadal variability. Continents bound the E- W Atlantic, both air/sea interaction and continental monsoon effects are important in causing the variability.

Fig. 2.14b. The equatorial and South Atlantic currents.

The Indian Ocean. The Indian Ocean circulation subjects to strong seasonal variations of the monsoon wind forcing, which is southwest monsoon (SWM) during boreal summer and northeast monsoon (NEM) during boreal winter. Therefore, its annual mean picture does not make much sense in the north Indian Ocean (north of 10S). In the south IO (south of 10S), however, circulation is not much affected by the monsoon. The STG exists in the south IO throughout the year as in the other two oceans. Associate with the STG are the SEC, East Madagascar Current and Mozambique Current which join the Agulhas Current (WBC), South Indian Ocean current (adjacent to ACC). In the southeast IO along the west Australian coast, Leeuwin current flows southward against the prevailing southeasterly wind. Fig 2.15a.

During the Northeast Monsoon season (Fig. 2.15b): the NEC (also called the Northeast Monsoon Current) flows westward, the Somali Current flows southward, the Equatorial Counter Current flows eastward, and the South Java Current flows southeastward. In the Bay of Bengal and the Arabian Sea, the East India Coastal Current (EICC) and West Indian Coastal Current flow along the Indian coasts with the coasts to their right (counter clockwise).

During the Southwest Monsoon season: the Somali Current, which is fed by the east African Coastal Current (EACC), flows northward (acting as the WBC). The Southwest Monsoon Current is in opposite direction as the wintertime NEC. The South Java Current

25 also reverses direction to flow northwestward (rather than southeastward) (Fig. 12.15a). South of 10S, the circulation remains the similar. The EICC and WICC also reverse directions.

Fig. 2.15. Schematic diagrams showing surface currents over the Indian Ocean, which exhibits strong seasonal variations. a) Southwest monsoon season (July- August); b) Northeast monsoon season (Jan- Feb).

During spring and fall, the equatorial surface jets – the Wyrtki jets, flow eastward along the equator. They are swift currents, attaining amplitudes of ~100cm/s (1m/s) during their peak months of May and November (Fig. 12.15c).

The EUC exists only during winter to early spring, when surface winds are easterlies as in the Pacific and Atlantic Oceans. During other seasons, it disappears unless during the Indian Ocean Dipole (IOD) mode years (will be discussed in later chapters), the EUC reappears during summer-fall.

Interacting with Pacific by the Indonesian Throughflow, and with Atlantic by the Agulhas Current.

WJ WJ

Fig. 2.15c. Bimonthly plots of sea surface salinity (SSS) and surface currents from ship drift observations.The swift eastward-flowing currents along the equator during spring and fall are referred to as the Wyrtki Jets (WJ), following C. Wyrtki’s work on ocean observations. From Han and McCreary (2001).

The Arctic Ocean. The Arctic Ocean belongs to a class of ocean basins know as Mediterranean seas. A Mediterranean Sea is defined as a sea that only has limited communication with the major ocean basins (the Pacific, Atlantic, and Indian), where the circulation is dominated by thermohaline forcing.

E-P >0 E-P<0 O2

Pycnocline

Marine life Marine life Devoid below Salinity increases: Pycnocline. concentration basin Dilution basin.

Fig 2.16: A schematic diagram showing the circulation in a Mediterranean Sea.

(a) E-P>0, salinity increases and thus density increases. Heavy surface water sinks to the bottom. Produce frequent deep water renewal--oxygen available for marine life. 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 ocean 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 deep water volume is large, deep water 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.

Precipitation is low in the Polar area but significant in the subpolar regions, which are dominated by the West Wind Drift and its associated high storm frequency. Precipitation over the Siberia and the resulting river run-off, estimated to be ~0.2Sv (Sverdrup; 1 Sv=106 m3/s) total, contribute to polar mass budget. [Others precipitate on ice and export out to the North Atlantic so that they do not contribute to the mass budget in the Arctic.] Since evaporation over ice is relatively low compared to water, the Arctic Mediterranean Sea is a dilution basin: Its outflow is fresher on average than its inflow. See figure 2.17 for surface circulation pattern and 2.18 for sea ice coverage and export.

Figure 2.17. The Arctic circulation.

Figure 2.18. Sea ice extent and export.

Thermohaline circulation--concept. The circulation patterns discussed above in the Pacific, Atlantic and Indian Oceans 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 its size shrinking poleward and westward with the increase of 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’s oceans (such as the Mediterranean Sea in the North Atlantic; Red Sea and Persian Gulf in the Indian Ocean).

Thermohaline circulation is the movement of water when its density changes due to changes of temperature and/or of salinity. Examples: change of salinity: the Arctic Ocean and the Mediterranean Sea; change of temperature: the North Atlantic Ocean.

(i) The Arctic Ocean case: Fresher river water due to precipitation exceeds evaporation drives the outflow and thus exporting fresher water and sea ice into the North Atlantic at the surface, and the saltier North Atlantic water flows into the Arctic beneath the outflow. It is a dilution basin. (Salinity decrease causes density to 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 Indian Ocean are all concentration basins. Their E is much larger than P and thus salinity increases, causing density to increase near the surface. The heavier 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 ocean 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) The North Atlantic Ocean. The two cases discussed above represent density change due to ``salinity'' change: either decrease or increase. In the North Atlantic, the strong surface cooling can cool the sea surface temperature (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, being an important part of the global thermohaline circulation.

2.3. Water masses and T-S diagram. 2.3.1. Water masses A water mass is a 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 and mix with them. The total volume of a water mass is given by the sum of all its elements regardless of their location (Tomczak 1999).

The Atlantic water masses. (i) Abyssal water masses.--Formed by thermohaline process. There are five sources of Abyssal Atlantic water masses, classified (over simplification) as: the Antarctic Intermediate Water (AAIW), Antarctic Bottom Water (AABW), Mediterranean Overflow Water (MOW), Labrador Sea Water (LSW) and Nordic Sea Overflow Water (NSOW); the last three contribute to the NADW, which fills the depth range between 1000m and 4000m. Two of these have southern ocean origin, and the other three are originated from the North Atlantic.

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 2.19: y-z section at 25W for salinity in the Atlantic Ocean]

Figure 2.19: 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).

Southern source: AAIW and AABW. They flow into the North Atlantic 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 a source at the sea surface near the tip of South America where surface salinity is low. The AAIW can be tracked as a salinity minimum up to the subtropical gyre in the North Atlantic.

The other southern source for the Abyssal Atlantic Water is at the ocean bottom - the Antarctic Bottom Water (AABW). This is actually the 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 appears as a cold, low 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 the AABW. It upwells into the southward-flowing NADW layer above it (and hence does not reach the sea surface).

Northern Source for the Abyssal Atlantic Water: MOW, LSW, NSOW--form the 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. At the sea surface, water flows into the Mediterranean Sea from the Atlantic via the Strait of Gibraltar. Within the Mediterranean there is strong evaporation and cooling and thus production of dense water. This water flows out of the Mediterranean Sea into the North Atlantic at the bottom of the Strait. The 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 from the MOW.) The MOW at Gibraltar has a temperature of ~13C, salinity of ~38.45 psu, and density of 29.07 sigma theta (sigma = density in kg/m3 subtracts 1000). This is actually denser than the bottom water of the North Atlantic. As the MOW plume exits the Gibraltar, it moves to the north (boundary to the right, in the Kelvin wave sense) and down the slope. Vigorous mixing as it moves down the slope reduces its high salinity and density. It finally equilibrates at a depth of ~1000 m.

The LSW and NSOW. The two northern origins of the NADW are the Labrador Sea Water and Nordic Sea Overflow Water. The latter is also referred to as, by its three separate components resulting from overflows at three separate sites into the North Atlantic, the Greenland-Iceland-Norwegian Sea overflow water or GIN-Sea overflow. In undiluted form the NSOW is restricted to the immediate vicinity of the Greenland- Iceland-Scotland Ridge. It forms the lower part of the NADW.

The Labrador Sea Water (LSW) is formed in the western Labrador Sea through convection to about 1500-2000 meters in late winter. It 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 vertical direction (Fig. 2.20).

Within both the subpolar and subtropical gyres it is marked by an oxygen maximum in the vertical direction. Within both gyres it is also marked by a thickness maximum resulting from its convective source. Figure 2.20.

Figure 2.20. Zonal salinity section at 47N of the Atlantic Ocean.

The 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 it to as the ``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 call it Nordic Sea Overflow Water (NSOW) here. The Greenland Sea convection usually reaches intermediate depths, and is the water is colder and denser than the convection in the Labrador Sea, hence producing the denser part of the NADW.

Outflow of the dense NSOW occurs at 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 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 waters. This modifies the properties of the NSOW while it enters the North Atlantic Ocean.

The deep part of the NADW, i.e. the NSOW, is the lower branch of the Atlantic Meridional Overturning Circulation (AMOC) and fills the western North Atlantic through the lower part of the Deep Western Boundary Current. The DWBC is under the Gulf Stream, and it is a core layer of higher oxygen. The DWBC was predicted in a theory by Stommel and Arons, and its presence was detected through the earliest use of deep- tracked floats.

The DWBC flows to the South Atlantic. One feature of the DWBC is that is exhibits re- circulations to the north on its offshore side, which spreads the DWBC waters into the ocean interior. Transports for the various components of the AMOC (wind vs buoyancy driven components) have been computed using the RAPID observations. On average, the AMOC strength is 15~20 Sv. Part of the NADW flows out of the Atlantic, joins the ACC and becomes an important part of the global thermohaline ocean circulation (Fig. 2.21).

Fig. 2.21. Schematic

diagram showing

the global thermohali

ne circulation

The Subpolar Mode water (SPMW). The two northern sources of the NADW are from the surface waters of the Gulf Stream and North Atlantic Current, which flow northward, eastward and then northward again in the subpolar region (Fig. 2.14). The surface waters cool and freshen along their paths toward 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, analogous to the Subtropical Mode Water of the subtropical gyres discussed below. The SPMW temperature ranges from ~14C near the North Atlantic Current, to 8C where the SPMW enters the Norwegian Sea and to 4C where the SPMW enters the Labrador Sea.

Ventilation of the subtropical gyre: Water masses in the thermocline and surface layer: thermohaline and wind effect. [Thermocline: a subsurface region of tropical- subtropical ocean where vertical gradients of temperature is the largest.] Thermohaline effect. As sea surface temperature (SST) generally decreases poleward (Fig. 2.22), 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 (clockwise) 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 in the eastern boundary currents, are usually quite small, generally the surface waters slide down. This process is referred to as subduction. Consequently, thermocline waters in the tropics have the same density as the surface water in the subtropics. The subtropical gyre is thus ``ventilated''. It is often referred to as the “ventilated thermocline”. Theories and dynamics for the ventilated thermocline will be introduced in ATOC 5061 (spring 2020).

Figure 2.22. Annual mean global SST.

The Subtropical Mode Water (STMW or ``Eighteen Degree Water''). The isopycnal bowl (Fig. 2.23) immediately south of the Gulf Stream created by the Gulf Stream itself and its recirculation south of the Gulf Stream is a site of winter convection. Convection is favored here for two reasons: The isopycnal bowl has reduced vertical stratification compared to other regions simply because of the bowl, and heat losses to the atmosphere in this region are large because of the conjunction of the warm Gulf Stream waters with the cold, dry air blowing off the North American continent. The convected water mass is called ``Eighteen Degree Water'' because of its dominant temperature is ~18C. There is a STMW for each of the subtropical gyres' western boundary currents. The Eighteen Degree Water spreads to the entire region of the western subtropical gyre via oceanic circulation. ``Mode'' means relatively large volume on a volumetric T-S diagram (introduced below).

Fig. 2.23. Meridional section of density in the Pacific Ocean.

Wind effect: The 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. The Central Water is the general name for the entire thermocline layer. 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 direction. 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, spread along a continuous temperature/salinity range), and Subtropical Underwater (salinity maximum arising from subduction of the very high central subtropical gyre waters beneath the fresher waters 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

37 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 16C-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 Oyashio current. Because they are quite fresh, they appear as a salinity minimum in the subtropical gyre. This intermediate layer is often referred to as 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 shown to flow to the central South Pacific.

The Indian Ocean. Abyssal water mass. The 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, and the PGW seldom gets out of the Arabian Sea and just mixes with the Arabian Sea water.

Water masses of the thermocline and surface. As in the other two oceans, the Indian Central Water (ICW) is a subtropical water mass formed and subducted in the Subtropical Gyre convergence region of the South Indian Ocean. Surface water masses: Bay of Bengal water--fresh due to river runoff and Arabian Sea water -- salty.

2.3.2. T-S diagram. A diagram of salinity as a function of temperature is called T-S diagram (Fig. 2.24). Water properties, such as temperature and salinity, are formed only when the water is at

38 the surface or in the mixed layer. Heating, cooling, rain, and evaporation all contribute to the property changes. 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 and 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 not conservative.

Figure 2.24. T-S diagram of observed profile at 9S of the Atlantic Ocean.