wind wave on trigonometric so do Waves and shallow water and Tsunami function on mathematic

 


        





  
In fluid dynamics, wind waves, or wind-generated waves, are surface waves that occur on the free surface of oceans, seas, lakes, rivers, and canals or even on small puddles and ponds. They result from the wind blowing over an area of fluid surface. Waves in the oceans can travel thousands of miles before reaching land. Wind waves range in size from small ripples, to waves over 100 ft (30 m) high.
When directly generated and affected by local winds, a wind wave system is called a wind sea. After the wind ceases to blow, wind waves are called swells. More generally, a swell consists of wind-generated waves that are not significantly affected by the local wind at that time. They have been generated elsewhere or some time ago. Wind waves in the ocean are called ocean surface waves.
Wind waves have a certain amount of randomness: subsequent waves differ in height, duration, and shape with limited predictability. They can be described as a stochastic process, in combination with the physics governing their generation, growth, propagation and decay—as well as governing the interdependence between flow quantities such as: the water surface movements, flow velocities and water pressure. The key statistics of wind waves (both seas and swells) in evolving sea states can be predicted with wind wave models. 

Although waves are usually considered in the water seas of Earth, the hydrocarbon seas of Titan may also have wind-driven waves 

The great majority of large breakers seen on a beach result from distant winds. Five factors influence the formation of the flow structures in wind waves:

• Wind speed or strength relative to wave speed- the wind must be moving faster than the wave crest for energy transfer
• The uninterrupted distance of open water over which the wind blows without significant change in direction (called the fetch)
• Width of area affected by fetch
• Wind duration - the time over which the wind has blown over a given area
• Water depth

All of these factors work together to determine the size of wind waves and the structures of the flows within:

• Wave height (from high trough to crest)
• Wave length (from crest to crest)
• Wave period (time interval between arrival of consecutive crests at a stationary point)
• Wave propagation direction

A fully developed sea has the maximum wave size theoretically possible for a wind of a specific strength, duration, and fetch. Further exposure to that specific wind could only cause a loss of energy due to the breaking of wave tops and formation of "whitecaps". Waves in a given area typically have a range of heights. For weather reporting and for scientific analysis of wind wave statistics, their characteristic height over a period of time is usually expressed as significant wave height. This figure represents an average height of the highest one-third of the waves in a given time period (usually chosen somewhere in the range from 20 minutes to twelve hours), or in a specific wave or storm system. The significant wave height is also the value a "trained observer" (e.g. from a ship's crew) would estimate from visual observation of a sea state. Given the variability of wave height, the largest individual waves are likely to be somewhat less than twice the reported significant wave height for a particular day or storm.
• Sources of wind wave generation: Sea water wave is generated by many kinds of disturbances such as Seismic events, gravity, and crossing wind. The generation of wind wave is initiated by the disturbances of cross wind field on the surface of the sea water. Two major Mechanisms of surface wave formation by winds (a.k.a.‘The Miles-Phillips Mechanism’) and other sources (ex. earthquakes) of wave formation can explain the generation of wind waves.
However, if one set a flat water surface (Beaufort Point,0) and abrupt cross wind flows on the surface of the water, then the generation of surface wind waves can be explained by following two mechanisms which initiated by normal pressure fluctuations of turbulent winds and parallel wind shear flows.
• The mechanism of the surface wave generation by winds
1) Starts from "Fluctuations of wind" (O.M.Phillips) : the wind wave formation on water surface by wind is started by a random distribution of normal pressure acting on the water from the wind. By the mechanism developed by O.M. Phillips (in 1957), the water surface is initially at rest and wave generation is started by adding turbulent wind flows and then, by the fluctuations of the wind, normal pressure acting on the water surface. From this pressure fluctuation arise normal and tangential stresses to the surface water, which generates wave behavior on the water surface. It is assumed that:-

1. The water originally at rest.
2. The water is not viscid.
3. The water is irrotational.
4. There are random distribution of normal pressure to the water surface from the turbulent wind.
5. Correlations between air and water motions are neglected.

2) starts from "wind shear forces" on the water surface (J.W.Miles, applied to mainly 2D deep water gravity waves) ; John W. Miles suggested a surface wave generation mechanism which is initiated by turbulent wind shear flows Ua(y), based on the inviscid Orr-Sommerfeld equation in 1957. He found the energy transfer from wind to water surface as a wave speed, c is proportional to the curvature of the velocity profile of wind Ua’’(y) at point where the mean wind speed is equal to the wave speed (Ua=c, where, Ua is the Mean turbulent wind speed). Since the wind profile Ua(y) is logarithmic to the water surface, the curvature Ua’’(y) have negative sign at the point of Ua=c. This relations show the wind flow transferring its kinetic energy to the water surface at their interface, and arises wave speed, c.
the growth-rate can be determined by the curvature of the winds ((d^2 Ua)/(dz^2 )) at the steering height (Ua (z=z_h)=c) for a given wind speed Ua {Assumptions; 1. 2D parallel shear flow, Ua(y) 2. incompressible, inviscid water / wind 3. irrotational water 4. slope of the displacement of surface is small}
Generally these wave formation mechanisms occur together on the ocean surface and arise wind waves and grows up to the fully developed waves.
For example,
If we suppose a very flat sea surface (Beaufort number, 0), and sudden wind flow blows steadily across the sea surface, physical wave generation process will be like;
1. Turbulent wind flows form random pressure fluctuations at the sea surface. Small waves with a few centimeters order of wavelengths are generated by the pressure fluctuations. (The Phillips mechanism)
2. The cross winds keep acting on the initially fluctuated sea surface causing the waves to become larger. As the waves grow, the pressure differences get larger causing the growth rate to increase. Finally the shear instability expedites the wave growth exponentially. (The Miles mechanism)
3. The interactions between the waves on the surface generate longer waves (Hasselmann et al., 1973) and the interaction will transfer wave energy from the shorter waves generated by the Miles mechanism to the waves have slightly lower frequencies than the frequency at the peak wave magnitudes, then finally the waves will be faster than the cross wind speed (Pierson & Moskowitz).

Conditions Necessary for a Fully Developed Sea at Given Wind Speeds, and the Parameters of the Resulting Waves

Wind Conditions			Wave Size
Wind Speed in One Direction	Fetch	Wind Duration	Average Height	Average Wavelength	Average Period and Speed
19 km/h (12 mph)	19 km (12 mi)	2 hr	0.27 m (0.89 ft)	8.5 m (28 ft)	3.0 sec 9.3 ft/sec
37 km/h (23 mph)	139 km (86 mi)	10 hr	1.5 m (4.9 ft)	33.8 m (111 ft)	5.7 sec 19.5 ft/sec
56 km/h (35 mph)	518 km (322 mi)	23 hr	4.1 m (13 ft)	76.5 m (251 ft)	8.6 sec 29.2 ft/sec
74 km/h (46 mph)	1,313 km (816 mi)	42 hr	8.5 m (28 ft)	136 m (446 ft)	11.4 sec 39.1 ft/sec
92 km/h (57 mph)	2,627 km (1,632 mi)	69 hr	14.8 m (49 ft)	212.2 m (696 ft)	14.3 sec 48.7 ft/sec

((NOTE: Most of the wave speeds calculated from the wave length divided by the period are proportional to sqrt (length). Thus, except for the shortest wave length, the waves follow the deep water theory described in the next section. The 28 ft long wave must be either in shallow water or between deep and shallow.))

   



Three different types of wind waves develop over time:

• Capillary waves, or ripples
• Seas
• Swells

Ripples appear on smooth water when the wind blows, but will die quickly if the wind stops. The restoring force that allows them to propagate is surface tension. Sea waves are larger-scale, often irregular motions that form under sustained winds. These waves tend to last much longer, even after the wind has died, and the restoring force that allows them to propagate is gravity. As waves propagate away from their area of origin, they naturally separate into groups of common direction and wavelength. The sets of waves formed in this way are known as swells.
Individual "rogue waves" (also called "freak waves", "monster waves", "killer waves", and "king waves") much higher than the other waves in the sea state can occur. In the case of the Draupner wave, its 25 m (82 ft) height was 2.2 times the significant wave height. Such waves are distinct from tides, caused by the Moon and Sun's gravitational pull, tsunamis that are caused by underwater earthquakes or landslides, and waves generated by underwater explosions or the fall of meteorites—all having far longer wavelengths than wind waves.
Yet, the largest ever recorded wind waves are common — not rogue — waves in extreme sea states. For example: 29.1 m (95 ft) high waves have been recorded on the RRS Discovery in a sea with 18.5 m (61 ft) significant wave height, so the highest wave is only 1.6 times the significant wave height. The biggest recorded by a buoy (as of 2011) was 32.3 m (106 ft) high during the 2007 typhoon Krosa near Taiwan.
Ocean waves can be classified based on: the disturbing force(s) that create(s) them; the extent to which the disturbing force(s) continue(s) to influence them after formation; the extent to which the restoring force(s) weaken(s) (or flatten) them; and their wavelength or period. Seismic Sea waves have a period of ~20 minutes, and speeds of 760 km/h (470 mph). Wind waves (deep-water waves) have a period of about 20 seconds.

^[14]

Wave type	Typical wavelength	Disturbing force	Restoring force
Capillary wave	< 2 cm	Wind	Surface tension
Wind wave	60–150 m (200–490 ft)	Wind over ocean	Gravity
Seiche	Large, variable; a function of basin size	Change in atmospheric pressure, storm surge	Gravity
Seismic sea wave (tsunami)	200 km (120 mi)	Faulting of sea floor, volcanic eruption, landslide	Gravity
Tide	Half the circumference of Earth	Gravitational attraction, rotation of Earth	Gravity

 
The speed of all ocean waves is controlled by gravity, wavelength, and water depth. Most characteristics of ocean waves depend on the relationship between their wavelength and water depth. Wavelength determines the size of the orbits of water molecules within a wave, but water depth determines the shape of the orbits. The paths of water molecules in a wind wave are circular only when the wave is traveling in deep water. A wave cannot "feel" the bottom when it moves through water deeper than half its wavelength because too little wave energy is contained in the small circles below that depth. Waves moving through water deeper than half their wavelength are known as deep-water waves. On the other hand, the orbits of water molecules in waves moving through shallow water are flattened by the proximity of the sea surface bottom. Waves in water shallower than 1/20 their original wavelength are known as shallow-water waves. Transitional waves travel through water deeper than 1/20 their original wavelength but shallower than half their original wavelength.
In general, the longer the wavelength, the faster the wave energy will move through the water. For deep-water waves, this relationship is represented with the following formula:

${\displaystyle C={L}/{T}}$

where C is speed (celerity), L is wavelength, and T is time, or period (in seconds).
The speed of a deep-water wave may also be approximated by:

${\displaystyle C={\sqrt {{gL}/{2\pi }}}}$

where g is the acceleration due to gravity, 9.8 meters (32 feet) per second squared. Because g and π (3.14) are constants, the equation can be reduced to:

${\displaystyle C=1.251{\sqrt {L}}}$

when C is measured in meters per second and L in meters. Note that in both instances that wave speed is proportional to wavelength.
The speed of shallow-water waves is described by a different equation that may be written as:

${\displaystyle C={\sqrt {gd}}=3.1{\sqrt {d}}}$

where C is speed (in meters per second), g is the acceleration due to gravity, and d is the depth of the water (in meters). The period of a wave remains unchanged regardless of the depth of water through which it is moving. As deep-water waves enter the shallows and feel the bottom, however, their speed is reduced and their crests "bunch up," so their wavelength shortens.


As waves travel from deep to shallow water, their shape alters (wave height increases, speed decreases, and length decreases as wave orbits become asymmetrical). This process is called shoaling.
Wave refraction is the process by which wave crests realign themselves as a result of decreasing water depths. Varying depths along a wave crest cause the crest to travel at different phase speeds, with those parts of the wave in deeper water moving faster than those in shallow water. This process continues until the crests become (nearly) parallel to the depth contours. Rays—lines normal to wave crests between which a fixed amount of energy flux is contained—converge on local shallows and shoals. Therefore, the wave energy between rays is concentrated as they converge, with a resulting increase in wave height.
Because these effects are related to a spatial variation in the phase speed, and because the phase speed also changes with the ambient current – due to the Doppler shift – the same effects of refraction and altering wave height also occur due to current variations. In the case of meeting an adverse current the wave steepens, i.e. its wave height increases while the wave length decreases, similar to the shoaling when the water depth decreases.


 
Some waves undergo a phenomenon called "breaking". A breaking wave is one whose base can no longer support its top, causing it to collapse. A wave breaks when it runs into shallow water, or when two wave systems oppose and combine forces. When the slope, or steepness ratio, of a wave is too great, breaking is inevitable.
Individual waves in deep water break when the wave steepness—the ratio of the wave height H to the wavelength λ—exceeds about 0.07, so for H > 0.07 λ. In shallow water, with the water depth small compared to the wavelength, the individual waves break when their wave height H is larger than 0.8 times the water depth h, that is H > 0.8 h. Waves can also break if the wind grows strong enough to blow the crest off the base of the wave.
Three main types of breaking waves are identified by surfers or surf lifesavers. Their varying characteristics make them more or less suitable for surfing, and present different dangers.

• Spilling, or rolling: these are the safest waves on which to surf. They can be found in most areas with relatively flat shorelines. They are the most common type of shorebreak
• Plunging, or dumping: these break suddenly and can "dump" swimmers—pushing them to the bottom with great force. These are the preferred waves for experienced surfers. Strong offshore winds and long wave periods can cause dumpers. They are often found where there is a sudden rise in the sea floor, such as a reef or sandbar.
• Surging: these may never actually break as they approach the water's edge, as the water below them is very deep. They tend to form on steep shorelines. These waves can knock swimmers over and drag them back into deeper water.

Wind waves are mechanical waves that propagate. along the interface between water and air; the restoring force is provided by gravity, and so they are often referred to as surface gravity waves. As the wind blows, pressure and friction perturb the equilibrium of the water surface and transfer energy from the air to the water, forming waves. The initial formation of waves by the wind is described in the theory of Phillips from 1957, and the subsequent growth of the small waves has been modeled by Miles, also in 1957.

Stokes drift in a deeper water wave (Animation)

Photograph of the water particle orbits under a – progressive and periodic – surface gravity wave in a wave flume. The wave conditions are: mean water depth d = 2.50 ft (0.76 m), wave height H = 0.339 ft (0.103 m), wavelength λ = 6.42 ft (1.96 m), period T = 1.12 s.

See also: Airy wave theory

In linear plane waves of one wavelength in deep water, parcels near the surface move not plainly up and down but in circular orbits: forward above and backward below (compared the wave propagation direction). As a result, the surface of the water forms not an exact sine wave, but more a trochoid with the sharper curves upwards—as modeled in trochoidal wave theory.
When waves propagate in shallow water, (where the depth is less than half the wavelength) the particle trajectories are compressed into ellipses.
In reality, for finite values of the wave amplitude (height), the particle paths do not form closed orbits; rather, after the passage of each crest, particles are displaced slightly from their previous positions, a phenomenon known as Stokes drift.^[22][23]
As the depth below the free surface increases, the radius of the circular motion decreases. At a depth equal to half the wavelength λ, the orbital movement has decayed to less than 5% of its value at the surface. The phase speed (also called the celerity) of a surface gravity wave is – for pure periodic wave motion of small-amplitude waves – well approximated by

${\displaystyle c={\sqrt {{\frac {g\lambda }{2\pi }}\tanh \left({\frac {2\pi d}{\lambda }}\right)}}}$

where

c = phase speed;

λ = wavelength;

d = water depth;

g = acceleration due to gravity at the Earth's surface.

In deep water, where ${\displaystyle d\geq {\frac {1}{2}}\lambda }$, so ${\displaystyle {\frac {2\pi d}{\lambda }}\geq \pi }$ and the hyperbolic tangent approaches ${\displaystyle 1}$, the speed ${\displaystyle c}$ approximates

${\displaystyle c_{\text{deep}}={\sqrt {\frac {g\lambda }{2\pi }}}.}$

In SI units, with ${\displaystyle c_{\text{deep}}}$ in m/s, ${\displaystyle c_{\text{deep}}\approx 1.25{\sqrt {\lambda }}}$, when ${\displaystyle \lambda }$ is measured in metres. This expression tells us that waves of different wavelengths travel at different speeds. The fastest waves in a storm are the ones with the longest wavelength. As a result, after a storm, the first waves to arrive on the coast are the long-wavelength swells.
For intermediate and shallow water, the Boussinesq equations are applicable, combining frequency dispersion and nonlinear effects. And in very shallow water, the shallow water equations can be used.
If the wavelength is very long compared to the water depth, the phase speed (by taking the limit of c when the wavelength approaches infinity) can be approximated by

${\displaystyle c_{\text{shallow}}=\lim _{\lambda \rightarrow \infty }c={\sqrt {gd}}.}$

On the other hand, for very short wavelengths, surface tension plays an important role and the phase speed of these gravity-capillary waves can (in deep water) be approximated by

${\displaystyle c_{\text{gravity-capillary}}={\sqrt {{\frac {g\lambda }{2\pi }}+{\frac {2\pi S}{\rho \lambda }}}}}$

where

S = surface tension of the air-water interface;

${\displaystyle \rho }$ = density of the water.

When several wave trains are present, as is always the case in nature, the waves form groups. In deep water the groups travel at a group velocity which is half of the phase speed. Following a single wave in a group one can see the wave appearing at the back of the group, growing and finally disappearing at the front of the group.
As the water depth ${\displaystyle d}$ decreases towards the coast, this will have an effect: wave height changes due to wave shoaling and refraction. As the wave height increases, the wave may become unstable when the crest of the wave moves faster than the trough. This causes surf, a breaking of the waves.
The movement of wind waves can be captured by wave energy devices. The energy density (per unit area) of regular sinusoidal waves depends on the water density ${\displaystyle \rho }$, gravity acceleration ${\displaystyle g}$ and the wave height ${\displaystyle H}$ (which, for regular waves, is equal to twice the amplitude, ${\displaystyle a}$):

${\displaystyle E={\frac {1}{8}}\rho gH^{2}={\frac {1}{2}}\rho ga^{2}.}$

The velocity of propagation of this energy is the group velocity.

Wind wave models

The image shows the global distribution of wind speed and wave height as observed by NASA's TOPEX/Poseidon's dual-frequency radar altimeter from October 3 to October 12, 1992. Simultaneous observations of wind speed and wave height are helping scientists to predict ocean waves. Wind speed is determined by the strength of the radar signal after it has bounced off the ocean surface and returned to the satellite. A calm sea serves as a good reflector and returns a strong signal; a rough sea tends to scatter the signals and returns a weak pulse. Wave height is determined by the shape of the return radar pulse. A calm sea with low waves returns a condensed pulse whereas a rough sea with high waves returns a stretched pulse. Comparing the two images above shows a high degree of correlation between wind speed and wave height. The strongest winds (33.6 mph; 54.1 km/h) and highest waves are found in the Southern Ocean. The weakest winds—shown as areas of magenta and dark blue—are generally found in the tropical Oceans.

 

 

 Wind wave model

Surfers are very interested in the wave forecasts. There are many websites that provide predictions of the surf quality for the upcoming days and weeks. Wind wave models are driven by more general weather models that predict the winds and pressures over the oceans, seas and lakes.
Wind wave models are also an important part of examining the impact of shore protection and beach nourishment proposals. For many beach areas there is only patchy information about the wave climate, therefore estimating the effect of wind waves is important for managing littoral environments.

Seismic signals

Main article: microseism

Ocean water waves generate land seismic waves that propagate hundreds of kilometers into the land. These seismic signals usually have the period of 6 ± 2 seconds. Such recordings were first reported and understood in about 1900.
There are two types of seismic "ocean waves". The primary waves are generated in shallow waters by direct water wave-land interaction and have the same period as the water waves (10 to 16 seconds). The more powerful secondary waves are generated by the superposition of ocean waves of equal period traveling in opposite directions, thus generating standing gravity waves – with an associated pressure oscillation at half the period, which is not diminishing with depth. The theory for microseism generation by standing waves was provided by Michael Longuet-Higgins in 1950, after in 1941 Pierre Bernard suggested this relation with standing waves on the basis of observations.

Internal waves

Internal waves can form at the boundary between water layers of different densities. These sub-surface waves are called internal waves. As is the case with ocean waves at the air-ocean interface, internal waves possess troughs, crests, wavelength, and period. Internal waves move very slowly because the density difference between the joined media is very small. Internal waves occur in the ocean at the base of the pycnocline, especially at the bottom edge of a steep thermocline. The wave height of internal waves may be greater than 30 meters (98 feet), causing the pycnocline to undulate slowly through a considerable depth. Their wavelength often exceeds 0.8 kilometres (0.50 mi) and their periods are typically 5 to 8 minutes. Internal waves are generated by wind energy, tidal energy, and ocean currents. Surface manifestations of internal waves have been photographed from space.
Internal waves may mix nutrients into surface water and trigger plankton blooms. They can also affect submarines and oil platforms.


A tsunami (plural: tsunamis or tsunami; from Japanese: 津波, lit. "harbour wave"; English pronunciation: /tsuːˈnɑːmi/) , also known as a seismic sea wave, is a series of waves in a water body caused by the displacement of a large volume of water, generally in an ocean or a large lake. Earthquakes, volcanic eruptions and other underwater explosions (including detonations of underwater nuclear devices), landslides, glacier calvings, meteorite impacts and other disturbances above or below water all have the potential to generate a tsunami. Unlike normal ocean waves which are generated by wind or tides which are generated by the gravitational pull of the Moon and Sun, a tsunami is generated by the displacement of water.
Tsunami waves do not resemble normal undersea currents or sea waves, because their wavelength is far longer. Rather than appearing as a breaking wave, a tsunami may instead initially resemble a rapidly rising tide, and for this reason they are often referred to as tidal waves, although this usage is not favoured by the scientific community because tsunamis are not tidal in nature. Tsunamis generally consist of a series of waves with periods ranging from minutes to hours, arriving in a so-called "internal wave train". Wave heights of tens of metres can be generated by large events. Although the impact of tsunamis is limited to coastal areas, their destructive power can be enormous and they can affect entire ocean basins; the 2004 Indian Ocean tsunami was among the deadliest natural disasters in human history with at least 230,000 people killed or missing in 14 countries bordering the Indian Ocean.
Greek historian Thucydides suggested in his late-5th century BC History of the Peloponnesian War, that tsunamis were related to submarine earthquakes, but the understanding of a tsunami's nature remained slim until the 20th century and much remains unknown. Major areas of current research include trying to determine why some large earthquakes do not generate tsunamis while other smaller ones do; trying to accurately forecast the passage of tsunamis across the oceans; and also to forecast how tsunami waves interact with specific shorelines. 

Tsunami

The term tsunami, meaning "harbour wave" in literal translation, comes from the Japanese 津波, composed of the two kanji 津 (tsu) meaning "harbour" and 波 (nami), meaning "wave". (For the plural, one can either follow ordinary English practice and add an s, or use an invariable plural as in the Japanese.) While not entirely accurate, as tsunami are not restricted to harbours, tsunami is currently the term most widely accepted by geologists and oceanographers.

Tidal wave

Tsunami aftermath in Aceh, Indonesia, December 2004.

Tsunami are sometimes referred to as tidal waves. This once-popular term derives from the most common appearance of tsunami, which is that of an extraordinarily high tidal bore. Tsunami and tides both produce waves of water that move inland, but in the case of tsunami the inland movement of water may be much greater, giving the impression of an incredibly high and forceful tide. In recent years, the term "tidal wave" has fallen out of favour, especially in the scientific community, because tsunami actually have nothing to do with tides, which are produced by the gravitational pull of the moon and sun rather than the displacement of water. Although the meanings of "tidal" include "resembling" or "having the form or character of" the tides, use of the term tidal wave is discouraged by geologists and oceanographers.

Seismic sea wave

The term seismic sea wave also is used to refer to the phenomenon, because the waves most often are generated by seismic activity such as earthquakes. Prior to the rise of the use of the term "tsunami" in English-speaking countries, scientists generally encouraged the use of the term "seismic sea wave" rather than the inaccurate term "tidal wave." However, like "tsunami," "seismic sea wave" is not a completely accurate term, as forces other than earthquakes – including underwater landslides, volcanic eruptions, underwater explosions, land or ice slumping into the ocean, meteorite impacts, or even the weather when the atmospheric pressure changes very rapidly – can generate such waves by displacing water.

History

See also: List of historic tsunamis

Lisbon earthquake and tsunami in 1755

While Japan may have the longest recorded history of tsunamis, the sheer destruction caused by the 2004 Indian Ocean earthquake and tsunami event mark it as the most devastating of its kind in modern times, killing around 230,000 people. The Sumatran region is not unused to tsunamis either, with earthquakes of varying magnitudes regularly occurring off the coast of the island.
Tsunamis are an often underestimated hazard in the Mediterranean Sea and parts of Europe. Of historical and current (with regard to risk assumptions) importance are the 1755 Lisbon earthquake and tsunami (which was caused by the Azores–Gibraltar Transform Fault), the 1783 Calabrian earthquakes, each causing several tens of thousands of deaths and the 1908 Messina earthquake and tsunami. The tsunami claimed more than 123,000 lives in Sicily and Calabria and is among the most deadly natural disasters in modern Europe. The Storegga Slide in the Norwegian sea and some examples of tsunamis affecting the British Isles refer to landslide and meteotsunamis predominantly and less to earthquake-induced waves.
As early as 426 BC the Greek historian Thucydides inquired in his book History of the Peloponnesian War about the causes of tsunami, and was the first to argue that ocean earthquakes must be the cause.

"The cause, in my opinion, of this phenomenon must be sought in the earthquake. At the point where its shock has been the most violent the sea is driven back, and suddenly recoiling with redoubled force, causes the inundation. Without an earthquake I do not see how such an accident could happen."^[18]

The Roman historian Ammianus Marcellinus (Res Gestae 26.10.15-19) described the typical sequence of a tsunami, including an incipient earthquake, the sudden retreat of the sea and a following gigantic wave, after the 365 AD tsunami devastated Alexandria.

Generation mechanisms

The principal generation mechanism (or cause) of a tsunami is the displacement of a substantial volume of water or perturbation of the sea. This displacement of water is usually attributed to either earthquakes, landslides, volcanic eruptions, glacier calvings or more rarely by meteorites and nuclear tests. The waves formed in this way are then sustained by gravity. Tides do not play any part in the generation of tsunamis.

Seismicity

Tsunami can be generated when the sea floor abruptly deforms and vertically displaces the overlying water. Tectonic earthquakes are a particular kind of earthquake that are associated with the Earth's crustal deformation; when these earthquakes occur beneath the sea, the water above the deformed area is displaced from its equilibrium position. More specifically, a tsunami can be generated when thrust faults associated with convergent or destructive plate boundaries move abruptly, resulting in water displacement, owing to the vertical component of movement involved. Movement on normal (extensional) faults can also cause displacement of the seabed, but only the largest of such events (typically related to flexure in the outer trench swell) cause enough displacement to give rise to a significant tsunami, such as the 1977 Sumba and 1933 Sanriku events.

• 

  Drawing of tectonic plate boundary before earthquake
  
• 

  Overriding plate bulges under strain, causing tectonic uplift.
  
• 

  Plate slips, causing subsidence and releasing energy into water.
  
• 

  The energy released produces tsunami waves.
  

Tsunamis have a small amplitude (wave height) offshore, and a very long wavelength (often hundreds of kilometres long, whereas normal ocean waves have a wavelength of only 30 or 40 metres), which is why they generally pass unnoticed at sea, forming only a slight swell usually about 300 millimetres (12 in) above the normal sea surface. They grow in height when they reach shallower water, in a wave shoaling process described below. A tsunami can occur in any tidal state and even at low tide can still inundate coastal areas.
On April 1, 1946, a magnitude-7.8 (Richter Scale) earthquake occurred near the Aleutian Islands, Alaska. It generated a tsunami which inundated Hilo on the island of Hawai'i with a 14-metre high (46 ft) surge. The area where the earthquake occurred is where the Pacific Ocean floor is subducting (or being pushed downwards) under Alaska.
Examples of tsunami originating at locations away from convergent boundaries include Storegga about 8,000 years ago, Grand Banks 1929, Papua New Guinea 1998 (Tappin, 2001). The Grand Banks and Papua New Guinea tsunamis came from earthquakes which destabilised sediments, causing them to flow into the ocean and generate a tsunami. They dissipated before travelling transoceanic distances.
The cause of the Storegga sediment failure is unknown. Possibilities include an overloading of the sediments, an earthquake or a release of gas hydrates (methane etc.).
The 1960 Valdivia earthquake (M_w 9.5), 1964 Alaska earthquake (M_w 9.2), 2004 Indian Ocean earthquake (M_w 9.2), and 2011 Tōhoku earthquake (M_w9.0) are recent examples of powerful megathrust earthquakes that generated tsunamis (known as teletsunamis) that can cross entire oceans. Smaller (M_w 4.2) earthquakes in Japan can trigger tsunamis (called local and regional tsunamis) that can only devastate nearby coasts, but can do so in only a few minutes.

Landslides

In the 1950s, it was discovered that larger tsunamis than had previously been believed possible could be caused by giant submarine landslides. These rapidly displace large water volumes, as energy transfers to the water at a rate faster than the water can absorb. Their existence was confirmed in 1958, when a giant landslide in Lituya Bay, Alaska, caused the highest wave ever recorded, which had a height of 524 metres (over 1700 feet). The wave did not travel far, as it struck land almost immediately. Two people fishing in the bay were killed, but another boat managed to ride the wave.
Another landslide-tsunami event occurred in 1963 when a massive landslide from Monte Toc entered the Vajont Dam in Italy. The resulting wave surged over the 262 m (860 ft) high dam by 250 metres (820 ft) and destroyed several towns. Around 2,000 people died Scientists named these waves megatsunamis.
Some geologists claim that large landslides from volcanic islands, e.g. Cumbre Vieja on La Palma in the Canary Islands, may be able to generate megatsunamis that can cross oceans, but this is disputed by many others.
In general, landslides generate displacements mainly in the shallower parts of the coastline, and there is conjecture about the nature of large landslides that enter water. This has been shown to lead to effect water in enclosed bays and lakes, but a landslide large enough to cause a transoceanic tsunami has not occurred within recorded history. Susceptible locations are believed to be the Big Island of Hawaii, Fogo in the Cape Verde Islands, La Reunion in the Indian Ocean, and Cumbre Vieja on the island of La Palma in the Canary Islands; along with other volcanic ocean islands. This is because large masses of relatively unconsolidated volcanic material occurs on the flanks and in some cases detachment planes are believed to be developing. However, there is growing controversy about how dangerous these slopes actually are.

Meteotsunamis

Some meteorological conditions, especially rapid changes in barometric pressure, as seen with the passing of a front, can displace bodies of water enough to cause trains of waves with wavelengths comparable to seismic tsunami, but usually with lower energies. These are essentially dynamically equivalent to seismic tsunami, the only differences being that meteotsunami lack the transoceanic reach of significant seismic tsunami, and that the force that displaces the water is sustained over some length of time such that meteotsunami can't be modelled as having been caused instantaneously. In spite of their lower energies, on shorelines where they can be amplified by resonance they are sometimes powerful enough to cause localised damage and potential for loss of life. They have been documented in many places, including the Great Lakes, the Aegean Sea, the English Channel, and the Balearic Islands, where they are common enough to have a local name, rissaga. In Sicily they are called marubbio and in Nagasaki Bay they are called abiki. Some examples of destructive meteotsunami include 31 March 1979 at Nagasaki and 15 June 2006 at Menorca, the latter causing damage in the tens of millions of euros.
Meteotsunami should not be confused with storm surges, which are local increases in sea level associated with the low barometric pressure of passing tropical cyclones, nor should they be confused with setup, the temporary local raising of sea level caused by strong on-shore winds. Storm surges and setup are also dangerous causes of coastal flooding in severe weather but their dynamics are completely unrelated to tsunami waves.They are unable to propagate beyond their sources, as waves do.

Man-made or triggered tsunamis

See also: Tsunami bomb

There have been studies of the potential of induction of and at least one actual attempt to create tsunami waves as a tectonic weapon.
In World War II, the New Zealand Military Forces initiated Project Seal, which attempted to create small tsunamis with explosives in the area of today's Shakespear Regional Park; the attempt failed.
There has been considerable speculation on the possibility of using nuclear weapons to cause tsunamis near to an enemy coastline. Even during World War II consideration of the idea using conventional explosives was explored. Nuclear testing in the Pacific Proving Ground by the United States seemed to generate poor results. Operation Crossroads fired two 20 kilotonnes of TNT (84 TJ) bombs, one in the air and one underwater, above and below the shallow (50 m (160 ft)) waters of the Bikini Atoll lagoon. Fired about 6 km (3.7 mi) from the nearest island, the waves there were no higher than 3–4 m (9.8–13.1 ft) upon reaching the shoreline. Other underwater tests, mainly Hardtack I/Wahoo (deep water) and Hardtack I/Umbrella (shallow water) confirmed the results. Analysis of the effects of shallow and deep underwater explosions indicate that the energy of the explosions doesn't easily generate the kind of deep, all-ocean waveforms which are tsunamis; most of the energy creates steam, causes vertical fountains above the water, and creates compressional waveforms. Tsunamis are hallmarked by permanent large vertical displacements of very large volumes of water which don't occur in explosions.

Characteristics

When the wave enters shallow water, it slows down and its amplitude (height) increases.

The wave further slows and amplifies as it hits land. Only the largest waves crest.

Tsunamis cause damage by two mechanisms: the smashing force of a wall of water travelling at high speed, and the destructive power of a large volume of water draining off the land and carrying a large amount of debris with it, even with waves that do not appear to be large.
While everyday wind waves have a wavelength (from crest to crest) of about 100 metres (330 ft) and a height of roughly 2 metres (6.6 ft), a tsunami in the deep ocean has a much larger wavelength of up to 200 kilometres (120 mi). Such a wave travels at well over 800 kilometres per hour (500 mph), but owing to the enormous wavelength the wave oscillation at any given point takes 20 or 30 minutes to complete a cycle and has an amplitude of only about 1 metre (3.3 ft). This makes tsunamis difficult to detect over deep water, where ships are unable to feel their passage.
The velocity of a tsunami can be calculated by obtaining the square root of the depth of the water in metres multiplied by the acceleration due to gravity (approximated to 10 m sec²). For example, if the Pacific Ocean is considered to have a depth of 5000 metres, the velocity of a tsunami would be the square root of √5000 x 10 = √50000 = ~224 metres per second (735 feet per second), which equates to a speed of ~806 kilometres per hour or about 500 miles per hour. This formula is the same as used for calculating the velocity of shallow waves, because a tsunami behaves like a shallow wave as it peak to peak value reaches from the floor of the ocean to the surface.
The reason for the Japanese name "harbour wave" is that sometimes a village's fishermen would sail out, and encounter no unusual waves while out at sea fishing, and come back to land to find their village devastated by a huge wave.
As the tsunami approaches the coast and the waters become shallow, wave shoaling compresses the wave and its speed decreases below 80 kilometres per hour (50 mph). Its wavelength diminishes to less than 20 kilometres (12 mi) and its amplitude grows enormously. Since the wave still has the same very long period, the tsunami may take minutes to reach full height. Except for the very largest tsunamis, the approaching wave does not break, but rather appears like a fast-moving tidal bore. Open bays and coastlines adjacent to very deep water may shape the tsunami further into a step-like wave with a steep-breaking front.
When the tsunami's wave peak reaches the shore, the resulting temporary rise in sea level is termed run up. Run up is measured in metres above a reference sea level. A large tsunami may feature multiple waves arriving over a period of hours, with significant time between the wave crests. The first wave to reach the shore may not have the highest run up.
About 80% of tsunamis occur in the Pacific Ocean, but they are possible wherever there are large bodies of water, including lakes. They are caused by earthquakes, landslides, volcanic explosions, glacier calvings, and bolides.

Drawback

An illustration of the rhythmic "drawback" of surface water associated with a wave. It follows that a very large drawback may herald the arrival of a very large wave.

All waves have a positive and negative peak, i.e. a ridge and a trough. In the case of a propagating wave like a tsunami, either may be the first to arrive. If the first part to arrive at shore is the ridge, a massive breaking wave or sudden flooding will be the first effect noticed on land. However, if the first part to arrive is a trough, a drawback will occur as the shoreline recedes dramatically, exposing normally submerged areas. Drawback can exceed hundreds of metres, and people unaware of the danger sometimes remain near the shore to satisfy their curiosity or to collect fish from the exposed seabed.
A typical wave period for a damaging tsunami is about 12 minutes. This means that if the drawback phase is the first part of the wave to arrive, the sea will recede, with areas well below sea level exposed after 3 minutes. During the next 6 minutes the tsunami wave trough builds into a ridge, and during this time the sea is filled in and destruction occurs on land. During the next 6 minutes, the tsunami wave changes from a ridge to a trough, causing flood waters to drain and drawback to occur again. This may sweep victims and debris some distance from land. The process repeats as the next wave arrives.

Scales of intensity and magnitude

As with earthquakes, several attempts have been made to set up scales of tsunami intensity or magnitude to allow comparison between different events.

Intensity scales

The first scales used routinely to measure the intensity of tsunami were the Sieberg-Ambraseys scale, used in the Mediterranean Sea and the Imamura-Iida intensity scale, used in the Pacific Ocean. The latter scale was modified by Soloviev, who calculated the Tsunami intensity I according to the formula

${\displaystyle \,{\mathit {I}}={\frac {1}{2}}+\log _{2}{\mathit {H}}_{av}}$

where ${\displaystyle {\mathit {H}}_{av}}$ is the average wave height along the nearest coast. This scale, known as the Soloviev-Imamura tsunami intensity scale, is used in the global tsunami catalogues compiled by the NGDC/NOAA and the Novosibirsk Tsunami Laboratory as the main parameter for the size of the tsunami.
In 2013, following the intensively studied tsunamis in 2004 and 2011, a new 12 point scale was proposed, the Integrated Tsunami Intensity Scale (ITIS-2012), intended to match as closely as possible to the modified ESI2007 and EMS earthquake intensity scales.

Magnitude scales

The first scale that genuinely calculated a magnitude for a tsunami, rather than an intensity at a particular location was the ML scale proposed by Murty & Loomis based on the potential energy.Difficulties in calculating the potential energy of the tsunami mean that this scale is rarely used. Abe introduced the tsunami magnitude scale ${\displaystyle {\mathit {M}}_{t}}$, calculated from,

${\displaystyle \,{\mathit {M}}_{t}={a}\log h+{b}\log R={\mathit {D}}}$

where h is the maximum tsunami-wave amplitude (in m) measured by a tide gauge at a distance R from the epicentre, a, b and D are constants used to make the M_t scale match as closely as possible with the moment magnitude scale.

Warnings and predictions

See also: Tsunami warning system

Tsunami warning sign

Drawbacks can serve as a brief warning. People who observe drawback (many survivors report an accompanying sucking sound), can survive only if they immediately run for high ground or seek the upper floors of nearby buildings. In 2004, ten-year-old Tilly Smith of Surrey, England, was on Maikhao beach in Phuket, Thailand with her parents and sister, and having learned about tsunamis recently in school, told her family that a tsunami might be imminent. Her parents warned others minutes before the wave arrived, saving dozens of lives. She credited her geography teacher, Andrew Kearney.
In the 2004 Indian Ocean tsunami drawback was not reported on the African coast or any other east-facing coasts that it reached. This was because the wave moved downwards on the eastern side of the fault line and upwards on the western side. The western pulse hit coastal Africa and other western areas.
A tsunami cannot be precisely predicted, even if the magnitude and location of an earthquake is known. Geologists, oceanographers, and seismologists analyse each earthquake and based on many factors may or may not issue a tsunami warning. However, there are some warning signs of an impending tsunami, and automated systems can provide warnings immediately after an earthquake in time to save lives. One of the most successful systems uses bottom pressure sensors, attached to buoys, which constantly monitor the pressure of the overlying water column.
Regions with a high tsunami risk typically use tsunami warning systems to warn the population before the wave reaches land. On the west coast of the United States, which is prone to Pacific Ocean tsunami, warning signs indicate evacuation routes. In Japan, the community is well-educated about earthquakes and tsunamis, and along the Japanese shorelines the tsunami warning signs are reminders of the natural hazards together with a network of warning sirens, typically at the top of the cliff of surroundings hills.
The Pacific Tsunami Warning System is based in Honolulu, Hawaiʻi. It monitors Pacific Ocean seismic activity. A sufficiently large earthquake magnitude and other information triggers a tsunami warning. While the subduction zones around the Pacific are seismically active, not all earthquakes generate tsunami. Computers assist in analysing the tsunami risk of every earthquake that occurs in the Pacific Ocean and the adjoining land masses.


As a direct result of the Indian Ocean tsunami, a re-appraisal of the tsunami threat for all coastal areas is being undertaken by national governments and the United Nations Disaster Mitigation Committee. A tsunami warning system is being installed in the Indian Ocean.

One of the deep water buoys used in the DART tsunami warning system

Computer models can predict tsunami arrival, usually within minutes of the arrival time. Bottom pressure sensors can relay information in real time. Based on these pressure readings and other seismic information and the seafloor's shape (bathymetry) and coastal topography, the models estimate the amplitude and surge height of the approaching tsunami. All Pacific Rim countries collaborate in the Tsunami Warning System and most regularly practise evacuation and other procedures. In Japan, such preparation is mandatory for government, local authorities, emergency services and the population.
Some zoologists hypothesise that some animal species have an ability to sense subsonic Rayleigh waves from an earthquake or a tsunami. If correct, monitoring their behaviour could provide advance warning of earthquakes, tsunami etc. However, the evidence is controversial and is not widely accepted. There are unsubstantiated claims about the Lisbon quake that some animals escaped to higher ground, while many other animals in the same areas drowned. The phenomenon was also noted by media sources in Sri Lanka in the 2004 Indian Ocean earthquake. It is possible that certain animals (e.g., elephants) may have heard the sounds of the tsunami as it approached the coast. The elephants' reaction was to move away from the approaching noise. By contrast, some humans went to the shore to investigate and many drowned as a result.
Along the United States west coast, in addition to sirens, warnings are sent on television and radio via the National Weather Service, using the Emergency Alert System.

Forecast of tsunami attack probability

Kunihiko Shimazaki (University of Tokyo), a member of Earthquake Research committee of The Headquarters for Earthquake Research Promotion of Japanese government, mentioned the plan to public announcement of tsunami attack probability forecast at Japan National Press Club on 12 May 2011. The forecast includes tsunami height, attack area and occurrence probability within 100 years ahead. The forecast would integrate the scientific knowledge of recent interdisciplinarity and aftermath of the 2011 Tōhoku earthquake and tsunami. As the plan, announcement will be available from 2014.

Mitigation

See also: Tsunami barrier

A seawall at Tsu, Japan

In some tsunami-prone countries earthquake engineering measures have been taken to reduce the damage caused onshore.
Japan, where tsunami science and response measures first began following a disaster in 1896, has produced ever-more elaborate countermeasures and response plans. The country has built many tsunami walls of up to 12 metres (39 ft) high to protect populated coastal areas. Other localities have built floodgates of up to 15.5 metres (51 ft) high and channels to redirect the water from incoming tsunami. However, their effectiveness has been questioned, as tsunami often overtop the barriers.
The Fukushima Daiichi nuclear disaster was directly triggered by the 2011 Tōhoku earthquake and tsunami, when waves exceeded the height of the plant's sea wall. Iwate Prefecture, which is an area at high risk from tsunami, had tsunami barriers walls totalling 25 kilometres (16 mi) long at coastal towns. The 2011 tsunami toppled more than 50% of the walls and caused catastrophic damage.
The Okushiri, Hokkaidō tsunami which struck Okushiri Island of Hokkaidō within two to five minutes of the earthquake on July 12, 1993 created waves as much as 30 metres (100 ft) tall—as high as a 10-story building. The port town of Aonae was completely surrounded by a tsunami wall, but the waves washed right over the wall and destroyed all the wood-framed structures in the area. The wall may have succeeded in slowing down and moderating the height of the tsunami, but it did not prevent major destruction and loss of life.

Waves and shallow water

When waves travel into areas of shallow water, they begin to be affected by the ocean bottom. The free orbital motion of the water is disrupted, and water particles in orbital motion no longer return to their original position. As the water becomes shallower, the swell becomes higher and steeper, ultimately assuming the familiar sharp-crested wave shape. After the wave breaks, it becomes a wave of translation and erosion of the ocean bottom intensifies

In fluid dynamics, the Boussinesq approximation for water waves is an approximation valid for weakly non-linear and fairly long waves. The approximation is named after Joseph Boussinesq, who first derived them in response to the observation by John Scott Russell of the wave of translation (also known as solitary wave or soliton). The 1872 paper of Boussinesq introduces the equations now known as the Boussinesq equations.
The Boussinesq approximation for water waves takes into account the vertical structure of the horizontal and vertical flow velocity. This results in non-linear partial differential equations, called Boussinesq-type equations, which incorporate frequency dispersion (as opposite to the shallow water equations, which are not frequency-dispersive). In coastal engineering, Boussinesq-type equations are frequently used in computer models for the simulation of water waves in shallow seas and harbours.
While the Boussinesq approximation is applicable to fairly long waves – that is, when the wavelength is large compared to the water depth – the Stokes expansion is more appropriate for short waves (when the wavelength is of the same order as the water depth, or shorter).


  

proof of Pythagoras calculation in dynamic function robot

Pythagorean theorem reads in a right-angled triangle the hypotenuse squared same applies to the number of squares of the other sides. In general, if the triangle ABC is right-angled in C then the Pythagorean theorem can be stated AB ^ 2 = AC + BC ^ 2 ^ 2.
The Pythagorean Theorem is a very famous theorem. This theorem will be used in calculating the broad flat wake. Besides being used in the calculation of the flat wake, calculations in 3 dimensions or the other will also often use the Pythagorean theorem. Many books write this theorem as c ^ 2 = a ^ 2 + b ^ 2. Where c is the hypotenuse.
Proof of this theorem is very diverse. Very many ways to prove Pythagoras's theorem. Here will be given some proof of the Pythagorean theorem. From the very basic evidence until proof is quite complicated. Most evidence of the Pythagorean theorem is the development of core evidence (evidence base). 


I. The first evidence :


Supplied four right-angled triangles. Look at the picture above. 4 triangles above are similar triangles. Have the sides a, b and c. and side c is the hypotenuse of the triangle. The third triangle side are the result of the rotation of 90, 180 and 270 degrees of the first triangle.
The area of each triangle is \ frac {ab} {2}. So spacious 4 triangles are 2ab.
The triangles be set so that membentung square with side c as shown below. 




Note the image of the arrangement 4 the triangle. The images form a square with sides c. and in it there is a small square. The long side of the small square is (b-a).
Directly we can determine the extent of the huge square, ie c ^ 2. And indirectly, a large square area with side c is equal to the area of 4 triangles plus a small square area having a side (b-a). Thus obtained,

c ^ 2 = 2ab + (b-a) ^ 2
c ^ 2 = 2ab + b ^ a ^ 2 + 2-2ab
c ^ 2 = b ^ 2 + a ^ 2   


II. the second proof  




 Note the picture. The picture is a picture of 2 square. Large square is a square that has a side length, and a small square has a side length that is b.
Large square area that certainly is a ^ 2. And a small square area is b ^ 2. So spacious wake above is b ^ 2 + a ^ 2  




Both the square we combine. And we created a line so that as in the picture. C side becomes the hypotenuse of the triangle. then we cut into triangles. and we move it to the top and right side as shown below.


Square area with a side c of the course is c ^ 2. Because the two square at the outset is equal to one large square with sides c above, then certainly an area of 2 square equal to the area's first large square with the c side.
thus, c ^ 2 = b ^ 2 + a ^ 2

III. The third proof  



The picture is a picture of a trapezoid made up of three triangles. The spacious trapezoid is \ frac {1} {2} (a + b) (a + b). searched using the trapezoidal area formula. That is half multiplied by the number of times the high side of the parallel trapezoid. Seek broad flat wake above can also use an extensive number of triangles (see picture). that is

\ Frac {1} {2} ab + \ frac {1} {2} ab + \ frac {1} {2} c ^ 2.

Size counts is fixed. Namely the trapezium shape. so it should be both widely sought by the different steps that should be the same. Retrieved,

\ Frac {1} {2} (a + b) (a + b) = \ frac {1} {2} ab + \ frac {1} {2} ab + \ frac {1} {2} c ^ 2
\ Frac {1} {2} (a ^ 2 + 2ab + b ^ 2) = ab + \ frac {1} {2} c ^ 2
\ Frac {1} {2} a ^ 2 + ab + \ frac {1} {2} b ^ 2 = ab + \ frac {1} {2} c ^ 2
a ^ 2 + b ^ 2 = c ^ 2  


MACAM-MACAM BILANGAN TRIPEL PYTHAGORAS

1. (3,4,5)
2. (5,12,13)
3. (7,24,25)
4. (8,15,17)
5. (9,40,41)
6. (11,60,61)
7. (12,35,37)
8. (13,84,85)
9. (15,112,113)
10. (16,63,65)
11. (17,144,145)
12. (19,180,181)
13. (20,21,29)
14. (20,99,101)
15. (21,220,221)
16. (23,264,265)
17. (24,143,145)
18. (25,312,313)
19. (27,364,365)
20. (28,45,53)
21. (28,195,197)
22. (29,420,421)
23. (31,480,481)
24. (32,255,257)
25. (33,56,65)
26. (33,544,545)
27. (35,612,613)
28. (36,77,85)
29. (36,323,325)
30. (37,684,685)
31. (39,80,89)
32. (39,760,761)
33. (40,399,401)
34. (41,840,841)
35. (43,924,925)
36. (44,117,125)
37. (44,483,485)
38. (48,55,73)
39. (48,575,577)
40. (51,140,149)
41. (52,165,173)
42. (52,675,677)
43. (56,783,785)
44. (57,176,185)
45. (60,91,109)
46. (60,221,229)
47. (60,899,901)
48. (65,72,97)
49. (68,285,293)
50. (69,260,269)
51. (75,308,317)
52. (76,357,365)
53. (84,187,205)
54. (84,437,445)
55. (85,132,157)
56. (87,416,425)
57. (88,105,137)
58. (92,525,533)
59. (93,476,485)
60. (95,168,193)
61. (96,247,265)
62. (100,621,629)
63. (104,153,185)
64. (105,208,233)
65. (105,608,617)
66. (108,725,733)
67. (111,680,689)
68. (115,252,277)
69. (116,837,845)
70. (119,120,169)
71. (120,209,241)
72. (120,391,409)
73. (123,836,845)
74. (124,957,965)
75. (129,920,929)
76. (132,475,493)
77. (133,156,205)
78. (135,352,377)
79. (136,273,305)
80. (140,171,221)
81. (145,408,433)
82. (152,345,377)
83. (155,468,493)
84. (156,667,685)
85. (160,231,281)
86. (161,240,289)
87. (165,532,557)
88. (168,425,457)
89. (168,775,793)
90. (175,288,337)
91. (180,299,349)
92. (184,513,545)
93. (185,672,697)
94. (189,340,389)
95. (195,748,773)
96. (200,609,641)
97. (203,396,445)
98. (204,253,325)
99. (205,828,853)
100. (207,224,305)
101. (215,912,937)
102. (216,713,745)
103. (217,456,505)
104. (220,459,509)
105. (225,272,353)
106. (228,325,397)
107. (231,520,569)
108. (232,825,857)
109. (240,551,601)
110. (248,945,977)
111. (252,275,373)
112. (259,660,709)
113. (260,651,701)
114. (261,380,461)
115. (273,736,785)
116. (276,493,565)
117. (279,440,521)
118. (280,351,449)
119. (280,759,809)
120. (287,816,865)
121. (297,304,425)
122. (300,589,661)
123. (301,900,949)
124. (308,435,533)
125. (315,572,653)
126. (319,360,481)
127. (333,644,725)
128. (336,377,505)
129. (336,527,625)
130. (341,420,541)
131. (348,805,877)
132. (364,627,725)
133. (368,465,593)
134. (369,800,881)
135. (372,925,997)
136. (385,552,673)
137. (387,884,965)
138. (396,403,565)
139. (400,561,689)
140. (407,624,745)
141. (420,851,949)
142. (429,460,629)
143. (429,700,821)
144. (432,665,793)
145. (451,780,901)
146. (455,528,697)
147. (464,777,905)
148. (468,595,757)
149. (473,864,985)
150. (481,600,769)
151. (504,703,865)
152. (533,756,925)
153. (540,629,829)
154. (555,572,797)
155. (580,741,941)
156. (615,728,953)
157. (616,663,905) 

Pythagoras theory to determine the position of robot 

at the moment of making a robot with a special feature heavily dependent according to the needs. One of the sensors used in robots are the rotary encoder
Rotary encoder: a digital sensor used to detect displacement or movement of the robot. Along with the development of robots, FPGA development shows a positive response, so scientists began trying to apply it to the world of robots. One example application is to determine the position of the robot. calculating the position of the robot with FPGA using the rotary encoder.
FPGA (Field Programmable Gate Array) that is programmable major device that is composed of independent logic modules that can be configured through the canals of programmable routing. FPGA is a type IC HDL (Hardware Description Language).

In general, the inner architecture of FPGA IC consists
on three main elements, namely Input / Output Block (IOB), Configurable Logic Block (CLB) and programmable interconnect. Input Output Blocks (IOB), as an interface between the external and internal pins of the device use rlogic .Configure Logic Blocks (CLB), part of which will process all forms of logic circuit and Programmable Interconnect, the part that connects the CLB CLB one with the other
















B. Rotary encoder
Rotary encoder is an electromechanical device that can monitor the movement and position. Rotary encoder generally uses optical sensors to produce a series of pulses which translates into motion, position, and direction. So that the angular position of a shaft rotating object can be processed into information in the form of digital code by a rotary encoder
Rotary encoder
composed of a thin disk that has holes in the circular disk. LED is placed on one side of the disc so that
the light will be heading to the disc. On the other hand, a photo-transistor is placed so that the photo-transistor can detect light from the LED opposite. When the position of the disk results in light of the LED can achieve photo-transistor through the existing holes, then the photo-transistor will experience saturation and will produce a square wave pulse. The more the pulse train generated in one round determines
The accuracy rotary encoder.

       





    








A. System Block Diagram
Block diagram of the system, manually controlled robot (driven), resulting in the movement, rotary encoder, serves to detect the movement. The resulting output rotary encoder
   1, rotary encoder
    2, and a rotary encoder
    3 in the form of pulses that represent the position of the x-axis (x), and the movement of the x-axis angle (θ). De2 by FPGA output data respectively
rotary encoder is processed to be calculated movement of each rotary encoder then proceeds in the form of a position (x, y) is displayed on the LCD


     




   

     




         


B. Determine the position (x, y) robot

Pythagorean theorem
Round rotary encoder used as a reference as position
the x-axis robot, when the robot walks then it will
generating robot movement angle to the axis x (θ). So with the known position of the x-axis and the angle thetanya, then the position of the y-axis robot can be known by using the Pythagorean theorem. Here is a triangle theorem of Pythagoras.   

Rotary encoder adalah divais elektromekanik yang dapat memonitor gerakan dan posisi. Rotary encoder umumnya menggunakan sensor optik untuk menghasilkan serial pulsa yang dapat diartikan menjadi gerakan, posisi, dan arah. Sehingga posisi sudut suatu poros benda berputar dapat diolah menjadi informasi berupa kode digital oleh rotary encoder untuk diteruskan oleh rangkaian kendali. Rotary encoder umumnya digunakan pada pengendalian robot, motor drive, dsb.

Rotary encoder tersusun dari suatu piringan tipis yang memiliki lubang-lubang pada bagian lingkaran piringan. LED ditempatkan pada salah satu sisi piringan sehingga cahaya akan menuju ke piringan. Di sisi yang lain suatu photo-transistor diletakkan sehingga photo-transistor ini dapat mendeteksi cahaya dari LED yang berseberangan. Piringan tipis tadi dikopel dengan poros motor, atau divais berputar lainnya yang ingin kita ketahui posisinya, sehingga ketika motor berputar piringan juga akan ikut berputar. Apabila posisi piringan mengakibatkan cahaya dari LED dapat mencapai photo-transistor melalui lubang-lubang yang ada, maka photo-transistor akan mengalami saturasi dan akan menghasilkan suatu pulsa gelombang persegi. Gambar 1 menunjukkan bagan skematik sederhana dari rotary encoder. Semakin banyak deretan pulsa yang dihasilkan pada satu putaran menentukan akurasi rotary encoder tersebut, akibatnya semakin banyak jumlah lubang yang dapat dibuat pada piringan menentukan akurasi rotary encoder tersebut.

Gambar 1. Blok penyusun rotary encoder

Rangkaian penghasil pulsa (Gambar 2) yang digunakan umumnya memiliki output yang berubah dari +5V menjadi 0.5V ketika cahaya diblok oleh piringan dan ketika diteruskan ke photo-transistor. Karena divais ini umumnya bekerja dekat dengan motor DC maka banyak noise yang timbul sehingga biasanya output akan dimasukkan ke low-pass filter dahulu. Apabila low-pass filter digunakan, frekuensi cut-off yang dipakai umumnya ditentukan oleh jumlah slot yang ada pada piringan dan seberapa cepat piringan tersebut berputar, dinyatakan dengan:

(1)

Dimana f_c adalah frekuensi cut-off filter, s_w adalah kecepatan piringan dan n adalah jumlah slot pada piringan.

Gambar 2. Rangkaian tipikal penghasil pulsa pada rotary encoder

Terdapat dua jenis rotary encoder yang digunakan, Absolute rotary encoder dan incremental rotary encoder. Masing-masing rotary encoder ini akan dipaparkan pada bagian berikutnya.

ABSOLUTE ROTARY ENCODER

Absolute encoder menggunakan piringan dan sinyal optik yang diatur sedemikian sehingga dapat menghasilkan kode digital untuk menyatakan sejumlah posisi tertentu dari poros yang dihubungkan padanya. Piringan yang digunakan untuk absolut encoder tersusun dari segmen-segmen cincin konsentris yang dimulai dari bagian tengah piringan ke arah tepi luar piringan yang jumlah segmennya selalu dua kali jumlah segmen cincin sebelumnya. Cincin pertama di bagian paling dalam memiliki satu segmen transparan dan satu segmen gelap, cincin kedua memiliki dua segmen transparan dan dua segmen gelap, dan seterusnya hingga cincin terluar. Sebagai contoh apabila absolut encoder memiliki 16 cincin konsentris maka cincin terluarnya akan memiliki 32767 segmen. Gambar 3 menunjukkan pola cincin pada piringan absolut encoder yang memiliki 16 cincin.

Gambar 3. Contoh susunan pola 16 cincin konsentris pada absolut encoder

Karena setiap cincin pada piringan absolute encoder memiliki jumlah segmen kelipatan dua dari cincin sebelumnya, maka susunan ini akan membentuk suatu sistem biner. Untuk menghasilkan sistem biner pada susunan cincin maka diperlukan pasangan LED dan photo-transistor sebanyak jumlah cincin yang ada pada absolut encoder tersebut.

Gambar 4. Contoh piringan dengan 10 cincin dan 10 LED – photo-transistor untuk membentuk sistem biner 10 bit.

Sistem biner yang untuk menginterpretasi posisi yang diberikan oleh absolute encoder dapat menggunakan kode gray atau kode biner biasa, tergantung dari pola cincin yang digunakan. Untuk lebih jelas, kita lihat contoh absolut encoder yang hanya tersusun dari 4 buah cincin untuk membentuk kode 4 bit. Apabila encoder ini dihubungkan pada poros, maka photo-transistor akan mengeluarkan sinyal persegi sesuai dengan susunan cincin yang digunakan. Gambar 5 dan 6 menunjukkan contoh perbedaan diagram keluaran untuk absolute encoder tipe gray code dan tipe binary code.

Gambar 5. Contoh diagram keluaran absolut encoder 4-bit tipe gray code

Dengan absolute encoder 4-bit ini maka kita akan mendapatkan 16 informasi posisi yang berbeda yang masing-masing dinyatakan dengan kode biner atau kode gray tertentu. Tabel 1 menyatakan posisi dan output biner yang bersesuaian untuk absolut encoder 4-bit. Dengan membaca output biner yang dihasilkan maka posisi dari poros yang kita ukur dapat kita ketahui untuk diteruskan ke rangkaian pengendali. Semakin banyak bit yang kita pakai maka posisi yang dapat kita peroleh akan semakin banyak.

Gambar 6. Contoh diagram keluaran absolut encoder 4-bit tipe binary code

Tabel 1. Output biner dan posisi yang bersesuaian pada absolute encoder 4-bit

INCREMENTAL ROTARY ENCODER

Incremental encoder terdiri dari dua track atau single track dan dua sensor yang disebut channel A dan B (Gambar 7). Ketika poros berputar, deretan pulsa akan muncul di masing-masing channel pada frekuensi yang proporsional dengan kecepatan putar sedangkan hubungan fasa antara channel A dan B menghasilkan arah putaran. Dengan menghitung jumlah pulsa yang terjadi terhadap resolusi piringan maka putaran dapat diukur. Untuk mengetahui arah putaran, dengan mengetahui channel mana yang leading terhadap channel satunya dapat kita tentukan arah putaran yang terjadi karena kedua channel tersebut akan selalu berbeda fasa seperempat putaran (quadrature signal). Seringkali terdapat output channel ketiga, disebut INDEX, yang menghasilkan satu pulsa per putaran berguna untuk menghitung jumlah putaran yang terjadi.

Gambar 7. susunan piringan untuk incremental encoder

Contoh pola diagram keluaran dari suatu incremental encoder ditunjukkan pada Gambar 8. Resolusi keluaran dari sinyal quadrature A dan B dapat dibuat beberapa macam, yaitu 1X, 2X dan 4X. Resolusi 1X hanya memberikan pulsa tunggal untuk setiap siklus salah satu sinya A atau B, sedangkan resolusi 4X memberikan pulsa setiap transisi pada kedua sinyal A dan B menjadi empat kali resolusi 1X. Arah putaran dapat ditentukan melalui level salah satu sinyal selama transisi terhadap sinyal yang kedua. Pada contoh resolusi 1X, A = arah bawah dengan B = 1 menunjukkan arah putaran searah jarum jam, sebaliknya B = arah bawah dengan A = 1 menunjukkan arah berlawanan jarum jam.

Gambar 8. Contoh pola keluaran incremental encoder

Gambar 9. output dan arah putaran pada resolusi yang berbeda-beda

Pada incremental encoder, beberapa cara dapat digunakan untuk menentukan kecepatan yang diamati dari sinyal pulsa yang dihasilkan. Diantaranya adalah menggunakan frequencymeter dan periodimeter.

Cara yang sederhana untuk menentukan kecepatan dapat dengan frequencymeter, yakni menghitung jumlah pulsa dari encoder, n, pada selang waktu yang tetap, T, yang merupakan periode loop kecepatan (Gambar 10). Apabila α adalah sudut antara pulsa encoder, maka sudut putaran pada suatu periode adalah:

(2)

Sehingga kecepatan putar akan kita dapatkan sebagai:

(3)

Kelemahan yang muncul pada cara ini adalah pada setiap periode sudut α_f yang didapat merupakan kelipatan integer dari α. Ini akan dapat menghasilkan quantification error pada kecepatan yang ingin diukur.

Gambar 10. Sinyal keluaran encoder untuk pengukuran kecepatan dengan frequencymeter

Cara yang lain adalah dengan menggunakan periodimeter. Dengan cara ini kita akan mengukur kecepatan tidak lagi dengan menghitung jumlah pulsa encoder tetapi dengan menghitung clock frekuensi tinggi (HF Clock) untuk sebuah pulsa dari encoder yaitu mengukur periode pulsa dari encoder (Gambar 11). Apabila α_p adalah sudut dari pulsa encoder, t adalah periode dari HF clock, dan n adalah jumlah pulsa HF yang terhitung pada counter. Maka waktu untuk sebuah pulsa encoder, T_p,  adalah:

(4)

Sehingga kecepatan yang akan kita ukur dapat kita peroleh dengan:

(5)

Seperti halnya pada frequencymeter, disini juga muncul quantification error karena waktu T_p akan selalu merupakan perkalian integer dengan t.

Gambar 11. Pengukuran kecepatan dengan menggunakan Periodimeter