When an earthquake begins, a quiet room can change within seconds. A hanging lamp swings, windows rattle and the floor moves underfoot. The movement reaches us at the surface, but it usually begins when rock slips along a fault at some depth inside the Earth.
The forces that move tectonic plates act for years, centuries or longer. They slowly change the stress within rocks. A fault may remain locked during this time because friction resists movement. Rock around the locked part bends or changes shape, rather like a loaded spring, although real rock also cracks and deforms permanently.
When a locked part can no longer hold, it slips suddenly. The rupture starts at one place and spreads across part of the fault. Some of the stored elastic energy leaves the source as vibrations called seismic waves. These waves travel through the Earth and along its surface, causing the ground to move as they pass.
Sensitive instruments record that motion. Different waves travel at different speeds and follow different paths, so they reach a recording station at different times. By comparing records from several stations, scientists can estimate where the rupture began, how deep it was and how large the event was. The same waves also help us study parts of the Earth's interior that humans cannot reach directly.
An earthquake's size does not tell us exactly what will happen in every place. A town close to the rupture may shake more strongly than a distant town, but distance is only one control. Depth, the direction in which the rupture spreads, the local ground and the strength of buildings all affect the result.
Some earthquakes begin beneath the sea. If a shallow marine rupture suddenly raises or lowers a large area of the seabed, it also moves the water above. That displacement can start a tsunami. The waves may remain low while crossing deep water, then slow down and grow as they reach shallower coastal water.
This connected sequence is the foundation of the topic. Changing stress can lead to sudden fault slip, and the slip sends out seismic waves. The waves create uneven shaking and carry information about their source. Some kinds of seabed displacement can also create a tsunami. We can now develop each part without losing sight of the whole process.
How slow deformation becomes sudden rupture
Stress, strain and a locked fault
Stress describes force acting within rock over an area. Compression pushes material together, tension pulls it apart and shear tries to move adjoining parts sideways. Strain is the change in shape, size or position that this stress produces. Stress is the cause within the material; strain is the response.
Rock can respond in several ways. A small change may be elastic, which means that the rock can recover much of the change when the stress falls. Under other conditions, rock can flow slowly, develop permanent deformation or break. Temperature, pressure, rock type, fluids and the rate at which stress changes all influence its behaviour.
Geologists use fault for a break, or a band of broken rock, where material on the two sides has shifted relative to each other. It is usually a surface or broken zone within rock, not a wide open crack. Some faults creep slowly. Other parts remain locked because rough surfaces and friction resist slip even while tectonic motion continues around them.
Stress can change around a locked fault, and the nearby rock can store elastic strain. Eventually, the applied stress may overcome friction and the strength of the locked part. Slip then begins and can spread quickly along the fault. This moving break is the rupture.
The rupture uses energy in several ways. It moves and breaks rock, produces heat through friction and sends seismic waves away from the source. Only part of the accumulated strain is elastic, and an earthquake does not release all the tectonic stress in a region. Permanent displacement remains across the fault, while continuing plate motion can load the fault system again.
Elastic rebound and permanent change
Elastic rebound explains how slow deformation can end in sudden movement. Before rupture, rocks beside a locked fault may bend elastically. When the fault slips, some of that distortion relaxes and the rocks rebound towards a less strained shape.
The two sides do not return to their original positions. The earthquake leaves a permanent offset, and the fault does not become whole again. This is why the spring comparison has limits: it helps us picture stored elastic energy, but a fault is a complex zone of rock with friction, cracking and lasting deformation.
Rupture also changes stress around the part that slipped. Some nearby areas lose stress, while others receive more. This redistribution helps explain why earthquakes often occur in sequences and why one rupture can affect neighbouring fault segments without providing a reliable timetable for the next event.
Where an earthquake begins and how a fault moves
Hypocentre, epicentre and rupture area
Rupture starts below the surface at a location called the hypocentre or focus. The vertical depth of this location is the focal depth. If we move straight upward from the hypocentre to the ground, the point reached is the epicentre.
These two points make an earthquake easy to place on a map, but the actual source can be much larger. After starting at the hypocentre, rupture may spread across a broad area of the fault. Different patches can move by different amounts, and the rupture may send more wave energy in one direction than another.
Greatest shaking or damage need not occur at the epicentre. A settlement may lie closer to another part of the rupture, or it may stand on ground that amplifies particular waves. The hypocentre is only the starting point; it is not the point from which all the energy of a large earthquake is released at once.
A surface rupture forms if fault displacement reaches the ground. Its trace may cut across roads, streams and fields. Many earthquakes do not break the surface, and a surface rupture does not have to pass through the epicentre.
Normal, reverse and sideways movement
Faults move in ways that reflect the stress acting on them. When tension stretches the crust, one block can move down relative to another along an inclined fault. This is a normal fault, common in rifts and other regions where the crust extends.
Compression can push one block up over another. This produces a reverse fault. A reverse fault with a relatively low-angle fault plane is called a thrust fault. Large thrust surfaces occur in fold-and-thrust mountain belts and at subduction interfaces.
Shear stress can make blocks slide mainly sideways past one another. This forms a strike-slip fault. A fault that combines vertical and sideways movement has oblique slip. Real faults do not have to fit one pure category along their whole length.
On an inclined fault, the block above the fault plane is the hanging wall, and the block below is the footwall. The hanging wall moves down in ideal normal faulting and up in ideal reverse faulting. These terms are less useful for a nearly vertical strike-slip fault, where horizontal movement dominates.
Fault type helps us connect a rupture to its tectonic setting, but it does not determine the earthquake's magnitude or local damage by itself. Rupture area, average slip, rock strength, depth and wave radiation also matter.
Earthquakes have several sources and several sequence patterns
Most earthquakes come from tectonic faulting. They cluster along plate boundaries, where plates separate, converge or slide past one another. They also occur within plates when the regional stress reactivates an old fault, rift or other zone of weakness.
Volcanic systems produce earthquakes when moving magma and changing pressure fracture or shift nearby rock. Mines or natural cavities can produce small, shallow events when a roof collapses. Explosions create seismic signals as well, although their source process differs from tectonic rupture.
Human activity can sometimes alter stress or fluid pressure enough to trigger movement on a fault that is already under stress. Fluid injection or withdrawal, mining and the filling of a large reservoir can contribute when the geological conditions are suitable. The Koyna–Warna region in India is an important case of reservoir-associated seismicity. This does not mean that every reservoir produces earthquakes; the response depends on existing faults, stress, rock properties and water-pressure changes.
Earthquakes commonly form sequences. Seismologists call the largest event in a related sequence the mainshock. Smaller events that follow nearby are aftershocks, which occur as the fault region adjusts to changed stress. Aftershocks generally become less frequent with time, though an individual aftershock cannot be timed exactly in advance.
A smaller event receives the name foreshock only after a larger nearby earthquake follows it. At the moment the smaller event occurs, there is no dependable way to know whether it will remain an ordinary earthquake or become a foreshock in retrospect. A swarm is a cluster in which no single event clearly dominates as the mainshock.
These patterns help scientists estimate probabilities and follow an active sequence. They do not permit an exact short-term prediction of the time, place and magnitude of a future earthquake. A probability forecast, a long-term hazard assessment and a warning issued after rupture has begun answer different questions.
Earthquake early warning begins after rupture as well. Sensors close to the source detect the first waves and send rapid information towards places that stronger shaking has not yet reached. The useful warning time depends on distance, detection speed and communication speed. A place close to the rupture may receive little or no warning. The system warns that shaking from an earthquake already under way may soon arrive; it does not predict the earthquake before it starts.
Seismic waves carry energy away from the fault
Fault slip disturbs the surrounding rock. That disturbance travels outward as elastic waves while the rock particles themselves move only around their positions. The wave carries energy; it does not carry a mass of rock from the fault to a distant city.
Some waves travel through the Earth's interior and are called body waves. Others remain concentrated near the surface and are called surface waves. Their different movements, speeds and paths explain both the order in which they arrive and the kind of ground motion they produce.
P waves: compression along the path
A P wave repeatedly compresses and expands material in the same direction in which the wave travels. Imagine gently pushing and releasing one end of a long spring: each coil moves back and forth, while the disturbance moves along the spring. The comparison shows the direction of motion, though rock inside the Earth behaves more complexly than a spring.
P means primary because these waves normally reach a station first. They can travel through solids, liquids and gases because all three can undergo compression. Their speed and direction change when the stiffness and density of the material change.
S waves: shear across the path
An S wave moves particles at right angles to the direction of travel. The motion may be side to side or up and down, depending on the wave. S means secondary because these waves arrive after P waves along comparable paths.
S waves require a material that can provide a restoring force against shear. Solids can do this, while liquids and gases cannot sustain a travelling shear wave. S waves therefore pass through solid crust and mantle but do not pass through the liquid outer core.
Love and Rayleigh waves near the surface
Body waves reaching the surface can help generate surface waves. Love waves move the ground mainly from side to side in a horizontal direction. Rayleigh waves produce a rolling motion with both vertical and horizontal components, often along an elliptical path.
Surface waves usually arrive after direct P and S waves and can have large amplitudes near the ground. They may also last longer. Still, no wave name is always the most destructive. A structure responds to the amplitude, direction, frequency and duration of the motion, as well as to the ground beneath it.
Why speed depends on more than density
Seismic-wave speed depends on the material's elastic stiffness as well as its density. Stiffer material resists deformation more strongly and can transmit a disturbance rapidly. Greater density adds inertia and can oppose rapid acceleration. The balance of these properties determines the speed, so density alone cannot tell us whether a wave will become faster or slower.
P waves move faster than S waves through the same broad solid region because compression and shear involve different elastic responses. There is no single fixed P-to-S speed ratio for every rock and depth. What remains dependable for basic interpretation is their arrival order and their different ability to pass through fluids.
Instruments turn ground motion into evidence
A seismometer senses ground motion. The complete system that detects, times and records the motion is often called a seismograph. The record it produces is a seismogram. Everyday use sometimes overlaps the first two terms, but the seismogram is always the trace or digital record that analysts examine.
A station may record three components of motion, usually one vertical and two horizontal. The first small arrival can be a P wave, followed by S waves and later surface waves. Wave amplitude, frequency and duration add more information about the source and the ground through which the waves travelled.
One station cannot describe the whole event. Stations around the source record the earthquake from several directions and at several distances. Analysts combine these arrivals with models of how waves move through the Earth, then estimate the event's origin time, location, depth and magnitude.
Early automatic values often change as more records arrive and specialists review the waveforms. A location or magnitude is therefore a measured estimate, not an unchangeable number discovered in one step.
Finding the source from arrival times
P waves travel faster than S waves, so a more distant station generally records a larger gap between their arrivals. A travel-time model uses that interval to estimate how far away the source lies. One station provides a possible distance but cannot show the direction to the earthquake.
On a simple map, all possible epicentres at that distance form a circle around the station. A second station adds another distance circle, and a third or further station helps identify the common area. People often call this triangulation, but the method uses intersecting distance constraints rather than measured angles.
Real networks use many arrivals instead of relying on three perfect circles. Wave speeds vary with depth, stations sit at different elevations and each picked arrival has uncertainty. A computer adjusts the origin time, latitude, longitude and depth until calculated arrivals fit the observations as well as possible.
Depth is harder to determine than a surface position. Nearby stations and waves that reflect near the surface can improve the estimate. Sparse station coverage or an unsuitable Earth model can leave greater uncertainty, which is another reason that published event locations may be revised.
Waves also reveal the Earth's interior
When a wave reaches a boundary between materials with different properties, part of it may reflect and part may bend, or refract, as it enters the next material. Comparing many travel paths reveals major internal boundaries and helps estimate the properties of the layers between them.
The liquid outer core produces a clear example. Direct S waves vanish beyond an angular distance of about 103 degrees from the epicentre because S waves cannot travel through that liquid layer. P waves enter the outer core, but the strong change in wave speed bends their paths enough to create an approximate direct P-wave shadow zone from about 103 to 140 degrees.
These angles are measured at the Earth's centre between the epicentre and the station. They are rounded classroom limits, not perfectly sharp borders. Reflected, converted and core-travelling phases can still appear within areas where the simplest direct wave is absent.
The missing direct S waves support the conclusion that the outer core is liquid. The sharp bending of P waves shows that the material changes greatly at the boundary of the core. This application shows how earthquakes help us investigate the interior. The complete model of the Earth's layers requires other evidence as well.
Magnitude measures the source; intensity describes a place
People often use size, strength and damage as if they meant the same thing. Seismology separates them. Magnitude gives the size of the rupture as a source event. Intensity records the shaking and observed effects at a particular place.
An earthquake receives one representative magnitude for a chosen method, although different methods or later analysis may produce slightly different estimates. The same earthquake produces many intensities because every place has a different position, ground condition and built environment. Lines joining places with similar intensity can form an irregular pattern rather than neat circles around the epicentre.
Why “Richter scale” is not a universal label
The scale popularly called the Richter scale is local magnitude, written ML. Its original form used particular instruments to measure nearby earthquakes in southern California. Similar local-magnitude methods remain useful in suitable regions and size ranges, but ML becomes less effective for very large events because the measured waves can saturate. Saturation means that the scale stops increasing enough to represent the full growth of the source.
Modern reports use several magnitude methods according to the available waves and the event. Moment magnitude, written Mw, is especially useful for large earthquakes. It comes from seismic moment. This physical measure combines the rock's rigidity, the surface area that ruptured and the average movement across it.
A larger rupture area or greater average slip usually raises the seismic moment. Moment magnitude converts this very large physical value into a manageable logarithmic number. It therefore connects reported source size to the dimensions and movement of the rupture more directly than an amplitude reading from one instrument.
Magnitude scales are logarithmic. On traditional amplitude-based scales, an increase of one whole magnitude corresponds to ten times the corrected wave amplitude. It is associated with roughly 32 times as much energy. These ratios compare source measures; they do not mean that the larger earthquake must cause ten or thirty-two times as much damage in a particular place.
Magnitude is not a ruler that begins at zero and ends at ten. Instruments can record very small events with negative magnitudes, and the scales have no formally fixed maximum. The finite size and strength of the Earth place physical limits on how large a natural earthquake can become, but the scale itself does not stop at a printed endpoint.
Why intensity changes from one place to another
Intensity records what people, objects, buildings and the ground experience at a location. Descriptive intensity scales arrange these effects in ordered levels, often written with Roman numerals. Reports from many places can be combined with instrument records to map how shaking varied.
Distance matters, but the relevant distance may be to the nearest part of a large rupture rather than only to the epicentre. A shallow earthquake often sends stronger motion to nearby ground than a similar deeper event. The direction in which rupture spreads can concentrate wave energy towards some areas, a process called directivity.
Duration and frequency also matter. A short, rapid motion affects structures differently from a long sequence rich in slower oscillations. Two places at similar distances can therefore record different shaking even before we consider their buildings.
Why the same earthquake affects places differently
Local ground can change the motion
Hard rock, loose sediment and deep sedimentary basins respond differently to arriving waves. Soft material can slow seismic waves and amplify some frequencies. It can also prolong motion by trapping and reflecting energy within a basin.
This behaviour is called site response. It explains why a neighbourhood on thick alluvium may shake differently from a nearby rocky ridge. The effect is selective: the ground does not amplify every frequency by the same amount.
Every structure also has natural periods at which it tends to sway. Resonance occurs when strong frequencies in the ground motion lie close to those natural periods, allowing motion to build. Low, stiff buildings and tall, flexible buildings usually respond to different frequency ranges. Design, materials, maintenance and construction quality can be as important as height.
Hazard, exposure and vulnerability
The possible shaking at a place is part of its earthquake hazard. Exposure refers to the people, buildings, roads and services present there. Vulnerability describes how readily they can be harmed. Risk grows from their interaction.
A powerful earthquake in a sparsely occupied area may produce limited human loss, while a smaller event near vulnerable buildings can be disastrous. Magnitude therefore has no fixed relationship with local damage. Depth, distance, rupture direction, site conditions, exposure and vulnerability must be considered together.
Effects that shaking can trigger
Strong shaking can damage buildings and bridges, break pipes and electrical lines, disrupt transport and start fires. Surface rupture can directly offset anything crossing a fault. These effects can interact, so the failure of one service may make another failure more serious.
Shaking can also destabilise steep slopes. Fractured or water-weakened material may move as rockfall, debris flow or landslide. Mountain roads and river valleys can be blocked, and a landslide dam may create a later flood hazard.
Liquefaction can occur when strong shaking affects loose sediment filled with water. Repeated motion raises pressure in the water between grains and reduces the effective contact that lets the grains support weight. The sediment temporarily loses much of its strength, allowing the ground to settle, spread sideways or eject water and sand.
The soil does not melt. Liquefaction needs a suitable combination of loose sediment, high groundwater and sufficient shaking. Young river deposits, deltas, reclaimed land and some coastal sediments may be susceptible, while dense or dry ground may not be.
These processes explain how an earthquake can become a wider disaster. Detailed preparedness, building design, emergency response and recovery belong to disaster management. Geography supplies the causal map of the source, the ground and the exposed landscape.
Earthquake depth and plate setting create broad patterns
Earthquakes do not occur at every depth. Geologists commonly group them as shallow, from the surface to about 70 kilometres; intermediate, from about 70 to 300 kilometres; and deep, from about 300 to 700 kilometres. These rounded limits are useful conventions, not boundaries between fixed Earth layers.
Shallow earthquakes occur in all major tectonic settings. Ridges and continental rifts produce shallow normal-fault earthquakes as plates or crustal blocks separate. Transform boundaries produce mainly shallow strike-slip earthquakes. Continental collisions and faults within plate interiors also produce mostly crustal events.
Subduction zones show the widest range. Shallow earthquakes occur near the trench and on the plate interface. Intermediate and deep earthquakes form within the cold lithospheric slab as it descends beneath the overriding plate. Their inclined pattern traces the slab towards depths of roughly 700 kilometres. Broad, locked subduction interfaces can also produce the largest earthquakes known.
Continental collision does not create the same oceanic trench-and-slab pattern everywhere. Compression spreads across thrusts and other faults within thickened crust. This produces a broad seismic belt, as seen across the Himalayan region.
Most earthquakes follow the global network of plate boundaries, but plate interiors are not earthquake-free. Regional stress can reactivate inherited faults and ancient rifts. Such events may be infrequent, yet they can be damaging because communities may have little recent experience and old continental crust can transmit waves efficiently.
Depth affects surface shaking without deciding the outcome alone. For a similar source size, a shallow rupture often produces stronger motion close to the source because the waves travel a shorter path. A deep earthquake may be felt across a wide area, and local ground can still alter the motion at each place.
India's seismic geography
India brings several tectonic settings together. The northern edge of the Indian Plate continues to converge with Eurasia. To the east, the collision region curves into the complex Indo-Burma and Andaman system. Within the plate, inherited rifts and faults remain capable of renewed movement.
The Himalayan arc and north-east India
Along the Himalayas, part of the Indian Plate underthrusts the mountain belt while crust shortens across several major structures. Locked fault sections can accumulate strain over broad areas. Earthquakes in the arc can affect both steep mountain slopes and the sediment-filled plains to the south.
North-east India lies where the Himalayan collision, the Shillong region and the Indo-Burma transition meet. Several structures and directions of movement overlap within a relatively small area. It is more accurate to treat this as a complex interaction zone than to explain every earthquake with one simple boundary line.
Andaman–Nicobar, Kachchh and the peninsula
The Andaman and Nicobar Islands lie beside the Sunda–Andaman subduction system. The plate contact produces earthquakes, and separate events also form inside the descending slab or the plate above it. A large, shallow thrust rupture here can move the seabed and generate a tsunami.
Kachchh is an intraplate region with ancient rift-related structures. The 2001 Bhuj earthquake showed that faults inside a plate can still produce a major event. It also showed why the phrase “stable plate interior” describes lower average activity, not complete safety.
Much of peninsular India is less active than the Himalayan and Andaman margins, yet old faults can reactivate under the present stress field. The Koyna–Warna case also shows that changes related to a reservoir can influence seismicity where suitable faults and fluid pathways already exist.
Thick sediments add another layer to Indian seismic geography. The Indo-Gangetic Plain and other river, delta and coastal basins can modify incoming waves. Some water-saturated sediments may also liquefy when shaking is strong enough.
What India's national zones mean
As of August 2026, India's applicable national framework follows IS 1893 (Part 1):2016 and uses Seismic Zones II, III, IV and V. Zone V is the highest broad hazard category, and Zone II is the lowest. A proposed revised zonation was withdrawn in March 2026, so a proposed Zone VI is not part of the applicable framework at this date.
The national map supports broad design and hazard assessment. It does not predict where or when the next earthquake will occur. A zone boundary also does not make the ground change abruptly at a line on the map.
Places within one zone may stand on very different rock, sediment, slopes and groundwater conditions. Microzonation studies these local differences at a finer scale. Site investigations go finer still. National zonation, microzonation and site assessment therefore answer related questions at different map scales.
The zone label also does not measure the complete risk of a city or building. Exposure and vulnerability can vary sharply within the same hazard zone. Any use of the national framework must keep this difference between broad hazard and local risk clear.
A tsunami begins by displacing water
A tsunami is a series of very long water waves created by the rapid displacement of a water body. Astronomical tides do not generate it, so the name “tidal wave” gives the wrong cause. Earthquakes generate most ocean-crossing tsunamis, but only when the source moves enough water quickly.
Why some marine earthquakes generate tsunamis
A large, shallow earthquake on a marine thrust fault can lift or lower a broad area of the seabed. The water surface above also rises or falls. Gravity then acts on this displaced water, and waves spread away from the source.
Subduction megathrusts can be especially effective because they may rupture a large, shallow, gently inclined surface and produce substantial vertical displacement. The size of the earthquake matters, but no magnitude number guarantees a tsunami. Rupture depth, area, fault orientation, slip direction and the amount and speed of vertical seabed movement all influence the result.
A predominantly strike-slip earthquake moves the seabed mainly sideways and usually displaces less water directly. A vertical component or an earthquake-triggered submarine landslide can still generate waves. This is why neither the phrase “underwater earthquake” nor a single magnitude threshold gives a complete test.
Submarine landslides can push water aside even without a very large fault rupture. Volcanic collapse or explosion can do the same, usually across a more limited region. A very large impact is another rare physical source.
How a tsunami crosses the ocean and changes near shore
Wind waves mainly disturb the upper part of the sea. A tsunami has such a long wavelength that its motion involves the water column from the surface towards the seabed. Its wavelength is far greater than the ocean depth. Wave science therefore classifies it as a shallow-water wave, even while it crosses a deep ocean basin.
For this kind of wave, speed depends mainly on water depth and rises roughly with the square root of that depth. The tsunami therefore moves fastest across deep ocean and slows as the water becomes shallower. Since real ocean depth changes along every route, different parts of the wave can travel at different speeds and bend around seabed features.
In deep water, a tsunami can move very fast while remaining low in height. Its energy is spread through a very long wave, so speed does not mean that a giant wall of water crosses the open ocean. A ship may pass over the gradual rise and fall without noticing it, although instruments can record the change.
Shoaling, run-up and inundation
As the tsunami enters shallower water, it slows and its wavelength shortens. The water can build to a greater height. This transformation is called shoaling.
The coast does not receive a simple copy of the offshore wave. Submarine ridges, channels and the continental shelf refract or redirect it. Bays can concentrate energy, while headlands, islands and reefs can alter the paths. Beach slope, river mouths, coastal buildings and the tide level at arrival further change the water movement.
At the shore, the sea may rise rapidly as a surge, form a bore or produce breaking waves. In some cases, the water first withdraws because a trough reaches the coast before a crest. That withdrawal is not universal, so its absence does not mean that no tsunami is coming.
Run-up is the highest vertical level that the water reaches above a reference sea level. Inundation is the horizontal distance or land area that the water floods. A steep coast may record high run-up but limited inland reach, while a low coastal plain may experience wide inundation.
A tsunami arrives as a series of waves. The first need not be the largest, because later waves may carry more energy or interact differently with the coast. Reflection within a bay or basin can also prolong changes in water level.
Warning begins after the source event starts
Earthquake prediction would require reliable advance knowledge of an event's time, place and magnitude. No method can provide that exact short-term prediction. A tsunami warning system works differently because it reacts rapidly after a possible source event has begun.
Seismic records first help estimate the earthquake's location, depth, magnitude and faulting. Models then test whether the source could have displaced the sea. Deep-ocean pressure measurements and coastal sea-level gauges can confirm and track the waves, while updated models estimate arrival times and possible coastal levels. Communication systems must then deliver the warning to the places that may be affected.
Every stage takes time and contains uncertainty. A distant coast may have enough time for measurement and modelling, while a nearby coast may receive strong shaking and a tsunami within a much shorter interval. Warning is therefore rapid detection, assessment and communication, not a prediction made before rupture.
Natural observations also have limits. Strong or long coastal shaking can indicate a nearby source, and an unusual rapid rise or fall of the sea can indicate that a tsunami is arriving. The water does not always withdraw first, and a person far from the source may feel little shaking. Detailed warning response and evacuation belong to disaster management; the geographical lesson is that source distance controls the time available.
The Indian Ocean example
The 2004 Sumatra–Andaman earthquake shows the full source-to-coast chain. A very large, shallow thrust rupture, with a moment magnitude of 9.1, spread along the Sunda–Andaman subduction system. Vertical seabed movement displaced the overlying water and sent tsunami waves across the Indian Ocean.
Magnitude alone does not explain that outcome. The shallow marine setting, the long rupture area, thrust movement and vertical displacement combined to form an efficient source. The waves then travelled through different ocean depths before shelves, bays, islands and coastlines shaped local run-up and inundation.
Two broad source regions matter especially for India. The Sunda–Andaman system can affect the islands and coasts around the Bay of Bengal, while the Makran subduction system north of the Arabian Sea can send waves towards western shores. Both the mainland coasts and the island territories therefore form part of India's tsunami geography, though their exposure differs by source and coastline.
An earthquake in either region does not automatically create a damaging tsunami. The source mechanism and seabed displacement must be assessed first, then sea-level observations and propagation models refine the picture. A coast may face weak local shaking yet still receive waves from a distant source.
Bringing the whole process together
Earthquake geography begins with slow deformation and sudden rupture. Stress changes around a fault, rock strains and a locked section may eventually slip. The hypocentre marks where rupture starts, while the rupture itself spreads across a finite fault area and sends several kinds of seismic waves outward.
P waves compress material and travel through solids and fluids. S waves shear solid material and stop at the liquid outer core. Surface waves move along the ground. Their timed arrivals help locate earthquakes, and their reflection, refraction and absence along certain paths reveal the layered interior.
Magnitude describes the source. Intensity describes the effects at a place. Depth, distance from the rupture, directivity, duration, frequency, local ground, buildings, exposure and vulnerability explain why one earthquake creates an uneven map of shaking and loss.
Plate settings organise the broad pattern. Ridges, transforms, subduction zones, continental collisions and inherited intraplate faults produce different distributions of depth and fault motion. India contains collision, subduction and intraplate settings, while its national seismic zones give only a broad hazard framework.
A tsunami needs rapid water displacement. A large, shallow marine thrust rupture can supply it, though many underwater earthquakes do not. The resulting long waves travel quickly but may remain low in deep water. They slow, shorten and often grow near shore, where seabed and coastline shape their final reach.
The same causal chain can guide any unfamiliar event. First locate the tectonic setting, then identify the rupture and depth. Read the waves and source magnitude, and examine the conditions that control local intensity. For a marine event, ask whether the rupture displaced enough water to begin a tsunami. Each step adds a necessary part of the explanation without turning one number or one map label into the whole story.