When a fault ruptures and releases stored elastic energy, that energy doesn't simply stay at the earthquake's source — it radiates outward through the Earth as seismic waves, traveling thousands of kilometers in minutes and shaking the ground along the way. These waves are the reason an earthquake in Chile can be recorded on instruments in Norway, and why a rupture 10 km underground can level buildings on the surface.
Understanding seismic waves is fundamental to nearly every aspect of earthquake science. The speed, direction, and behavior of these waves determine how seismologists locate earthquakes, how engineers design earthquake-resistant buildings, and how scientists have mapped Earth's hidden interior — from crust to core — without ever directly observing it. Seismic waves also explain why two cities equidistant from an earthquake can experience vastly different levels of damage, depending on the rock and soil beneath them.
This article covers the physics, behavior, and real-world consequences of each wave type, from the first tremor you feel to the rolling motion that can topple structures.
What Are Seismic Waves?
Seismic waves are waves of energy that propagate through Earth's interior and along its surface, generated by the sudden release of strain energy when rock fractures along a fault. The process is analogous to the waves that radiate outward when a stone is dropped in water, though seismic waves are three-dimensional and travel through solid rock at speeds of several kilometers per second.
The point within the Earth where the fault first ruptures is the hypocenter (or focus). The point on the surface directly above the hypocenter is the epicenter. Seismic waves radiate outward from the hypocenter in all directions, and their properties — speed, amplitude, frequency, and particle motion — determine what people feel during an earthquake and what kind of damage occurs.
Seismic waves divide into two fundamental categories: body waves, which travel through the Earth's interior, and surface waves, which are confined to the vicinity of Earth's surface. Each category contains distinct wave types with different physical characteristics.
Body Waves
Body waves propagate through the interior of the Earth in all directions from the earthquake source. They are the first waves to arrive at seismic stations and at locations distant from the epicenter. The two types of body waves — P-waves and S-waves — differ in their speed, particle motion, and ability to travel through different materials.
P-Waves (Primary Waves)
P-waves, or primary waves, are the fastest seismic waves and always the first to arrive at a recording station — hence the name "primary." They are also called compressional waves or longitudinal waves because the particle motion is parallel to the direction of wave propagation, creating alternating zones of compression and dilation (expansion) in the material through which they travel.
The motion resembles that of a Slinky toy pushed and pulled along its length: as the wave passes, material is alternately squeezed together and pulled apart in the direction the wave is traveling. This push-pull motion means P-waves temporarily change the volume of the material they pass through.
Key properties of P-waves:
- Speed: Approximately 6-8 km/s through crustal rock (varies with rock type and depth). Velocity increases with depth due to increasing pressure and density, reaching roughly 13.7 km/s at the base of the mantle. Through the liquid outer core, P-wave velocity drops to about 8 km/s before increasing again to ~11 km/s in the solid inner core.
- Particle motion: Parallel to the direction of propagation (compressional/longitudinal).
- Medium: P-waves can travel through solids, liquids, and gases — the only seismic wave type capable of traversing all three states of matter. This is because compressional deformation can occur in any material.
- Human perception: P-waves arriving at the surface are often perceived as a sudden, sharp jolt or a loud bang, particularly for nearby earthquakes. In some cases, P-waves are heard as a deep rumbling sound because their frequencies overlap with the audible range.
- Damage potential: Generally lower than other wave types due to smaller amplitudes, though near-field P-waves from large earthquakes can cause significant vertical shaking.
S-Waves (Secondary Waves)
S-waves, or secondary waves, arrive after P-waves and travel at roughly 55-60% of P-wave speed. They are also called shear waves or transverse waves because the particle motion is perpendicular to the direction of propagation. Material moves side-to-side (or up-and-down) as the wave passes, shearing the rock without changing its volume.
The motion is similar to the wave produced by shaking a rope side-to-side: the wave travels along the rope's length while each point on the rope moves perpendicular to that direction.
Key properties of S-waves:
- Speed: Approximately 3.5-4.5 km/s through crustal rock. Like P-waves, S-wave velocity increases with depth through the mantle, reaching about 7.3 km/s near the core-mantle boundary.
- Particle motion: Perpendicular to the direction of propagation (transverse/shear). S-waves can be further divided into SH waves (horizontal particle motion) and SV waves (vertical particle motion).
- Medium: S-waves can travel only through solids. Liquids and gases cannot sustain shear stress — they flow rather than deform elastically when subjected to transverse forces. This single property has profound implications for understanding Earth's structure.
- Human perception: S-waves produce the strong side-to-side or rolling shaking typically associated with earthquakes. The onset of S-wave shaking is usually felt as a sudden intensification following the initial P-wave jolt.
- Damage potential: Substantially higher than P-waves. The shearing motion is particularly destructive to buildings and structures because it induces lateral forces that buildings must resist.
The S-P Time Lag
Because P-waves travel faster than S-waves, they progressively separate as they move away from the earthquake source. At any given location, the time difference between the P-wave arrival and S-wave arrival — called the S-P interval or S-P time lag — is directly proportional to the distance from the earthquake.
As a rough rule of thumb, multiplying the S-P time in seconds by 8 gives an approximate distance in kilometers. For example, an S-P interval of 10 seconds indicates a distance of approximately 80 km. This principle is the foundation of earthquake location using triangulation from multiple seismic stations.
For people near an earthquake, the S-P interval provides a natural warning. If you feel a sharp jolt (P-wave) followed by a pause before strong shaking begins (S-wave), the length of that pause indicates how far away the rupture is. A long pause means the earthquake is distant and the strong shaking may be moderate; a very short pause — or no perceptible gap — means the earthquake is nearby and shaking could be severe.
Surface Waves
Surface waves are generated when body waves interact with Earth's free surface (the ground-air boundary). They are confined to the shallow subsurface and travel along the surface like ripples on a pond. Surface waves travel slower than body waves but typically carry much larger amplitudes and persist for longer durations, making them the primary cause of earthquake damage in most situations.
Two types of surface waves are recognized: Love waves and Rayleigh waves, named after the scientists who first mathematically described them.
Love Waves
Love waves, predicted mathematically by British mathematician A.E.H. Love in 1911, are surface waves that produce purely horizontal, side-to-side motion perpendicular to the direction of wave propagation. There is no vertical component to Love wave motion.
Love waves require a velocity structure in which a low-velocity surface layer overlies higher-velocity material — a condition met virtually everywhere on Earth's surface. They can be thought of as horizontally polarized shear waves (SH waves) trapped in the surface layer by constructive interference.
Key properties of Love waves:
- Speed: Typically 2-4.5 km/s, faster than Rayleigh waves but slower than S-waves. The velocity depends on frequency — long-period (low-frequency) Love waves travel faster than short-period (high-frequency) Love waves, a phenomenon called dispersion that allows seismologists to probe Earth's layered structure.
- Particle motion: Horizontal, perpendicular to the direction of wave travel (side-to-side).
- Damage potential: Very high. The horizontal shearing motion is extremely destructive to building foundations and causes lateral displacement of structures. Love waves are often the most damaging wave type for multi-story buildings.
Rayleigh Waves
Rayleigh waves, predicted by British physicist Lord Rayleigh in 1885, produce an elliptical retrograde rolling motion. As a Rayleigh wave passes, particles at the surface trace an elliptical path — moving up, then forward in the direction of wave travel, then down, then backward, creating a rolling motion similar to ocean swells. The amplitude decreases rapidly with depth.
Key properties of Rayleigh waves:
- Speed: Approximately 2-4 km/s, slightly slower than Love waves of the same period. Rayleigh waves are also dispersive.
- Particle motion: Retrograde elliptical in the vertical plane containing the direction of propagation. At the surface, the motion is backward (retrograde) relative to the wave's travel direction.
- Damage potential: Very high. The rolling motion can be visible to the naked eye during large earthquakes — observers have described seeing the ground moving in visible waves. Rayleigh waves are particularly damaging to long structures like bridges, pipelines, and rail lines.
Why Surface Waves Cause the Most Damage
Surface waves are typically the most destructive component of earthquake shaking for three reasons:
- Larger amplitudes: Surface waves carry more energy concentrated in the shallow subsurface, producing ground motions with larger amplitudes than the body waves from the same earthquake.
- Longer duration: Surface waves persist much longer than body wave arrivals. While P-waves and S-waves are relatively brief pulses, surface waves can produce sustained shaking for tens of seconds to several minutes, depending on earthquake size and distance.
- Lower frequencies: Surface waves are enriched in low frequencies (long periods), which resonate with the natural frequencies of many buildings and structures. Mid-rise to high-rise buildings (3-20 stories) are particularly vulnerable because their natural oscillation periods (roughly 0.3-2 seconds) can match the dominant periods of surface waves.
Seismic Wave Comparison
| Property | P-Waves | S-Waves | Love Waves | Rayleigh Waves |
|---|---|---|---|---|
| Also called | Primary, compressional, longitudinal | Secondary, shear, transverse | — | — |
| Type | Body wave | Body wave | Surface wave | Surface wave |
| Typical speed (crust) | 6-8 km/s | 3.5-4.5 km/s | 2-4.5 km/s | 2-4 km/s |
| Particle motion | Parallel to propagation (push-pull) | Perpendicular to propagation (side-to-side) | Horizontal, perpendicular to propagation | Retrograde elliptical (rolling) |
| Travels through | Solids, liquids, gases | Solids only | Near-surface solids | Near-surface solids |
| Arrival order | First | Second | Third (typically) | Fourth (typically) |
| Amplitude | Smallest | Moderate | Large | Large |
| Duration | Brief | Brief to moderate | Long | Longest |
| Damage potential | Low to moderate | Moderate to high | Very high | Very high |
Seismic Wave Velocities Through Earth's Layers
The speed at which seismic waves travel depends on the elastic properties and density of the material. As waves descend into the Earth, they encounter different compositions, pressures, and temperatures that alter their velocities. These velocity changes create the reflections and refractions that allow seismologists to map Earth's internal structure.
| Earth Layer | Depth Range | P-Wave Velocity | S-Wave Velocity | Key Features |
|---|---|---|---|---|
| Upper Crust | 0-15 km | 5.8-6.5 km/s | 3.2-3.7 km/s | Varies widely with rock type (granite, sediment) |
| Lower Crust | 15-35 km (continental) | 6.5-7.2 km/s | 3.6-4.0 km/s | Mohorovičić discontinuity at base |
| Upper Mantle | 35-410 km | 7.8-9.0 km/s | 4.4-5.0 km/s | Low-velocity zone (asthenosphere) at ~100-200 km |
| Transition Zone | 410-660 km | 9.0-10.3 km/s | 5.0-5.6 km/s | Mineral phase transitions (olivine to wadsleyite to ringwoodite) |
| Lower Mantle | 660-2,891 km | 10.3-13.7 km/s | 5.6-7.3 km/s | Steady velocity increase with depth |
| Outer Core | 2,891-5,150 km | 8.0-10.4 km/s | 0 (liquid) | S-waves cannot penetrate; P-waves slow dramatically |
| Inner Core | 5,150-6,371 km | 11.0-11.3 km/s | ~3.6 km/s | Solid iron-nickel; S-waves detected (PKJKP phase) |
Data sources: Preliminary Reference Earth Model (PREM, Dziewonski & Anderson, 1981); International Association of Seismology and Physics of the Earth's Interior (IASPEI) 1991 model.
Shadow Zones: How Waves Revealed Earth's Interior
One of seismology's greatest achievements is the discovery of Earth's layered internal structure through the observation of shadow zones — regions on the far side of the planet from an earthquake where certain wave types are absent or anomalously weak.
The S-Wave Shadow Zone
When an earthquake occurs, S-waves radiate outward through the mantle and are recorded by seismic stations up to an angular distance of about 104° from the epicenter. Beyond 104°, S-waves disappear entirely — they are not recorded at any station on the opposite side of the Earth. This S-wave shadow zone extends from 104° to 180° (the point directly opposite the earthquake, called the antipode).
The explanation, first proposed by Richard Dixon Oldham in 1906 and refined by Beno Gutenberg in 1914, is definitive: S-waves cannot travel through liquid, and the Earth has a liquid outer core. S-waves traveling at steep angles toward the core encounter the core-mantle boundary at a depth of approximately 2,891 km and are absorbed or reflected — they cannot pass through. Only stations within 104° of the epicenter receive direct S-waves traveling entirely through the solid mantle.
This observation was the first seismological proof of the liquid outer core and remains one of the most elegant demonstrations of how seismic waves reveal Earth's hidden structure.
The P-Wave Shadow Zone
P-waves, which can travel through both solids and liquids, do penetrate the outer core — but their path is dramatically altered. When P-waves enter the liquid outer core, they slow from about 13.7 km/s to 8 km/s, causing them to refract (bend) sharply. This refraction creates a P-wave shadow zone between approximately 104° and 140° from the epicenter, where direct P-waves are absent.
Beyond 140°, P-waves re-emerge because waves that entered the core at steep angles are refracted back toward the surface on the far side. Additionally, faint P-wave arrivals were eventually detected within the shadow zone itself — arrivals too early to be explained by waves diffracted around the core. In 1936, Danish seismologist Inge Lehmann proposed that these arrivals were caused by P-waves reflecting off the boundary of a solid inner core at a depth of approximately 5,150 km. The inner core, confirmed by subsequent studies, produces these reflected phases (called PKiKP and PKIKP) that partially fill the shadow zone.
Cross-section of Earth showing P-wave and S-wave shadow zones
Data: Earthquake source at surface, P-waves recorded 0-104° and beyond 140°, P-wave shadow zone 104-140°, S-waves recorded 0-104° only, S-wave shadow zone 104-180°. Core-mantle boundary at 2,891 km depth, inner-outer core boundary at 5,150 km depth. Show ray paths bending through mantle and core layers.
Reading a Seismogram
A seismogram is the recorded trace of ground motion at a seismic station during an earthquake. Understanding how to read a seismogram reveals the physics of earthquake waves in action.
Annotated seismogram diagram showing labeled P-wave, S-wave, Love wave, and Rayleigh wave arrivals
Data: Time axis (horizontal) in seconds, amplitude axis (vertical). P-wave arrival: first, small amplitude, high frequency, sharp onset. S-wave arrival: second, larger amplitude, moderate frequency, onset marked. Love wave arrival: large amplitude horizontal motion, lower frequency, extended duration. Rayleigh wave arrival: largest amplitude, lowest frequency, longest duration, gradual onset. S-P time interval labeled. Noise baseline shown before P arrival. Example for a moderate earthquake at approximately 100 km distance.
Identifying Wave Arrivals
For an earthquake at moderate distance (roughly 100-500 km), a typical three-component seismogram (recording vertical, north-south, and east-west motion) shows the following sequence:
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Background noise: Before the earthquake waves arrive, the seismogram shows low-level ambient noise from ocean waves (microseisms), wind, human activity, and other sources.
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P-wave arrival: The first seismic signal appears as a sharp onset above the noise level. On the vertical component, P-waves are usually most prominent because their push-pull motion has a strong vertical component for waves arriving from below. P-wave motion is typically high-frequency and relatively small in amplitude.
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S-wave arrival: After a quiet interval (the S-P time), the S-wave arrives with a noticeable increase in amplitude, particularly on the horizontal components. The onset is often abrupt, and the frequency content is generally lower than the P-wave.
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Surface wave arrivals: Following the S-wave, surface waves arrive with progressively increasing amplitude, lower frequency, and much longer duration. Love waves (horizontal motion) and Rayleigh waves (elliptical motion visible on both vertical and horizontal components) are often intermingled but can sometimes be distinguished by their polarization.
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Coda: After the main surface wave train passes, the seismogram gradually returns to background noise levels as scattered energy reverberates through Earth's crust. This "coda" can persist for minutes after a large earthquake.
Using the S-P Interval
The S-P time interval is the simplest tool for estimating distance to an earthquake. The interval increases linearly with distance at a rate of approximately 1 second per 8 km for crustal earthquakes. Standard travel-time curves published by seismological agencies provide precise distance estimates from S-P intervals for any depth and distance combination.
At seismic observatories, analysts (and increasingly, automated algorithms) pick P and S arrival times from seismograms, compute S-P intervals, and use these with data from other stations to triangulate the earthquake's location. This process is the operational foundation of earthquake monitoring as described in seismology and earthquake detection.
Site Effects: Why Local Geology Matters
The intensity of earthquake shaking at any location depends not only on the earthquake's magnitude and distance but also critically on the local geological conditions. Seismic waves can be dramatically amplified — or dampened — by the properties of the soil and rock beneath a site, a phenomenon known as site effects or site amplification.
The Physics of Amplification
When seismic waves travel from hard bedrock into soft, unconsolidated sediments (such as clay, silt, sand, or artificial fill), two things happen simultaneously:
- Velocity decrease: Wave speed drops dramatically — from perhaps 3-4 km/s in bedrock to 200-500 m/s in soft soil.
- Amplitude increase: To conserve energy, the wave amplitude must increase as velocity decreases. The amplification factor can range from 2 to 10 or more, depending on the impedance contrast (the product of velocity and density difference between layers).
Additionally, soft sediment layers can trap seismic energy, producing resonance when the natural frequency of the sediment column matches the frequency of the incoming waves. This resonance effect can amplify shaking at specific frequencies by enormous factors — precisely the frequencies most dangerous to buildings of certain heights.
Mexico City, 1985
The September 19, 1985, Michoacán earthquake (M8.0) provided the most dramatic illustration of site effects in modern seismology. The earthquake's epicenter was approximately 350 km from Mexico City, yet destruction in the capital was catastrophic — over 10,000 people died and more than 400 buildings collapsed or were severely damaged.
The devastation was concentrated in the historic city center, built on the soft clay sediments of the ancient Lake Texcoco bed. These lake sediments, with shear-wave velocities as low as 40-80 m/s, amplified seismic waves by factors of 8-50 compared to bedrock sites at the city's edge. The dominant resonance period of the lake sediments (~2 seconds) closely matched the natural oscillation period of 6-15 story buildings, which explains why mid-rise structures suffered disproportionate damage while shorter and taller buildings often survived.
Recordings from instruments in the lake zone showed peak ground accelerations roughly five times larger than those at nearby bedrock stations. The earthquake led directly to major revisions in Mexico City's building code and the installation of one of the world's first earthquake early warning systems (SASMEX).
San Francisco Marina District, 1989
During the October 17, 1989, Loma Prieta earthquake (M6.9), the Marina District of San Francisco experienced severe damage and fires, even though the epicenter was approximately 100 km to the south. The district was built largely on artificial fill placed after the 1906 earthquake — ironically, much of it was rubble from buildings destroyed in the 1906 disaster.
This loose fill amplified seismic waves significantly compared to bedrock areas of the city just blocks away. Additionally, liquefaction occurred in the saturated fill, causing the ground to lose bearing capacity and buildings to settle, tilt, and collapse. The Marina District experience starkly demonstrated that earthquake damage maps cannot be drawn based on distance alone — local geology dominates the outcome.
Amplification in Modern Seismic Hazard Analysis
Today, site effects are explicitly incorporated into seismic hazard assessments and building codes. The National Earthquake Hazards Reduction Program (NEHRP) classifies sites into categories from A (hard rock) to E (soft soil) based on the average shear-wave velocity in the upper 30 meters (Vs30). Each site class carries specific amplification factors applied to design ground motions.
USGS ShakeMaps — which show estimated shaking intensity across a region after an earthquake — incorporate site amplification using geological maps and Vs30 models to account for the effects of local geology. This produces more accurate shaking estimates than simple distance-based models and helps emergency responders identify areas likely to have experienced the worst damage.
How Seismic Waves Reveal Earth's Structure
Beyond their destructive power, seismic waves are humanity's primary tool for investigating Earth's deep interior. Just as medical CT scans use X-rays that pass through the body to create images of internal organs, seismologists use earthquake waves that pass through the planet to create images of its internal structure — a technique called seismic tomography.
Refraction and Reflection
When seismic waves encounter a boundary between materials with different elastic properties (such as the crust-mantle boundary or the core-mantle boundary), they are refracted (bent) and reflected — the same phenomena that affect light passing through glass. The angle and degree of refraction depend on the velocity contrast across the boundary, governed by Snell's Law.
These refractions and reflections create characteristic patterns in seismic wave arrivals at stations around the globe. By analyzing thousands of earthquake recordings, seismologists in the early 20th century identified the major internal boundaries of the Earth:
- Mohorovičić discontinuity (Moho): Discovered in 1909 by Croatian seismologist Andrija Mohorovičić, this boundary between Earth's crust and mantle is marked by a sharp increase in P-wave velocity from about 6.5-7.2 km/s to 7.8-8.1 km/s. The Moho lies at approximately 5-10 km depth beneath the oceans and 30-70 km beneath continents.
- Core-mantle boundary: At 2,891 km depth, P-wave velocity drops from 13.7 km/s to 8 km/s, and S-waves cease entirely — marking the transition from the solid silicate mantle to the liquid iron outer core.
- Inner-outer core boundary: At 5,150 km depth, P-wave velocity increases again, and very faint S-waves (PKJKP phases) have been detected, indicating the inner core is solid.
Seismic Tomography
Modern seismic tomography uses vast databases of wave travel times from thousands of earthquakes recorded at thousands of stations to construct three-dimensional velocity models of Earth's interior. Regions where waves travel faster than average are interpreted as cooler, denser material (such as subducting tectonic plates plunging into the mantle), while regions with slower velocities suggest hotter, less dense material (such as mantle plumes rising beneath hotspots like Hawaii or Iceland).
Tomographic imaging has revealed structures invisible to other methods: deep subducted slabs extending to the core-mantle boundary, large low-shear-velocity provinces (LLSVPs) — two continent-sized anomalies at the base of the mantle beneath Africa and the Pacific — and the complex geometry of mantle convection driving plate tectonics at the surface.