Introduction

This chapter surveys earthquake phenomena by classifying event types, causal mechanisms, observational descriptors, measurement methods, and applied practices that constitute seismology and earthquake-hazard science. The classification distinguishes event sequences (foreshock–mainshock–aftershock), rupture styles and tectonic settings (e.g., blind thrusts, doublets, interplate and intraplate events, megathrusts), special categories defined by triggering or rupture behavior (remotely triggered, slow-slip, submarine, supershear), and cascading or clustering phenomena (tsunamis, earthquake swarms). Together these categories frame hazard assessment, monitoring, and mitigation strategies.

Event sequences are organized around the mainshock—the largest shock in a localized series. Smaller earthquakes preceding a mainshock are termed foreshocks; those that follow are aftershocks, which typically cluster spatially around the rupture zone and decrease in frequency with time as the crust readjusts. Doublet sequences, in which two comparable-magnitude ruptures occur in close succession, complicate stress transfer and emergency response.

Rupture style and tectonic setting strongly influence magnitude, recurrence, and impact. Interplate earthquakes occur at plate boundaries (transforms, subduction zones, collisional margins) and often concentrate strain release along well-defined interfaces; intraplate earthquakes originate within plate interiors and commonly display different recurrence behavior and deeper stress localization. Megathrust events on subduction interfaces can attain the largest magnitudes and frequently produce widespread coastal effects, including tsunamis. Blind thrusts are compressional faults that do not break the surface, increasing the potential for unexpected damage because surface rupture is absent.

Some earthquakes are distinguished by their trigger or rupture speed. Remotely triggered events are faults that fail in response to passing seismic waves from distant large earthquakes, indicating long-range stress sensitivity. Slow earthquakes, including slow-slip events and very-low-frequency tremor, release strain over extended intervals with little high-frequency shaking. Submarine earthquakes occur beneath the seafloor and, when they produce significant vertical displacement, are the primary source of tsunamis. Supershear ruptures propagate along a fault at speeds exceeding the shear-wave velocity of the surrounding rock, concentrating energy and altering ground-motion patterns.

Cascading hazards extend the geographic and temporal footprint of seismic events. Tsunamis—ocean waves generated by abrupt vertical seafloor displacement—can propagate across ocean basins and threaten coastlines well beyond the seismic epicenter, with megathrust faults being a predominant source. Earthquake swarms consist of numerous similar-magnitude events without a single dominant mainshock; such swarms are often associated with magmatic or fluid migration and require distinct forecasting and response approaches.

Mechanistically, earthquakes arise from sudden fault slip driven by tectonic stress accumulation, but seismicity also results from volcanic processes (magma movement, dike intrusion, caldera collapse) and from anthropogenic activities (reservoir filling, mining, hydrocarbon operations, wastewater injection), the latter typically producing shallow, operation-correlated events.

Fundamental observational descriptors and wave physics underpin event location and interpretation. The hypocenter (focus) marks the subsurface initiation point of rupture and is defined by latitude, longitude, and depth; the epicenter is its surface projection. Epicentral distance—measured between an observing station and the epicenter—enters travel-time analyses. Seismic shadow zones, where particular seismic phases are absent or attenuated, yield constraints on Earth's internal structure (for example, S waves do not traverse the liquid outer core).

Seismic waves provide diagnostic information: P waves (compressional) travel fastest and arrive first at instruments, enabling rapid location and early-warning actions; S waves (shear) travel only through solids, arrive later, and generally carry larger amplitudes responsible for much structural damage. Analysis of P- and S-wave arrivals and amplitudes is central to determining hypocenters, focal mechanisms, and energy release.

Instrumentation and quantification link observations to hazard models. Seismometers record ground motion across frequency bands and form the primary dataset for locating and characterizing events. Magnitude scales (logarithmic measures such as local and moment magnitude) quantify the energy released and facilitate global comparison; intensity scales describe spatially variable shaking effects and damage at specific sites, providing complementary input for engineering and emergency planning.

Prediction and forecasting practice distinguish precise, deterministic predictions from probabilistic forecasts of likelihood over given time windows. Coordinating bodies (for example, formal committees for earthquake prediction and forecasting) integrate research, monitoring, and communication to advance short-term early warning, improve probabilistic seismic-hazard assessments, and guide public and institutional response.

Analytical and applied methods support interpretation and mitigation. Shear-wave splitting reveals anisotropy and fault-aligned fabrics; the Adams–Williamson relation links seismic velocity to density in Earth models; Flinn–Engdahl regionalizations standardize catalog reporting; seismites preserve paleoseismic shaking in sediments; and earthquake engineering translates seismological knowledge into design and retrofit criteria. Collectively, these observational, theoretical, and applied tools constitute modern seismology and its contributions to reducing earthquake risk.

Section A — Aftershocks

Aftershocks are earthquakes that follow a larger mainshock and occur in the same general area as part of the ensuing seismic sequence; they are usually smaller than the initiating event and commence immediately after the mainshock, persisting for time spans that may range from seconds to years while generally decreasing in frequency. Spatially, aftershocks concentrate around the primary rupture—often on the same fault segment or adjacent patches—so their epicentral distribution commonly delineates the lateral and depth extent of the main rupture and highlights nearby structurally weakened zones.

Mechanically, aftershocks result from stress redistribution caused by the mainshock: both permanent changes in the static stress field and transient dynamic stress transfer from seismic waves can bring neighboring fault patches closer to failure and trigger subsequent slips. Empirically, aftershock sequences display a characteristic temporal decay in occurrence described by Omori-type laws (rate ∝ 1/(t + c)) and a magnitude-frequency distribution consistent with the Gutenberg–Richter relation, implying many small aftershocks and relatively few large ones.

From a hazard and scientific perspective, aftershocks can produce additional damaging ground motion, aggravate structural failure, and induce secondary hazards such as landslides, liquefaction, or—when occurring offshore—contribute to tsunami risk, thereby influencing emergency response and short-term seismic risk assessments. Dense seismic monitoring of aftershock sequences enables precise relocation of events, mapping of the rupture plane, refinement of focal mechanisms and stress-drop estimates, and improved characterization of regional fault geometry and seismic hazard.

B. Blind-thrust earthquakes

Blind-thrust earthquakes occur on low-angle compressional (thrust/reverse) faults whose fault planes do not break the Earth’s surface, so they lack an obvious scarp or surface rupture directly above the fault trace. These faults typically dip at relatively low angles (commonly less than ~45°) and form in compressional tectonic settings such as fold-and-thrust belts, foreland basins, accretionary wedges and intra-continental contractional provinces; they are frequently located beneath sedimentary basins and under pre-existing anticlines or folds.

Rupture of blind thrusts is generally confined to shallow crustal depths within the seismogenic layer (commonly the upper ~0–20 km). Because slip is arrested before reaching the surface, seismic energy and displacement can be concentrated at depth, producing strong near-source ground motions even for moderate magnitudes. In sedimentary basins this effect is often amplified by site response and by efficient forward-directivity of rupture, resulting in elevated spectral accelerations and prolonged shaking durations.

Although the primary fault plane remains concealed, cumulative blind-thrust activity commonly produces secondary geomorphic and stratigraphic signatures rather than discrete scarps. Repeated events generate broad folding, warping, uplifted anticlines, subtle warps in Quaternary deposits, drainage deflections and localized tilting; such fault-bend or fault-propagation folds and growth strata provide the principal surface clues to buried thrusts and permit paleoseismic reconstruction where scarps are absent.

The lack of a surface trace complicates hazard assessment because blind thrusts are difficult to recognize by conventional field mapping, increasing epistemic uncertainty in seismic-source models. Their shallow location beneath urbanized sedimentary basins can concentrate shaking and damage locally, as demonstrated by historic metropolitan events (notably the 1994 Mw 6.7 Northridge earthquake). Consequently, hazard analyses must explicitly consider buried thrust sources and basin-specific amplification when estimating ground-motion and risk.

Detection and characterization of blind thrusts therefore relies on indirect geophysical, geodetic and subsurface techniques: seismic reflection and refraction imaging, dense seismicity and aftershock mapping, focal-mechanism and waveform inversion, microseismic monitoring, InSAR, continuous and campaign GPS, and borehole data. Stratigraphic folding and growth deposits can furnish paleoseismic evidence and, when precisely dated, allow estimation of slip rates, recurrence intervals and long-term seismic potential despite the absence of modern surface rupture.

From an engineering and planning perspective, managing the risk from blind-thrust faults requires probabilistic seismic-hazard analyses that incorporate subsurface fault geometries, improved pre-development subsurface imaging, basin-calibrated ground-motion models, stringent building standards for strong shaking, and targeted retrofitting in known basin settings. Over many earthquake cycles, blind-thrust deformation produces identifiable landscape and stratigraphic records—uplifted terraces, basin-margin anticlines and folded sequences—that, when integrated with geochronology, provide a means to quantify the long-term behaviour of these cryptic but potentially destructive seismic sources.

Cryoseisms

Cryoseisms are seismic events produced by the abrupt mechanical failure of frozen ground, ice bodies or water-saturated sediments when thermally or hydraulically induced stresses exceed material strength. Thermal contraction and expansion, growth and rupture of ice lenses, frost heave, hydrofracture associated with freezing pore water, and sudden collapse during thawing are the primary processes that generate an impulsive release of elastic energy and perceptible ground vibration.

These events are extremely shallow in origin, confined to the active layer above permafrost or within seasonally frozen soils, and typically involve relatively small volumes of material—from a few cubic metres to, in some cases, thousands of cubic metres. Cryoseisms are most common in cold-climate settings: Arctic and sub-Arctic regions, extensive permafrost zones, high mountain (subalpine and alpine) areas, and continental regions with strong winter freezes. River and coastal ice covers can also produce cryoseismic activity when ice jams or rapid temperature changes compromise ice integrity.

Cryoseismicity is strongly seasonal, concentrating in late autumn through early spring and often triggered by rapid temperature drops, abrupt refreezing after thaw, intense nocturnal radiative cooling, or cold snaps following warm spells. Rapid changes in ground moisture and anthropogenic actions that alter thermal or hydraulic conditions (for example, rapid reservoir drawdown or dewatering) can likewise precipitate shallow fractures. Seismically, cryoseisms are characterized by impulsive, short-duration signals with energy concentrated at higher frequencies than typical tectonic earthquakes; they commonly appear as very shallow events and may also register as airborne pressure waves or audible “booms.”

Distinguishing cryoseisms from tectonic earthquakes relies on multiple lines of evidence: very shallow focal depth, pronounced seasonality tied to freeze–thaw cycles, absence of rupture propagation and sustained aftershock sequences, high-frequency seismic content, and temporal coincidence with meteorological changes. Field indicators—such as surface cracking, broken ice, localized subsidence or displaced frozen soils—further support a cryoseismic interpretation.

Although generally smaller in energy than tectonic earthquakes, cryoseisms can produce damaging local effects: loud noises, localized ground shaking, cracking of foundations, damage to brittle building elements, displacement or collapse of road and pipeline surfaces, and rockfalls on frozen slopes. Effective identification and risk assessment combine seismic and acoustic monitoring with meteorological observations (rate and timing of temperature change), field inspection, and mapping of ground thermal regimes, permafrost distribution, ground-ice content and soil saturation. Mitigation measures in susceptible areas include designing foundations and infrastructure to accommodate shallow ground movement, ensuring adequate drainage to reduce pre-freeze saturation, insulating or elevating structures above the active layer in permafrost terrain, and integrating weather and ground-temperature monitoring into maintenance and planning to anticipate periods of elevated cryoseismic risk.

Section D — Deep-focus (plutonic) and doublet earthquakes

Deep-focus (plutonic) earthquakes are defined by hypocentres deeper than 70 kilometres (≈43 mi), placing their origins beneath the brittle upper crust and commonly within the transition from lower crust to upper mantle associated with subducting slabs. At these depths, elevated pressures and temperatures alter rupture behaviour and the generation of seismic waves compared with shallow events, so that the mechanics of faulting and the frequency content of radiated energy differ from crustal earthquakes.

Because rupture occurs far below the surface, deep-focus earthquakes rarely produce clear surface faulting; nevertheless, their seismic waves can be transmitted efficiently through the denser, more continuous mantle, allowing perceptible shaking over much larger distances. Depth therefore influences observed intensity patterns, attenuation rates and the spatial extent of felt effects, complicating simple correlations between local shaking and source magnitude.

Doublet earthquakes are sequences in which two or more principal shocks of comparable magnitude occur separated in time, rather than a single dominant mainshock followed by smaller aftershocks. Such multiplicity challenges standard aftershock classification and hazard estimation because successive large events can reconfigure stress fields, create overlapping damage zones, and necessitate separate cataloging and emergency responses.

Depth-based categories (e.g., deep-focus) and temporal/magnitude-based categories (e.g., doublet) capture different dimensions of seismicity—source depth versus the timing and relative size of events—and both are critical for interpreting tectonic setting, assessing seismic risk, designing monitoring strategies and conducting post-event analyses.

E. Earthquake swarms

An earthquake swarm is a temporally concentrated cluster of seismic events occurring within a restricted geographic volume and distinguished by the absence of a single, clearly dominant mainshock. Hypocenters in swarms typically delineate a coherent spatial pattern—such as a fault segment, fracture network or magmatic conduit—and the elevated rate of events can persist from hours to several months, standing out from the background seismicity of the region.

Unlike classic mainshock–aftershock sequences, swarms are characterized by numerous events of comparable magnitude rather than by one large event followed by declining aftershocks; this difference alters how energy release and short‑term hazard are interpreted. Causation is multifactorial and may include magmatic or hydrothermal processes beneath volcanoes, migration of natural or induced fluids that modify pore pressure and promote slip, tectonic stress changes, and anthropogenic actions such as reservoir filling or fluid injection. Determining the operative mechanism is therefore central to assessing the likelihood of continued activity and its potential consequences.

Seismological investigation of swarms depends on dense monitoring and advanced analysis—precise relocation of hypocenters, focal‑mechanism solutions, temporal clustering metrics and waveform‑similarity methods—to resolve the spatial geometry, depth distribution and evolving stress or fluid processes that drive the sequence. From a risk perspective, swarms can produce substantial local impacts because many moderate earthquakes in close succession may produce cumulative damage; swarms beneath or near volcanic systems also commonly serve as indicators of changing volcanic unrest and require integrated monitoring and response.

Statistically, swarm sequences often deviate from the standard Gutenberg–Richter magnitude distribution and the Omori decay law that describe typical aftershock populations. Consequently, forecasting swarm behavior benefits from tailored statistical and physical models that incorporate the specific temporal clustering, magnitude‑frequency characteristics and underlying physical drivers of the episode.

F. Foreshocks

Foreshocks are smaller earthquakes that precede a larger seismic event (the mainshock) and are linked to it both in space and time; by definition a foreshock has a lower magnitude than the subsequent mainshock. The interval between a foreshock and its mainshock is highly variable—ranging from seconds to much longer periods—so an earlier quake cannot be labeled a foreshock until the larger rupture occurs. Spatially, foreshocks typically cluster within or adjacent to the same fault zone or rupture patch that later fails, reflecting local stress concentration and transfer toward the forthcoming rupture area.

Because the foreshock designation depends on the occurrence of a later, larger event, identification is inherently retrospective: individual small earthquakes cannot be distinguished from ordinary background seismicity in real time without the benefit of hindsight. Nonetheless, sequences of foreshocks are seismologically informative; analyses of waveform similarity, focal mechanisms, and spatial-temporal clustering can illuminate evolving stress conditions, nucleation processes, and changes in fault coupling associated with an impending rupture, even though such signals are not deterministic predictors.

Foreshock sequences are distinct from aftershocks—events that follow a mainshock—and from earthquake swarms, which comprise clustered activity without a single dominant event; the defining characteristic of a foreshock is its temporal precedence combined with spatial linkage to a later, larger mainshock. In terms of hazard and forecasting, foreshock activity can raise the probability of an imminent large earthquake but cannot be relied upon as a consistent early-warning indicator: many foreshock-like occurrences do not culminate in a mainshock, and many mainshocks occur without identifiable foreshocks, making foreshocks a probabilistic, not deterministic, element of seismic risk assessment.

Section H — Harmonic tremor

Harmonic tremor is a sustained, quasi-continuous seismic and acoustic phenomenon associated with active volcanic systems. Unlike discrete tectonic earthquakes, which appear as isolated impulsive events, tremor is characterized by prolonged, rhythmic energy release visible as continuous low-frequency signals in both seismic and infrasound records.

The primary physical mechanisms generating tremor are oscillatory processes linked to magma and volatiles: resonant or pulsatory flow of magma within conduits and the rapid exsolution or venting of volcanic gases. Either mechanism, or their interaction, can maintain the persistent oscillations that define a tremor signal.

Tremor has both subsurface and atmospheric expressions. Seismometers record the crustal propagation of long-duration, low-frequency seismic energy, while infrasound arrays detect pressure waves produced by surface or venting processes. Combined seismic and infrasonic observations therefore offer complementary perspectives on conduit dynamics, open-vent activity, and shallow magmatic processes.

Spatially, tremor is most commonly observed at conduit systems, magma chamber margins, fumaroles and vents of active volcanoes, making it a useful spatial indicator of where magma transport and gas release are occurring. Temporally it typically persists continuously for minutes to hours or longer, distinguishing it from episodic volcanic earthquakes and enabling near-real-time monitoring of evolving eruptive behavior.

In volcanic surveillance and hazard assessment, detection of harmonic tremor is an important diagnostic for magma movement and gas venting and is routinely used to inform eruption forecasts, plume detection and alert-level decisions. Robust interpretation, however, requires integration with other datasets—ground deformation, gas-emission measurements, seismic swarms and petrological or geodetic evidence—to constrain the underlying physical processes.

Importantly, the occurrence of tremor does not by itself constitute a definitive eruption precursor. Tremor-like signals can sometimes arise from non-magmatic sources, and accurate attribution depends on detailed analysis of signal attributes (frequency content, duration, spatial distribution) in concert with independent geological and geophysical observations.

Section I — Types of Earthquake

Earthquakes may be classified according to their origin as induced, interplate, or intraplate events, each reflecting distinct processes and spatial patterns. Induced seismicity consists of relatively small, shallow tremors directly attributable to human interventions—for example fluid injection, extraction, reservoir impoundment, or mining—that perturb the local stress and strain regime of the crust. These events commonly occur in close temporal and geographic association with industrial operations and illustrate how anthropogenic activities can alter seismic risk independent of long‑term tectonic loading.

Interplate earthquakes arise at the margins between tectonic plates, where convergent, divergent, and transform motions concentrate deformation along major faults and subduction interfaces. Because plate boundary kinematics govern rupture style and focal depth, interplate seismicity produces well‑defined belts of frequent activity, from shallow strike‑slip faults at transform limits to deeper thrusting within subduction zones. In contrast, intraplate earthquakes occur within the interiors of plates, typically reflecting reactivation of ancient structures or transmission of far‑field stresses; they are less frequent, more spatially dispersed, and therefore less predictable based on plate boundary maps alone.

These distinctions have direct implications for monitoring, hazard assessment, and mitigation. Induced events demand operational surveillance and often allow attribution to specific human actions; interplate earthquakes are addressed through probabilistic models tied to plate boundary deformation; intraplate events require attention to buried, relict faults and broader stress transfer, complicating hazard forecasts in regions distant from plate margins. Understanding the causal regime—anthropogenic, boundary‑driven, or intraplate—is therefore essential for tailoring observation strategies and risk‑reduction measures.

M. Megathrust earthquakes

Megathrust earthquakes occur on the plate interface of subduction zones, where one lithospheric plate descends beneath another at destructive convergent margins. The term designates ruptures along the principal fault that separates the overriding and subducting plates, typically located at shallow depths near the trench and extending downdip beneath the forearc.

Mechanically, these events result from failure of the locked portion of the subduction interface after prolonged elastic strain accumulation as plates converge. When that stored strain is released by sudden slip, very large seismic moment is produced because rupture can span great lengths along strike and considerable width down-dip, generating strong shaking over broad coastal and inland regions.

The structural setting of megathrusts includes characteristic subduction-zone elements—an oceanic trench at the front, an accretionary prism or wedge on the overriding plate, a forearc basin, and an inland volcanic arc—reflecting the fault’s position between the descending slab and overriding lithosphere. Because rupture often deforms the seafloor, these earthquakes are prime sources of large tsunamis; they also induce intense ground motion, coastal uplift or subsidence, slope failures and landslides, and persistent modifications to coastal geomorphology.

Owing to the extensive fault area and prolonged strain buildup, megathrusts rank among the planet’s largest seismic events and thus constitute the highest-magnitude hazard at convergent margins. Repeated megathrust activity influences long-term geodynamics and landscape evolution by driving continental-margin development and mountain-building, recycling oceanic lithosphere into the mantle, and sustaining arc magmatism and volcanism.

Remotely triggered earthquakes (Section R)

Remotely triggered earthquakes are seismic events induced by a separate, typically large, earthquake but occurring well outside that event’s immediate aftershock zone. They can appear on faults and in crustal regions that were already near critical stress and may be stimulated at regional to continental distances, far beyond the area affected by the mainshock’s static stress changes.

The dominant driver of remote triggering is transient dynamic stress perturbation from passing seismic waves (body and surface waves), which can instantaneously modify stress and pore pressure on distant faults. In many cases the perturbation initiates secondary, aseismic processes—such as fluid diffusion or slow slip—that delay failure by hours to days (or longer) before seismicity is observed. Consequently, triggered events may occur synchronously with the wave train or only after a latency controlled by these slower mechanisms.

Susceptibility to remote triggering is greatest where faults, fractures or hydrothermal systems are already close to failure, where pore-fluid pressures are elevated, or where structural weaknesses (e.g., fault intersections, volcanic fields) exist. Triggered seismicity commonly appears as increased rates of small-magnitude earthquakes (microseismicity), although moderate-sized events can also be produced; the pattern and rate change in seismicity are key diagnostics of triggering.

Identifying remote triggering requires demonstrating that activity lies outside the mainshock’s aftershock region, is temporally linked to the seismic-wave passage or subsequent processes, and exceeds background seismicity levels by statistical tests. Analytical tools include rate-statistics and clustering analyses, waveform correlation, and geodetic measurements (InSAR, GPS) that can detect remote aseismic deformation. Understanding remote triggering improves knowledge of long-range stress interactions, the role of fluids and slow slip in nucleation, and coupling among tectonic, volcanic and hydrothermal systems—insights that are important for refining seismic-hazard assessments and fault-mechanical models.

Section S — Selected earthquake types

Slow earthquakes are events in which tectonic strain is released over hours to months rather than the seconds-to-minutes typical of ordinary quakes. Their ruptures propagate at very low speeds and often involve aseismic slip that produces little or no high‑frequency shaking; they are therefore most readily observed through geodetic measurements (continuous GPS, strainmeters) and by long‑period or low‑frequency seismic signals. Because strain is discharged gradually, slow events alter the timing and distribution of stress on adjacent faults and thus have important ramifications for the temporal clustering and assessment of seismic hazard.

Submarine earthquakes occur on the seafloor beneath oceans or other large water bodies and are commonly associated with plate‑boundary structures such as subduction zones, transform faults and mid‑ocean ridges. In a marine setting rupture geometry, wave propagation and surface expression differ from onshore earthquakes: seafloor displacement can directly set the overlying water column in motion and generate tsunamis, and observing these events relies on a combination of land seismic networks, ocean‑bottom seismometers and hydroacoustic or tsunami‑monitoring systems. Their coastal and offshore context therefore modifies both hazard character and instrumentation requirements.

Supershear earthquakes are distinguished by rupture fronts that travel faster than the surrounding shear‑wave speed, producing concentrated, coherent wavefronts analogous to a sonic boom. This supersonic‑like propagation concentrates radiated energy into a narrow, Mach‑cone pattern perpendicular to the fault, producing strong directional ground motions and amplified shaking at specific azimuths. Recognizing supershear behavior is important for interpreting high‑frequency seismic records, understanding rupture dynamics, and anticipating spatially variable damage patterns.

Strike‑slip earthquakes result from lateral shear on near‑vertical faults, where adjacent crustal blocks move horizontally past one another in left‑ or right‑lateral sense. They commonly produce linear surface ruptures and pronounced horizontal offsets of geomorphic features and infrastructure, and are characteristic of transform boundaries and many crustal fault systems. Compared with dip‑slip events, strike‑slip ruptures typically generate little net vertical displacement, which influences patterns of surface deformation and related hazard exposure.

T. Tsunami earthquakes

Tsunami earthquakes are events in which the tsunami produced is anomalously large relative to the earthquake size inferred from conventional short‑period seismic measurements. In these cases, magnitude estimates that rely on high‑frequency body or surface waves systematically understate the seismic moment and therefore the tsunami potential, because the source radiates comparatively little energy at short periods.

Mechanically, tsunami earthquakes involve unusually slow rupture and a dominance of long‑period, low‑frequency energy. Large coseismic slip is often concentrated very close to the trench or on the shallow portion of a subduction interface, so modest high‑frequency radiation can nonetheless produce substantial vertical seafloor displacement. Such behaviour is most commonly observed at shallow subduction zones and in accretionary prisms, or wherever a sedimentary wedge can accommodate large, slow slip near the seafloor.

Because short‑period amplitudes are muted, traditional magnitude scales (e.g., mb and other short‑period regional measures) can greatly underestimate the true moment; long‑period analysis, moment magnitude (Mw) and full moment‑tensor or finite‑fault inversions are required to capture the low‑frequency source and the amount of slip relevant for tsunami generation. Robust assessment therefore depends on integrating long‑period seismic data with geodetic observations (coastal and seafloor GPS), real‑time tsunami measurements (tide gauges, DART buoys) and forward tsunami modeling to predict local run‑up.

The key to tsunami efficiency is the transfer of displacement into the overlying water column: vertical coseismic motion or collapse of shallow slope sediments produces disproportionately large tsunamis even when felt shaking is weak. After generation, regional bathymetry and coastal geometry strongly modulate amplitudes; shallow continental shelves, submarine canyons and funnelled bays can greatly amplify run‑up for the same seafloor displacement.

From a hazard perspective, tsunami earthquakes pose a particular risk of underwarning for nearby coasts because ground shaking may be weak or brief and rapid short‑period magnitude estimates may be misleading; local arrival times are short, so reliance on felt shaking as an indicator is unsafe. Impacts therefore tend to be severe and concentrated along proximate coastlines and within the same ocean basin, where low‑frequency energy couples most efficiently into the water column. Recognizing the tsunami‑earthquake phenomenon is essential for earthquake and tsunami science, for designing early‑warning algorithms that incorporate long‑period and geodetic signals, and for coastal planning that acknowledges that weak shaking or low short‑period magnitudes do not preclude large tsunami hazard.

Volcano-tectonic earthquakes

Volcano-tectonic (VT) earthquakes are seismic events produced by brittle failure of crustal rock in response to stress changes caused by magmatic movement rather than by regional plate-boundary loading alone. They occur when magma is emplaced, intrudes as dikes or sills, pressurizes conduits, or is withdrawn from storage, each of which modifies the local stress field sufficiently to cause slip on existing fractures or to create new fractures.

Mechanically, magma injection imposes tensile and shear perturbations that favor dike propagation and high-frequency brittle cracking, whereas magma withdrawal can produce extensional collapse, subsidence or downward stress transfer that likewise provokes seismic failure. Because VT events record brittle rupture, their waveforms are typically impulsive with clear P- and S-wave arrivals, strong high-frequency content and relatively short codas. This contrasts with long-period earthquakes and volcanic tremor, which are dominated by emergent onsets and lower frequencies associated with fluid resonance.

VT seismicity is generally shallow, concentrated in the upper crust (commonly within the upper ~15 km), and tends to be small-to-moderate in magnitude—frequently below about M 5—although larger earthquakes may occur when substantial fault slip or caldera collapse accompanies magmatic processes. Hypocenter distributions often trace the geometry of the volcanic plumbing system: vertical dike swarms, lateral sills, conduits and chamber margins can be delineated by mapped VT hypocenters, providing a three-dimensional view of intrusion pathways and stress concentration zones beneath a volcano.

Temporally, VT events commonly cluster into swarms as magma propagates, stalls or interacts with heterogeneous host rock. Changes in swarm behaviour—rates, hypocenter migration (upward or outward), or the onset of new clusters—are important indicators of evolving magmatic activity. Because VT earthquakes directly record the brittle response to magma movement, they are integral to eruption forecasting when interpreted alongside deformation, gas and thermal observations: increases in VT rates, systematic migration of hypocenters and evolving focal mechanisms can signal magma ascent and elevated eruption likelihood.

Local structural controls strongly modulate VT occurrence. Pre-existing faults, joints and rock anisotropies determine where magma-induced stresses produce failure; intrusion can reactivate faults or, in the case of withdrawal, initiate collapse structures such as grabens or ring faults. Magma-driven stress changes also produce measurable ground deformation (inflation/deflation). Integrating seismic patterns with geodetic data (GPS, InSAR) allows estimation of intrusion geometry, magma volume changes and the stress evolution that drives VT sequences.

VT earthquakes pose hazards independent of eruptive output. Their ground shaking can damage infrastructure, trigger rockfalls and slope failures, and in submarine settings may destabilize slopes and generate landslides or tsunamis. Distinguishing VT events from regional tectonic earthquakes and fluid-driven seismicity relies on contextual and waveform criteria: location beneath or adjacent to volcanic edifices, impulsive high-frequency signals, focal mechanisms consistent with brittle failure, and corroborating evidence for magmatic movement or deformation.

Accurate interpretation of VT seismicity requires dense seismic networks and integrated analyses: precise hypocenter locations, focal-mechanism solutions, moment-tensor or waveform inversions, and swarm-statistics all help resolve intrusion geometries (e.g., planar dikes versus spherical sources), discriminate injection from withdrawal processes, and quantify magma transport. Such multidisciplinary monitoring is essential for characterizing subsurface magmatic processes and for assessing volcanic hazard.