Quaking or shaking of the earth is a common phenomenon undoubtedly known to humans from the earliest times. Before the development of strong-motion accelerometers that can measure peak ground speed and acceleration directly, the intensity of the earth-shaking was estimated based on the observed effects, as categorized on various seismic intensity scales. Only in the last century has the source of such shaking been identified as ruptures in the Earth’s crust, with the intensity of shaking at any locality dependent not only on the local ground conditions but also on the strength or magnitude of the rupture, and on its distance.
Subsequent scales (see seismic magnitude scales) have retained a key feature, where each unit represents a ten-fold difference in the amplitude of the ground shaking and a 32-fold difference in energy. Subsequent scales are also adjusted to have approximately the same numeric value within the limits of the scale.
Although the mass media commonly reports earthquake magnitudes as “Richter magnitude” or “Richter scale”, standard practice by most seismological authorities is to express an earthquake’s strength on the moment magnitude scale, which is based on the actual energy released by an earthquake.
Measuring and locating earthquakes
Seismic waves travel through the Earth’s interior and can be recorded by seismometers at great distances. The surface wave magnitude was developed as a means to measure remote earthquakes and to improve the accuracy for larger events. The moment magnitude scale not only measures the amplitude of the shock but also takes into account the seismic moment (total rupture area, average slip of the fault, and rigidity of the rock). The Japan Meteorological Agency seismic intensity scale, the Medvedev–Sponheuer–Karnik scale, and the Mercalli intensity scale are based on the observed effects and are related to the intensity of shaking.
Every tremor produces different types of seismic waves, which travel through rock with different velocities:
- Longitudinal P-waves (shock- or pressure waves)
- Transverse S-waves (both body waves)
- Surface waves – (Rayleigh and Love waves)
Propagation velocity of the seismic waves through solid rock ranges from approx. 3 km/s (1.9 mi/s) up to 13 km/s (8.1 mi/s), depending on the density and elasticity of the medium. In the Earth’s interior, the shock- or P-waves travel much faster than the S-waves (approx. relation 1.7:1). The differences in travel time from the epicenter to the observatory are a measure of the distance and can be used to image both sources of quakes and structures within the Earth. Also, the depth of the hypocenter can be computed roughly.
In the upper crust, P-waves travel in the range 2–3 km (1.2–1.9 mi) per second (or lower) in soils and unconsolidated sediments, increasing to 3–6 km (1.9–3.7 mi) per second in solid rock. In the lower crust, they travel at about 6–7 km (3.7–4.3 mi) per second; the velocity increases within the deep mantle to about 13 km (8.1 mi) per second. The velocity of S-waves ranges from 2–3 km (1.2–1.9 mi) per second in light sediments and 4–5 km (2.5–3.1 mi) per second in the Earth’s crust up to 7 km (4.3 mi) per second in the deep mantle. As a consequence, the first waves of a distant earthquake arrive at an observatory via the Earth’s mantle.
On average, the kilometer distance to the earthquake is the number of seconds between the P- and S-wave times 8. Slight deviations are caused by inhomogeneities of subsurface structure.
S-waves and later arriving surface waves do most of the damage compared to P-waves. P-waves squeeze and expand the material in the same direction they are traveling, whereas S-waves shake the ground up and down and back and forth.
Earthquakes are not only categorized by their magnitude but also by the place where they occur. The world is divided into 754 Flinn–Engdahl regions (F-E regions), which are based on political and geographical boundaries as well as seismic activity. More active zones are divided into smaller F-E regions whereas less active zones belong to larger F-E regions.
Standard reporting of earthquakes includes its magnitude, date and time of occurrence, geographic coordinates of its epicenter, depth of the epicenter, geographical region, distances to population centers, location uncertainty, several parameters that are included in USGS earthquake reports (number of stations reporting, number of observations, etc.), and a unique event ID.
地震是地球表层或表层下的振动所造成的地面震动，可由自然现象如地壳运动、火山活动及陨石撞击引起，亦可由人为活动如地下核试验造成，不过历史上主要的灾害性地震都由地壳的突然运动所造成。地震的影响力涵盖岩石圈及水圈 ── 当地震发生时，可能会连带引发地表断裂、大地震动、土壤液化、山崩、余震、海啸、甚至是火山活动，并影响人类的生存及活动。
地震产生的原因是因为地壳在板块运动的过程中累积应力，当地壳无法继续累积应力时，地壳会破裂，释放出地震波，使地面发生震动，地震可由地震仪透过对地震波的观察来量测，地震震级表示地震所释放出来的能量大小，地震烈度指地震在该地点造成的震动程度，地震的发生处称为震源，其投影至地表的位置为震中。
并非世界上所有的地区都会发生地震。地震与火山分布一样，主要集中在板块相互作用的地区。目前世界上主要分为三个频繁发生地震的“地震带”：环太平洋地震带（占80%）、从地中海一路向东延伸至喜马拉雅山区和印尼的欧亚地震带、位于各大洋洋中脊的洋中脊地震带。并不是所有地震都发生在以上三个地震带，另外有一小部分大地震发生在板块内部，主要集中在大的活动断层带及其附近地区。
地震波
根据弹性回跳理论，造成地震的原因是岩石中断层的破裂。当断层破裂时，两侧的岩体会相对移动并释放出累积的能量。虽然其中大部分的能量都在克服摩擦力中损失为热能，但是剩下的部分则转换为动能，并以弹性波的形式散发出去，这些波称为地震波。地震波是地震的直接表现，因此，研究地震波的到来时间、大小、振动方式等，就可以了解一个地震的发生时间、大小、发生机制等，进而研究地震。
在地球物理学上，由于地震波具备物理上实体波的特性，因此，地震波在穿越不同介质时，便有机会发生折射、反射及全反射。当许多波叠加在一起时，还有机会发生共振，并产生驻波。换句话说，研究地震波，除了了解地震本身外，还可以一窥地球内部堂奥。因为地球很大，挖深井等直接方法研究内部构造效果有限，因此分析地震波是目前人类最常用的地球物理方法。
地震波主要分为三种：实体波、面波和尾波。
地震仪纪录下的地震波，红线是先到来的P波，绿线是较晚的S波。
地震波是地震震源瞬间散发能量初方式，当地球物质在实体波经过时，可能以三维方式（上下、左右、前后）震动。如果不同质点间的震动方向属于（相对于波速方向的）前后震动，代表震波以前后压缩、纵波的方式向外传递，这种一密一疏的震波称为“Ｐ波”。P代表主要（Primary）或压缩（Pressure）。由于Ｐ波的传播来自于在传播方向上施加压力，而地球内部几乎不可压缩，因此P波很容易通过介质传递能量。事实上，Ｐ波是所有地震波里最快的波，因此也会是地震仪第一个记录到的波。因为压缩力在固体、液体中都能存在，因此Ｐ波能在固体和液体中传播。
还有一种实体波到来的较晚，称为“S波”。S波中的S代表次要（Secondary）或剪力（Shear）。在S波的行进过程中，不同于P波，质点会在上下或左右方向震动、以横波的方式前进。因为液体无法忍受剪切，所以S波不能通过液体（例如外地核），Ｐ波则可。Ｓ波的波速约为Ｐ波的0.58倍，振幅约为P波的1.4倍。由于当地震波从地底来到地表时，S波的震动方向平行于地表的分量较多，较容易水平拉扯建筑物，而一般建筑垂直耐震能力较强，水平耐震能力较弱，故S波经常是造成地震破坏的主因。
由于接近地表的地层地震波速率较低。因此，再进地表处发生的地震，很容易把能量送进地表的低速层内，这些蓄积的能量波称为“陷波”。当累积的陷波彼此干涉，倘若发生建设性干涉，便有机会使地层共振，使能量沿地表传播。面波传递速度较S波慢一些。P波及S波干涉的面波为瑞利波（Rayleigh Wave），又称为地滚波，粒子运动方式类似海浪，在垂直面上，粒子呈逆时针椭圆形振动，震动振幅一样会随深度增加而减少。由S波相互干涉的面波为勒夫波（Love Wave），振动只发生在水平方向上，没有垂直分量，差别是侧向震动振幅会随深度增加而减少。
在近距离地震纪录（小于200千米）中，在S波后方的波包并非面波，而是尾波。地球内部虽然大致是均匀的，但小部分有不均匀的质点分布，越靠近地表越多（例如断层或岩石裂痕）。当震波向外传播时，这些不均匀或散射质点或与震波作用，产生散射现象。此散射波在纪录中会形成尾波。尾波的长短与震波耗散为热能的程度有关。例如月球因为刚性较低，耗散低，故尾波时间长。尾波如同地震图上异质性所留下的“指纹”，研究尾波，可以促进对一地地质结构之了解。