A Study of Recent Earthquakes Part 24

Earthquake. Disturbed Area in Sq. Miles.

Indian 1,750,000 j.a.panese 330,000 Neapolitan 39,200 Charleston 2,800,000 Riviera 219,000 Andalusian 174,000 Hereford 98,000 Inverness 33,000

Here we see that the Charleston earthquake was perceptible over a greater area than the Indian earthquake, while the Neapolitan earthquake was inferior to that of Hereford in this respect. The explanation of course is that the boundaries of the disturbed areas are isoseismal lines corresponding to different degrees of intensity, the inhabitants of Great Britain and the United States being evidently more sensitive to weak tremors, or more observant, than those of Italy, Spain, or Central Asia. The only disturbed areas that are bounded by isoseismals of the same intensity are the two last. Very roughly, then, we may say that the intensity of the Hereford earthquake was three times as great as that of the Inverness earthquake.

POSITION OF THE EPICENTRE.

One of the first objects in the investigation of an earthquake is to determine the position and form of the epicentre. In a few rare cases, as in the j.a.panese and Indian earthquakes, when the fault-scarp is left protruding at the surface, only careful mapping is required to ascertain both data. But, in the great majority of earthquakes, the fault-slip dies out before reaching the surface and the position of the epicentre is then inferred by methods depending chiefly on the time of occurrence or on the direction or intensity of the shock.

At first sight, methods that involve the time of occurrence at different places seem to be of considerable promise. No scientific instruments are so widely diffused as clocks and watches; but, on the other hand, few are so carelessly adjusted. It is the exception, rather than the rule, to find a time-record accurate to the nearest minute; and, as small errors in the time may be of consequence, methods depending on this element of the earthquake are seldom employed. If, however, the number of observations is large for the size of the disturbed area, the construction of coseismal lines may define approximately the position of the epicentre. In the Hereford earthquake of 1896, the centre of the innermost coseismal line (Fig.

62) is close to the region lying between the two epicentres.

The method of locating the epicentre by means of the intersection of two or more lines of direction of the shock was first suggested by Mich.e.l.l in 1760,[79] and has been employed by Mallet in investigating the Neapolitan earthquake, by Professors Taramelli and Mercalli in their studies of the Andalusian and Riviera earthquakes, as well as by other seismologists. The diversity of apparent directions at one and the same place caused its temporary neglect, until Professor Omori showed in 1894 that the mean of a large number of measurements gives a trustworthy result (p. 19). His interesting observations should reinstate the method to its former place among the more valuable instruments at the disposal of the seismologist.

No observations, however, are at present so valuable for the purpose in view as those made on the intensity of the shock. For many years, it has been the custom to regard the epicentre as coincident with the area of greatest damage to buildings; and, when the area is small, the a.s.sumption cannot be much in error. It is of course merely a rough way of obtaining a result that is generally given more accurately by means of isoseismal lines; but there are exceptional cases, such as the Neapolitan and Ischian earthquakes, when the destruction wrought by the earthquake furnishes evidence of the greater value.

A single isoseismal accurately drawn not only gives the position of the epicentre with some approach to exactness, but also by the direction of its longer axis determines that of the originating fault.

When two or three such lines can be traced, the relative position supplies in addition the hade of the fault (p. 219). The successful application of the method requires, it is true, a large number of observations, and these cannot as a rule be obtained except in districts that are somewhat thickly and uniformly populated, such as those surrounding the cities of Hereford and Inverness. In the Charleston earthquake, also, the position and form of the epicentres were deduced from the trend of isoseismal lines based on the damage to railway-lines and various structures within a spa.r.s.ely inhabited meizoseismal area.

In a few cases, of which the Indian earthquake may be regarded as typical, a fourth method has recently been found of service. The numerous after-shocks which follow a great earthquake originate for the most part within the seismic focus of the latter; and, as they usually disturb a very small area, it is not difficult to ascertain approximately the positions of their epicentres. Some, as in the Inverness after-shocks of 1901, result from slips in the very margin of the princ.i.p.al focus; but, as a rule, the seat of their activity tends to contract towards a central region of the focus. Bearing in mind, then, that some of the succeeding shocks originate at and beyond the confines of the focus, and that others may be sympathetic shocks precipitated by the sudden change of stress, it follows that the shifting epicentres of the true after-shocks map out, in part at any rate, the epicentral area of the princ.i.p.al earthquake.

DEPTH OF THE SEISMIC FOCUS.

It is much to be regretted that we have no satisfactory method of determining so interesting an element as the depth of the seismic focus. That it amounts to but a few miles at the most is certain from the limited areas within which slight shocks are felt or disastrous ones exhibit their maximum effects. Nor can we suppose that the rocks at very great depths are capable of offering the prolonged resistance and sudden collapse under stress that are necessary for the production of an earthquake.

The problem is evidently beyond our present powers of solution, and its interest is therefore mainly historical. All the known methods are vitiated by our ignorance of the refractive powers of the rocks traversed by the earth-waves. But, even if this ignorance could be replaced by knowledge, most of the methods suggested are open to objection. Falb's method, depending on the time-interval between the initial epochs of the sound and shock, is of more than doubtful value.

Dutton's, based on the rate of change of surface-intensity, is difficult to apply, and in any case gives only an inferior limit to the depth. Time-observations have been employed, especially in New Zealand; but the uncertainty in selecting throughout the same phase of the movement, and the large errors in the estimated depth resulting from small errors in the time-records, are at present most serious objections. There remains the method devised by Mallet, and, though he claimed for it an exaggerated accuracy, it still, in my opinion, holds the field against all its successors. When carefully applied, as it has been by Mallet himself, by Johnston-Lavis and Mercalli, we probably obtain at least some conception of the depth of the seismic focus.

Professor Omori and Mr. K. Hirata have recently[80] lessened the chief difficulty in the application of Mallet's method. They have deduced the angle of emergence from the vertical and horizontal components of the motion as registered by seismographs, instead of from the inclination of fissures in damaged walls. In two recent earthquakes recorded at Miyako in j.a.pan, they find the angle of emergence to be 7.2 and 9 respectively, the corresponding depths of the foci being 5.6 and 9.3 miles. These are probably the most accurate estimates that we possess, and it will be noted that they differ little from the mean values obtained for the Neapolitan, Andalusian, and Riviera earthquakes--namely, 6.6, 7.6, and 10.8 miles.

NATURE OF THE SHOCK

In one respect, the earthquakes described above fail to represent the progress of modern seismology. They furnish no diagrams made by accurately constructed seismographs within their disturbed areas. The curve reproduced in Fig. 36, as already pointed out, is no exception to this statement. For another reason, the records that were obtained in j.a.pan of the earthquake of 1891 are trustworthy for little more than the short-period initial vibrations; for, owing to the pa.s.sage of the surface-waves, visible in and near the meizoseismal area, the j.a.panese seismographs registered the tilting of the ground rather than the elastic vibrations that traversed the earth's crust.

Notwithstanding this defect, personal impressions of an earthquake-shock give a fairly accurate, if incomplete, idea of its nature. Nearly all observers placed under favourable conditions agree that an earthquake begins with a deep rumbling sound, accompanied, after the first second or two, by a faint tremor which gradually, and sometimes rapidly, increases in strength until it merges into the shock proper, which consists of several or many vibrations of larger amplitude and longer period, and during which the attendant sound is generally at its loudest; the earthquake dying away, as it began, with tremors and a low rumbling sound.

[Ill.u.s.tration: FIG. 79.--Seismographic Record of Tokio Earthquake of 1894. (_Omori._)]

The vibrations that produce the sensible shock are by no means all that are present during an earthquake. The Indian earthquake, for instance, seemed to last about three or four minutes at Midnapur; but the movements of the bubble of a level showed that the ground continued to oscillate for at least five minutes longer (p. 280). Many of these unfelt waves are rendered manifest by seismographs, although there are still others that elude registration either from the extreme shortness or the great length of their periods.

In Fig. 79 is shown the princ.i.p.al part of a diagram obtained at Tokio during the j.a.panese earthquake of June 20th, 1894 (p. 18), the curve representing the N.E.-S.W. component of the horizontal motion during the first 25 seconds of the record. The instrument employed is one specially designed for registering strong earthquakes, and is unaffected by very minute tremors. Those which formed the commencement of this earthquake lasted for about 10 seconds, as shown by ordinary seismographs, and the vibrations had attained a range of a few millimetres before they affected the instrument in question. For the first 2-1/2 seconds, they occurred at the rate of four or five a second. The motion then suddenly became violent, and the ground was displaced 37 mm. in one direction, followed by a return movement of 73 mm., and this again by one of 42 mm., the complete period of the oscillation being 1.8 seconds. The succeeding vibrations were of smaller amplitude and generally of shorter period for a minute and a half, then dying out during the last three minutes as almost imperceptible waves with a period of two or more seconds.[81]

Though incomplete in some respects, this diagram ill.u.s.trates clearly the division of the earthquake-motion into three stages--namely, the preliminary tremors, the princ.i.p.al portion or most active part of an earthquake, and the end-portion or gradually evanescent slow undulations. In all three stages, however, both tremors and slow undulations may be present; and, as the latter, owing to their long period, are more or less insensible to human beings, the ripples of the final stage give the impression of a tremulous termination as described above. The duration of each stage varies considerably in different earthquakes. Thus, in a valuable study of 27 earthquakes recorded at Miyako, in j.a.pan, during the years 1896-98, Messrs. Omori and Hirata show[82] that the duration of the preliminary stage varies from 0 to 26 seconds, with an average of about 10 seconds; that of the princ.i.p.al portion from 0.7 to 26 seconds, also with an average of about 10 seconds; and that of the end portion from 28 and 105 seconds, with an average of about one minute. The total apparent duration, however, depends on the instrument employed; one of the earthquakes, that of April 23rd, 1898, disturbing the seismograph at Miyako for two minutes; while, at Tokio, a horizontal pendulum designed by Professor Omori oscillated for at least two hours. The periods of both ripples and slow undulations, again, vary from one earthquake to another; but it is worthy of notice that the average period of the undulations is almost constant in all three stages of the motion, being 1.1, 1.3, and 1.3 seconds, respectively, for the east-west component of the horizontal motion, and 1.0 second throughout for the north-south component. For the ripples, the average period is .08 second in the preliminary stage, .10 second in the princ.i.p.al portion, and .08 second again in the end portion; those of the princ.i.p.al portion being slightly larger in amplitude, as well as longer in period, than the ripples of the first and third stages.

SOUND-PHENOMENA.

Besides the ripples already mentioned, there are others of still smaller amplitude and shorter period that are sensible, but as a rule only just sensible, to us as sounds. All the known evidence points to the extraordinary lowness of the earthquake-sound. According to some observers, it seems as if close to their lower limit of audibility; while others, however intently they may listen, are unable to hear the slightest noise. In other words, the most rapid vibrations present in an earthquake do not recur at a rate of much more than about 30 to 50 per second; or, if they do, they are not strong enough to impress the human ear.

To most observers, the sound seems to increase and decrease in intensity with the shock, and so gradually and smoothly does this change take place that the sound is frequently mistaken for that of an underground train approaching the observer's house, pa.s.sing beneath it, and receding in the opposite direction. Some persons, especially if situated within the meizoseismal area, hear also loud crashes in the midst of the rumbling sound and simultaneously with the strongest vibrations. At a moderate distance, say from 30 to 40 miles, the sound becomes more harsh and grating while the shock is felt; and, at a greater distance, even this change disappears, and nothing is heard but an almost monotonous sound like the low roll of distant thunder.

The explanation of this is that the sound-vibrations are of different periods and varying amplitude, and the limiting vibrations tend to become inaudible with increasing distance, the lower on account of their long period, the higher owing to their small amplitude.

The magnitude of the sound-area depends, even more than that of the disturbed area, on the personal equation of the observers. The lower limit of audibility varies not only in different individuals, but also in different races. In Great Britain, it is doubtful whether an earthquake ever occurs unaccompanied by sound; and in the meizoseismal area the noise is heard by nearly all observers. With Italians, the average lower limit of audibility is higher than with the Anglo-Saxon race; slight shocks frequently occur without noticeable sound, but with strong ones, the larger number of observers is sure to include one or more capable of hearing the rumbling noise. The j.a.panese are, however, seldom affected by the most rapid earthquake-vibrations, and the strongest shocks may be unattended by any recorded sound. The result is manifest in the size of the sound-area in different countries. In the Hereford earthquake, the sound-area contained 70,000 square miles; in the Neapolitan earthquake, about 3,300 square miles; while, in j.a.panese earthquakes, the sound is rarely heard more than a few miles from the epicentre.

Another effect of this personal equation of the observers is that the sound-vibrations apparently outrace those of longer period. The Italians, for instance, generally hear the sound that precedes the shock, and more rarely the weaker sound that follows it. In j.a.pan, only the earlier sound-vibrations, if any, seem to be audible. In Great Britain, on the contrary, the fore-sound is perceptible to four, and the after-sound to three, out of every five observers; and these proportions are maintained roughly to considerable distances from the epicentre. It follows, therefore, that the sound-vibrations and those which const.i.tute the shock must travel with nearly, if not quite, the same velocity; and that the greater duration of the sound is due either to the prolongation of the initial movement or to the overlapping of the princ.i.p.al focus by the sound-focus. Neither alternative can be regarded as improbable, but observations made on British earthquakes point to the latter explanation as the true one.

It will be sufficient to refer to two phenomena in support of this statement. In the first place, the percentage of observers who hear the fore-sound varies with the direction from the epicentre. Thus, during the Inverness earthquake of 1901, the majority of observers in Aberdeenshire regarded the sound as beginning and ending with the shock; while, in counties lying more nearly along the course of the great fault, the sound was generally heard both before and after the shock (p. 253). In this case, then, the initial and concluding sound vibrations must have come chiefly from the margins of the seismic focus; and those from the margin nearest to an observer would be more sensible than those from the farther margin. Again, in slight earthquakes, such as the Cornwall earthquake of April 1, 1898,[83] the curves of equal sound intensity, while their axes are parallel to those of the isoseismal lines, are displaced laterally with respect to these curves, owing to the arrival of the strongest sound-vibrations from the upper margin of an inclined seismic focus.

When a fault-slip occurs, the displacement is obviously greatest in the central region, and dies out gradually towards the margins of the focus. The phenomena described above show that the evanescent displacement within these margins generate sound-vibrations only; and that the greater slip within the central region produces also the more important vibrations that compose the shock. As the former are perceptible over a limited district, while the latter may be felt through half a continent, it is clear that the sound-area should bear no fixed relation in point of size to the disturbed area, but should be comparatively greater for a slight shock than for a strong one.

VELOCITY OF THE EARTH-WAVES.

If we consider only the earthquakes here described, we see at once how great is the diversity in the estimated velocity of the earth-waves.

On the one hand, we have a value as high as 5.2 kms. per sec. for the Charleston earthquake, and, at the other end of the scale, a value of 0.9 km. per sec. for the Hereford earthquake. Between them, and equally trustworthy, lie the estimates of 3.0 km. per sec. for the Indian earthquake, and 2.1 kms. per sec. for the j.a.panese earthquake and its immediate successors.

It is difficult to account entirely for such discordance. Errors of observation may be responsible for a small part of the differences.

The initial strength of the disturbance appears to have some effect, and the nature of the rocks traversed must be a factor of consequence when the distances in question are not very great. In the j.a.panese and Hereford earthquakes, all three may have combined to produce the divergent results, the distance in these cases being only 275 and 142 kms. respectively.

In the Indian and Charleston earthquakes, the distances are much greater (1944 and 1487 kms.), and the variety of rocks traversed must tend to give a truer average. In the former, the result obtained (3.0 kms. per sec.) agrees so closely with the velocity of the long-period undulations of distant earthquakes as to suggest that it was these waves that were timed at the stations west of Calcutta and disturbed the magnetographs at Bombay.[84]

Omitting, then, the Indian estimate, we find that, for the j.a.panese and Charleston earthquakes, the velocity increases with the distance as measured along the surface. To a certain extent, such a result might have been expected, had we a.s.sumed the earthquake-waves to travel along the chords joining the focus to very distant places of observation.

The wave-paths that penetrate the earth are straight lines, however, only when the conditions that determine the velocity are uniform throughout, and such uniformity we have no reason to expect. From what we know of the earth's interior, there can, indeed, be little doubt that the velocity of earthquake-waves increases with the depth below the surface, and that the wave-paths in consequence are curved lines with their convexity downwards. It would be out of place to state more than the princ.i.p.al result of the recent investigations by Dr. A.

Schmidt[85] and Prof. P. Rudzki[86] on this subject. These are based on the a.s.sumptions that the velocity increases with the depth below the surface, and that it is always the same at the same depth. From the focus of the earthquake, wave-paths diverge in all directions. Those which start horizontally curve upwards, and intersect the surface of the earth in a circle dividing the whole surface into two areas of very unequal size. Within the small area, the surface-velocity is infinite at the epicentre, and decreases outwards until it is least on the boundary-circle. In the larger region beyond, the surface-velocity increases with the distance from the epicentre, until, at the antipodes of that point, it is again infinite. But, as the depth of the focus is always slight compared with the radius of the earth, the small circular area surrounding the epicentre is practically negligible, and we may regard the surface-velocity of the waves that traverse the body of the earth as a quant.i.ty that continually increases with the distance from the epicentre.

How fully this interesting theoretical result has been confirmed is well shown in Mr. Oldham's recent and very valuable investigation on the propagation of earthquake-motion to great distances.[87] A study of the records of the Indian earthquake revealed the existence of three series of waves, the first two consisting in all probability of longitudinal and transversal waves travelling through the body of the earth, and the third of undulations spreading over its surface (pp.

282-285). Extending his inquiries to ten other earthquakes originating in six different centres, Mr. Oldham distinguishes the same three phases in their movements; the third phase being the most constantly recorded, the second less so, while the first phase is the most frequently absent. With the exception of a few very divergent records, the initial times of these phases and the maximum epoch of the third phase are plotted on the accompanying diagram (Fig. 80), in which distances from the epicentre in degrees of arc are represented along the horizontal line and the time-interval in minutes along the perpendicular line. The dots near the two lower curves refer to the records of the heavily weighted Italian instruments, and the crosses to those of the light horizontal pendulums, which respond somewhat irregularly to the motion of the first two phases (p. 282). In the third phase, there is less divergence between the indications of the two cla.s.ses of instruments, and dots are used in each case for the initial, and crosses for the maximum epoch.

[Ill.u.s.tration: FIG. 80.--Time-curves of princ.i.p.al epochs of earthquake-waves of distant origin. (_Oldham._)]

Of the smoothed curves drawn between these series of points, those marked A, B, and C represent the time-curves of the beginnings of the first, second, and third phases respectively, while D is the time-curve for the maximum of the third phase.

The concavity of the two lower lines towards the horizontal base-line shows that the surface-velocity of the corresponding waves increases rapidly with the distance, far more so than would be possible with rectilinear motion. The rates at which these waves travel through the earth therefore increase with the depth, and the wave-paths must in consequence be curved lines convex towards the centre of the earth.

If the time-curves A and B were continued backwards to the origin, their inclinations at that point to the horizontal line give the initial velocities of the corresponding waves, which prove to be about 5 and 3 kms. per sec. respectively. Now, according to recent experiments made by Mr. H. Nagaoka on the elastic constants of rocks,[88] the mean velocity of seven archaean rocks is 5.1 kms. per sec. for the longitudinal waves, and 2.8 kms. per sec. for the transversal waves--values which agree so closely with those obtained for the first two series of earthquake-waves as to leave little doubt with regard to their character.

The other time-curves, C and D, corresponding to the initial and maximum epochs of the third phase, are practically straight lines.

Some of the records are slightly discordant for the average curve, especially for the initial epoch; but it is often difficult to define the commencement of this phase with precision. At any rate, the observations show no distinct sign of an increase in the surface-velocity of these waves with the distance from the origin. It may therefore be concluded that they travel along the surface with velocities which are practically constant for each individual earthquake, the largest waves at the rate of about 2.9 kms. per sec., and the advance waves with a velocity of about 3.3 kms. per sec., rising occasionally to over 4.0 kms. per sec.

STRUCTURAL CHANGES IN THE EPICENTRAL AREA.

Changes of elevation have long been known as accompaniments of great earthquakes, though many of the earlier observations and measurements left much to be desired in accuracy and completeness. The j.a.panese earthquake of 1891, however, placed the reality of such movements beyond doubt, and revealed the existence of a fault-scarp, with a height in one place of 18 or 20 feet, and a length of at least 40, if not of 70, miles. In the Indian earthquake of 1897, the fault-scarps were shorter, though more p.r.o.nounced in character, the largest known (the Chedrang fault) being about 12 miles long, and having a maximum throw at the surface of 35 feet. In some other recent earthquakes, also, remarkable fault-scarps have been developed. After the great shocks felt in Eastern Greece on April 20th and 27th, 1894, a fissure was traced for a distance of about 34 miles, running in an east-south-east and west-north-west direction through the epicentral district, and varying in width from an inch or two to more than three yards. That it was a fault, and not an ordinary fissure, was evident from its great length, its uniform direction, and its independence of geological structure. The throw was generally small, in no place exceeding five feet.[89] Again, in British Baluchistan, after the severe earthquake of December 20th, 1892, a fresh crack was observed in the ground running for several miles in a straight line parallel to the axis of the Khojak range. It coincided almost exactly with a line of springs, and was clearly produced by a fresh slip along an old line of fault, for before the earthquake it had the appearance of an old road, and the natives a.s.sert that the ground has always cracked along this line with every severe shock. In 1892, the change in relative height of the two sides of the fault was small, in one place where it was measured being only two inches.[90]