refraction statics 2

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2Near-Surface Topography

and Geology

2.1 INTRODUCTION

In Section 1.1, I quoted the following definition of sta-tic corrections from Sheriff (1991):

Corrections applied to seismic data to compensatefor the effects of variations in elevation, weathering

thickness, weathering velocity, or reference to a datum.The objective is to determine the reflection arrival timeswhich would have been observed if all measurementshad been made on a (usually) flat plane with no weath-ering or low-velocity material present.

In this chapter, the near surface itselfits topography orelevation profileand the thickness and velocity of theweathered layer are examined and discussed.

I stated in Section 1.3 that near-surface information isoften undersampled and that interpretive steps must betaken to interpolate values between control, or observa-tional, points. Consequently, the decisions taken or the

interpretive judgments made about the near surface canhave a significant impact on the quality of the final sec-tions and may also affect the time relief on a particularreflector or refractor. However, these effects can bereduced with the careful use of residual static correc-tions, although in most cases the success of thisapproach requires that the original datum static correc-tions be a good first approximation. In addition, thequality of the field data, and hence the final section,often depends on the nature of the near surface. The loss

of quality can vary from minimal to almost total wherecomplexities in the near surface are such that little if anycoherent reflected signals from a source are recorded atreceiver locations.

In any discussion of the near surface, it is normallyassumed that this term refers to a survey on land or per-haps the transition zone. This is not necessarily the case

because near-surface problems also occur just below thewater bottom, especially in areas of relatively rapiddeposition such as at the mouths and deltas of majorriver systems (e.g., Schatz and Lindblade, 1986; Monk etal., 1994a, b). Also, highly irregular water-bottom topog-raphy, such as submarine canyons and trenches (e.g.,Dent, 1983; Berryhill, 1986; Yilmaz, 1987) also requiretime corrections, although Section 6.2 shows that theseare not true static corrections. In addition, the waterlayer itself is not homogeneous. It contains layers of dif-fering temperature and salinity that can vary spatiallyand as a function of calendar time.

The present chapter looks first at the relationship of

datum static corrections and a simple near-surfacemodel, in which the only lateral changes are in the ele-vation of the surface. This is followed by a description ofthe weathered layer, from both a seismic and geologicstandpoint, including the various ways in which thislayer is formed and what effect it has on seismic data.The next section deals with near-surface irregularities andlooks at various types of topography and their charac-teristics as they relate to the definition of a near-surfacemodel for datum static corrections. This is followed by

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the effect of seasonal changes such as temperature andrainfall on the near surface, with emphasis on theweathered layer. The last section in this chapter dealswith near-surface velocities, both for compressional (P)and shear (S) waves, for both consolidated and uncon-solidated material, together with velocities for water, ice,

and permafrost.

2.2 TOPOGRAPHIC VARIATIONS

Although static corrections are needed to compensatefor all types of differences in near-surface conditions,often the first parameter to be considered is variations insurface elevation. However, even when the surface pro-file is flat or smooth, large variations in datum static cor-rections or traveltimes in the near surface can occur insome areas as a result of lateral near-surface changes. Itis important to remember that the final reflection time

section is really an intermediate product; the depth sectionis the final product. In some surveys, the final compen-sation for the effects of the near surface can be per-formed at the depth-conversion stage, accepting that thefinal time section still contains near-surface time anom-alies, which will normally be limited to long-wave-length (low-spatial-frequency) features.

To illustrate the concept of datum static correctionsand its implications for variations in topography, a sim-ple depth model is shown in Figure 2-1a, which shows asurface escarpment with a relief of 80 m. A simple fieldexperiment, with a collocated point source and receiver,would generate a surface-referenced reflection time pro-

file as in Figure 2-1b, in which the times from the surfaceto each of the reflectors and back to the surface areshown. With respect to the surface, there is an increasein two-way traveltime to the deeper reflector at the eastend of the profile. This false or apparent structure is relat-ed to the choice of reference surface, in this case, theearths surface.

This dependence is demonstrated by looking at thesame information plotted as a depth profile. To convertfrom time to depth, the relevant velocity information isrequired and can be obtained either from well informa-tion or from velocity information extracted from theseismic data. In the model experiment, the stackingvelocity down to the second reflector can be estimated ifadditional data are acquired at longer offsets. The rmsvelocity (approximately equal to the stacking velocity)computed using the standard Dix equation (Dix, 1955)gives velocities with respect to the surface of about 2460m/s on the west end of the profile and about 2220 m/son the east end. These velocities and the time profileshown in Figure 2-1b produce the depth profile shownin Figure 2-1c, which has been converted to absolute

depth with inclusion of the elevation profile of the refer-ence surface (earths surface). These velocities givedepths of 251 m below sea level on the west end of theprofile and 257 m on the east end, as compared to a cor-rect depth of 250 m for the whole line.

The error in the reflector depth and its lack of hori-zontality occurred because the stacking (rms) velocitywas used on its own for the depth conversion. The cor-rect approach would be to partition the stacking veloci-ty into two discrete layers with interval velocities of1600 and 2500 m/s. This assumes that the boundarybetween the two layers is recognized. Thus, it is possibleto perform the mapping from the surface and still obtaina correct depth profile, providing adequate velocityinformation is available.

10 Static Corrections for Seismic Reflection Surveys

1600 m/s

2500 m/s

3000 m/s

(a)

100

0

-100

-200

-300

West East

Elevation(m)

(c)100

0

-100

-200

-300

Elevation(m)

(b)0.0

0.2

0.4

0.1

0.3Two-waytime(s)

Fig. 2-1. Topographic model with an 80-m escarpment:(a) velocity versus depth model; (b) two-way reflectiontime profile, referenced to the surface; (c) depth profilecomputed from the reflection times and stacking velocityinformation.


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If static corrections are computed to a datum at sealevel using a near-surface velocity of 1600 m/s, theresulting two-way time profile (Figure 2-2a) shows thedeep reflector to be flat (in time). If, however, an incor-rect near-surface velocity is used, the time section nolonger shows a flat reflector. An example is shown inFigure 2-2b in which the near-surface velocity used is1400 m/s, 12.5% lower than the correct value of 1600m/s. For these parameters, the time relief on the sectionis about 14 ms (two-way time) and the absolute reflec-tion times are incorrect by about 2 ms for the low-eleva-tion portion of the line on the west. It is still possible toarrive at the correct depth picture if good estimates ofthe stacking velocity are obtained from the base of thenear-surface layer and the reflector.

Stacking velocities and their associated referencedatum are discussed in Sections 3.2 and 6.5. The simple

approach of using a space-invariant time-depth relation-ship, however, would translate the above time anomalyinto a depth anomaly with a false relief of about 16 m.Figure 2-2c shows the effect when the near-surface veloc-ity is 1800 m/s, 12.5% faster than the correct velocity of1600 m/s. This shows a time relief of about 11 ms in the

opposite direction of that in Figure 2-2b.Figures 2-3 and 2-4 are seismic sections from an areawhere the near-surface elevation profile is close to thatdepicted in the model of Figure 2-1a. Both sections areinitial stacks, but in Figure 2-3, no datum static correc-tions have been applied. The shallow event in Figure2-3, at a time of about 0.2 s at SP 10 and 0.3 s at SP 90, iscontinuous and almost a mirror image of the surfaceprofile and should be compared to the model data inFigure 2-1b. At this shallow reflection time, the range ofoffsets used to produce the stack is limited and the sec-tion is almost a near-trace display. This implies thatreflections are not mis-stacked due to inadequate datum

static corrections, and if there is sufficient signal-to-noiseratio and the reflections can be picked, a reflection timeprofile can be obtained. However, converting to depthstill requires adequate velocity information.

In the case of deeper events, however, there is a dif-ferent conclusion. For example, the event at 0.8 s at SP10appears to break up at SP 42 and then become continu-ous again from SP 54 at about 0.9 s. Deeper events alsoshow a discrete break near SP 53, such as the event atabout 1.3 s. The break-up of these eventsoften called acycle skip, in which an en echelon pattern is formedisdue to a mis-stack of the data as a result of significantlydifferent traveltimes for the various offsets being

stacked together. I include a more detailed explanationof this phenomenon in Section 6.4.3.

Thus, the section without datum static corrections(Figure 2-3) may be acceptable for shallow targets, butthe stack is unfocused or degraded and thus not accept-able for deeper events in which a large range of offsetsare included in the stack. The section with datum staticcorrections applied (Figure 2-4) shows these events to becontinuous and to have a better stack response overall,as demonstrated by the improved signal-to-noise ratio.This section also shows that events dip to the right,which may indicate genuine dip or be due to an incor-rect near-surface velocity for the computation of thedatum static corrections (similar to the effects in Figures2-2b and c), or a combination of the two.

2.3 WEATHERED LAYER

The weathered layer is a term that has a slightly differ-ent meaning to geologists and geophysicists and shouldthus be differentiated as either seismic weathering or

Chapter 2Near-Surface Topography and Geology 11

West East(a)

0.0

0.1

0.2

Two-waytime(s)

(b)0.0

0.1

0.2

Two-waytime(s)

(c)0.0

0.1

0.2

Two-waytime(s)

Fig. 2-2. Deep reflector from model data in Figure 2-1with datum static corrections applied using correctionalvelocities of (a) 1600 m/s, (b) 1400 m/s, and (c) 1800 m/s.


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geologic weathering. Seismic weathering is defined bySheriff (1991) as follows:

Anear-surface low-velocity layer, usually the portionwhere air rather than water fills the pore spaces of rocksand unconsolidated earth. Seismic weathering is usual-ly different from geologic weathering (the result of rockdecomposition). The term LVL (low-velocity layer) isoften used for the seismic weathering. Frequently the

base of the weathering is the water table. Sometimes theweathering velocity is gradational, sometimes it issharply layered. Weathering velocities are typically 500to 800 m/s (although weathering velocity may be 150m/s for the first few cm) compared to subweatheringvelocities of 1500 m/s or greater. Weathering thicknessis calculated from uphole-survey data and from refrac-tion first breaks.

Other names proposed for the seismic weatheredlayer that differentiate it from the geologic definitioninclude surface correction zone (McDermott, 1931) andaerated layer for the shallow part (Lester, 1932; Zirbel,1940; Lyons, 1946). However, it is now usually referredto as the weathered layer or the low-velocity layer (LVL)as indicated in Sheriffs definition. When two distinctnear-surface low-velocity layers are present, it is some-

times called double layer weathering; occasionally, theweathered layer, for the purpose of deriving datum sta-tic corrections, encompasses several layers. The base of

the weathered layer normally occurs where there is achange to a layer of appreciably higher velocity orwhere it no longer changes rapidly with depth. Thisboundary is sometimes coincident with the water tableor the base of geologic weathering.

In most cases, the seismic weathered layer is thickerthan the weathered layer as defined by the geologist.Geologic weathering is usually defined as the disintegra-tion or decomposition of rocks in situ. The term weath-ering refers to the atmosphere, but weather is in factonly one of several factors that cause the phenomenon.The weathering process in effect prepares rocks forsubsequent erosion and transportation by wind orwater in the form of rivers and glaciers. The processnormally occurs at the surface and works its waydownward; this is especially true for unconsolidatedsediments such as sands and clays. The processes thatcause the disintegration or decomposition of rocks areeither physical or chemical; they can also be classifiedin terms of water, gases, and organisms. Further detailscan be found in engineering geology and geomorphol-ogy textbooks (e.g., Ollier, 1984; Beavis, 1985; Rahn,1986).


Fig. 2-3. Initial stack without datum static corrections of a line containing a near-surface topographic feature with about80 m of relief.

SP 1 12 3624 48 60 72 84 96

0

50

100

Elevation(m)


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Physical processes involved in the weatheringprocess include thermal expansion, crystal growth, andliving organisms. Because rocks are generally poor con-ductors of heat, atmospheric temperature changes causeexpansion or contraction of the outer part of the rockmass down to a depth of a few centimeters below thesurface; the resulting stresses often produce cracking

which, in time, leads to disintegration. Also, the growthand development of crystals can exert sufficient pres-sure so that some rocks are eventually shattered. Themost common example of this is the formation of icefrom water, which is accompanied by an expansion involume of about 10%. The cycle is initiated with waterpenetrating small cracks in the rock; many cycles ofthawing and freezing can occur during a single season,leading to the break up of the rock. Another example issalt crystallization, which can cause disintegration in

some areas. The effects of living organisms include theroots of trees, which can widen preexisting cracks, lead-ing to disintegration. Burrowing creatures create tunnelsunderground, mainly working on the partly disintegrat-ed rock or soil, which may then be affected by otherweathering agents to produce further disintegration anddecomposition.

Decomposition or chemical weathering of rocksincludes action by water and weak acids, even thoughmany rock-forming minerals are close to being chemi-cally inert. This action partially or totally alters the rockso that its structure is weakened or its volume changed,either by an increase in size or by formation of cavities.A few minerals are soluble in water, such as rock salt,and others take up water, such as anhydrite. Oxides,such as iron oxide, are formed by the action of water andoxygen, much as rust forms on untreated metal when


Fig. 2-4. Initial stack with datum static corrections of a line containing a near-surface topographic feature with about 80m of relief (compare with Figure 2-3).

SP 1 12 3624 48 60 72 84 96

-50

-25

0

Datums

tatic

correction(ms)

0

50

100

Elevation(m)


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exposed to the atmosphere. Many minerals are solublein weak acids such as carbonic acid, formed from waterand carbon dioxide, which acts on rocks such as lime-stone to produce weathered surfaces. Another exampleof a weak acid is the humic acid in water, which occurswidely, especially in tropical areas.

As stated above, seismic weathering is composed oflow-velocity material (see Section 2.6 for velocities)and has a thickness varying from a few centimeters to50 m or more. It can be extremely irregular and varyrapidly in both lateral and vertical directions, lithology,thickness, density, and velocity (see Section 2.4 fordetails). The weathered layer normally has a very highrate of energy absorption and often acts as a scattererof energy.

Tullos and Reid (1969) measured an attenuation of13.1 dB per cycle, or Q = 2 (since attenuation in dB percycle is about 27/Q), in the top 3 m (10 ft) of a loam,sand, and clay in the near surface of the Gulf Coast

region. Rice et al. (1991) reported Q = 1 or less for the top0.3 m (1 ft) for four different types of near-surface con-ditions, Q 2 for the depth range of 0.30.9 m (13 ft)and Q 5 for the depth range 0.92 m (36 ft). Meissnerand Theilen (1984) reported many values of Q < 10. Forshear-wave data, Meissner et al. (1985) noted Q valuesof 5 for loam, 6.5 for Tertiary mudstone, 8 for diluvialsand, and 20 for alluvial silt (see Kudo and Shima, 1981);for peat, the attenuation was severe for P-waves but notfor S-waves. Stmpel et al. (1984) observed the absorp-tion in S-waves to be lower than in P-waves for partial-ly saturated sands and boulder clays.

The weathered layer usually acts as a high-cut (low-

pass) filter (e.g., Meekes et al., 1990), such that it is oftendifficult to obtain high-frequency data unless specialsteps are taken. These steps include placing the sourcebelow the weathered layer, as is done routinely formany dynamite surveys, and in some instances, placingthe receivers below the weathered layer as well (e.g.,Gaby and Solari, 1948; Pullin et al., 1987; Rice et al.,1991). The characteristics of the weathered layer alsoaffect the type of ground roll propagation that occurs.However, if a deep source is used and is located belowthe weathered layer, the amount of ground roll generat-ed is usually substantially reduced.

If the interface between the base of the weatheredlayer and the underlying unweathered layer corre-sponds to a significant change in acoustic impedance, itmay act as a multiple generator or be the interfaceresponsible for the generation of ghost reflections (vanMelle and Weatherburn, 1953). A large reflection coeffi-cient of 0.7 to 0.8 was observed by Levin (1962) at themudwater interface in Lake Maracaibo, Venezuela, andJones et al. (1958) reported coefficients of 0.85 to 0.95at a lake in the United States.

In areas where the weathered layer changes rapidlyin the horizontal direction, the traveltimes of reflectionor refraction data may be sufficiently different for thevarious components in the source or receiver array thatthe downgoing or upcoming signal is significantlyattenuated. Examples of such intra-array raypath prob-

lems are described in Section 2.4.2.In some dynamite surveys, several shots are occa-sionally located in the same hole. Once a hole is drilled,however, the near surface is no longer the same due tothe disturbance of the material surrounding the holeand to the invasion of the drilling fluid. This implies thatestimates of the near-surface velocity from an upholesurvey no longer measure the near surface in its originalcondition, a point discussed further in Chapter 4 (in par-ticular, see Section 4.4.3). Similarly, once a charge hasbeen detonated, the material in the vicinity of the chargeis further changed. For example, Lyons (1946) noted thatthis slowed the velocity of the material near the shot and

observed traveltime differences of 24 ms, and occa-sionally up to 68 ms, as a result of this disturbance.

2.4 NEAR-SURFACE IRREGULARITIES

Various parameters are associated with the near sur-face, including the surface elevation or water-bottomtopography, the weathered layer or LVL, and in someareas, one or more subweathered layers. A simple near-surface model was used in Section 2.2 (Figures 2-1 and2-2) to show the relationship between changes in eleva-tion and the resulting datum static corrections; Figures

2-3 and 2-4 demonstrated the impact of this on real data.Recall from Section 2.3 that the weathered layer is usu-ally that part of the near surface where air rather thanwater fills the pore spaces of rocks and unconsolidatedearth. Its base is normally defined as the depth where achange to an appreciably higher velocity occurs orwhere the velocity stabilizes, which sometimes coin-cides with the water table.

Variations in the physical properties of the weatheredlayer, in both the horizontal and vertical directions, andconditions under which the subweathered layer shouldbe included in the computation of datum static correc-tions both need to be analyzed by the geophysicist. Iftaken to extremes, this includes anomalous layers atgreat depths below the surface, such as gas pockets andgas hydrates at depths in excess of 1000 m. Section 6.2covers ways in which such deep-seated anomalies aretreated, which do not normally fit the definition of a sta-tic correction because a simple time shift is no longerassociated with a particular surface location.

It was suggested by Thralls and Mossman (1952) thatthe type of datum static correction technique required



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for an area depends on the types of near-surface forma-tions that are present. They broadly divided these intoeither youthful or mature topography. Youthful topogra-

phy is characterized by active vertical erosion (Figure2-5), whereas in mature topography (Figure 2-6), the sur-face gives no real indication of the variations in the nearsurface. Examples include old channels or glacial val-leys that have been infilled with recent sediments. Insome areas, mature topography has been elevated orrejuvenated, and if followed by erosion, both types oftopography will be present. Most near-surface cross-sec-tions can be broadly classified into one or the other ofthese categories, or a combination of the two.

If static corrections are to be computed to a flat refer-ence plane or datum (which is the usual practice, asindicated by the definition in Chapter 1), then the near

surface often includes consolidated material, althoughthis depends on the datum plane elevation. Examples ofthis are shown in Figure 2-5 (under the hills) and inFigure 2-6 (on either side of the old channel). It is gener-ally not sufficient to consider just the weathered layer;information about the subweathered layers, at leastdown to the reference plane, must be included. Thus,datum static corrections and weathering corrections (thetime through the weathered layer) are generally not thesame; the former usually includes time through consol-idated layers as well as the time through the weatheredlayer. I discuss the component parts of datum static cor-rections in more detail in Sections 3.1 and 3.5.

The regional distribution of the LVL (Lyons, 1946) canbe broadly classified as follows:

1. areas that are approximately uniform;2. areas where the LVL and other anomalous layers

are thicker on hills and thinner in valleys, such asthose tied to the water table;

3. areas where the LVL is thicker in valleys andthinner on the hills, implying more alluvial fill,for example; and

4. areas where the LVL is distributed almost

randomly.I show later in this section that the velocity and thick-ness of weathered and subweathered layers may changeslowly, or rapidly and erratically, in both the horizontaland vertical directions.

Seismic acquisition is undertaken in many countriesaround the world encompassing a great variety of topo-graphic settings. Several examples of near-surface mod-els are given in the following sections and are used toillustrate specific problems with the definition of thenear surface and subsequent computation of datum sta-tic corrections. This is not intended to be a complete list

of all possible types of topography, but more a represen-tative set that is more detailed than the two types men-tioned earlier. In addition to examples of youthful andmature topography, examples are given of sand dunetopography, areas where the weathered layer is highlyirregular (such as swamps and large river deltas), per-mafrost topography, and mountain front topography.

In the case of marine surveys, the other near-surfacefeature that should be considered is the water layer. It isgenerally assumed to be homogeneous, apart fromlarge-scale variations in velocity with depth and tem-perature. In some areas, however, the water layer con-tains layers of locally different temperature and salinity.

These can occur, for example, near the mouth of a largeriver system, where the salinity decreases due to theinflux of freshwater. In some areas, salt is leached outfrom the seabed, causing local changes in salinity. Someareas are influenced by major ocean currents which cancause appreciable changes in water temperature. Thesevarious factors are also likely to vary as a function of cal-endar time, a topic discussed in Section 2.5. The impactof changes in temperature and salinity on water veloci-ty is described in Section 2.6.3.


Elevation(m)

150

100

50

00 5 km

V1

V2

V3

V4

V5

Weathered layer

Fig. 2-5. Generalized cross-section illustrating youthfultopography, containing a weathered layer and five sub-weathered layers; vertical exaggeration is 75.

350

Elevation(m)

300

250

200

V1

V2

V3

V4

Weathered layer

0 4 km

Fig. 2-6. Generalized cross-section illustrating maturetopography, containing a weathered layer and four sub-weathered layers; vertical exaggeration is 50 (afterThralls and Mossman, 1952).


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2.4.1 Sand Dune Topography

Sand dunes are formed by wind-blown sand andoccur mainly in arid regions; they vary from an insignif-icant height up to about 200 m. The dips on the wind-ward slope of a dune can be as much as 1520, but areoften much smaller than this, whereas on the slip faceslope they are about 30, which approximates the maxi-mum stable slope that can be obtained by loose sand. Asdunes increase in height, some compaction occurs, suchthat the velocity near the base of the dune and awayfrom the edge is likely to be higher than it is close to thesurface. Because dunes are formed by the wind, theirposition can change with time, which may involve ahorizontal movement of many meters per year. Thisimplies that their elevation profile and near-surface

characteristics are calendar time variant (see Section 2.5).In many areas, gravel plains or sabkhas form the baseof the dunes themselves (e.g., Robinson and Al-Husseini, 1982; Friedmann, 1988). A refractor, typicallywith a smoother profile than the surface topography, ispresent below the sand in many sand dune areas andcan be used to estimate traveltimes in the near surfaceusing the techniques described in Chapter 5. Additionalnear-surface complications may arise as a result ofanomalous zones in the area, such as high-velocitystringers (thin high-velocity layers within low-velocityformations). Thus, the near-surface model of sand dunetopography can be summarized as typically having a

broad band surface profile (low to high spatial frequen-cies), with the base of the weathered layer predomi-nantly low frequency.

A generalized cross-section of an 8-km line showingdunes with a vertical relief of 6090 m is shown inFigure 2-7 and is illustrative of the topography of theRub Al-Khali, Saudi Arabia. In this example, the dunesare about 1 km wide and their base is seen to increase inelevation by about 5 m toward the right end of the pro-file, equivalent to a dip of less than 0.1.

To examine for the effects of compaction, a crossplot

is generated of traveltime through many dunes in thearea against the elevation of the dune surface above thesabkha or dune base. The values can be estimated fromuphole surveys or from arrival times from a refractorclose to the elevation of the sabkha. Figure 2-8 providesan example that encompasses observations from a largenumber of dunes. It shows that the velocity generallyincreases as the thickness of the dune increases. Forthicknesses of a few meters, the velocity ranges fromabout 150 m/s to more than 1500 m/s; for thicknesses of


200

100

Elevation(m)

0 2 km

0

Fig. 2-7. Elevation profile of sand dune topography fromthe Rub Al-Khali, Saudi Arabia; vertical exaggeration is20 (after Robinson and Al-Husseini, 1982).

Sandth

ickness(m)

0

25

50

75

100

One-way traveltime (ms)

0 50 100 150

Fig. 2-8. Crossplot of one-way traveltime versus sanddune elevation above sabkha (after Robinson and Al-Husseini, 1982).


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about 2060 m, the interval velocity is about 700 m/s,which increases to about 950 m/s for greater thickness-es. The limitations of a sand compaction curve are thatthe velocity often varies from dune to dune and fromone part of the dune to another, such as from the wind-ward side to the slip face. However, a sand compaction

curve is often used successfully as a first approximationin the computation of datum static corrections.To compute datum static corrections, a definition of

the velocities and thicknesses of the layers forming thenear surface is required. The sand compaction curve isone method that can be used to define the near-surfacevelocities, with the thickness of the sand assumed to bethe elevation above the sabkha or sand base. However,this requires an adequate compilation of values beforethe method can be initiated. As previously stated, trav-eltimes are normally estimated from uphole or refrac-tion surveys. In both cases, the height and complexity ofthe dunes may influence the particular approach to be

used; for example, it is often difficult to position upholedrilling rigs near the tops of dunes, and it is awkward todrill in soft sand. Also, if the refractor is at a considerabledepth below the surface, there are some constraints onthe refraction approach used (see Chapter 5). Otherappropriate methods to obtain near-surface informationare discussed in Section 3.8, including refraction-basedmethods with uphole control and the use of a near-sur-face reflection crew, designed to obtain shallow reflec-tions that can be tied to deep uphole surveys.

2.4.2 Highly Irregular Weathered Layer

Near-surface topography can be classified as havinga highly irregular weathered layer when rapid changesoccur in near-surface velocity and/or in thickness; inmost cases, this is associated with recent uncompactedsediments. Examples include swamps, bogs, andmarshes (especially those with marsh gas), recentdeposits near the mouths of large rivers, some loessdeposits, karst limestone, alluvial fans or stream chan-nels, and mature topography where irregular bedrockhas been infilled with sediment. In these areas, the near-surface weathered layer changes rapidly along the linein both thickness and average velocity.

An example is shown in Figure 2-9 in which consid-erable variations take place over distances of less than100 m in both weathering thickness (Figure 2-9a) andaverage weathering velocity (Figure 2-9b). The velocityprofile shows a rapid lateral change from about 1000 to200 m/s over a distance of about 200 m where some ofthe low-velocity segments are associated with swampyterrain.

In such areas, there can be large variations in datumstatic corrections (tens of milliseconds) between adjacent

source (and receiver) locations for group intervals assmall as 20 m, even though the elevation changes arecomparatively small. Connelly et al. (1991) quoted dif-ferences of 040 ms for adjacent traces in an area charac-terized by a significant thickness of Quaternary glacial

till. Lynch (1986) observed short- and long-wavelengthstatic problems in the range of 2080 ms as a result of alow-velocity (about 300 m/s) biomass layer beneathlakes in Alberta, Canada.

Intra-array static corrections can become sufficientlylarge in these areas for the downgoing or upcoming sig-nals to be appreciably attenuated, unless the arrays areshort (or contoured if the elevation profile is the cause ofthe irregularity). It is generally better to preserve the sig-nal and record the unwanted noise rather than attenuateboth the signal and the coherent noise (discussed furtherat the end of this section and in Section 6.6.1).

This type of near-surface model can be summarizedas one in which both the surface profile and the base ofthe weathered layer have a broad band of spatial-fre-quency components. In the case of swampy areas, how-ever, the surface profile is normally low frequency.

Swamps are characterized as wet and saturated withmoisture but not usually covered with standing water.Mangrove swamps, however, are tidally submerged,salt-tolerant, coastal woodlands, with deposition oftenfrom both inland rivers and the sea and generally com-


0

Elevation(m)

(a)

400 m

120

80

40

0

0

1000

500

Velocity(m/s)

(b)

Fig. 2-9. Near-surface profiles illustrating highly irregularweathered layer: (a) surface elevation and base of weath-ered layer; vertical exaggeration is 5; (b) average weath-ering velocity (computed from uphole survey times).


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posed of soft organic-rich sandy to clayey mud. Swampsalso occur in the Arctic tundra, especially in the summermonths when the frozen surface soil above the per-mafrost starts to thaw, where they are called muskeg.Pullin et al. (1987) buried phones beneath the muskeg,which varied in thickness from 5 to 10 m, to minimize

the effects of the large static variations and to improvedata quality. Swamps and recent deposits near themouths of large rivers are likely to have biogenic gaspresent, formed in place by the anaerobic decomposi-tion of organic matter. This can have a significant effecton the velocity, lowering it appreciably below the veloc-ity of sound in air (see Section 2.6).

A high rate of sedimentation is normally associatedwith large present-day delta systems, such as theMississippi and Yellow rivers. For example, it is esti-mated that the sediment deposited near the mouth ofthe Yellow River exceeds 1 billion tons per year(Milliman et al., 1985). Shallow gas is often present in

such areas, which (as stated above) significantly reducesthe velocity. For example, Tinkle et al. (1988) reportedMississippi Delta muds in which the first 30 m (100 ft)had a velocity less than 300 m/s (1000 ft/s), with thelowest observed interval velocity less than 120 m/s (400ft/s). These zones of abnormally low velocities are usu-ally localized and can often be seen on geotechnical sur-veys targeted at the top few hundred meters below thewater bottom. Anderson and Bryant (1987), for example,showed maps of the occurrence of such gas anomaliesfor much of the shallow waters off the Texas andLouisiana coasts.

Delta front features include mudflow gullies and

lobes, collapsed depressions, and intergully platforms,which are typically quite variable in both spatial andvertical directions. In some cases, these anomalies arecalled mud lumps (Fulton, 1980). Near-surface studiesoffshore from the Mississippi Delta by Meeder et al.(1988) indicated the presence of short-wavelength (high-frequency) variations in excess of 100 ms. Monk et al.(1994a, b) reported trace-to-trace traveltime differencesapproaching 100 ms in the West Delta area of the Gulf ofMexico. They also noted that the water-bottom reflectiv-ity varied from 0.4 to 0.4 due to the hard water bottombeing overlain by soft mud in some places. Many otherareas are described in the literature, including Cooper etal. (1979) on the Aleutian Basin in the Bering Sea. Inaddition, a variety of case histories have been presentedand published as proceedings at Offshore TechnologyConferences (OTC).

Anomalously low velocities have been reported frommany lakes, which are also associated with the presenceof shallow gas. Levin (1962) reported mud velocities aslow as 90 m/s (300 ft/s) in Lake Maracaibo, less thanone-third of the velocity of sound in air. Rapid changes

in velocity are observed in such areas, both in verticaland horizontal directions. For example, at a lake in theUnited States, Jones et al. (1958) noted that the velocitychanged by a factor of two in a horizontal distance of 8m (25 ft) and that the velocities were very low, at 75170m/s (250550 ft/s). (More examples of low velocities are

given in Section 2.6.2, and datum static corrections inmarine areas are discussed in Section 3.6.)An example of loess topography in China was given

by Herd (1990), who noted that the loess deposit can beas thick as 300 m with rapidly changing elevation andthickness, such that intra-array static corrections are sig-nificant. The near-surface velocities range from 600 m/sabove the water table, where the deposits are in a drystate, to about 1200 m/s at the water table. In this area,the loess is thought to be an eolian sand (loess is gener-ally considered to be fine-grained loam with grainscoarser than clay but finer than sand).

Steeples et al. (1988) described irregular bedrock

under a comparatively smooth surface profile at a depthof 48 m below the surface. The time profile, mapped bya high-frequency reflection technique, indicated a varia-tion of more than 15 ms (one-way time) in a distance of40 m. Along with the low velocities in this part of theTexas panhandle, this translates into a depth variation ofabout 4 m. If we assume that the bedrock has a muchfaster velocity than the weathered layer, this scale oftime variation again indicates the need for short arraylengths to minimize signal attenuation by intra-arraystatic differences. Many other irregular bedrocktopographies are reported in the literature (e.g., Hunteret al., 1984; Miller et al., 1989; Miller and Steeples, 1990;

Steeples et al., 1990).A further example, from the San Joaquin Valley in

California, was described by Gaby and Solari (1948) inwhich the surface is peat with zones of sand and sandyclay. The lateral variation is such that a 9-m (30-ft) holecould be all peat in one location, but 15 m (50 ft) away atthe next location, it could be all sandy clay with little orno peat present. The near-surface velocities in the areavary from 130 m/s (425 ft/s) at one location to about 600m/s (2000 ft/s) at a neighboring location. The large vari-ations observed in refraction arrival times were attrib-uted to near-surface velocity changes and not to depthchanges, on the assumption that the water table is flat.This was subsequently confirmed during a later part ofthe survey when the detectors were buried to a depth of8.5 m (28 ft), below most of the near-surface anomalies.

The effect of intra-array static corrections wasdemonstrated by Berni and Roever (1989) who com-pared array outputs with and without the application ofsuch corrections. The primary data set was recordedwith a group interval of 0.25 m in a high-velocity karstlimestone area of the Paris Basin, where caves are pre-



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sent. These comparisons showed significant improve-ments in the output with the intra-array static correc-tions applied. Steeples et al. (1990) used synthetic data todemonstrate signal distortion for arrays of variouslengths using static corrections derived from a shallowhigh-resolution reflection profile in west Texas. Arrays

are designed to attenuate coherent noise and therebyimprove the signal-to-noise ratio. In areas where signif-icant intra-array static corrections occur, it is generallybetter to preserve the signal and record the unwantednoise rather than attenuate the coherent noise and thesignal with the use of long arrays. This assumes that thenoise can be attenuated by subsequent data processingtechniques (see Section 6.6.1).

The computation of datum static corrections requiresa description of the near surface. In rapidly changing sit-uations such as those described in this section, discreteinformation is required at each source and receiver loca-tion. Appropriate methods (discussed in Section 3.8)

include refraction-based methods and high-resolutionshallow reflection surveys. The more standard approachof interpolation between uphole locations is unlikely tosample the near surface sufficiently. This was demon-strated by the profile in Figure 2-9b where uphole sur-veys would have to be located at an uneconomical spac-ing of less than 200 m to sample the changes in weath-ering velocity adequately.

2.4.3 Youthful Topography

Youthful topography is characterized by active verti-cal erosion, as shown in Figure 2-5. In this example from

the Paris Basin, the weathered layer is only a few metersthick. Immediately below the weathered layer, however,there are a succession of layered rock units whose effectsshould be understood before datum static corrections arecomputed. This is because the layers change laterally inboth thickness and velocity and contribute to the long-wavelength component of the datum static corrections. Ifthese variations are not taken into account, the resultingincorrect near-surface corrections could lead to false timestructures. I show in Sections 6.3 and 6.6.3, however, thatthe common-midpoint (CMP) stack response is primari-ly dependent on the high-spatial-frequency (short-wave-length) component of the datum static corrections. Thismeans that a simpler near-surface model (without thelateral changes in the deeper layers) may be adequate forthis component of the static corrections.

Hanot and Renoux (1991) described near-surfaceproblems associated with Upper Cretaceous chalks inthe Paris Basin. The chalk is subdivided into three layerswith velocities of about 2700, 5000, and 3300 m/s, andthese vary in thickness and velocity along the line (seeSection 4.4.4 for more examples). Examples of youthful

topography from Wyoming were documented byThralls and Mossman (1952), and Selem (1955) showedexamples from several lines in southern Italy.

This type of near-surface model can be summarizedas one in which both the surface profile and the base ofthe weathered layer are broad band in a spatial-frequen-

cy context. In addition, the subweathered layers must bedefined adequately for the computation of datum staticcorrections. If the subweathered layers are less complex(e.g., if just one homogeneous formation is present), thenthe definition of the near surface is relatively simple.

To compute datum static corrections, a description ofthe near surface is required; in the case of youthfultopography, this means a definition of the weatheredand subweathered layers. Appropriate methods are dis-cussed in Section 3.8; these include refraction-basedmethods and uphole surveys. The refraction data areused for the weathered layer and possibly for the defin-ition of the time through one or more of the subweath-

ered layers. The interpolation between the uphole sur-veys can be based on refraction information or some-times on shallow reflection information.

2.4.4 Mature Topography

In mature topography, the surface profile gives no realindication of the variations in the near surface. This isbecause old topography, such as river channels or glacialvalleys, are infilled with sediments. The example shownin Figure 2-6 depicts an old channel about 8 km wide and100 m deep. The dimensions of such features vary wide-ly from area to area, with depths from tens of meters to

several hundred meters. Osterhoudt (1946) reported anancient Mississippi river channel observed to be a steep-sided valley more than 180 m (600 ft) deep with a widthof 8 km (5 miles). Thralls and Mossman (1952) describeda feature from Kansas over 100 m deep with a width ofmany kilometers. Examples of much smaller features,with channel widths of 160300 m and depths of 2040m, were documented by Haeni (1986). In some cases,such features may be isolated but many extend overappreciable distances. Their areal extent can be mappedwith information from a grid of lines as shown byKotcher et al. (1984), who described deep erosional val-leys filled with glacial drift in northern Michigan.

If the channels are deep, then it is likely that there willbe a compaction effect on the channel-fill velocity, in asimilar way to sand thickness velocities referred to inSection 2.4.1. This type of near-surface model is one inwhich the surface profile bears little relationship to thesubsurface bedrock, and either may have a smooth orrugged profile.

In describing the near surface of mature topographyfor computation of datum static corrections, an ade-



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quate definition of the weathered and subweathered lay-ers is required. In some areas, methods (discussed inSection 3.8) must accommodate the great depths in theold topography. These include refraction-based methodsfor the weathered layer and the definition of any channelor equivalent features, as well as uphole surveys.

However, a long recording spread may be required if therefractors are deep (see Chapter 5). If the channel is broadwith respect to the spreadlength, a detailed descriptionof the long-wavelength component of the datum staticcorrections is required to obtain the correct structuralrelief on the final section. This requirement may not benecessary to achieve an adequate CMPstack response, asthis depends on the high-frequency (short-wavelength)component (see Sections 6.3 and 6.6.3).

2.4.5 Permafrost Topography

Permafrost can be simply defined as permanently

frozen soil or rock (Sheriff, 1991); it is more preciselydefined as a thermal condition of soil and rock in whichthe temperature stays below 0C for at least two consec-utive winters and the intervening summer. This timeframe is a minimum condition, as some permafrost islikely to be thousands of years old. At such tempera-tures, the water in the pore spaces of rock freezes, result-ing in a significant increase in the host rock velocity (seeSection 2.6.4).

Permafrost is found over large areas of land in theArctic, Antarctic, and adjoining regions (e.g., Black,1954), where it also extends beneath lakes, seas, andoceans. Permafrost thickness, away from large bodies of

water, is about 3060 m multiplied by the magnitude ofthe mean annual temperature of the ground surface (indegrees Celcius) and can be as much as 600 m (Taylorand Judge, 1985; Specht et al., 1986). Its thicknessdecreases offshore as a result of the warming effect ofthe sea. The base of the permafrost is governed by theaverage thermal conditions over a long period; it is at adepth where the influence of the cold surface conditionsare balanced by the normal increase of temperature withdepth. The position of the top of the permafrost, how-ever, is controlled more by annual seasonal fluctuationsin temperature.

Offshore, the upper boundary of the permafrost canalso vary appreciably. Rogers and Morack (1980) indi-cated a change in depth from 100 to 22 m beneath theocean bottom when going from 14 to 17 km offshore,although no direct correlation between permafrostdepth and distance offshore was noted. Several profilesfrom the Beaufort shelf were shown by Palacky andStephens (1992) in water depths ranging from about 10to 50 m. As a result of the anomalously high velocity ofpermafrost, there is an associatedpull-up in seismic time

that can amount to hundreds of milliseconds. A modelfrom Prudhoe Bay, Alaska, is shown in Figure 2-10, inwhich the permafrost has a velocity of 3660 m/s and the

equivalent unfrozen sediments offshore a velocity of1980 m/s. This model shows that a thickness of 610 m ofpermafrost (Figure 2-10a) results in a pull-up of about280 ms (Figure 2-10b).

The apparent dip on the seismic line shown in Figure2-11 varies from about 500 ms in the shallow part of thesection (1.01.5 s) to about 400 ms deeper in the section.A large percentage of this apparent dip is caused byanomalously high near-surface velocities at the southend of the line which are associated with a thick per-mafrost layer beneath the onshore part of the line whichthins offshore to the north. This line thus represents amajor difference between the apparent structure on atwo-way time section and a depth section.

Booker et al. (1976) showed a localized permafrostmelt problem associated with a lake in the CanadianArctic where there is a sag in the time section of morethan 150 ms (see Figure 7-64); this sag corresponds tolocations beneath the lake where less high-velocity per-mafrost occurs. Todd et al. (1991) indicated that thethickness of permafrost below a lake can be significant-ly less than in adjoining areas.


0

1

2

3

Depth

(km)

(a)Coast

1980 m/sPermafrost3660 m/s

Sedimentary section2410 m/s

(b)2.0

2.2

2.4

Two-

waytime(s)

Fig. 2-10. Permafrost transition zone model showingapparent time dip on seismic reflections: (a) velocity ver-sus depth model; (b) two-way reflection time profile (afterSpecht et al., 1986).


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In this type of near-surface model, the key element isto know the thickness and lateral extent of the per-mafrost, which can be considered to be the subweath-ered layer. The base of the permafrost is predominantlylow frequency, apart from a few places where localizedhigher heat flows result in thin spots, often called taliks.

In some areas, localized melt zones occur within the per-mafrost. Many other terms, such as frost boils, ice poly-gons, ice wedges,peat rings, andpingos are associated withnear-surface permafrost features.

To compute datum static corrections, we need a defi-nition of the permafrost. Appropriate methods (dis-cussed in Section 3.8) include refraction-based methods(although the high velocity of the permafrost normallyprecludes their use in mapping its base), methods asso-ciated with the dispersion properties of a thick near-sur-

face layer, and other geophysical (nonseismic) tech-niques. The base of the permafrost is a transition zone,not normally associated with a discrete change inacoustic impedance, such that mapping it with a reflec-tion survey is often not possible. In most cases, the long-wavelength variations in permafrost thickness are

accounted for in time-to-depth conversion; the per-mafrost profile and its velocity are needed for thisprocess.

2.4.6 Mountain Front Topography

An example of a generalized cross-section from thefolded mountain front in southwestern Iran is shown inFigure 2-12, in which the formation boundaries are basedon surface observations by a field geologist; this is an


Fig 2-11. A 15-km seismic line illustrating two-way traveltime pull-up on the onshore segment of an onshoreoffshorepermafrost line.

3.0


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interpretive procedure, especially in areas of such com-plex geology. The weathered layer itself is often thin, per-haps only a few meters thick. Beneath the weatheredlayer, the various rock units may have velocities similarto or appreciably different from one another.

The concept of static corrections, with its assumption

of vertical raypaths in near-surface layers, is likely to bea very poor approximation in such areas. This problemis minimized if the velocities are slow with respect to theformations deeper in the section and do not vary appre-ciably from unit to unit (see Section 6.2). However, inareas of mountain front topography, the velocities aremore likely to be fast than slow. The near-surface profilecan be summarized as one in which the main element isthe complex geology, together with a thin weatheredlayer at the surface.

To compute datum static corrections for complexgeology may involve geologic mapping and/or geo-physical methods. Approaches (discussed in Section 3.8)

include refraction- and uphole-based methods, whichare used to compute the velocity of the geologic units inthe area. Many of the refraction interpretation methodsare unlikely to be applicable, however, because of thesteep dips that are present in such areas.

2.5 TIME-VARIANT CHANGESIN NEAR-SURFACE LAYERS

Some parameters associated with the near surfacecan change with the seasons or with calendar time. This

results from such factors as temperature changes, rain-fall, tidal effects, ice movement, wind, recent erosionand deposition, earthquakes, and human activities. Ifthe results of a survey are analyzed on an areal basis andthe observations were not obtained at the same time orduring the same season, the differences noted caninclude geologic and seasonal or time-variant changes.Careful analysis is then required to differentiatebetween these effects.

The recognition of any seasonal or time-variantchange is very important for monitoring or time-lapsesurveys, when a survey is repeated after many monthsor even years. The objective is to analyze changes in the

subsurface response at the target zone as a function oftime (e.g., Greaves and Fulp, 1987; King, Dunlop, et al.,1988; Nur, 1989; Johnston, 1989; Matthews, 1992; Archeret al., 1993; Johnstad et al., 1993; Justice et al., 1993). Ifseasonal or time-variant near-surface changes are notrecognized and accounted for, they may obscure subtlechanges at the target or objective level. Thus, it is neces-sary in some areas to note when the near-surface infor-mation was obtained and whether, for example, theground was water-logged. These observations should

be archived with the primary data. Various causes oftime-variant near-surface changes are discussed in thissection, along with the types of areas that are affectedand the magnitude of such changes.

I noted in Section 2.4 that layering may exist withinthe water layer due to local differences in temperatureand salinity and can change over time. These changesmay be seasonal or caused by temperature variations,

tides, or currents. In areas where ocean currents result inrelatively warm and cold bodies of water merging, therecan be rapid changes in average seawater temperature.In a deep-water survey, changes in velocity due to thistemperature change may be large enough to require dif-ferential static shifts to be used to adjust the data to a ref-erence water velocity.

2.5.1 Temperature

The effect of temperature includes the local freezingand thawing of the near surface in Arctic tundra orlakes; the unfrozen near surface has a velocity signifi-

cantly lower than when it is frozen (see Section 2.6.4 forvelocity comparisons). The velocity in ice and per-mafrost also increases as the temperature decreases (seeSections 2.6.3 and 2.6.4). Changes in temperature alsohave an impact on the velocity in both freshwater andseawater, the latter also being affected by changes insalinity. The velocity increases by several meters per sec-ond per degree Celcius for a rise in temperature [seeTable 2-3 and equation (2.12)].

2.5.2 Precipitation

The main effect of rainfall on near-surface layers is

the resulting change in the water table elevation, whichcan be significant in areas with discrete wet and dry sea-sons. The difference in velocity depends on whether thesection of the near surface is water wet or dry, and it canbe a factor as large as 4 (see Section 2.6.1). For example,if the water table elevation changes by 3 m and the dryand wet near-surface velocities are 500 and 1800 m/s,respectively, then the resulting difference in two-waytime through the near surface is about 9 ms. Birkelo et al.(1987) mapped the top of a saturated zone at a depth of


3

2

1

0

-1Elevation(km)

0 4 km

Fig. 2-12. Near-surface profile illustrating rugged moun-tain front topography from southwestern Iran; no verticalexaggeration.


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about 3 m with a shallow high-frequency survey andobserved an increase in time of a few milliseconds as aresult of lowering this zone by about 0.4 m. Clymer et al.(1986) noted variations of 56 ms in the time through thenear-surface layers in winter versus summer due toshallow moisture changes. Another effect of precipita-

tion is the accumulation of snow, especially in polarregions, which not only changes the elevation but alsothe velocity of the underlying snow as a result of com-paction (see Section 2.6.3 for more details).

2.5.3 Tides

Tidal variations change the absolute water depth andhence the time through the water layer. In nearly allmarine surveys, however, datum is at mean sea level,and no corrections are made for the water layer so thatno allowance is made for the state of the tide. Time shifts(static corrections) that are routinely applied compen-

sate for the source and receiver depths (see Section 3.6).In the case of 2-D marine surveys, intersecting lines mis-tie if there is a significant change in sea level between thetwo lines. A variation of 3 m, for example, results in atwo-way time shift of 4 ms; this would normally beaccommodated during the interpretation phase of thesurvey (e.g., Ramsden, 1985).

The state of the tides can be very important in transi-tion zone surveys. For example, Hildebrant (1990)reported a tidal range of 9 m (30 ft) on a tidal mudflatsurvey in Cook Inlet, Alaska. This meant that a carefultidal record was kept so that the water depth wasknown at each shot instant for each receiver location.

(See Figures 3-17 and 3-18 for examples of time shiftsdue to tidal changes.)

In the case of 3-D surveys, however, it is common fordata from two (or more) boat tracks to be gatheredtogether into one common-midpoint (CMP) or subsur-face bin. If the different boat tracks occur at differentstates of the tide, there will be a time shift between thedifferent components of the data sets; if this is notremoved prior to any multichannel processing, it canintroduce artifacts, from anf-k process for example (seeSection 6.6.1), and it can also degrade the final stackresponse. Water temperature variations mentioned ear-lier can have a similar impact on the data in a subsurfacebin. In addition, a change in water velocity leads to achange in the multiple period of multiples associatedwith the water layer so that different multiple periodic-ities may be present within a subsurface bin.

Tidal variations in different parts of the world rangefrom several centimeters to several meters in large lakes,enclosed seas, and the open ocean, with the mostextreme case in the Bay of Fundy, Canada, where thetidal range is up to 15 m. The occurrence of a large tidal

range generally depends on the shape of bays and adja-cent sea, and there are many areas where the range is inexcess of 3 m.

Another factor that should be considered, especiallyfor surveys with a significant high-frequency content, isthe short- to medium-period changes in sea level associ-

ated with ocean swell. Nontrivial changes are most like-ly where there is a significant distance to the shorelinetoward the source of the prevailing winds or storms.

2.5.4 Ice Movement

The main example of ice movement is the passage ofglaciers down from higher elevations to the coast, wherethey break off from the land mass and form icebergs.The normal effect of this on the near surface is a slightchange of elevation over time. The effect on datum stat-ic corrections is normally small because the elevationchanges are small and the velocity of sound in ice is

comparatively fast (see Section 2.6.3). In some cases,moraines are also carried down by the glacier, resultingin a change over time in the near-surface compositionfrom ice to rock debris. The icebergs scour the sea bot-tom in some areas, giving rise to minor changes in waterdepth, but these are likely to be minimal.

2.5.5 Wind

Near-surface changes associated with wind includethe location and elevation of sand dunes. They canchange position with time, and as pointed out in Section

2.4.1, this can involve a horizontal movement of manymeters per year. Thus, the elevation profile can changeas a function of time. It is also likely that the near-surfacevelocity will also change because the velocity is affectedby compaction and any recent sand deposits will changethe compaction curve. The changes are likely to besmall, but they may become significant in a few areas orwhere comparisons are made of surveys conducted overmany years.

2.5.6 Recent Erosion and Deposition

Most of the changes referred to above take place overrelatively short time periods, especially when comparedwith the geologic time scale. Some geologic changes,such as recent erosion and deposition, may have animpact on the near surface when comparisons are madeover a period of months or years. For example, coastlineerosion and landslides can have a significant effect onsurface elevation. Furthermore, the leaching of salt intothe water may alter the salinity sufficiently such that theseawater velocity is altered (see Section 2.6.3).



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Recent deposition, such as near the mouth of a majorriver system, produces changes in the near surface thatcan be observed over short periods of time. For example,Antoine (1975) observed that some gas pockets movedmore than 30 m (100 ft) in a year and that there werechanges in topography apparently caused by upwelling

gases. Son-Hindmarsh et al. (1984) observed changes inwater-bottom topography in the Mississippi Delta mud-flow area as a result of mudflow activity. Both of theseprocesses can result in significant changes in datum sta-tic corrections as a function of calendar time. This isbecause the velocities in these environments can be verylow, especially for aerated or gaseous mud which oftenhas a velocity less than that of air (see Section 2.6.2).

Sand waves on the water bottom have also beenobserved to migrate with time. Whitney et al. (1979)cited reports indicating rates of meters per month. Theyalso observed a small movement in Cook Inlet, Alaska,where the sand waves were 214 m high with wave-

lengths of 50600 m.

2.5.7 Volcanic Activity and Earthquakes

Volcanic activity and earthquakes can have a signifi-cant impact on the near surface. For example, thedestruction of part of Mount St. Helens in the state ofWashington in 1980 caused a sudden major change in theelevation profile of the mountain and rapidly depositeda thick layer of volcanic ash and debris on the near sur-face of the surrounding countryside. The main shockwaves or aftershocks of earthquakes can cause land-slides, eroding one area and depositing it in another area.

The magnitude and location of some earthquakes is suchthat the surface elevation is changed, perhaps by asmuch as several meters in the case of a major fault.

2.5.8 Human Activity

The last time-variant topic is the various changes thatresult from human activities. These include local fea-tures such as road and rail cuts and embankments, quar-ries, dumps, major construction projects, and channeldredging in rivers and harbors. This construction activ-ity can cause major changes both in near-surface veloci-ty and surface elevation. Because such features are oftenlocalized and well defined, their extent must be consid-ered in any subsequent interpolation procedure. Thisshould ensure that the anomalous near surface is local-ized and that simple linear interpolation between con-trol points is unlikely to be appropriate. A major geo-graphic feature that occurs as a result of human activityis the draining and reclamation of large land areas, suchas the polders in The Netherlands. Over time, the near-surface velocity of each polder has changed consider-

ably as a result of compaction and drying out. Thisresults in discrete changes in velocity at the boundariesbetween polders.

2.6 NEAR-SURFACE VELOCITIES

The computation of datum static corrections requiresa near-surface model that includes the thicknesses andvelocities of the layers present. The mechanics of howdatum static corrections are computed are discussed inChapter 3, which shows that near-surface velocities aswell as those below the weathered layer are needed. Therange of velocities this encompasses is large, from about100 to 7000 m/s. To put this velocity range into perspec-tive, it can be compared to the velocity of sound in airand water, which have typical values of about 340 and1500 m/s, respectively.

Water velocity is significantly reduced when air (or

gas) bubbles are present and increased when in a frozenstate as ice. Similarly, sedimentary rocks have increasedvelocity when the temperature is below 0C for longperiods and the water in the pore spaces freezes (per-mafrost). The velocity of the weathered layer is general-ly less than the subweathered layers below it; however,when the near surface is ice or permafrost, there is typi-cally a velocity inversion at the base of the near-surfacelayer. Some velocities are shown to vary appreciablywith temperature or water saturation and can changewith the seasons of the year. Anisotropy is also a factor,especially variations between the velocity parallel ver-sus perpendicular to the bedding planes.

The velocity of sound in air (about 330350 m/s) isoften considered to be near the low-velocity limit forseismic velocities; this is actually an incorrect assump-tion because velocities in the near surface have beenmeasured down to about one-sixth of this value. Ricker(1977) suggested an even lower velocity in the LVL, not-ing that it may drop as low as 30 m/s (100 ft/s). Thevelocity of sound in air is given here only for complete-ness; it is not needed to compute datum static correc-tions. If static corrections are estimated without consid-ering the physics of the situation, it is possible that airvelocity would be used (incorrectly) to correct downto datum for a receiver located in the center of a bridgeor where datum is above the surface. The velocity ofsound in air (V) in meters per second is given by

V 331.5 + 0.607C, (2.1)

where C is the temperature (in C) (Sheriff, 1991).Many authors have published information on veloci-

ties, most indicating a considerable range of velocitiesassociated with the near surface and especially the



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weathered layer. Press (1966) produced a compilationfrom many sources, giving priority to recent measure-ments on well-described common rocks. Drawing onthis work, I will now discuss near-surface velocities forthe following materials:

1. consolidated rocks;2. unconsolidated material;3. mud;4. water, snow, and ice; and5. permafrost.

The section ends with some general points on aniso-tropy.

2.6.1 Velocities in Consolidated Rocksand Unconsolidated Sediments

Velocity ranges are listed in Table 2-1 for several

igneous, sedimentary, and metamorphic rocks for com-pressional waves and, in some cases, shear waves. Mostreferences in the literature are for velocities at a depthbelow the surface; these are generally higher than for thesame rock type close to the surface, and it is the surfacevelocities that are required in the computation of datumstatic corrections. The ratio between compressional (VP)and shear (VS) wave velocities in consolidated rocks istypically in the range of 1.52.0+ (Pickett, 1963; Gregory,1976; Tatham, 1982; Domenico, 1984; Castagna et al.,1985; Tatham and McCormack, 1991).

Velocities in unconsolidated sediments are listed inTable 2-2, which includes data from Press (1966) and

several other authors who observed a large velocityrange for many sediments. Velocities in unconsolidatedmaterials typically depend on water saturation and arealso related to compaction and thickness (such as insand dunes). For near-surface unsaturated rocks inArizona, New Mexico, and California, Watkins et al.(1972) established the following relationships:

= 0.175 ln VP + 1.56 (2.2)

and

= 0.516VP 1.815, (2.3)

where is the fractional porosity, VP is the compres-sional-wave velocity (in m/s), and is the bulk density(in g/cm3). Gardener (1984) noted that the velocity of acrumbling cliff may be as low as 10% of the fresh rockvelocity and that underwater, the velocity in a faultedzone can be 50% of the adjacent fresh rock velocity.

Many authors have shown that water saturation is asignificant factor influencing compressional velocity;the velocity increases by a factor of about 1.25 whengoing from a dry to a water-saturated state (e.g., White

and Sengbush, 1953; Hamilton, 1971, 1976; Domenico,1976; Gregory, 1976; Al-Husseini et al., 1981; Wiest andEdelmann, 1984). This increase in water saturation andvelocity often coincides with the water table. Water sat-uration, however, has little or no effect on the shear-wave velocity. Hamilton (1971, 1976) showed that thelongitudinal (compressional) and transverse velocitiesincrease appreciably with differential pressure in drysands, but that pressure has a minimal effect on thevelocity of brine-saturated sands.

The ratio between compressional- and shear-wavevelocities has a large range in unconsolidated near-sur-face layers, with values from about 1.3 to 10.0. Ericksonet al. (1968) observed ratios up to 6, and cited otherauthors who had obtained values of 7. Lash (1980) mea-sured a range from 6.2 to 10.0 in the top 30 m (100 ft) ina well near Houston, with the ratio decreasing to 5.3 forthe depth range 6090 m (200300 ft). A compilation ofVP/VS ratios against depth was shown by Dohr andJanle (1980) (also Tatham and McCormack, 1991), whichincludes many values from the large compilation byMolotova and Vassilev (1960) and indicates values upto 6 close to the surface. Table 2-3 presents a compilation

from many authors for various sediments showing awide range of values. The ratio normally increases sig-nificantly at the water table as a result of the increase inVP and minimal change in VS. For example, White andSengbush (1953) observed the ratio to increase from 2 to3.5 at the water table, Anno (1983, 1987) from 1.3 to 3.3,Wiest and Edelmann (1984) from 4.5 to 8, and Lawton(1990) from 2.0 to 8.3.

From a large number of measurements, Hamilton(1976) derived a regression curve, nearly independent of


Table 2-1. Seismic Velocities in Igneous, Sedimentary,

and Metamorphic Rocks.a

Material Velocityb(km/s)

VP VS

Anhydrite 4.15.0 2.672.99

Basalt 5.066.4 2.723.21

Chalk 2.14.2Dolomite 3.56.9

Gneiss 3.57.5

Granite 4.86.0 2.873.23

Gypsum 2.03.5

Limestone 1.77.0

Marble 3.756.94 2.023.86

Salt 4.46.5

Sandstone 1.44.3

Sandstoneshale 2.14.5

Shale and slate 2.34.7

aData from Press (1966).

bRanges forVPorVSshould not be used to compute VP/VSbecause the

extremes in each range are not necessarily for the same sample.


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moisture content, relating the shear-wave velocity VS (inm/s) for sands at depths of 0.112 m to depth D (in m).This is given as

VS = 128D0.28. (2.4)

For silt-clays down to a depth of 36 m, the empiricalrelationship is

VS = 116 + 4.65D. (2.5)

2.6.2 Velocities in Muds

Lester (1932) observed that the velocity of the weath-ered layer is generally in the range of 610760 m/s(20002500 ft/s), but very low velocities in the near sur-face (down to about half the velocity of sound in air) arealso seen, with one value quoted at 170 m/s (550 ft/s).This observation was equated to the work of Wood(1955), first published in 1930. He computed the veloci-


Table 2-2. Seismic Velocities in Unconsolidated Sediments.

Material Velocity (m/s) References

VP Vs

Aerated layer 170 Lester (1932)

Aerated layer 150 Lyons (1946)Alluvial silt 102270 Hamilton (1976)

Alluvium 5002000 Press (1966)Alluvium 225400 Miller et al. (1989)

Ash and lapilli 500 Watkins et al. (1972)

Basaltic cinders 300 Watkins et al. (1972)

Bay mud 90 Hamilton (1976)

Bentonite talus 85260 Watkins et al. (1972)

Clay 11002500 Press (1966)

Clay 3002500 100600 Milkereit et al. (1986)

silty clay 137174 Hamilton (1976)

wet 120359 Hamilton (1976)

Diatomite talus 80200 Watkins et al. (1972)

Diluvium 7001800 Press (1966)

Embankments and fill 400 Press (1966)

Glacial

sand and gravel (unsaturated) 380500 Press (1966)sand and gravel (saturated) 1670 Press (1966)

till (unsaturated) 4301040 Press (1966)

till (saturated) 1730 Press (1966)

Lignite 4001800 180280 Milkereit et al. (1986)

Loam 8001800 Press (1966)

Loess 300600 Press (1966)

Rhyolitic flows 260 Watkins et al. (1972)

Rhyolitic pumice 460 Watkins et al. (1972)

Sand 4002800 200500 Milkereit et al. (1986)

calcareous 800 Press (1966)

consolidated 610820 300410 Al-Husseini et al. (1981)

fine 85567 Hamilton (1976)

loose 430460 190240 Al-Husseini et al. (1981)

loose 2002000 Press (1966)

above water table 1000 400 Press (1966)

below water table 1800 500 Press (1966)

medium 53475 Hamilton (1976)

wet 7501500 Press (1966)

Soil 110200 Press (1966)

Subwater-bottom muds 451500 Jones et al. (1958, 1964),

Levin (1962), May et al.

(1988), Tinkle et al. (1988)

Volcanic ash 160 Watkins et al. (1972)

Weathered layer 300900 Press (1966)

Weathered layer 130610 Gaby and Solari (1948)


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ty of a fluid that included a certain volume of air orother gas (as bubbles) that does not interact with thefluid as

, (2.6)

where 1

is the density of air (in g/cm3), 2

is the densi-ty of fluid, x is the proportion of air by volume, E1 is theelasticity (bulk modulus) of air (in dynes/cm2), and E2is the elasticity of fluid (Wood, 1955, 361).

Equation (2.6) is based on the standard relationshipfor the velocity (V) of a compressional wave in a fluid,given by

(2.7)

or

, (2.8)

where Kis the bulk modulus, is the density of the fluid(in g/cm3), and is the compressibility (reciprocal ofbulk modulus, in cm2/dyne). Equation (2.6) can beexpressed in terms of the compressibility of the air andfluid as

, (2.9)

where 1 is the compressibility of air and 2 is the com-pressibility of fluid. For comparison, the compressional-wave velocity in a sediment that has structural strengthis given by

Vx x x x

=+ [ ] + [ ]

1

1 11 2 1 2

1 2

( ) ( )

/

V=

11 2

/

VK

=

1 2/

VE E

xE x E x x=

+ [ ] + [ ]

1 2

2 1 1 2

1 2

1 1( ) ( )

/


Table 2-3. Ratio of Compressional-Wave Velocity (VP) to Shear-Wave Velocity (VS).

Material VP/Vs Referencesa

Boulder clay 4.48.9 Stmpel et al. (1984)

Chalk 2.4 White and Sengbush (1953)

Chalk 1.812.4 Molotova and Vassilev (1960)

Clay 3.04.2 Milkereit et al. (1986)

Clay 1.78.5 Molotova and Vassilev (1960)

argillaceous 1.555.0 Molotova and Vassilev (1960)

wet 4.513.7 Molotova and Vassilev (1960)

Consolidated sediments 2.25.0 Wiest and Edelmann (1984)

Gravel 4.57.0 Fromm et al. (1985)

Gravel 3.05.0 Molotova and Vassilev (1960)

Lignite 2.26.4 Milkereit et al. (1986)

Loess 3.0 Molotova and Vassilev (1960)

Sand 2.05.5 Milkereit et al. (1986)

loose 1.82.3 White and Sengbush (1953)

below water table 2.73.0 White and Sengbush (1953)

fine-grained wet sand 1.72.1 Molotova and Vassilev (1960)

fine sand/silt 5.18.7 Beeston and McEvilly (1977)

dry and partly saturated 1.43.9 Stmpel et al. (1984)coarse sand and fine gravel 1.42.7 Stmpel et al. (1984)

saturated sand 4.510.0 Stmpel et al. (1984)

sand and pea gravel 3.6 Beeston and McEvilly (1977)

sandy clay 3.44.5 Fromm et al. (1985)

sandy clay 1.742.4 Molotova and Vassilev (1960)

water-saturated, clean sandstone 3.4 Castagna et al. (1985)sand and silt (alluvium) 1.9 Harris and Street (1991)

sand, gravel, and clay 2.84.6 Harris and Street (1991)

Shale 2.5 Molotova and Vassilev (1960)

shaly sand 2.7 Fromm et al. (1985)

noncalcareous 5.0 Castagna et al. (1985)Soil 1.72.0 Molotova and Vassilev (1960)

Unconsolidated sediments 2.58 Wiest and Edelmann (1984)

a In some cases, values were computed from a figure in the reference.


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, (2.10)

where G is the sediment rigidity modulus and is thedensity of sediment.

Lester (1932) suggested that some types of near-sur-face sediments, such as Gulf Coastgumbos (sticky mate-rials such as wet clay) could be treated as fluids. Figure2-13a shows the relationship between the compression-al-wave velocity and the proportion of air in the gumbo,ranging from 107 to101 (with extension to unity wherethe velocity becomes the air velocity). The figure usesthe parameter values given by Lester (1932): a fluid den-sity of 1.9 g/cm3, an air density of 0.0012 g/cm3, andelasticities of fluid and air of 5.58 1010 and 1.2 106dynes/cm2, respectively. Thus, the value quoted earlierof 170 m/s (550 ft/s) corresponds to a value of about0.002 parts per volume (0.2%) of air. It should also be

noted that as more air is introduced, the velocitydecreases even further, to well below 100 m/s, when themixture becomes afroth.

The velocity decreases because air bubbles intro-duced into the fluid increase the compressibility appre-ciably as compared to bubble-free fluid, which results ina significant reduction in the velocity, as indicated byequation (2.7). The suggestion by Ricker (1977) thatvelocities down to 30 m/s (100 ft/s) may occur wasbased on the premise that the near surface had the elas-ticity of air and the density of earth. Schubel andSchiemer (1973) observed an increase in compressibilityof about 100 when comparing turbid sediments with

clear sediments of similar grain size. For comparisonwith the gumboair velocity relationship, Figure 2-13bshows a similar relationship for water using a density of1.0 g/cm3 and an elasticity of 2.25 x 1010 dynes/cm2.

Silberman (1957), Anderson and Hampton (1980a, b),and Domenico (1982) supported the formula proposedby Wood (1955) [equation (2.6)], but stated that it wasonly valid if the frequency of the sound waves wasbelow the resonant frequency of the bubbles. This is nor-mally the case for seismic exploration, as a resonant fre-quency of 200 Hz corresponds to a large and impracticalbubble radius of 23 cm and lower resonant frequenciescorrespond to even larger bubbles. Mallock (1910) hadshown earlier that if froth were present, there was anenormous reduction in the velocity related to the incom-pressibility of liquids.

As stated above, the velocity defined by equations(2.6) and (2.9) assumes a mixture of air and fluid; insome cases, the sediment can be expected to have somestructural strength. Anderson and Hampton (1980b)suggested that an alternative approach was to assumethat the structural strength of a saturated sediment wasunmodified by the addition of gas. The faster velocities

given by this approach could be considered as an upperthreshold, whereas those computed by a fluid equationwould give the lower threshold. The velocities they

computed for a 1% gas content were 359 m/s for clay,524 m/s for silt, 826 m/s for fine sand, and 1039 m/s forcoarse sand. These can be compared to the velocities forthe gas-free sediments which are 1488, 1552, 1749, and1907 m/s, respectively. Thus, the gas-charged sedimentshave velocities that are 2454% of the gas-free sedi-ments, substantially higher than 5% for the fluid case,where 1% gas results in a decrease in velocity from 1714to 80 m/s.

In many cases, the gas is likely to be biogenic, formedin place by the anaerobic decomposition of organic mat-ter (see Section 2.4.2). These anomalously low velocitieshave been reported from many areas. In LakeMaracaibo, Levin (1962) reported mud velocities downto 90 m/s (300 ft/s). In a lake in northeastern UnitedStates, Jones et al. (1958) reported velocities in the rangeof 75170 m/s (250550 ft/s); similar velocities wereobserved by Bobber (1959) near Orlando, Florida. Joneset al. (1964) observed velocities down to 45120 m/s(150400 ft/s) in soft river mud. In the Mississippi Delta,May et al. (1988) observed velocities in the range of3001500 m/s (10005000 ft/s), and Tinkle et al. (1988)noted that the first 30 m (100 ft) of mud had velocities

VK G

= +

4 3

1 2/

/


0

500

1000

1500

2000

Velocity(m/s)

10-7

10-6

10-5

10-4

10-3

10-2

10-1

1

(a)

(b)

Proportion of air in mixture

Air velocity

Fig. 2-13. Plots of compressional-wave velocities offluidair mixtures expressed as a function of the propor-tion of air in the mixture (by volume): (a) mud; (b) water(after Lester, 1932;Wood, 1955).


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lower than 300 m/s (1000 ft/s), with the lowest less than120 m/s (400 ft/s), measured over a 3-m (10-ft) interval.Edrington and Calloway (1984) observed a mud veloci-ty of 800 m/s, 30 m below the mud line. A humus layerabout 15 cm thick in a rubber plantation area was report-ed by Singh (1983) to have a velocity as low as 55 m/s.

These reports indicate a considerable range of veloci-ties, and it must be remembered that the velocities canchange rapidly over short distances (see Section 2.4.2)because gas pockets, for example, are often localized. Inaddition, because velocities are often very low, a com-paratively small change in velocity or thickness canhave a dramatic effect on the traveltime through thelayer.

2.6.3 Velocities in Water, Snow, and Ice

The velocity of sound in water is often quoted at 1500m/s, although it varies, for water at the surface, fromabout 1400 m/s for freshwater at 0C to about 1550 m/sfor seawater at 35C. Various formulas have been pro-

posed to relate the velocity to variations in temperature,salinity, and depth. Wood (1955) gives the followingexpression for the velocity of sound in seawater (at617C):

V= 1410 + 4.21T 0.037T2 + 1.14S, (2.11)

where Vis the velocity (in m/s), Tis the temperature (inC), and S is the salinity (in parts per thousand, ppt).

Sheriff (1991) quotes a similar relationship but whichincludes a factor for depth:

V= 1449.2 + 4.6T 0.055T2 + 0.0003T3+ (1.34 0.010T)(S 35) + 0.016Z, (2.12)

where Z is the depth (in m). Surface velocities comput-ed from equation (2.12) for freshwater and seawaterwith a salinity of 35 ppt are listed in Table 2-4 for a rangeof temperatures. Surface velocities are also showngraphically in Figure 2-14 for a range of temperaturesand salinities. A more precise formula, as used in U.S.Naval Oceanographic Office publications (e.g., Bialek,1966), is given by Wilson (1960a, b).

Below the surface, water temperature decreases with

depth and is on the order of 810C at 1000 m depth, 4Cat 2000 m, and 3C at 3000 m (Ewing and Worzel, 1948).This results in an overall decrease in the velocity ofsound in water with depth, but this is offset by anincrease in velocity resulting from increased pressure(depth factor). The net effect is a decrease of about 30m/s down to a depth of about 1250 m, at which pointthe velocity increases with further depth. The preciseminimum velocity and its associated depth vary in dif-ferent locations around the world. This velocity mini-


Table 2-4. Seismic Velocities at the Surface

in Freshwater and Seawater.a

Temperature Velocity (m/s)

(C) Freshwater Seawater

(35 ppt salinity)

0.0 1402 1449

2.5 1414 14605.0 1426 1471

7.5 1436 1481

10.0 1447 1490

12.5 1456 1499

15.0 1465 1507

17.5 1474 1514

20.0 1482 1522

22.5 1489 1528

25.0 1496 1535

27.5 1503 1540

30.0 1509 1546

32.5 1515 1551

35.0 1521 1556

aData from Sheriff (1991).

Temperature(C)

1540

1520

1500

1480

1460

1440

1420

30

20

10

0

0 10 20 30 40

Salinity (parts per thousand)

Fig. 2-14. Surface water velocity as a function of temper-ature and salinity (after Sheriff, 1991).


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mum gives rise to the SOFAR channel (Ewing andWorzel, 1948). Any layering within the water layer gives

rise to local changes in velocity which are caused byvariations in temperature, salinity, or both.In snow-covered areas, the near-surface layer is

calledfirn, which is composed (from the top down) ofsnow, compacted snow, ice, and finally, compacted ice ata depth of about 40100 m. The density changes fromabout 0.4 g/cm3 close to the surface to about 0.9 g/cm3for compacted ice; a representative profile of densityversus depth is shown in Figure 2-15. In any specificarea, the values may differ slightly from this profile,depending primarily on weather conditions over thepreceding years.

The velocity of sound in snow is dependent on its

density. Robin (1958) gave the following relationshipbetween firn density (in g/cm3) and velocity VP(in m/s), under the assumption that the temperaturecoefficients of ice and firn are identical:

, (2.13)

where t is the mean annual near-surface firn tempera-ture (in C). An alternative relationship was given by

Kohnen (1972) as

, (2.14)

where (z) is the density at depthz (in g/cm3), V(z) is thevelocity at depthz (in km/s), and Vice is the velocity inice (in km/s).

A representative plot of firn velocity versus depth isshown in Figure 2-16 for both compressional and shearwaves. In any given area, the velocity versus depth pro-file will be similar to that shown. However, the velocitiesclose to the surface are likely to be different and canrange from about 500 to 1750 m/s for compressional

waves, depending on near-surface characteristics. Thevelocity range at depth is less because the variations inice properties are much less than for uncompactedsnow.

The velocity of sound in various forms of ice isshown in Table 2-5; typical compressional-wave veloci-ties are in excess of 3000 m/s and at or above 3900 m/sfor compacted ice. The velocity in ice increases as thetemperature decreases by about 2.3 m/s/C for com-pressional waves and by 1.1 m/s/C for shear waves

zzV V

( )

= +

0 915 12 25

1 22

1

..

( ) .ice

=

+2

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