Temporal signiﬁcance of sequence boundaries · 2015. 10. 6. · variable, whichever that is among the eustasy, tectonics or sedimentation controls. Types A and B surfaces merge

ELSEVIER Sedimentary Geology 121 (1998) 157–178

Temporal significance of sequence boundaries

Octavian Catuneanu a,Ł, Andrew J. Willis b, Andrew D. Miall c

a Department of Geology, Rhodes University, Grahamstown 6140, South Africab Rigel Energy Corporation, 1900, 255 5th Ave. S.W., Calgary, Alberta T2P 3G6, Canada

c Department of Geology, University of Toronto, Toronto, Ontario M5S 3B1 Canada

Received 15 May 1997; accepted 22 June 1998

Abstract

This paper analyses the temporal significance of stratigraphic surfaces bounding the marine portions of the depositionalsequence, genetic stratigraphic sequence and transgressive–regressive sequence. These bounding surfaces, known asthe ‘correlative conformity’ (c.c.), ‘maximum flooding surface’ (MFS) and ‘conformable transgressive surface’ (CTS),respectively, may either be defined on the basis of stratal stacking patterns (which we call ‘type A surfaces’), or on thebasis of water-depth changes and relative sea-level changes (which we call ‘type B surfaces’). The type A MFS andCTS are time lines in a depositional-dip section, corresponding to the turnaround points from shoreline transgression toregression and vice versa. They separate prograding (coarsening-upward) from retrograding (fining-upward) geometries,with a timing determined by the interplay between the rates of sedimentation and relative sea-level rise in the shorelinearea. The timing of type A MFS and CTS is not affected by the offshore variations in sedimentation and subsidencerates, but it is only controlled by the shoreline movements and the associated facies shifts. The type A c.c. separatesrapidly prograding and offlapping forced regressive strata from the overlying lower rate prograding and aggrading normalregressive strata. This surface is diachronous, younger basinward, with the rate of offshore sediment transport. The timingof the type A c.c. in the shoreline area corresponds to the end of relative sea-level fall, but it develops under relativesea-level rise conditions offshore. The timing of the type B MFS and CTS depends on the offshore variations in thesedimentation and subsidence rates. These surfaces, defined on the basis of bathymetric changes, become younger andolder seaward, respectively, tending to merge together offshore. The type B c.c. marks the end of relative sea-level fallin any point along a depositional-dip section. It is diachronous, older basinward, as its timing depends on the offshorevariations in subsidence rates. The diachroneity of type B surfaces reaches a quarter of the period of the highest frequencyvariable, whichever that is among the eustasy, tectonics or sedimentation controls. Types A and B surfaces merge togetherin the shoreline area, but they become temporally divergent offshore. Deepening-upward and shallowing-upward faciesshould not be confused with transgressive and regressive systems tracts. The latter are strictly controlled by the shorelinemovements, which determine the direction of facies shifts and the stratal stacking patterns. 1998 Elsevier Science B.V.All rights reserved.

Keywords: sequences; systems tracts; bounding surfaces; diachroneity

Ł Corresponding author. E-mail: [email protected]

0037-0738/98/$ – see front matter c 1998 Elsevier Science B.V. All rights reserved.PII: S 0 0 3 7 - 0 7 3 8 ( 9 8 ) 0 0 0 8 4 - 0

158 O. Catuneanu et al. / Sedimentary Geology 121 (1998) 157–178

1. Introduction

A controversial topic in modern stratigraphyis the assessment of the relationship between se-quence stratigraphy and chronostratigraphy. Are thesequence-bounding surfaces time lines, i.e. gener-ated at the same time everywhere within the areaof occurrence? The answer to this question is ofparamount importance for stratigraphic correlation,and although this problem has been around for sometime (Miall, 1991, 1994), an agreement is yet to bereached. Part of the problem derives from the wayconcepts are defined and used, often with contradic-tory meanings, as we present in this introduction.

The various types of sequences and boundingsurfaces are illustrated in Fig. 1. The depositionalsequence (Jervey, 1988; Posamentier et al., 1988;Van Wagoner et al., 1990; Haq, 1991; Vail et al.,1991; Hunt and Tucker, 1992) is defined in relation-ship to the relative sea-level (base-level) curve, and itis bounded by the subaerial unconformity (SU) andits marine correlative conformity (c.c.). The timingof the SU is generally related to the stage of base-level fall (Fig. 1; Posamentier et al., 1988; Hunt andTucker, 1992; Embry, 1995), whereas the c.c. wasinitially considered to form during early sea-levelfall (Posamentier et al., 1988) or at the beginning ofthe sea-level fall (Posamentier et al., 1992), whichwas later revised to the end of relative fall (Fig. 1;Hunt and Tucker, 1992; Helland-Hansen and Mar-tinsen, 1996). The depositional sequence comprisesfour systems tracts with distinct stratal stacking pat-terns (Figs. 1 and 2): the highstand systems tract

Fig. 1. Types of sequences, bounding surfaces and systems tracts, defined in relationship to the relative sea-level and transgressive–regressive curves. The relative sea-level depends on the combined effect of eustasy and tectonics, whereas the generation of transgressiveand regressive facies depends on the combined effect of relative sea-level changes and sedimentation. The depositional sequenceboundary (i.e., subaerial unconformity and its marine correlative conformity) is generated at the end of relative sea-level (base-level) fall(Hunt and Tucker, 1992; Helland-Hansen and Martinsen, 1996). The genetic stratigraphic sequence boundary (i.e., maximum floodingsurface) is taken at the top of marine and nonmarine transgressive facies (Galloway, 1989). The T–R sequence boundary is taken at thetop of marine regressive facies (i.e., the conformable transgressive surface, Embry, 1995). Within nonmarine facies, the T–R sequenceis arbitrarily chosen to coincide with the depositional sequence in spite of the fact that the RST extends above the SU, due to thedifficulty in field recognition of the CTS-correlative. In special circumstances (i.e., short LST stages and strong erosion associatedwith the ravinement surface), the most basinward portion of the nonmarine LST may not be preserved and thus the CST could bemapped in the continuation of the SU. Note that normal regressive facies accumulate in the earliest (LST) and latest (HST) stagesof relative sea-level rise, due to sedimentation outpacing the low rates of relative rise. We take here the LST to be equivalent to the‘lowstand prograding-wedge systems tract’ of Hunt and Tucker (1992). Abbreviations: DS D depositional sequence; GS D geneticstratigraphic sequence; T–R D transgressive–regressive sequence; LST D lowstand systems tract; TST D transgressive systems tract; HSTD highstand systems tract; FSST D falling stage systems tract; RST D regressive systems tract; SU D subaerial unconformity; c.c. D

(HST) forms during late relative rise, when the sed-imentation rate exceeds the rate of relative rise inthe shoreline area (normal regression); the fallingstage systems tract (FSST) forms during relativefall (forced regression); the lowstand systems tract(LST) forms during early relative rise, when the sed-imentation rate exceeds the rate of relative rise in theshoreline area (normal regression); and the transgres-sive systems tract (TST) which forms when the rateof relative sea-level rise in the shoreline area exceedsthe sedimentation rate. The former three systemstracts (HST, FSST and LST) form together a progra-dational package known as a regressive systems tract(RST; Embry and Johannessen, 1992). A RST fol-lowed by a TST form together a genetic stratigraphicsequence (Fig. 1; Galloway, 1989), bounded by max-imum flooding surfaces (MFS). The combination ofa TST followed by a RST gives the transgressive–regressive (T–R) sequence (Fig. 1; Embry and Jo-hannessen, 1992; Embry, 1993, 1995), bounded byconformable transgressive surfaces (CTS) in the ma-rine portion of the basin. The nonmarine correlativeof the CTS is either unidentifiable within the fluvialsuccession overlying the SU, or eroded by the ravine-ment surface. In either case, the SU was convenientlychosen to represent the T–R sequence boundary inthe nonmarine succession (Embry, 1993, 1995).

We focus our analysis on bounding surfaces de-veloped within marine successions, i.e. the c.c., CTSand MFS.

The c.c. may be defined in three ways, whichallow for different temporal interpretations.

(1) It is a surface taken by definition as a time

O. Catuneanu et al. / Sedimentary Geology 121 (1998) 157–178 159

correlative conformity; CTS D conformable transgressive surface; CTS-c D CTS-correlative (i.e., the nonmarine correlative of the marineCTS); MFS D maximum flooding surface; R D ravinement surface; BSFR D basal surface of forced regression; IV D incised valley;(A) D creation of accommodation space (base-level rise); NR D normal (sediment supply-driven) regression; FR D forced (base-levelfall-driven) regression.



line, which begins at the basinward termination ofthe SU and extends throughout the conformable ma-rine succession (Jervey, 1988; Embry, 1995). Thetiming of this conformable surface was related tovarious portions of the sea-level or relative sea-levelcurves, finally settling at the end of relative fallin the shoreline area, i.e. the depositional surfacewhich existed at the end of the forced regression ofthe shoreline (Embry, 1995: “The subaerial uncon-formity is developed and migrates seaward duringbase-level fall and reaches its maximum extent at theend of the fall : : : , the depositional surface in themarine realm at this time of change from base-levelfall to base-level rise is the correlative conformity”,portrayed as a time line in his fig. 1). The time linesignificance of such surface is of course valid along adepositional dip section, as varying subsidence ratesalong the depositional strike may offset the transitionbetween base-level fall and base-level rise along theshoreline;

(2) It is a surface defined on the basis of stratalstacking patterns, separating the offlapping forced re-gressive lobes from the overlying aggradational LST(Fig. 2; Haq, 1991: “a change from rapidly prograd-ing parasequences to aggradational parasequences”).This definition implies a diachronous c.c., youngerbasinward, with a diachroneity rate that matches therate of offshore sediment transport (Fig. 3). Thetransport rate of terrigenous sediments along thedepositional dip within a marine basin varies from10�1–100 m=s in the case of low gradient marine sys-tems, to 101–102 m=s in the case of turbiditic flowsassociated with higher gradients (Reading, 1996).For this reason, the c.c. is represented with a higherdiachroneity within the terrigenous progradationalwedge relative to the deeper marine basin where ittops the submarine fan deposits (Fig. 3).

(3) It is the surface that marks the end of rel-ative sea-level fall within the marine basin (Posa-mentier and Allen, 1993: “eustasy and sea-floorsubsidence=uplift determine the timing of sequencebounding surfaces”). This definition also implies adiachronous c.c., as the relative sea-level partly de-

Fig. 2. Systems tracts of the depositional sequence, defined on the basis of stratal stacking patterns. The sinusoidal curves illustraterelative sea-level changes in the shoreline area, which may be different in terms of rates and sign from the coeval relative sea-levelchanges occurring farther offshore.

pends on varying subsidence rates across the basin.We investigate in this paper the diachroneity rate ofthis type of c.c.

The CTS represents the marine T–R sequenceboundary, and only a systems tract boundary in theview of the depositional sequence model (Fig. 1).It may be defined either: (1) on the basis of stratalstacking patterns, as a conformable surface whichseparates regressive strata (progradational pattern)below from transgressive strata (retrogradational pat-tern) above (Embry, 1993, 1995); or (2) on thebasis of bathymetric (water-depth) changes, as a con-formable surface recording the start of a deepeningepisode, i.e. formed when the water-depth reachesthe shallowest peak (Embry, 1993).

Although these two definitions are consideredequivalent, they allow different temporal signifi-cances for the CTS. The former relates to theshoreline movements and the associated changes instacking patterns, which brings the CTS to a timeline in a depositional dip section, independent ofthe offshore variations in sedimentation and sub-sidence rates, as there is only one point in timewhere the shoreline is at its most basinward position.The sediment supplied from the onshore during theshoreline regression generates a coarsening upwardmarine succession related to the basinward faciesshift, sharply overlain by much finer transgressivestrata as the coarse terrigenous sediments are trappedwithin the shoreline systems during transgression(Fig. 3). This provides a lithological criterion to pin-point the CTS position in outcrops or subsurface logs(Embry, 1993; fig. 5 in Catuneanu et al., 1997).

The second definition implies a diachronous CTS,as the water-depth changes depend on varying sedi-mentation and subsidence rates across the basin. Inthis light, it is recognized that the CTS is youngerin areas with higher sedimentation rates, where thetransition from shallowing to deepening may occurlater, although this diachroneity is considered mi-nor (Embry, 1995). We investigate in this paper thediachroneity rate of this type of CTS.

The MFS may also be defined in two ways: (1)

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Fig. 3. Wheeler diagram illustrating bounding surfaces defined on stratal stacking patterns. Zone A D low-rate diachroneity at the top of the condensed section, which equalsthe rate of offshore sediment transport. Zone B D higher-rate diachroneity at the top of the condensed section, which equals the rate of shoreline regression.


on the basis of stratal stacking patterns, marking thechange from transgressive strata below to regressivestrata above (Galloway, 1989); or (2) on the basisof bathymetric (water-depth) changes, being formedwhen the water reaches the deepest peak (Naish andKamp, 1997).

Again, these two approaches define bounding sur-faces which are not necessarily superimposed. Inthe former approach, the MFS is associated withthe condensed section separating retrograding fa-cies, below, from prograding facies above. Ideally, itcorresponds to the time line coeval to the momentin time where the shoreline is at its most land-ward position within a given depositional dip section(Fig. 3). In this case, the MFS separates retrogradingfrom prograding stratal patterns (downlap surface)irrespective of the variations in sedimentation andsubsidence rates along the depositional dip. Even so,a certain diachroneity exists along the depositionalstrike, as variations in sedimentation and subsidencerates determine temporally offset transitions fromtransgression to regression along the shoreline (Gilland Cobban, 1973; Martinsen and Helland-Hansen,1995). In practice, it is very difficult to pinpoint thetime line surface within the condensed section, andthe more readily recognizable base of the overlyingterrigenous progradational wedge (limit between thecondensed section and the overlying terrigenous pro-grading facies in Fig. 3) may be approximated asthe downlap surface. This ‘MFS’ is of course di-achronous, with the rates of offshore sediment trans-port (zone A in Fig. 3) or shoreline=sedimentarylobes progradation (zone B in Fig. 3), which canbe emphasized using volcanic ash layers as timemarkers (Ito and O’Hara, 1994).

The latter definition implies a diachronous MFSalong both depositional dip and strike sections, asthe basin bathymetry depends on varying sedimen-tation and subsidence rates. As noted by Naish andKamp (1997), the maximum water depth (i.e., theirMFS) often occurs within the lower part of theHST (normal regressive) progradational wedge. Thusthe boundary between prograding and retrogradinggeometries (downlap surface) will correspond to aphysical surface, recognizable on the basis of stratalstacking patterns, whereas the MFS is unknowablelithologically and can only be identified using forampaleobathymetry.

We apply end-member boundary conditions tosurfaces formed as a result of the complex interplaybetween eustasy, subsidence and sedimentation (i.e.,CTS and MFS defined on the basis of water-depthchanges), as well as to surfaces formed as a resultof the interaction between eustasy and tectonics (i.e.,the c.c. formed at the end of relative fall, and itscounterpart, the surface marking the start of relativefall). The temporal significance of these surfaces willbe compared with the timing of surfaces defined onthe basis of stratal stacking patterns.

2. Controlling factors on relative sea-levelchanges and water-depth changes

The water-depth changes depend on the interplaybetween eustasy, tectonics (which combined givethe relative sea-level changes) and sedimentation(Fig. 4). Sediment compaction may also interfere inthis process by creating additional accommodationspace (base-level rise effect), but it has exactly thesame consequences as the tectonic subsidence andtherefore we incorporate compaction within ‘tec-tonics’ in our quantitative modelling. The causes,magnitudes and rates of change of these variableshave been described at length by Galloway (1989)and need not be reiterated here. It is sufficient tonote that all three variables can attain comparablerates of change and thus have equal potential to in-fluence the generation of CTS and MFS. A positiverate of change of the water-depth (i.e., deepening)will result in upward-deepening facies (UDF), anda negative rate (i.e., shallowing) in upward-shallow-ing facies (USF). The boundary between UDF andoverlying USF (and vice versa: MFS and CTS re-spectively) therefore occurs at the point where therate of water-depth change is equal to zero. Similarly,the correlative conformity portion of the depositionalsequence boundary (c.c.), as well as the surfacemarking the start of relative fall, form when the rateof relative sea-level changes equals zero.

As sequence-bounding surfaces form only wherethe rate of water-depth or relative sea-level changesis equal to zero, it is implicit that the interactingvariables must be of the same order of magnitudeto periodically balance one another and meet thiscondition. If the magnitude of one of the variables


Fig. 4. Diagrammatic illustration of two ways in which therate of water-depth change (W) can be obtained from the threeprimary variables eustasy (E), tectonics (T) and sedimentation(S) via subsidiary derived variables. Reducing the three primaryvariables to combinations of two proxies facilitates the construc-tion of simple geometrical models which still retain the effectsof all three, rather than the usual approach of eliminating onevariable. Case 1 combines the rate of tectonic subsidence=uplift(T) and the rate of sedimentation (S) to define the rate of move-ment of the depositional surface (D), which when added to therate of eustatic change (E) gives the rate of water-depth change(W). Case 2 combines the rates of eustatic change (E) and tec-tonic subsidence=uplift (T) to define the rate of relative sea-levelchange (R), which when added to the sedimentation rate (S) alsogives the rate of water-depth change (W).

is always much larger than the others, then theeffect of the smaller variables will be effectivelysuppressed. For example, if the rate of subsidenceis always greater than the combination between eu-stasy and sedimentation, continuous relative sea-level rise and transgression occur and no sequence-bounding surfaces will form. An example of suchdisparity between subsidence and eustatic rates isdescribed from the Banda Arc by Fortuin and deSmet (1991).

Where the rates of change of the variables arein the same range, one will normally vary witha higher frequency than the others. The higher-fre-quency variable will be the ‘driving force’ behind thehigh-frequency changes in the stratigraphic record.It is commonly assumed that eustasy is the higher-frequency variable (Posamentier and James, 1993),but this need not be so. It is quite conceivable thateustasy could maintain a uniform rate of change overthe duration of several episodes of subsidence anduplift. In this case tectonic subsidence=uplift wouldbe the driving force behind sequence formation, asit the case for instance with the Late Cretaceoussequences of the western Canada foreland basin(Catuneanu et al., 1997).

To facilitate simpler graphical presentation of theresults, the rate of water-depth change (W) can beobtained from the three variables rate of eustaticchange (E), rate of tectonic subsidence (T) and rateof sedimentation (S), via derived proxy variables inthree ways, two of which make geological sense(Fig. 4).

2.1. Case 1

If we combine the rate of tectonic subsidence (T)with the rate of sedimentation (S), we arrive at aderivative variable (D) reflecting the rate of move-ment of the depositional surface. The sum of the rateof movement of the depositional surface (D) and therate of eustatic change (E) gives the rate of change ofthe water-depth (W), which controls the stratigraphicpattern of UDF and USF as described above.

2.2. Case 2

We can also arrive at the rate of change of thewater-depth (W) in a second way. The rates ofeustatic change (E) and tectonic subsidence (T) canbe combined into a variable reflecting the rate ofrelative sea-level change (R) (Posamentier et al.,1988). By combining this derived variable (R) withthe primary variable rate of sedimentation (S), wearrive at the rate of water-depth change (W). Ifthe rate of relative sea-level rise exceeds the rateof sedimentation (S), then (W) will be positive, acontinuous UDF succession will be deposited, andno bounding surfaces will form. If the reverse is


true, i.e. the rate of sedimentation (S) is greaterthan the rate of relative sea-level rise, then (R) willbe negative and continuous shallowing (‘normal’regression in the shoreline area) will occur.

3. Timing of bounding surfaces controlled bywater-depth changes

3.1. Two-dimensional model

To illustrate the effect that subsidence and sedi-mentation rates have on the temporal formation ofCTS and MFS defined on the basis of bathymetricchanges, we have constructed a simple two-dimen-sional geometrical basin model applied to an openmarine shelf setting, which we will refer to as ProfileA. We have chosen to model the case where eustasyis the highest-frequency variable and subsidence andsedimentation are grouped together as the variable(D) (Case 1 in Fig. 4), since it facilitates comparisonwith the depositional sequence model of Posamen-tier et al. (1988). The input values used for thevariable rates of change are obtained from the litera-ture (Pitman, 1978; Pitman and Golovchenko, 1983;Angevine, 1989; Galloway, 1989; Jordan and Flem-ings, 1991; Macdonald, 1991; Frostick and Steel,1993).

Fig. 5. The eustatic variation considered in the model. Eustasy is assumed to vary sinusoidally with an amplitude of 10 m and period of2 Ma. Only half the eustatic cycle (180º phase or 1 Ma) from highstand to lowstand is shown. The rate of eustatic change at each 0.125Ma time increment, given by the differential of the sine curve, is shown on the left.

The assumptions of the model are as follows:(1) Eustasy varies sinusoidally with an amplitude

of 10 m and period of 2 Ma (Fig. 5). For the sake ofbrevity, we have illustrated only half of the eustaticcycle, from highstand to lowstand. The rate of eu-static fall increases from zero at highstand (0 Ma) toa maximum of 15.7 m=Ma at the inflexion point (0.5Ma), and then decreases to zero at lowstand (1 Ma).

(2) The modelled portion of the basin is 200 kmacross, and the tectonic subsidence rate (T) is con-stant at any particular point, but increases basinwardfrom 20 m=Ma at the proximal end of the profileto 40 m=Ma at the distal end. This is similar to thesimple divergent margin models of Pitman (1978),Angevine (1989) and Jordan and Flemings (1991).

(3) Rather than assuming an unrealistic constantsedimentation rate (S) across the basin, we assumethat (S) decreases from 15 m=Ma at the proximal endof the basin profile to 5 m=Ma at the distal end. Thisreflects the tendency of coarser-grained sedimentsto be trapped close to the shoreline. A consequenceof this is that since the sedimentation rate at anypoint along the basin profile is a function of itsdistance from the shoreline, it must therefore varyin time as the shoreline progrades and retreats. It isthis feedback loop between water-depth changes andsedimentation that causes progradation and retrogra-dation (Demarest and Kraft, 1987). Incorporation of


Fig. 6. Computation of the rate of facies shift as a function of the rate of water-depth change and gradient of the depositional surface.This phenomenon links facies progradation–retrogradation and water-depth changes.

this induced facies shift into the model would ne-cessitate recalculating the sedimentation rate at eachpoint along the profile at each model time step.The magnitude of the sedimentation rate shift is afunction of the rate of water-depth change and thegradient of the depositional surface (Fig. 6). By as-suming an extreme shelf gradient of 1º and usingour highest rate of water shallowing (20.7 m=Ma),we calculate that the maximum lateral sedimentationrate shift is 148 m per 0.125 Ma (one time step). Atthe scale of the modelled basin profile (200 km) thisis insignificant, so we have ignored it and assumedthe sedimentation rate (S) to be constant at any givenpoint through time.

Combining the rates of tectonic subsidence (T)and sedimentation (S) at each point along the basinprofile gives the rate of movement of the depositionalsurface (D). As we have demonstrated above that thelateral migration of the sedimentation rate (S) isessentially negligible at the scale of the model, therate (D) remains constant at any particular pointon the profile through the course of the eustatichalf-cycle. The value of (D) does vary spatially,reflecting the reality of differential subsidence andsediment supply.

The model was advanced in increments of 0.125Ma. For each incremental time step, by adding therate of eustatic change (E) to the rate of deposi-tional surface movement (D) at each point alongthe profile, we calculate the rate of the water-depthchange (W) across the basin. The result is a graphic

output (Fig. 7) showing which portions of Profile Aare undergoing water-depth shallowing (W negative),and which water-depth deepening (W positive). Theboundary between these two zones marks the pointat which water-depth is stationary, which as notedabove is the condition for a sequence-bounding sur-face to form. This will be a CTS where it tops USF,and a MFS where it tops UDF.

3.2. Model results

The model is started at eustatic highstand, wherethe rate of eustatic change (E) is zero. The rate ofwater-depth change (W) at this time is thus equal tothe rate of movement of the depositional surface (D)and is positive along the entire length of the profile(Fig. 7). This water-depth deepening results in UDFthroughout Profile A.

The successive incremental time steps of themodel through a 1 Ma eustatic half-cycle from high-stand to lowstand are shown in Fig. 7. At each timestep, the distance along the profile at which the rateof water-depth change (W) equals zero is indicated;this is the point at which a sequence-bounding sur-face is formed, separating areas of coeval depositionof upward-deepening and upward-shallowing facies.The bounding surface generated between time steps1 and 5 is referred to as a MFS because it separatesUDF, below, from the overlying USF. The boundingsurface generated starting with time step 5 is a CTSas it separates USF from the overlying UDF.


As the rate of eustatic fall (E) increases fromtime step 1 to 5 (the eustatic inflexion point), aprogressively larger value of net subsidence (D) isrequired to balance it and maintain the stationarywater-depth condition under which the MFS forms.This results in the formation of the MFS movingbasinward through time in the direction of increasing(D). The MFS is thus younger offshore than it istowards the basin margin.

Time step 5 (0.5 Ma) is the inflexion point onthe falling limb of the sinusoidal eustatic curve andrepresents the maximum rate of eustatic fall. Thisis balanced by (D) at a distance of 71.3 km alongthe profile (Fig. 7). Basinward of this point, the rateof subsidence of the depositional surface (D) alwaysexceeds the maximum rate of eustatic fall (E), and aMFS is not formed.

During time steps 5 to 9 the rate of eustaticfall decreases to zero. Continued subsidence (D)results in a water-depth deepening and UDF. As therate of eustatic fall decreases, it is balanced by aprogressively lower value of (D) and the boundary(in this case the CTS) thus moves toward the basinmargin. The CTS is therefore older offshore than itis toward the basin margin (Fig. 7).

3.3. Strike variability

To further illustrate the diachroneity of the bound-ing surfaces we have added two other basin profiles(B and C) to the model (Fig. 8). These representdip-sections across the same basin at 50 and 100km along strike from Profile A. Profiles B and Care assigned slightly different values of subsidencerate (T) and sedimentation rate (S) to reflect the typeof strike-variability found in reality. All three mod-els were run through the same eustatic half-cycle,

Fig. 7. Timing of sequence-bounding surface formation alongbasin Profile A as a function of the interaction between eustasy(E), subsidence (T) and sedimentation (S). The input values of(T) and (S) are given at the top together with the resulting rangeof values of (D). The value of (E) comes from Fig. 5. The rateof water-depth change (W) is shown across the basin profileat nine incremental time steps through a eustatic half-cyclefrom highstand to lowstand. A sequence-bounding surface formswhere the rate of water-depth change (W) equals zero, andmigrates laterally with time as different combinations of thethree variables meet this criterion.



Fig. 9. Isochron maps of the basin containing Profiles A, Band C showing the timing of bounding surface formation: (1)the time of formation of the maximum flooding surface (formedduring time steps 0 to 5), and (2) the time of formation ofthe conformable transgressive surface (formed during time steps5 to 9). The duration of bounding surface formation (i.e., itsdiachroneity) during the considered half of the eustatic cyclewavelength is 0.5 Ma, or one quarter of the eustatic cycle period.

which allows the formation of bounding surfaces tobe depicted in a map view.

The graphical incremental time steps for Pro-files B and C are shown in Fig. 8. By using theoutput time–distance data for bounding surface for-mation point along each of the three profiles, wehave constructed isochron maps of sequence-bound-ing surface formation through the eustatic half-cycle(Fig. 9). The isochrons join points where boundingsurface formation was synchronous on the three pro-files and show its diachronous movement. The rateof movement varies with the rate of eustatic change.

Fig. 8. Incremental output of Profiles B and C model runs. These profiles are assigned different rates of subsidence (T) and sedimentation(S) to Profile A, but the same eustatic curve (E) shown in Fig. 5 was used. The different rates of movement of the depositional surface(D) in Profiles A, B and C result in along-strike variation in the timing of boundary formation.

The diachroneity (duration) of bounding surface for-mation along the 200 km of Profile B is 0.5 Ma,which is approaching the resolution of ammonitezonation in the Jurassic and Cretaceous.

It is noted that the USF generated during themodel run, outlined by a MFS at the base and aCTS at the top, does not extend across the entirebasin but it wedges out offshore where the ratesof relative subsidence of the depositional surface(Fig. 4) completely outpace the rates of eustatic falland therefore a continuous deepening of the sea takesplace (Figs. 7–9). This situation reinforces the factthat bounding surfaces may only be generated whenthe three controlling factors vary in the same range.

4. Timing of bounding surfaces controlled byrelative sea-level changes

4.1. Two-dimensional model

The data input for the numerical model presentedin the preceding section depicts a situation in whichthe subsidence rates are always greater than the ratesof eustatic fall. This is a case of continuous relativesea-level rise in which sedimentation, together withvarying rates of eustatic fall, represent key elementsin the deposition of UDF and USF, and implic-itly in the formation of CTS and MFS. However,no surfaces controlled by relative sea-level changes(i.e., the c.c. and its counterpart surface generated atthe start of relative sea-level fall) may form duringcontinuous relative sea-level rise (Fig. 1). For thesesurfaces to form, a transition from relative rise torelative fall and vice versa is required, which mayonly happen when eustasy and tectonics vary withinat least partially overlapping ranges. To model thissituation, we take two additional depositional-dipsections, referred to as Profiles D and E (Figs. 10and 11), which are parallel to each other and sepa-rated by a distance of 100 km measured along thedepositional strike. We use the same curve of eustaticvariation as for Profiles A–C (Fig. 5), this time in


Fig. 10. Diachronous formation of surfaces separating deposits accumulated under relative sea-level fall and rise conditions (left column),as well as deposits accumulated under water deepening and shallowing conditions (right column). The rates of tectonic subsidence forProfile D have been selected to partially overlap with the rates of eustatic fall (Fig. 5), to allow both relative sea-level fall and rise tomanifest within the modelled area. E D rate of eustatic change.


Fig. 11. Similar computations as in Fig. 10, only with a different set of rates for tectonic subsidence, to illustrate the possible strikevariability within the sedimentary basin.


interplay with lower rates of subsidence and sedi-mentation. The difference between the sets of inputdata selected for Profiles D and E consists in thesubsidence rates, being slightly greater for Profile E.

The model run, advancing at incremental timesteps of 0.125 Ma, is similar to the one described inSection 3.3. The interplay between partially overlap-ping eustatic and subsidence rates allows the mani-festation of both relative sea-level fall and rise duringthe considered eustatic half-cycle. At each time step,the distance along the profile at which the rate ofrelative sea-level change (R) equals zero is indicated;this is the point at which a sequence-bounding sur-face (c.c. or start of relative fall surface) is formed,separating areas of coeval relative fall and rise. Asthe point of stationary relative sea-level migrates intime within the basin, the c.c. (end of relative fall)and its counterpart (onset of relative fall) are found tobe diachronous, with a net diachroneity approaching0.5 Ma (a quarter of the eustatic cycle; Figs. 10 and11, left columns). By adding the effect of sedimenta-tion to the relative sea-level changes, the water-depthchanges as well as the timing of the CTS and MFScould be modelled as well (Figs. 10 and 11, rightcolumns). The slightly steeper slopes for the CTSand MFS relative to the c.c. and the start of relativefall surface (Figs. 10 and 11) indicate higher rates ofdiachroneity for the former, which is explained bythe additional control of sedimentation.

Fig. 12 illustrates the case of Profile D duringone and a half eustatic cycles, with the CTS, MFS,c.c. and SRFS (start of relative fall surface) curvesplotted from Fig. 10. The shoreline trends, as wellas the various types of systems tracts separated bythese bounding surfaces are also represented. Thebasinward increase in subsidence rates, parallelledby a decrease in sedimentation rates, imposes a limitin the seaward extent of the USF, assuming thatthe subsidence rate in the basin centre exceeds themaximum rate of eustatic fall. The basinward extentof bounding surfaces controlled by relative sea-levelchanges (c.c. and SRFS) reaches the point where thesubsidence rates start to exceed the maximum rate ofeustatic fall. Similarly, bounding surfaces controlledby water-depth changes (CTS and MFS) extend upto the point where the rate of relative subsidence ofthe depositional surface (D D T � S, Fig. 4) starts toexceed the maximum rate of eustatic fall.

It can be noted that we did not use the termsTST and RST in Fig. 12, but rather UDF and USF,although the second sets of definitions for CTS andMFS (see Section 1) would make the terms TSTand UDF, as well as RST and USF, to be equiva-lent. According to Embry (1995), the CTS may bediachronous because “transgression may begin laterin areas of high sediment supply”, which is exactlythe point made in Fig. 12. In this light, the TSTwould be equivalent to the UDF, and the RST withthe USF. On the other hand, transgressive (retrogra-dational) and regressive (progradational) stackingpatterns are controlled by the shoreline movements(Fig. 3), which does not allow for such equivalence(see also the discussion in Section 1).

4.2. Three-dimensional model

The model results from Profiles D and E havebeen plotted together in bloc diagrams (Figs. 13 and14) to illustrate different aspects of sequence strati-graphic diachroneity, such as the diachroneity ofbounding surfaces along the depositional strike anddepositional dip (Fig. 13, lower diagram), and thesimultaneous formation of different systems tractsduring discrete time steps (Fig. 14). The distancealong the depositional strike between Profiles D andE has been selected arbitrarily (100 km in this case),to support a realistic subsidence variability withinthe basin.

Our results suggest that the bounding surfacescontrolled by water-depth changes (CTS and MFS)always extend farther basinward than the surfacescontrolled by relative sea-level changes (c.c. andSRFS), due to the effect of sedimentation. As a re-sult, the LST and HST merge together beyond thewedging out point of the FSST. The FSST includesstrata accumulated under relative sea-level fall con-ditions, and bounded by diachronous surfaces (c.c.and SRFS) which merge together offshore (Figs. 10–14). This raises a recurrent sore issue of sequencestratigraphy: where is the place within the sequencestratigraphic framework of the deep marine gravityflow deposits? While the position of the basin floorsubmarine fans relative to the depositional sequenceboundary is a controversial issue (above the c.c.:Posamentier et al., 1988; Emery and Myers, 1996;or below the c.c.: Hunt and Tucker, 1992; Helland-


Fig. 12. Time–distance depositional-dip section based on the model results from Profile D, constructed for one and a half eustatic cycles.The different rates of shoreline regression illustrate forced (high rate) regressions and normal (low rate) regressions. Abbreviations: DS Ddepositional sequence; GS D genetic stratigraphic sequence; TR D transgressive–regressive sequence; LSTD lowstand systems tract; FSSTD falling stage systems tract; HST D highstand systems tract; USF D upward shallowing facies; UDF D upward deepening facies; CTS Dconformable transgressive surface; MFSD maximum flooding surface; SRFSD start of relative fall surface; c.c. D correlative conformity.

Hansen and Martinsen, 1996), they are generallyregarded as formed during the forced regression ofthe shoreline and therefore part of the FSST. How-ever, if we define the FSST boundaries on the basisof the interplay between eustasy and varying subsi-dence rates throughout the basin (Fig. 12), then thegravity flows could be initiated in the shallow marinearea subject to relative sea-level fall, but the sedimentwould be transported and deposited beyond the edgeof the FSST, within an area characterized by relativesea-level rise and accumulation of UDF (Fig. 13).

5. Conclusions

The timing of bounding surfaces may be assessedin relationship to the criteria employed to definethem. There are two main ways to define strati-graphic surfaces: (1) in relationship to the stratalstacking patterns, which we call ‘type A surfaces’,and (2) in relationship to the water-depth changesand relative sea-level changes, which we call ‘typeB surfaces’. The types A and B surfaces are repre-sented in Fig. 15.

Fig. 13. Block-diagrams illustrating strike and dip diachroneity of surfaces modelled in Figs. 10 and 11 (profiles D and E). The lowerdiagram shows a possible relationship between the source area for gravity flow deposits (S), which may be placed within a regionaffected by relative sea-level fall, and the associated submarine fan (SF) which may be placed in a region characterized by relativesea-level rise. Abbreviations: DS D depositional sequence; GS D genetic stratigraphic sequence; TR D transgressive–regressive sequence;LST D lowstand systems tract; FSST D falling stage systems tract; HST D highstand systems tract; USF D upward shallowing facies;UDF D upward deepening facies; CTS D conformable transgressive surface; MFS D maximum flooding surface; SRFS D start of relativefall surface; c.c. D correlative conformity.


Fig. 14. Block diagrams illustrating coeval deposition of different systems tracts, as derived from the model results for Profiles D and E(Figs. 10 and 11). For abbreviations see Fig. 13.

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Fig. 15. Temporal significance of types A and B surfaces. Type A surfaces depend on the shoreline movements (interplay between the rates of sedimentation and relativesea-level rise in the shoreline area), relative sea-level changes in the shoreline area, and rates of offshore sediment transport (not represented in this figure, but suggestedin Fig. 3). They do not depend on the offshore variations in sedimentation and subsidence rates. Type B surfaces depend on the offshore variation in sedimentation and=orsubsidence rates. Abbreviations: D D diachroneity rate; E D eustasy, T D tectonics; S D sedimentation; SU D subaerial unconformity; MRS D maximum regressive surface;MTS D maximum transgressive surface; f.u. D fining-upward; c.u. D coarsening-upward; d.u. D deepening-upward; s.u. D shallowing-upward; HST D highstand systemstract; FSST D falling stage systems tract; LST D lowstand systems tract; RST D regressive systems tract; TST D transgressive systems tract.


The timing of type A surfaces depends either onthe shoreline movements, which determine the direc-tion of facies shift and the prograding or retrogradinggeometries, or on the rate of offshore sediment trans-port. Type A CTS and MFS are time lines alongdepositional dip sections, marking the turnaroundpoints from shoreline regression to transgression andvice versa (Figs. 3 and 15). Their timing depends onthe interplay between the rates of sedimentation andrelative sea-level rise in the shoreline area, and it isnot affected by the offshore variations in the sedi-mentation and subsidence rates. Both type A CTSand MFS are potentially diachronous along the de-positional strike, as variations in sedimentation andsubsidence rates determine temporally offset transi-tions from regression to transgression and vice versaalong the shoreline. Type A c.c. (end of forced re-gression of the shoreline) and its counterpart (basalsurface of forced regression: onset of forced regres-sion of the shoreline) are diachronous along thedepositional dip, with a rate that matches the rateof offshore sediment transport (Fig. 3). They arealso not affected by the offshore variations in thesubsidence rates. Type A c.c. and basal surface offorced regression are potentially diachronous alongthe depositional strike as well, due to variations insubsidence rates along the shoreline. All type A sur-faces could be theoretically recognized on the basisof stratal stacking patterns (Figs. 2 and 3); also thismay often be very difficult at the outcrop scale.

Type B surfaces relate to bathymetric changes(CTS and MFS) or relative sea-level changes (c.c.and SRFS) throughout the basin, and are character-ized by much higher diachroneity rates than the typeA surfaces (Fig. 15). The timing of type B surfacesdepends on the offshore variations in the sedimenta-tion and subsidence rates, as modelled in Figs. 7, 8,10 and 11. The model results indicate that the degreeof diachroneity of bounding surface formation spansup to a quarter of the eustatic cycle period, whichis in agreement with the quarter-cycle phase shiftnoted by Angevine (1989) Christie-Blick (1991) andJordan and Flemings (1991). The cases describedabove involve a sinusoidal eustatic curve of higherfrequency than the other variables. The higher-fre-quency variable could equally be sedimentation (S)or subsidence (T), so it can be concluded that thedegree of diachroneity of a bounding surface reaches

up to a quarter of the period of the highest frequencyvariable.

Types A and B surfaces merge together in theshoreline area (Fig. 15). The differentiation betweenthe two becomes more important offshore, especiallyin regards to the gravity flow deposits (Fig. 15).For practical reasons, the use of type A surfaces forstratigraphic correlation seems to be a better choice,both from a chronostratigraphic point of view, andfrom the point of view of the field signatures ofsystems tracts and bounding surfaces.

The terms CTS and MFS are currently usedwith both architectural and bathymetric implicationswhich, as this paper demonstrates, are not equiv-alent. Alternative terminology is suggested for theboundaries between prograding and retrograding ge-ometries, i.e. ‘maximum regressive surfaces’ for thetop of prograding facies, and ‘maximum transgres-sive surfaces’ for the top of retrograding facies (Fig.15; Catuneanu, 1996; Helland-Hansen and Martinsen,1996).

Acknowledgements

Financial support during completion of this workwas provided by Rhodes University, NSERC Canada,University of Toronto, AAPG, Petrel Robertson,Wascana Energy, Petro-Canada, Union Pacific Re-sources and GSA. We thank Ashton Embry, Christo-pher Kendall and Nicholas Christie-Blick for com-ments and suggestions on previous versions of themanuscript. We also wish to thank reviewers WilliamHelland-Hansen, Tim Naish and John Howell forvaluable advice and constructive criticism; this paperwas very much improved as a result of their reviews.

References

Angevine, C.L., 1989. Relationship of eustatic oscillations toregressions and transgressions on passive continental margins.In: Price, R.A. (Ed.), Origin and Evolution of SedimentaryBasins and their Energy and Mineral Resources. Am. Geo-phys. Union, Geophys. Monogr. 48, 29–35.

Catuneanu, O., 1996. Reciprocal Architecture of Bearpaw andpost-Bearpaw Sequences, Late Cretaceous–Early Tertiary,Western Canada Basin. Ph.D. thesis, University of Toronto,Toronto, 301 pp.

Catuneanu, O., Sweet, A.R., Miall, A.D., 1997. Reciprocal ar-


chitecture of Bearpaw T–R sequences, uppermost Cretaceous,Western Canada Sedimentary Basin. Bull. Can. Pet. Geol. 45(1), 75–94.

Christie-Blick, N., 1991. Onlap, offlap, and the origin of unconfor-mity-bounded depositional sequences. Mar. Geol. 97, 35–56.

Demarest, J.M., Kraft, J.C., 1987. Stratigraphic record of Qua-ternary sea levels: implications for more ancient strata. In:Nummedal, D., Pilkey, O.H., Howard, J.D. (Eds.), Sea-LevelFluctuations and Coastal Evolution. Soc. Econ. Paleontol.Mineral. Spec. Publ. 41, 223–239.

Embry, A.F., 1993. Transgressive–regressive (T–R) sequenceanalysis of the Jurassic succession of the Sverdrup Basin,Canadian Arctic Archipelago. Can. J. Earth Sci. 30, 301–320.

Embry, A.F., 1995. Sequence boundaries and sequence hier-archies: problems and proposals. In: Steel, R.J., Felt, V.L.,Johannessen, E.P., Mathieu, C. (Eds.), Sequence Stratigraphyon the Northwest European Margin. Norwegian PetroleumSociety Spec. Publ. 5, Elsevier, Amsterdam, pp. 1–11.

Embry, A.F., Johannessen, E.P., 1992. T–R sequence stratigra-phy, facies analysis and reservoir distribution in the uppermostTriassic–Lower Jurassic succession, western Sverdrup Basin,Arctic Canada. In: Vørren, T.O., Bergsager, E., Dahl-Stamnes,O.A., Holter, E., Johansen, B., Lie, E., Lund, T.B. (Eds.), Arc-tic Geology and Petroleum Potential. Norwegian PetroleumSociety Spec. Publ. 2, Elsevier, Amsterdam, pp. 121–146.

Emery, D., Myers, K.J. (Eds.), 1996. Sequence Stratigraphy.Blackwell, Oxford, 297 pp.

Fortuin, A.R., de Smet, M.E.M., 1991. Rates and magnitudesof late Cenozoic vertical movements in the Indonesian BandaArc and the distinction of eustatic effects. In: Macdonald,D.I.M. (Ed.), Sedimentation, Tectonics and Eustasy: Sea-LevelChanges at Active Margins. Int. Assoc. Sedimentol. Spec.Publ. 12, 79–90.

Frostick, L.E., Steel, R.J., 1993. Tectonic control and signaturesin sedimentary successions. Int. Assoc. Sedimentol. Spec.Publ. 20, 520 pp.

Galloway, W.E., 1989. Genetic stratigraphic sequences in basinanalysis, I. Architecture and genesis of flooding-surfacebounded depositional units. Am. Assoc. Pet. Geol. Bull. 73,125–142.

Gill, J.R., Cobban, W.A., 1973. Stratigraphy and geologic his-tory of the Montana Group and equivalent rocks, Montana,Wyoming, and North and South Dakota. U.S. Geol. Surv.Prof. Pap. 776, 73 pp.

Haq, B.U., 1991. Sequence stratigraphy, sea-level change, andsignificance for the deep sea. In: Macdonald, D.I.M. (Ed.),Sedimentation, Tectonics and Eustasy: Sea-Level Changes atActive Margins. Int. Assoc. Sedimentol., Spec. Publ. 12, 3–39.

Helland-Hansen, W., Martinsen, O.J., 1996. Shoreline trajecto-ries and sequences: description of variable depositional-dipscenarios. J. Sediment. Res. 66 (4), 670–688.

Hunt, D., Tucker, M.E., 1992. Stranded parasequences and theforced regressive wedge systems tract: deposition during base-level fall. Sediment. Geol. 81, 1–9.

Ito, M., O’Hara, S., 1994. Diachronous evolution of systemstracts in a depositional sequence from the middle Pleistocenepalaeo-Tokyo Bay. Japan. Sedimentology 41, 677–697.

Jervey, M.T., 1988. Quantitative geological modeling of sili-ciclastic rock sequences and their seismic expression. In:

Wilgus, C.K., Hastings, B.S., Kendall, C.G.St.C., Posamen-tier, H.W., Ross, C.A., Van Wagoner, J.C. (Eds.), Sea-LevelChanges: An Integrated Approach. Soc. Econ. Paleontol. Min-eral. Spec. Publ. 42, 47–70.

Jordan, T.E., Flemings, P.B., 1991. Large-scale stratigraphic ar-chitecture, eustatic variation, and unsteady tectonism: a theo-retical approach. J. Geophys. Res. 96 (B4), 6681–6699.

Macdonald, D.I.M. (Ed.), 1991, Sedimentation, Tectonics andEustasy: Sea-Level Changes at Active Margins. Int. Assoc.Sedimentol. Spec. Publ. 12, 518 pp.

Martinsen, O.J., Helland-Hansen, W., 1995. Strike variabilityof clastic depositional systems: does it matter for sequencestratigraphic analysis? Geology 23, 439–442.

Miall, A.D., 1991. Stratigraphic sequences and their chronos-tratigraphic correlation. J. Sediment. Petrol. 61, 497–505.

Miall, A.D., 1994. Sequence stratigraphy and chronostratigraphy:problems of definition and precision in correlation, and theirimplications for global eustasy. Geosci. Can. 21, 1–26.

Naish, T., Kamp, P.J.J., 1997. Foraminiferal depth palaeoecologyof Late Pliocene shelf sequences and systems tracts, WanganuiBasin, New Zealand. Sediment. Geol. 110, 237–255.

Pitman, W.C., 1978. Relationship between sea level change andstratigraphic sequences of passive margins. Geol. Soc. Am.Bull. 89, 1389–1403.

Pitman, W.C., Golovchenko, X., 1983. The effect of sea levelchange on the shelf edge and slope of passive margins. In:Stanley, D.J., Morre, G.T. (Eds.), The Shelfbreak: Critical In-terface on Continental Margins. Soc. Econ. Paleontol. Mineral.Spec. Publ. 33, 41–58.

Posamentier, H.W., Allen, G.P., 1993. Siliciclastic sequencestratigraphic patterns in foreland ramp-type basins. Geology21, 455–458.

Posamentier, H.W., James, D.P., 1993. An overview of sequencestratigraphic concepts: uses and abuses. In: Posamentier, H.W.,Summerhayes, C.P., Haq, B.U., Allen, G.P. (Eds.), SequenceStratigraphy and Facies Associations. Int. Assoc. Sedimentol.Spec. Publ. 18, 3–18.

Posamentier, H.W., Jervey, M.T., Vail, P.R., 1988. Eustatic con-trols on clastic deposition I, conceptual framework. In: Wilgus,C.K., Hastings, B.S., Kendall, C.G.St.C., Posamentier, H.W.,Ross, C.A., Van Wagoner, J.C. (Eds.), Sea-Level Changes:An Integrated Approach. Soc. Econ. Paleontol. Mineral. Spec.Publ. 42, 109–124.

Posamentier, H.W., Allen, G.P., James, D.P., Tesson, M., 1992.Forced regressions in a sequence stratigraphic framework:concepts, examples and exploration significance. Am. Assoc.Pet. Geol. Bull. 76, 1687–1709.

Reading, H.G. (Ed.), 1996. Sedimentary Environments: Pro-cesses, Facies and Stratigraphy. Blackwell, Oxford, 688 pp.

Vail, P.R., Audemard, F., Bowman, S.A., Eisner, P.N., Perez-Cruz, C., 1991. The stratigraphic signatures of tectonism,eustasy and sedimentology — an overview. In: Einsele, G.,Ricken, W., Seilacher, A. (Eds.), Cycles and Events in Stratig-raphy. Springer-Verlag, New York, N.Y., pp. 611–659.

Van Wagoner, J.C., Mitchum, R.M., Campion K.M., Rahmanian,V.D., 1990. Siliciclastic sequence stratigraphy in well logs,cores, and outcrops: concepts for high-resolution correlationof time and facies. Am. Assoc. Pet. Geol. Methods Explor.Ser. 7, 55 pp.

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Temporal signiﬁcance of sequence boundaries · 2015. 10. 6. · variable, whichever that is among the eustasy, tectonics or sedimentation controls. Types A and B surfaces merge