Marine Synrift Sedimentation Ravnas & Steel AAPG 1998

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ABSTRACT

Marine rift basins represent a continuum rangingfrom mixed nonmarine/marine through shallowmarine to deep marine, or from partly emergentthrough partly submergent to completely submer-gent basin types. These rift basin types have stronglyvariable synrift sedimentary architectures because oftemporal changes in relative sea level, accommoda-tion creation, and sediment supply throughout therift cycle. Accommodation changes are controlledmainly by local basin-floor rotation, basinwide back-ground subsidence, and, to a lesser degree, byeustatic changes. Sediment supply determines howmuch of the accommodation is filled and in whatmanner, and is controlled by the distance to themain hinterland areas, and the size and sediment-yield potential of any local fault-block source area.

Marine siliciclastic synrift successions, whetherdominantly shallow or deep marine in nature, areclassified in terms of sediment supply as overfilled,balanced, underfilled, and starved. Sediment-over-filled and sediment-balanced infill types are charac-terized by a threefold sandstone-mudstone-sandstone synrift sediment-infill motif; the sedi-ment-underfilled type is represented by a two-foldconglomerate-sandstone-mudstone motif; and thesediment-starved type commonly is represented bya one-fold mudstone motif. The sequential develop-ment, linked depositional systems, and stratigraph-

ic signatures of the early synrift, the rift climax, andthe late synrift to early postrift stages vary signifi-cantly between these rift basin infill types, as dothe tectonic significance (timing of initiation andduration) of stratal surfaces, such as footwallunconformities, nondepositional hiatuses, andmarine condensed sections. The construction ofthe fourfold rift basin infill classification schemeprovides a first basis and a strong tool for predict-ing the distribution and geometry of synrift reser-voir and source rock types, despite the inherentvariability of the marine synrift infills.

INTRODUCTION

Intracratonic rift basins form by the stretchingand faulting of continental lithosphere (e.g.,McKenzie, 1978; Wernicke, 1985; Lister et al., 1986;Rosendahl, 1987; Kusznir et al., 1991; Ziegler,1992). Because rift basins form in a variety of set-tings, their sedimentary fill shows a great variability(e.g., Leeder, 1995). Ideas on synrift sediment archi-tecture have evolved from studies of nonmarine(e.g., Crossley, 1984; Frostick and Reid, 1987;Leeder and Gawthorpe, 1987; Morley, 1989;Lambiase, 1990) and marine rift basins (e.g., Surlyk,1978, 1989; Leeder and Gawthorpe, 1987; Prosser,1993; Gawthorpe et al., 1994). In addition to recog-nizing the asymmetric geometry of fault-controlledsedimentary packages, it has long been realized thatchanging lithological signatures and stacking pat-terns can reflect variation in fault-related subsi-dence rate, assuming a variation in the balancebetween sediment supply and subsidence rate. Clay-prone intervals (with time-equivalent conglomer-ates restricted to narrow belts adjacent to the faultscarps) in syntectonic successions are known torepresent periods of rapid differential subsidence

(Steel, 1988), whereas the intervals of coarse faciesextending farthest from the fault scarp relate toperiods of relative tectonic quiescence and minimalaccommodation creation (Blair, 1987).

Hamblin and Rust (1989) and Lambiase (1990)combined these ideas, proposing a threefold,fining- to coarsening-upward, sandstone-mudstone-sandstone lithology motif for lacustrine synrift suc-cessions. The basal and capping sandstones were

110 AAPG Bulletin, V. 82, No. 1 (January 1998), P. 110146.

Copyright 1998. The American Association of Petroleum Geologists. Allrights reserved.

1Manuscript received December 14, 1995; revised manuscript receivedDecember 9, 1996; final acceptance June 16, 1997.

2University of Bergen, Geological Institute, Allegaten 47, N-5007 Bergen,Norway. Present adress: Norske Conoco A.S., P.O. Box 288, N-4001Stavanger, Norway.

3University of Wyoming, Deptartment of Geology and Geophysics,Laramie, Wyoming 82071-3355.

We thank our fellow colleagues and research students in the Joule II:Integrated Basin Studies-Dynamics of the Norwegian Margin (IBS-DNM)project for stimulating cooperation and discussions. Constructive commentsby K. T. Biddle, C. K. Morley, T. H. Nilsen, and F. Surlyk on an earlierversion of the manuscript sharpened the final product and are greatlyappreciated. The drafting offices at the Geological Survey of Wyoming,Statoil, Norsk Hydro, and the Geological Institute, University of Bergen,created the illustrations. This work was funded by the Research Council ofNorway as part of the Joule II Research Programme (CEC Contract No.JOU2-CT92-0110).

For clarification, the term rift margin depositional system as used in thispaper refers to a geological feature that originates at the rift margin but mayextend out into the basin.

Architecture of Marine Rift-Basin Successions1

R. Ravns2 and R. J. Steel3


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argued to reflect the early and late synrift stages,respectively, whereas the intervening mudstoneswere related to the climax of rifting. The architec-tural styles of fluvial and arid synrift sedimentinfill probably reflect the depositional systemsresponse to the temporal variation in the riftbasins basin-floor and bounding fault-scarp slopes,rather then being related to any intrabasinal base

level, such as local lake level. Nevertheless, a three-fold sandstone-mudstone-sandstone synrift motifhas also been suggested for fluvial rift basin infills(Mack and Seeger, 1990; see also Alexander andLeeder, 1987).

In contrast to the architectural styles of nonma-rine synrift successions, workers have made fewerattempts to synthesize data on mixed nonma-rine/marine and marine synrift sediment architec-ture. Surlyk (1978, 1989) developed models forhanging-wall turbidite systems sourced from theadjacent footwall, and demonstrated a four-stageevolution of the turbidite systems with variablerates of sea level changes and extensional faulting.

Furthermore, Surlyks (1978, 1989) work resultedin a two-fold lithosome model for marine synriftsuccessions (Figure 1A). Subsequently, Prosser(1993) proposed a complementary, three-foldlithosome model for terminal marine rift basins,analogous to that suggested for lacustrine riftbasins (Figure 1B). These two models a re herepostulated to form type members in a series ofarchitectural patterns present in marine synriftsuccessions. Gawthorpe et al. (1994) examinedthe combined effects that varying tectonic stylesand subsidence rates, as well as glacioeustatic sealevel fluctuations, had on synrift sequence devel-opment in an extensional back-arc setting subject-ed to regional uplift.

Our understanding of rift basin formation and syn-rift sediment architecture has been aided by numeri-cal and analog structural and stratigraphic modeling.Such studies have highlighted the structural parame-ters (Barr, 1987a, b, 1991; Kusznir et al., 1991;Roberts et al., 1993a) and the onlap or offlap pat-terns expected to characterize the resultant synriftstratigraphy (Schlische and Olsen, 1990; Schlische,1991; Hardy, 1993; Roberts et al., 1993a; Waltham etal., 1993).

The aim of this paper is to further develop ourunderstanding of synrift sediment architecture ofmarine half-graben type rift basins, building on theearlier studies of Surlyk (1978, 1989), Prosser(1993), and Gawthorpe et al. (1994). Our overallcontribution is a synthesis of the controls on silici-clastic sequence architecture, with an emphasis onhow variable this architecture can be, in regionallyextended and subsiding marine terrains formedduring nonglacial periods. In contrast to previousstudies, our database from the northern North Sea

and western Portugal is wide ranging and based onfield studies, three-dimensional seismic coverage,and abundant well data. An additional feature wewish to emphasize is that the Middle JurassicEarlyCretaceous synrift successions studied herein havebeen derived from thick, poorly consolidated sedi-mentary terrains. Drainage development and sedi-mentary response to tilting and uplift thus are like-

ly to have been more rapid than for basementterrains, a feature at variance with assumptions inone of the most recent models (i.e., Prosser, 1993).In Prossers model, the sedimentary response isassumed to be delayed with respect to an extreme-ly rapid fault rate, leaving the rift climax relativelystarved of sediment. We also suggest that the roleof sediment supply, especially as enacted throughprerotational bathymetry and distance to main hin-terland areas, has been relatively neglected in previ-ous studies. Likewise, the importance of the shore-line position on the hanging wall with respect tolocation of the rotational fulcrum for the develop-ment of sediment supply has been underestimated

and not differentiated (see, however, Hardy, 1993;Roberts et al., 1993a).

RIFT BASIN DEVELOPMENT ANDARCHITECTURE

Synrift Subsidence

Synrift extension involves upper crust extensionaccommodated by faulting and stretching of lowercrust and lithospheric mantle by plastic deforma-tion (Kusznir et al., 1991; Kusznir and Ziegler,1992). Synrift subsidence is a state resultant fromthe balance of the following items.

(1) Isostatic adjustment of the crust to mechani-cal stretching of the lithosphere leads to the subsi-dence of the thinned crust (McKenzie, 1978;Cochran, 1983). This is accompanied by flexuraluplift or by downwarping of the rift zone, depend-ing on the depth of the lithospheric necking level(Kooi et al., 1992; Ziegler, 1992).

(2) Upwelling of the aesthenosphere into thespace created by mechanical stretching of the litho-sphere, which, together with thermal upward dis-placement of the aesthenosphere-lithosphereboundary, causes uplift of the rift zone (Turcotteand Emmermann, 1983; Ziegler, 1992).

According to Ziegler (1992), the structural style ofrift basin development is influenced by the thick-ness and thermal state of the crust and subcrustallithosphere at the onset of rifting; the amount ofcrustal extension and width over which it is dis-tributed ( factor); the mode of extension, whetherorthogonal or oblique; and the lithological composi-tion of prerift and synrift deposits. In addition, any

112 Marine Rift-Basin Architecture


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preexisting basement grain commonly exerts astrong control on the preferential development andorientation of synrift structures (see review inMorley, 1995). Changing crustal conditions, espe-cially temperature, can cause temporal changes inrift geometry during extension (Kusznir and Park,1987), such as from low- to high-angle fault geome-tries (Morley, 1989; Bosworth, 1992).

General Features of Rift Basins

An analysis of rifted terrains shows clearly thatthe fundamental morphological element is the faultblock or half graben formed within the hangingwall of major basin-bounding master faults (Figures2, 3) (Bosworth, 1985; Rosendahl et al., 1986;Morley, 1995). The rift basin may consist of a singlehalf graben or a series of half grabens, commonly

with pronounced variat ion in morphology topography along strike. The width of classic ear rift valleys ranges from 50 to 100 km. Twid th contr ast s wit h regional ly extended provinces, which may exceed 1000 km in widHalf-graben subbasins commonly have width1035 km, although widths of more than 50 may occur. Shallow detachments commonly pduce shorter wavelength fault blocks comparedfault blocks delineated by basement faults (eRattey and Hayward, 1993). The length of recbasin-bounding master faults is typically so3560 km, with individual fault strands at surfrarely exceeding 1520 km (Jackson and Wh1989). The basin-bounding master faults tendoverlap and switch polarity along the rift zoresulting in the formation of transfer and accomodation zones (Bosworth, 1985; Rosenda1987; Morley et al., 1990).

Ravns and Steel

(C) ACCOMMODATION DEVELOPMENT IN TIME AND SPACE(CHANGING RATE OF ROTATION / SUBSIDENCE / EUSTASY/SEDIMENT SUPPLY)

Spatial variations along fault lengths

Late synrift / tectonic quiescenceRift climaxEarly synrift

(A) SEDIMENT SUPPLY (BATHYMETRY + SOURCE AREA SIZE / LITHOLOGY)

INCREASING DISTANCE FROMLARGE SOURCE AREAS

MARGINAL &NONMARINE

INCREASING SUBMERGENCE DEEP-WATERSUBBASINS

SL

Hinterland sourceareas (large) Intrabasinal source areas (small)

(B) ACCOMMODATION (ROTATION + BACKGROUND SUBSIDENCE + FAULTSPACING + EUSTASY)

prerift surface

rotation

fulcrum

backgroundsubsidence

fault spacing

HANGIN

GWALL

(onlapo

finfill)

FOOTWALL

Sb

(zero net vertical movement)

rotationalpivot

Figure 2Controls on riftarchitecture. (A) Sedimensupply in relation tohalf-graben location,

bathymetry, source areasize, and lithology;(B) accommodationdevelopment, synrift tiltinand subsidence (Sb =

background subsidence);and (C) accommodationdevelopment in time andspace, and architecture ofthe synrift sediment fill.


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Basin-b

oundingmasterfault

F

H

HF

F

H

F

H

l

ll

lll

Subbasin-bound

ingmasterfault

Subbasin-bou

ndingmasterfault

Subbasin-bou

ndingmasterfault

Subbasin-b

oundingmasterfault

Subbasin

-bou

ndingm

aste

rfault

Subbas

in-boundingmasterfault

Sediment dispersal direction

Drainage divide

Accommodation space

Footwall

Hanging wallH

F

(A)

(B)

Figure 3(A) Half-grabenmorphology andrift-interior sedimentdispersal pattern.(B) Dip section throughrifted terrains showingaccommodation,intrabasinal drainage,and sediment dispersaldirection in (I) a solitaryhalf graben, (II) a seriesof half grabens withsimple basin-boundingmaster faults, and (III)a series of half grabens

with segmentedfault-complex orcollapsed footwallas basin-boundingstructures.


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Extension generates footwall uplift and adjacenthanging-wall collapse (Stein and Barrientos, 1985;Jackson et al., 1988; King et al., 1988; Stein et al.,1988). The throw on the subbasins boundary faultdecreases away from the central part of the faulttoward its tips (Figure 3A) (Walsh and Watterson,1988; Jackson and White, 1989), producing the typi-cal half-graben accommodation: wedge-shaped in

vertical section, and commonly scoop-shaped inplan view (Figure 3). The hanging-wall depositionalsinks may attain considerable depths; marine synriftstrata have thicknesses of up to 3 km (Surlyk, 1978).

A complicating feature of a rifted terrain with aseries of half grabens is that the basin-boundingmaster faults, which themselves do not moveuniformly along their length (Schwartz andCoppersmith, 1984; Wallace, 1985a), may or maynot move in concert across the region. Althoughinitiation, reactivation, migration, and cessation ofmovement on faults is likely to vary across the area,it is still unclear whether an entire rifted area expe-riences synchroneity or diachronism of its tectonic

maxima and minima. Dating and chronostratigraph-ic resolution is still too poor to determine whichalternative was more likely in the northern NorthSea, although there are indications of a weaklydiachronous rifting acme.

Another feature of rifted areas is their commonpulsed tectonic development. It has been arguedthat there are recurrent tectonic phases in manyextensional terrains, particularly well illustrated inthe Middle JurassicEarly Cretaceous infill historyof the northern North Sea rift basins (Underhill,1991a, b; Rattey and Hayward, 1993; Frseth et al.,1995), but also observed in other marine (Sellwoodand Netherwood, 1984; Purser et al., 1990) andnonmarine rifts (Lambiase, 1990). During MiddleJurass icEar ly Cretaceous rift ing in the NorthSeaNorwegian Sea, the successive rift phases com-monly had a duration of some 46 m.y., whereasthe entire rift episode lasted a few tens of m.y.

Unique Aspects of Rift Basins

One of the unique aspects of fault blocks or halfgrabens that affects their role as depositional sinksis the unusually steep slope gradient attainable bythe basin floor and the systematic variation of thisslope in time. This steep slope gradient arises as aresult of extensional rotation, which causes simul-taneous subsidence and uplift of downdip andupdip areas of the fault block, respectively. Rapidlyattained basin-floor tilt angles of 23, with longerterm cumulative tilt of up to 10, is not uncommonin the Late Jurassic subbasins of the northern NorthSea (Frseth et al., 1995). An important conse-quence of the steep local slopes of half grabens,

further enhanced by the tendency for accommotion creation to outpace sediment supply for muof the rift cycle, is the potential developmenprominent transgressive systems tracts.

A second unique aspect of rift basins, particuly marine rift basins, is the variability of the sment supply. Supply of sediment to any subbasisensitive both to hinterland distance and to the

and gradient of the local fault-block source ar(which themselves vary through the rift cycle). Tsediment-yield potential of areas updip of the fblock is likely to be less than the volume of acent depositional sinks; therefore, the riskmarine subbasins being underfilled is great.

The important implication of this spectectonostratigraphic style is that the creationaccommodation space is complex, varying in tiand space, because of background subsidence uplift, fault-block rotation and differential sudence, linked variable but commonly limited sment supply, and varying eustatic sea level. In mconventional sequence-stratigraphic models, in

grabens or other gently sloping basins sedimsupply and basin floor tectonics are assumed toconstant or uniformly varying in space; thereforis hardly surprising that sequence stratigraphdifficult to use as a predictive tool in analyzing srift basin fills (Steel, 1993). In a synrift half-grasituation there is updip erosion and generationsediment concomitant with, and only a short tance from, downdip accommodation creation sediment accumulation. The upslope-downslomigration of the shoreline zone generatedchanges in rotational style and rate and inf luenby changes in regional subsidence and eustatic level, then, in addition to the hinterland-basinwmovement of the rift-margin shoreline, largdetermines the regressive or transgressive tendcies of the infilling sedimentary systems.

CONTROLS ON RIFT-INFILL ARCHITECTURE

In this section, we evaluate the factors that ctrol the variability of synrift stratigraphy in marhalf grabens during a single rift phase and the pceding and succeeding tectonic quiescent stagThe resultant synrift stratigraphy appears to be ctrolled mainly by sediment supply in relation to tonic subsidence and uplift (or accommodatdevelopment).

Controls on Accommodation

Globally synchronous changes in sea level refrom changes in the volume of the ocean basand in the volume and distribution of wate

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them. For half grabens connected to the oceans,five possible causes are suggested for eustaticchanges in sea level (Table 1) (Allen and Allen,

1990). Except for glacio-eustasy and desiccationcrisis, eustatic sea level changes are of relativelyslow rates. Glacio-eustatically controlled accommo-dation development is commonly asymmetrical,resulting from lower rates of sea level fall than rise(Table 1), and produces rather short-term (10100k.y. time scale) variations in sea level. These short-term variations are attributed to Milankovitchcyclicity, and may result in the development ofhigh-frequency sequences (e.g., see Allen andAllen, 1990; Plint et al., 1992).

Structurally controlled accommodation resultsfrom the effects of background subsidence andfrom slip on basin-bounding (and intrabasinal)faults (Figure 2). Both tilting and vertical move-ments accompany slip on active normal faults.Vert ical movements involve hanging-wall subsi-dence and, commonly, footwall uplift (Barr, 1987a,b; Jackson et al., 1988; Jackson and White, 1989).These vertical movements decrease with distancefrom the fault, causing a tilting of the footwall andhanging-wall blocks (Figure 3). The rotational pivotdenotes the line that separates the areas subjectedto footwall uplift from areas experiencing hanging-wall subsidence as induced by fault-block tiltingalone. The superimposed background subsidence,which causes the enti re fault block to subside,results in an updip movement of the line of zeronet vertical movement and thus a narrowing of thezone experiencing synrotational uplift. The ful-crum is the line separating those areas of the faultblock subjected to uplift from areas subjected tosubsidence, taking into account the effects of bothtilting and background subsidence (Figure 2). Theupdip-downdip movement of the shoreline zone isdetermined predominantly by the combined effect

of the fulcrum position and the prerotational sealevel stand (Figure 4), in addition to any eustaticfluctuations.

If the rift basin includes a series of master faults,the dip section is a series of fault blocks or halfgrabens. The space created by fault-block rotation(Figure 2) depends upon the rate of rotation, totalextension, and initial master fault spacing (Barr1987a, b; Yielding, 1990), the initial fault dip andfault shape (planar, listric, or ramp-flat-rampgeometries) (e.g., Morley, 1989), as well as litho-sphere properties (ter Voorde and Cloetingh,1995). In continental rifts and proto-oceanictroughs, coseismic footwall uplift commonlyranges from 5 to 25% of the total displacementalong the fault (Table 2) (Jackson and McKenzie,1983; Stein et al., 1988; Jackson and White, 1989;Yielding and Roberts, 1992). During periods of seis-mic inactivity (interearthquake intervals), theupper crust deforms isostatically by flexuring (Kinget al., 1988; Stein et al., 1988; Weissel and Karner,1989; Kusznir et al., 1991). Hanging-wall deposi-tion and footwall erosion further influence the iso-static relaxation path following coseismic deforma-tion. Accordingly, a total footwall uplift of some50% of the hanging-wall subsidence may resultfrom coseismic and postseismic deformation (Steinand Barrientos, 1985; Roberts and Yielding, 1991).Footwall uplift, however, may be suppressed if thestretching factor is large, the initial fault spacing issmall (Barr, 1987a, b), or if the background subsi-dence is high.

The amount of stretching increases toward therift basin axis, and is commonly expressed by largerthrows on the basin-bounding faults and strongertilting of the basin floor (Roberts et al., 1993b).Fault spacing commonly shows a systematic vari-ance as well, and decreases toward the rift axis.This variance in fault-spacing possibly reflects the


Table 1. Eustatic Changes in Sea Level Showing Causes, Rates, and Magnitudes*

Mechanism Rate Magnitude (m)

Continuing lithospheric differentiation 0.20.4 mm/1000 yr 150250(0.00020.0004 mm/yr)

Changes in the volumetric capacity of the Not expected to exceedocean basins caused by sediment 1.0 mm/1000 yr influx or removal (0.001 mm/yr)

Changes in the volumetric capacity of the Maximum of 7.0 mm/1000 yr


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Ravns and Steel

(A)SLIGHTLYSUBMERGEDSUBBASINS

(B)

PARTLYEMERGEDSUBBASINS

(C)COMPLETELYSUBMERGED

SUBBASINS

C

C'

B

B'

A

A'

A A

A'

A'

B

B'

B

B'

C

C'

C

C'

ll

lll

l

l

l

ll

ll lll

lll

fulcrum

fulcrum

fulcrum

SLbelow

fulcrum

Expansionofsea

Transgressionand

storageinbackbarrierandcoastal

plains.

Mainlysed

iment-balancedsubbasinsdueto

potentialaxialsys

tem.

Han

ging-wallforced(andpossiblynormal)regression.

Blanketofhemipelagicmudstonesandpossible

submarinegravitationalresed

imentation.

Starvationorsubmarine

erosion

onsubmerged,updipareas

Forced

regression

above

fulcrum

possible

normal

regression

below

fulcrum

fulcrum

fulcrum

fulcrum

fulcrum

fulcrum

SLabovefulcrum

Subm

erged

wideningoffootwallisland

dueto

shoreline

progradation

footwall

uplift

fulcrum

Figure4

Bathymetryof(A)slightlysubmergedsubbasins,(B)partlyemergedsubbasins,and(C)completelysubmergedsubb

asins.Bathymetryat

theonsetofriftingisshownasblockdiagrams(rowI)

andverticalcross-sections(rowII)

.Riftclimaxbathymetryandsedim

entinfillpatternare

showninrowIII.No

tethedifferentdevelopmentofhan

ging-wallshorelinesintheslightlysubmergedandpartlyemergedsubb

asins.


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progressively attenuated lithosphere or higher heatflow beneath the more extended regions (Jacksonand White, 1989).

Slip on normal faults related to single large earth-quakes (Richter scale magnitude >6.5) (Jackson andWhite, 1989) ranges between 1.5 and 4 m, with themaximum reported slip on recent faults as large as 6 m(Fraser et al., 1964; Wallace, 1985a). Reported earth-quake recurrence intervals range from 500 to 7600 yr(Table 2), although large-magnitude earthquakes

sometimes are temporally clustered (Machette et al.,1991). After faulting has occurred at one locality,activity may shift to other segments or lineamentsand not return for thousands of years (Wallace andWhitney, 1984). Displacement rates on normal faultsin extensional terrains are from 0.01 to more than5.0 mm/yr (Table 2), with the values including possi-ble slip on successive small- and moderate-magnitudeearthquakes. Recurrence intervals (and thus sliprates) on recent faults vary laterally along fault


Table 2. Compilation of Displacement Rates on Normal Faults in Extensional Settings with References for FurtherData

Hanging-Wall Footwall EarthquakeRift Basin and Displacement Rates Subsidence Rates Uplift Rates RecurrenceReferences* (mm/yr) (mm/yr) (mm/yr) Intervals (yr)

Basin and Range ProvinceTeton fault, Wyoming (13) 0.451.6 7006000

Wasatch fault zone, Utah (46) 0.353.9 5004000average 2000c.r.i. 400

Stillwater front, Nevada (79) 0.260.5 28007600Owens Valley, California (10) >0.4


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zones, the earthquake frequency and magnitudebeing lower on fault segments located toward theends of the lineaments (Schwartz and Coppersmith,1984). Displacement rates vary laterally also alongindividual fault segments (Wallace, 1985a).

Spatial and temporal variations in style and therate of fault-block rotation, whether linear, nonlin-ear, or pulsating, influence the rate of accommoda-tion creation. During a single rift phase, the displace-ment rates increase during the early synrift substagetoward a maximum during the rift climax, and thenprogressively wane during the late synrift substage.Tectonic quiescent stages are prolonged periods oflow seismic activity occurring between recurrentrift phases and characterized by background subsi-dence and minimal fault-block rotation.

A comparison of publ ished da ta on ra te s ofeustatic sea level change with data on normal faultdisplacement rate shows clearly that, except duringglacial periods, structural control on accommoda-tion development by far exceeds eustatic control

(Tables 1, 2). Of the eustatic mechanisms, oglacio-eustasy appears to be of sufficiently high and magnitude to compete with accommodatchanges resulting from extensional faulting. Oteustatic mechanisms show rates and sometimalso magnitudes too low to leave a basinwimprint, and are most likely suppressed or masby higher tectonic subsidence rates in medial downdip reaches of the subsiding half grabeNevertheless, the eustatic signal resulting from other mechanisms (in combination with the strturally induced accommodation developmentlikely to have influenced the resultant relative level changes at (local) footwall and hanging-wshorelines.

Although it is difficult to disentangle the eustsignal in this type of setting (it may be cloudedvariations in sediment supply and suppressedstrong, local tectonics), some general pointsworth noting . Eustat ic sea level changes af fbroader areas on gently dipping hanging-w

Ravns and Steel

Synrift lithologies

Prerift lithologies

Highstand sea level

Lowstand sea levelLowstand wave base

Shoreline affecteeustatic sea levevariations

hsl

lslwb

hsl

lslhsl

lsl

hsl

lslwb

"broad" area affected by eustaticsea level variations

gentlehang

ing-wallslop

e

narrow area affectedby eustatic sea level

variations

steephang

ing-wa

llslope

hsl

lsl

wb

hsl

lslhsllsl

(A) (B)

Figure 5Areas affected by eustasy on (A) gently dipping slopes or hanging wall in shallow-marine half graband (B) steeply dipping slopes or hanging wall in deep-marine half grabens. Note that the fault-scarp successihave the steepest slopes. Fault-scarp shorelines thus are expected to be less influenced by eustasy.


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slopes compared to their steeper fault-scarpequivalents (Figure 5), especially if the hanging-wall shorelines are located at or close to the fault-block fulcrum. In addition, eustatic sea levelchanges show a greater influence during periodsof low rates of fault-related subsidence and basin-floor rotation; that is, during the early and latesynrift substages and intervening tectonic quies-

cent stages. If the rift basin development werecharacterized by stepwise rotation or a series ofrift phases, eustatic accommodation changesprobably would be more influential during theearlier phases because these would be represent-ed by the formation of shallower subbasins withgentler slopes, whereas later rift phases would becommonly dominated by deeper basins withsteeper gradients.

Surlyk et al. (1993) argued that a basinwide sealevel signal can be recognized in the large-scalecyclicity in marine synrift sediments of easternGreenland. The sea level changes may have been ofeustatic nature, but these workers emphasized thatthe changes more likely were a direct result of exten-sion. Should this hold true, it would argue strongly infavor of near-synchroneity of basinwide extension.

Cases of spatially variable accommodation resultingfrom Pleistocene glacio-eustasy in extensional set-tings in Greece have been discussed by Gawthorpeet al. (1994). They noted that close to the center ofthe normal faults, the effects of glacioeustatic fallsmay be canceled by high rates of hanging-wall subsi-dence. Away from the fault zones and toward thefault tips, however, relative sea level variations even-tually would be dominated by eustasy.


Prerift drainagerift margin

Synrift drainagerift margin and

interior

Rift-marginalAntecedent

axialRift-marginalAntecedenttransverse

transferzoneRift-marginAntecedenttransverse

Rift-margin or interior drainage

major incision andriver capture

Antecedentaxial

(backstepping)

Consequent-Juvenilehanging wall

Antecedenttransverse

transferzone

Rift-m

argindrainage

Major sediment entrypoint at cross-cutting fault

Rift-interior drainage

Consequent-Juvenilefault-scarp & hanging wallR

ift-marginald

rain

age

Antecedent transverse(major incision)

Redirected(axially)

Large footwall-sourced system(antecedent orconsequent)

ConsequentReversed

(A)

(B)

Figure 6Nomenclature, style,and development of drainagetypes. (A) Prerift to earliest synriftdrainage. (B) Synrift drainage.


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Sediment Supply and Drainage Development

In rift basins, extensional faulting exerts theprime control on morphology and topography, andhence on drainage (e.g., Jackson and Leeder, 1993;Leeder and Jackson, 1993). Additional factors thatinfluence the temporal and spatial variation in sedi-ment supply and drainage development or estab-

lishment include the following items.(1) Climate, which includes factors such as tem-perature, wind levels, and rainfall, and their season-al variations and fluctuations, exerts a major influ-ence on vegetation and biotype development,weathering processes, sediment supply, deposition-al environment, and the resultant lithofacies. Forlarge, elongate rifts (e.g., East African rift, RedSeaGulf of Suez rift, Rio Grande rift), climate mayvary considerably along the length of the rift basin.

(2) The distance of hinterland area to successivesubbasins determines the extent to which individu-al half grabens receive sediment from a rift-margindrainage or are dependent on supply from local

(fault-block) drainage alone.(3) The size of drainage catchment, which fol-

lows Hacks Law (Hack, 1957), depends largely onthe nature of the drainage itself, whether transverse(at high angle to basin-bounding faults), axial (atlow angle or paralleling the basin-bounding faults),antecedent or consequent, with consequentdrainage, including reversed, captured, or juveniledrainage (Figure 6). The dimensions and nature ofthe local (i.e., consequent or fault block) drainagecatchments depend largely on fault geometry andspacing. In general, the hanging-wall catchmentsare larger than those on the footwall, resulting ingreater sediment-yield potential for hanging-wallsuccessions.

(4) The half-graben morphology itself, especial-ly the gradient and relief of its slopes, the pres-ence, spacing, segmentation, and orientation ofany intrabasinal structures, and the presence oftransfer zones or crosscutting faults, influencessediment supply. The steeper fault-scarp slopesresult in a more rapid response and evolution ofthe footwall drainage systems. Spatial variations inresponse time between catchment generation andthe development of drainage systems and theircharacteristic transport, dispersal, and deposition-al processes would be expected. Moreover, seg-mentation of the footwall may shift the local fault-block drainage divide away from the originalbasin-bounding master fault (Figure 3). These pro-cesses, either on their own or in concert, result inan increase of the footwall catchment area at theexpense of the hanging-wall complement. In addi-tion, note that fault-scarp successions form mainlyas a response to fault-related subsidence along anddrainage development on the footwall of the

basin-bounding master fault. Hanging-wall succsions, however, record changing conditionsthe hanging wall, usually reflecting activity aon the basin-bounding master fault of the neiboring half graben. Accommodation zonestransfer faults determine the level to which invidual subbasins are filled before sediment trafer to the neighboring subbasin can oc

(Lambiase, 1990). The presence of hinterlandbasinward-dipping half-graben boundary fauwould be a major control on the progradatiopattern and sediment architecture related to trverse, rift-margin depositional systems

(5) The prerift substrate (sediment cover arock types) clearly influence the response tisediment-yield potential, and caliber of any lointrabasinal sediment source. Also importanwhether the rift infill contains volcanics.

Sediment flux in relation to drainage develment in rift basins has been discussed extensivRift subbasins receive their sediment from prodation of hinterland-derived, rift-margin deposit

al systems or from rift-interior sources.Rift-margin depositional systems usually rep

sent antecedent drainage systems, and mayeither axial or transverse (Figure 6). Axial systecommonly consist of relatively large fluvio-delsystems, which drain tilted subbasins upstreTransverse systems consist either of relatively smantecedent or of consequent hanging-wall and fowall fans, fan-deltas, deltas, and shorelines (Leeand Gawthorpe, 1987; Leeder and Jackson, 19Seger and Alexander, 1993; Gawthorpe et 1994), or large, antecedent fluvio-deltaic and shline systems traversing broad, submerged, platmal areas that border the rift basin proper (eHellem et al., 1986; Ravns and Bondevik, 1997

Rift-interior sources include footwall shopeninsulas, and islands (Figure 6B). Depositiosystems range from antecedent to consequhanging-wall and fault-scarp talus (collapse) conslope aprons, fan-deltas, and deltas, with the latwo possibly feeding submarine turbidite lobfans, or shorelines. Only consequent drainwould develop on footwall peninsulas or islaformed during rotation of a previously submerterrain. Upon reaching downdip (basinal) settinsediment delivered from the transverse systemay spread out laterally (Feretinos et al., 1988become redirected axially down the regional paslope or toward the center of individual, closubbasins (Papatheodorou and Ferentinos, 1993

Major sediment entry points form at the outleantecedent drainage systems (Leeder et al., 19Nelson et al., 1992), transfer zones (relay ram(Surlyk, 1977; Surlyk et al., 1981; Leeder Gawthorpe, 1987; Morley et al., 1990; Gawthoand Hurst, 1993; Morley, 1995), crosscutting fa

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(Cherry, 1993) or basinward of easily erodable sub-strate (unconsolidated sediments, etc.) (Leeder etal., 1991) (Figure 6B). To maintain the course oftransverse antecedent river systems during rotation-al faulting, high downcutting potential is impliedwith erosion rates larger than or equal to uplift rates,leading to river incision. Steep fault-scarp slopes maylack pronounced entry points; instead, sediment

may derive from a linear feeder systems or a series ofsmaller entry points (e.g., Surlyk, 1978).During a rift phase, subbasins located nearest

the hinterlands tend to trap most of the coarsematerial delivered by the rift-margin depositionalsystems. Subbasins located far from the hinterlandsreceive their sediment predominantly from rift-interior sources. When considered in a simplified,schematic manner, each rift phase is characterizedby fault-block rotation superimposed on the back-ground subsidence (Sb in Figure 2) (Barr 1987a, b;Jackson and White, 1989). If the prerotation eleva-tion of the terrain is equal to Sb, the volume of thefault block that is above sea level at the end of fault-

ing is equal to the adjacent submer ged sinks(assuming no along-strike variations). Decreasingthe prerotation elevation would diminish thesediment-yield potential of the fault blocks.Consequently, stretching of a terrain that was origi-nally close to or below sea level results in theformation of intrabasinal sediment sources that areof limited size compared to the volume of the adja-cent basinal areas (sinks). This scenario results inthe sediment-underfilled or clay-prone synrift signa-ture commonly observed in marine rift basins.Notably, if the prerotation water depth is great, syn-rotational footwall uplift would not be sufficient toproduce any subaerial emergence, and the conse-quent sediment-fill processes would be entirelysubmarine (Figure 4C).

Sediment-underfilled and sediment-starved halfgrabens are dependent on sediment supply fromother, distant sources to be completely filled, mostlikely the rift-margin hinterland. The subbasinsbecome filled in a successive manner, the rate ofprogradation of the rift-margin shorelines deter-mined by sediment supply, syndepositional accom-modation variations, depositional foundation(Helland-Hansen and Gjelberg, 1994), and waterdepth to accommodation zones or transfer faults(Lambiase, 1990). Other infill patterns are present incarbonate- and evaporite-prone rift basins. A discus-sion of such infill types is, however, beyond thescope of this study.

From these considerations, it is clear that theprerotation sea level stand and the size of and dis-tance to the hinterland are major controls on sedi-ment supply variations in rift basins (assuming simi-lar climate and substrate lithologies). The distanceto and sediment-yield potential of the hinterland

determines whether sediment from this source willbe delivered to any subbasin, and hence commonlyis the prime control on the degree to which any sub-basin will be sediment overfilled, sediment bal-anced, sediment underfilled, or sediment starved.Prerotation bathymetry determines the size of fault-block catchment areas and influences the timing ofsupply from such sources. These interrelated factors

are discussed further in following sections.

Influence of Prerift Bathymetry andBathymetry Development

Rift basins commonly consist of a series ofregionally tilted, elongate half grabens that tend tohave a lower basin floor elevation toward the axisof the rift basin (Figure 2). This tilting suggeststhat a continuum from essentially nonmarine orpartly submerged through mixed nonmarine/marine or partly emergent to completely sub-merged marine subbasins both across and along

the length of the rifted terrain may exist. Hence,one should expect a significant variability insediment-infill pattern.

At the landward end of marginal half grabens,marine incursions are likely in downdip areas dur-ing the rift climax only, with other substages char-acterized by erosion and nonmarine deposition(see Frostick and Reid, 1987; Leeder and Alexander,1987; Hamblin and Rust, 1989; and Lambiase,1990, among others, for discussions on the differ-ent nonmarine, synrift depositional models).

Subbasins with much of their basin floor close tosea level (Figure 4A) experience rapid drowning oftheir low-gradient downdip reaches and retreat ofthe rift-margin depositional systems (both axial andtransverse) during the early synrift stages.Rejuvenation of the rift topography may promoteprogradation of any local, hanging-wall or footwalltransverse system. However, this is counterbal-anced and, with time, suppressed by the increasingoverall subsidence, which results in further con-finement of coarse siliciclastics to near the bound-ary faults and to high up on the hanging wall.Marine influence in the succession reaches a maxi-mum during the rift climax stage. Progradation ofthe shoreline may occur locally at the mouths ofmajor sediment conduits. Increasing rates of ero-sion in the newly created tectonic uplands alongwith establishment of antecedent and consequentdrainage systems eventually result in renewedprogradation of the depositional systems duringthe late climax and late synrift substages. Providedthere is sufficient sediment available, individualsubbasins may be completely infilled during thelate synrift and tectonic quiescent or earlypostrift stages.



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If the fault-block emergence point (and henceany prerotational exposed areas) is located abovethe fulcrum (as in partly submerged subbasins),footwall uplift during the succeeding rift phase islikely to increase the width of any preexisting foot-wall peninsulas or islands, resulting in a progressiveincrease in the sediment-yield potential (Figure 4B).Any transverse drainage that existed during the pre-

ceding (prerift or tectonic quiescent) stage wouldhave a rapid and enhanced sedimentary responseto the new tectonic regime, expressed as prograda-tion and offlap of sedimentary lobes down thehanging wall. Peak downdip progradation dependson the uplift rate vs. the denudation rate, andoccurs either during or, more likely, sometime afterthe rift climax. Afterward, the hanging-wall shore-line retreats sourceward, that is, up the hangingwall itself. This response is clearly at variance withthe case described where the prerift shoreline zonewas located below the fulcrum.

Extension of a completely submerged terrain maytrigger gravitational collapse if there was some inher-

ited relief prior to the initiation of the rift phase. Ifthe prerotational water depth was relatively shallow(that is, if the background subsidence and the initialwater depth together were less than the footwalluplift), enhanced faulting and associated footwalluplift eventually would lead to the formation of foot-wall peninsulas or islands, providing a rift-interiorsource area (Dahl and Solli, 1993; Roberts et al.,1993a; Solli, 1995; Ravns and Bondevik, 1997).Subaerial erosion and the establishment of newlyformed, consequent, hanging-wall and footwalldrainage would develop some time after the updipreaches of the fault block have become emergent. Ifthis occurs during the early synrift stage, continuedfootwall uplift would lead to widening of the newlycreated footwall peninsulas or islands during the cli-max and early late-rift stages (Roberts et al., 1993a),accompanied by an increase in sediment supply andsustained downdip progradation during the climaxand early late synrotational stages (Ravns andBondevik, 1997).

If, however, footwall peninsulas or islands wereformed late in the rotational tilt stage, the exposedareas would be fairly narrow and, most likely, capa-ble of producing only limited amounts of sediment.Moreover, these limited catchment areas probablywould be surrounded by narrow and fairly steepshorelines. Hence, no pronounced shorelineprogradation is expected to occur.

The fault-scarp systems may show either a simplebackstepping or a forestepping-to-backsteppingstacking pattern. In the latter motif, the lowerforestepping part of the scarp represents theresponse to the newly created footwall uplands inthe early synrift stage. As increasing basinal subsi-dence rate outpaces the sediment supply, continued

retreat of any fault-scarp system would oc(e.g., Surlyk, 1989).

If the prerotational water depth is great (if sum of the background subsidence and the inwater depth together is larger than the footwuplift), footwall uplift cannot produce emergareas (Figure 4C). Instead, submarine shoals mform in the footwall of the boundary faults. S

subbasins may be draped by a veneer of hemipeic mudstones across intrabasinal highs, possiintercalated with submarine resedimentation pructs in depositional sinks (Underhill, 1991a,Deposition from the sediment gravity flows is prably initiated by gravitational instability along flank s of adjacent, intrabasinal highs, and deposit may in some cases reach considerathicknesses (Frseth et al., 1995). Sediment inbility may be tectonically triggered, and the retant deposits may increase in abundance, possforming the bulk of the accumulated sediment ding periods of high earthquake activity (the climstage). Updip areas may show considerable s

ment starvation, in particular where they are sjected to current or wave agitation. If sufficienshallow (i.e., usually above wave base), footwsubmarine erosional unconformities may form result of persistent wave or current erosion.

FACIES TRACTS

Marine synrift successions can be divided itwo main types: (1) nonmarine, paralic, and shallmarine infills (Figure 7A); and (2) deep-marinfills (Figure 7B). Deep-marine infills may inclshallow-marine deposits peripheral to emergfootwall highs and on submarine shoals. Tfacies tracts present in the shallow and demarine basin types outlined in the following stions should be regarded only as idealized modand do not take account of all possible facies aciations present in rift basins. Varying sedimyield and caliber, in particular, affect the variaty and architecture of the facies tracts (Orton Reading, 1993; Reading and Richards, 1994), asthe prevalent basinal regime, whether stowave, tidal, or f luvial influenced.

Shallow-Marine Facies Tracts

The shallow-marine facies tracts derived frthe hanging wall, the axial system, and the fscarp vary greatly (Figure 7A). The stacking tern of the resultant sediment geometries depeon the interplay of accommodation developmand type of accommodation in relation to sement supply.

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Hanging-wall dip-slope facies tracts includehanging-wall shorelines and associated backbarrierenvironments, deltas, and fan deltas (Figures 7A, 8),with fan deltas commonly represented by shallow-water, Hjul st rm type s (Col el la , 1988 a, 19 94 ;Massari and Colella, 1988; Gawthorpe and Colella,1990) and their offshore correlatives. Where back-barrier environments are volumetrically well rep-resented in the infill, characteristic features of the

transgressive segments are (1) unusually steep,seaward-dipping ravinement surfaces; (2) anunusually thick carpet of (transgressive) sandabove the ravinement surface; and (3) thick wedgesof backbarrier deposits, which also accumulate dur-ing the transgression, although below the ravine-ment surface. Onlap up the steep hanging-wallslopes tends to enhance the thickness of transgres-sive deposits. Figure 9 illustrates these points from


(B)

(A)

12

35

64

Regressive

Transgressive Tid

ali

nflu

enc

e

1

1 23

3

3

12

Figure 7Schematic, idealized presentations of synrift sediment infill models for (A) mixed nonmarine andshallow-marine rift basins and (B) deep-marine rift basins. The various styles of shallow-marine infills are basedon data from the Tarbert, Hugin, and Ling formations, axial parts of the Viking Graben, and the (point 1 in A) Ose-

berg, Hild, and Sleipner fields, and along the (point 2 in A) Alwyn, Brent, and Statfjord fields (see also Johannessenet al., 1995), and the (point 3 in A) Heather Formation and its intercalated sandstones on the rift-margin HordaPlatform, Oseberg and Brage fields. The deep-marine variant is based on data from the (point 1 in B) Brae Forma-tion, southern Viking Graben, (point 2 in B) Draupne Formation, Visund field, (point 3 in B) Munin sandstonemember, Statfjord North field, (point 4 in B) Kimmeridge Clay Formation and intercalated Ptarmigan and Magnussandstone members, Penguin half graben, (point 5 in B) Lusitanian Basin hanging-wall succession, and (point 6 inB) Draupne Formation, Oseberg, Brage, and Veslefrikk fields.


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(A)

(B)

Figure8

Correlationpanel

oftheVikingGroupinthe

Oseberg-Brageareaofthe

northernNorthSea(modi-

fiedfromRavnsandBonde-

vik,

1997).(A)Duringthe

Bathonian

Kimmeridgian,

threephasesoftiltin

g(tothe

east)andgeneralsubsidence

produced

three,

stacked

regressive-transg

ressive

sequences.(B)Thetransgres-

sivesegmentsforme

dduring

periodswithhightiltrates

andarecharacter

izedby

local,transversefault-scarp

andhanging-wallshoreline

andfandeltas.


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3/15-4

29/6-1

30/7-8

4

3

5

2

1

H

ildarea

S.Alwynarea

S

E

NW

1b

regressive

1atransgressive

Barrier&

shoreface

Coal-bearing

lagoon/deltaplain

Transgressive

sand

Regressive

Transgressiv

e

Ravinementsurface

Sequenceboundary

Maximumf

loodingsurface

GR

GR

GR

25m

Figure9

CorrelationpaneloftheTarbertFormationin

theHild-SouthAlwynareaofthenorthernNorthSea(modifiedfromR

n-

ningandSteel,1987).Fivesequences(numbered15

)de

velopedbyslighttiltingepisodesduringtheBajocianBathonian.

Note

the

seawardpartitionin

gofsedimentduringregressiona

ndlandwardpartitioning(storage

)duringtransgression.

Fartherw

est,

marineclaystonesar

epresentatthemaximumflooding

surfacelevels.


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Middle Jurassic (Brent Group) fault blocks in thenorthern North Sea, where the destruction ofshoreline barriers by steep-trajectory transgressionrepeatedly resulted in sediment being partitionedboth landward (backbarrier) and seaward (lowershoreface).

The footwall-derived fault-scarp facies tract com-monly consists of isolated to coalesced fan deltas

with interfan areas dominated by shoreline-paralicto shallow-marine environments (Figure 8) (e.g.,Surlyk, 1978, 1989; Leeder et al., 1991). Fan-deltatypes may show an evolutionary pattern fromshallow-water shelf type (Hjulstrm-type) throughshoal-water mouth-bar types (Postma and Drinia,1993) to Gilbert-type deeper water, gravitationallymodified fan deltas (Figure 7A) (Colella, 1988a, b,1994). Waning fault-related subsidence during the

late synrift stage may be accompanied by a chafrom an aggradational to a progradational stackpattern (Colella, 1994), possibly associated witchange from an alluvial fan to a braid plainthe feeder system (Gawthorpe et al., 199Alternatively, the late synrift substage is characized by renewed shoreline deposition.

Axial facies tracts include both fluvio-deltaic

shoreline systems (Figure 10). High constructdeltas (Orton, 1988) commonly form when basenergy (waves and tides) is low, whereas coaplains fronted by parallel beach ridges result frhigher wave energy. However, contrasting respoes to the basinal marine regime, and the devement of reflective or dissipative shorelines, are adependent on the sediment supply and the caliof the sediment delivered to the delta front (Ort

Ravns and Steel

12

9

30/9-8

GR

N

S

50

40

30

20

10

0

30/9-10

GR

30/9-3

GR

13

Starvedouter shelf

Inner shelf-lower shore

Backbarrietidallyinfluencedembaymen

Lowerdeltaplain

(m)

11

10

10

Mar ine muds tone outer - midd le shel f

Marine sandy siltstones middle - inner shelf

Marine sandstones

shoreface (wave dom.)

shoreface (mouth-bar / tide dom.)mouth bar or tidal/fluvial channels

Lower delta plaindeposits

(Tarbert Formation)

(Ness Formation)

Erosional unconformity

Time lines w/palynologycontrol point

Transgressive

Regressive

Figure 10Rift-axis parallel correlation panel of the Tarbert Formation in the Oseberg area, northern North Varying rates of extension produced four landward backstepping, regressive-transgressive shoreline prisms. regressive part formed during periods with lower tilt rates, whereas the transgressive part formed during peri

with higher tilt rates. Note the development of thick transgressive packages that show a higher degree of marinfluence upward in the succession, both individually and overall.


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Figure 11(A) Schematic basin-fill stratigraphy of the OxfordianKimmeridgian synrift succession of the Lusitani-an Basin, western Portugal. (B) Detailed architecture of synrift hanging-wall succession with submarine channelfill, disturbed slope sediments, and prograding Gilbert-type fan-delta sediments. Height of section is approximately150 m. (From Ravns et al., in press b.)


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Ravns and Steel

Figure12

Strike-ori

entedpaneloftheBraeFormationsubparalleltothemasterfaultalong

thewestsideoftheSouthVikingGrabenlessthan1km

outfromthisfault(fromTheriaultandSteel,1995).Thepanelshowsthehighlyvariablesedimentaryarchitectureofafootwall-derivedconglomerat-

icandsandysubmarinefansuccession.


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1988; Orton and Reading, 1993). The fluvio-deltaicsystems, whether originally transverse or axial, mayfeed basinal sediment gravity flows (Cockings etal., 1992), with these flows turning axial uponreaching downdip settings (Figure 7A).

Individual half grabens commonly form as scoop-shaped depressions along master fault zones. Anydowndip axial fluvio-deltaic system thus wouldprograde downslope from the sourceward end ofthe basin, but upslope basinward. The stacking pat-tern of the f luvio-deltaic system would vary as aresponse to such along-strike variations in thedepositional foundation.

The basinal marine regime may vary temporally,reflecting enhanced or subdued basin topographydeveloped in response to varying styles and rates ofextensional faulting. Ravns and Bondevik (1997)

suggested that tidal processes were enhanced by thedissection of rifted terrains into a series of elongatesubbasins (i.e., structurally controlled embaymentsor estuaries) during the rift phases. The funnel-shaped morphology of many mixed nonmarine/marine rift basins (e.g., Surlyk and Clemmensen,1983) probably provides ideal conditions for theamplification of originally weak tidal currents. Anextreme case of tidal influence strongly enhanced by

the weak rotation of narrow fault blocks in theMiddle Jurassic of the Hebrides (Mellere and Steel,1996), resulted in thick tidal deltaic and estuarinedeposits, possibly in tide-swept straits, where faultblocks were less than 8 km wide. The relative influ-ence of tides vs. waves probably is determined bythe water depth of individual subbasins (Ravns etal., in press a). Tidal currents may be particularlystrong over shallow accommodation zones, with theshallow water depth combined with a narrow sea-way accelerating the ti da l f low (Co ll ier andThompson, 1991).

In contrast, larger wavefetch can result in a pre-dominance of wave- and storm-influenced sedi-ments. The latter are probably more common dur-ing intervening tectonic quiescent stages whenfootwall islands are completely submerged, and a

larger, more open-marine embayment occupies therift basin (Ravns and Bondevik, 1997).

Deep-Marine Facies Tracts

Deep-marine facies tracts (Figure 7B) also varyspatially across the hanging wall, the axial, and thefault-scarp areas of half grabens.


Figure 13(A) Schematic dip section of a fault-scarp succession composed of sediments derived from a collapsedfootwall consisting of poorly lithified prerift lithologies. (B) Facies and stacking pattern of the fault-scarp wedge.(From Ravns and Steel, 1997.)

Base D

raupne Formation

W E

CU= coarseningupward

FU= fining upward

1

2

Cretaceous

CUFU

FU

FU

CUFU

FU

FU

CUFU

CUFU

FU

Heather

Formation

DraupneFormation

MINOR MAJOR

Grainsize

Increasingcongl./sand + shaleratio

Siltst./mudst.

Sandstone

Conglomerate

(A) (B) ?

Unconformity

Submarine unconf.

or condensed interval

Draupne clasticwedge

Major sequencesin well

2

1

150m

?

~5 - 6 km

BaseCretaceous

Visund FaultZone

Visun

dFaultB

lockwith

older

Jurassic

&Tria

ssic

1


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When updip areas of individual fault blocks aresubaerially exposed, narrow shoreline lithosomesform around the footwall islands (Surlyk, 1978,1984, 1989; Dahl and Solli, 1993). The steeperdepositional slopes of the deep-marine settingusually result in narrow facies belts, in more rapidtransitions between the individual (proximal-to-distal) facies tracts, and in the scarcity of either

alluvial or shoreline deposits (due to cannibaliza-tion into downdip sites). Shoreline and nonma-rine strata may be important only during the laterift and subsequent stages in subbasins becomingcompletely filled (Gury et al., 1986). Shorelinematerial may be transported to basinal settingsthrough localized channelized features, whereasthe remainder of the medial hanging-wall succes-sion is mud prone overall. Progradation of theshallow-marine and nonmarine feeder systemsinto deeper waters down on the hanging wallduring the late synrift stage may promote the for-mation of a shelf ramp or, notably, Gilbert-typedeltas in this setting (Figure 11). On completely

submerged fault blocks, the hanging-wall faciestracts are represented by mudstone blankets andgravitationally resedimented deposits.

The footwall-derived, fault-scarp facies tracts(Figures 12, 13) commonly consist of isolated col-lapse breccias, talus aprons, slope aprons, and gravi-tationally modified, deep-water cone deltas, fan-deltas, delta-front slope aprons, and canyon-fansystems (Surlyk, 1978, 1984, 1989; Ferentinos et al.,1988; Dart et al., 1994). Submarine aprons and fandeltas may coalesce to form shelf-slope systems withinterfan areas that are dominated by various shore-line to shallow-marine environments (Ferentinos etal., 1988; Leeder et al., 1991). One of the best docu-mented cases is that of the Brae oil field, which con-sists of resedimented deposits along the footwall ofthe southern Viking Graben. The great lateral vari-ability of the fan deposits along a short strike seg-ment near the fault scarp is shown in Figure 12.

The vertical organization of fault-scarp succes-sions characterized by a high supply of coarse silici-clastic sediments shows a development from initialsubmarine talus through a prograding slope apron toa coalescent fan delta to submarine fan, and then areturn to renewed mud-prone deposition, common-ly in the form of hemipelagic claystone blankets ormud-rich submarine fans (Surlyk, 1978, 1984; Turneret al., 1987; Cher ry, 1993; Surlyk et al., 1993).Provided sediment supply was sufficient and therewas complete filling of the original deep-marinebasin, an upward transition from slope apronsthrough aggradational fan-delta sequences to pro-gradational fan-delta sequences may develop(Gawthorpe et al., 1990).

The transverse depositional systems, in combina-tion with axial deltaic systems, may feed basinal

turbidite systems. Sediment gravity flows traing downslope at an angle to the basin axis mbe deflected, resulting in an overall along-atransport direction. Two different types of sment gravity flow elements may be present: bastepping and forestepping turbidite segme(Figure 14).

The backstepping turbidite segment appear

be dominated by a tendency to downdip shiftand an overall sourceward-stepping or foresteppto-backstepping stacking pattern. Sediment grty flow lobes tend to be isolated, embedded win thick basinal mudstones. Large-scale finiupward motifs reflect an upward decrease in volume of coarse-grained turbidites. The stackpattern is interpreted to be typical of early to synrift sediment gravity flow deposits (Surl1978, 1989; Surlyk et al., 1993; Ravns and St1997).

The forestepping turbidite segment is domied by sheetlike turbidites that show an ovebasinward-stepping stacking pattern. Initial

bidites infill basinal lows, whereas subsequflows onlap successively higher up on the haing-wall and fault-scarp slopes, in additionreaching progressively farther basinward. Tstacking pattern is interpreted to be charactetic for late synrift to tectonic quiescence or epostrift turbidites (Surlyk, 1989; Ravns and St1997). In some cases, the forestepping turbidare backed by an axially prograding submarislope to shallow-marine feeder system (Leinfeland Wilson, in press).

The development of a dissected, deep-marrift basin, with individual half grabens silledaccommodation zones, transfer faults, and fowall highs or shoals provides ideal conditionsthe development of a stratified water columpossibly leading to dysaerobic and anaerobic cditions in basinal reaches. Deep-marine rift basthus are favorable settings for the formatof high-quality marine source rocks, as exemfied by the Late Jurassicearliest CretaceoKimmeridge Clay and the Draupne and Speformations of the North SeaNorwegian Sea system (e.g., Dor, 1991).

SYNRIFT STRATIGRAPHY

Rift Basin Infill Types

Although prerift elevation, bathymetry, and bnal regime are important variables for the stratiphy of rift basins, as emphasized by the varied sigtures and stacking patterns expected in differbasinal settings, the factors of accommodation geration and sediment supply play the dominant r

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(Figure 15). Marine half grabens, whether shallow ordeep marine, may be grouped or classified accord-ing to the degree of balance attained between tec-tonically generated accommodation and sedimentsupply during infilling (Figure 15). Sediment-over-filled basins are characterized by an abundance ofsediment supply, complete infilling, and transfer ofexcess material to adjacent subbasins. Sediment-

balanced basins are typified by sediment supply(nearly) keeping pace with tectonic subsidence(when averaged over the entire synrift interval).Sediment-underfilled and sediment-starved basinsare characterized by a sediment supply that is toolow to keep pace with the tectonic subsidence, andhence by an inability to completely infill the accom-modation space. Note that the temporal variationthat an individual subbasin experiences as it sub-sides and evolves through a series of successive riftphases is also implicit in this classification.

The subdivision of rift basin infill types partly par-allels the morphological subdivision from marginalhalf grabens through partly submergent and partly

emergent to completely submerged rift basin types.This reflects the changing sediment yield of themain source areas for the different basin types. Rift-margin, sediment-overfilled, and sediment-balanced half grabens receive sediment from largehinterland areas. Rift-interior half grabens, in con-trast, located at distance from the rift-margin hinter-land, rely increasingly on supply from rift-interiorsources, thereby commonly showing a sediment-underfilled or sediment-starved synrift lithologymotif. Variation in sediment supply from rift-interiorand rift-margin sources is dependent on the size ofthe catchments, and how the catchments sedimentyield potential varies in time by the interaction ofrotational uplift/subsidence, basinwide subsidence,and eustatic sea level stand and its changes. As aresult, different infill patterns, and thus differentsynrift lithology motifs, result from the same tecton-ic scenario. Importantly, fault-block rotation resultin uplift and relative sea level fall in areas above thefault-block fulcrum, whereas areas downdip of thefulcrum are subject to continuous subsidence andconsequent relative sea level rise. For each rift basininfill type there are specific hanging-wall updip anddowndip signatures that reflect the varying accom-modation development and sediment supplythroughout the rift phase. Note that basinal succes-sions ultimately record the influence of hanging-wall, footwall, and axial drainage.

Early Synrift Signatures

On the hanging wall, the early synrift period ischaracterized by deposition of relatively coarse mate-rial derived either from a preexisting depositional

system (which may deliver its load simply into thenewly created deeper water) or from downslopeflushing of weathered detritus during the initial tilt-ing (Figure 15A). During this period, sediment sup-ply is generally not able to keep pace with theincreasing rate of accommodation, resulting inareal reduction of the drainage system and an over-all sourceward retreat of the depositional system.

Shoreface retreat during this transgression is likelyto contain the sediment yield from the shrinkingdrainage area within narrow backbarrier andcoastal plain zones. This scenario likely is commonin the early synrift stage, but assumes a partly sub-mergent fault block with its pretilt shoreline some-what downdip of the fulcrum (Figure 4A); how-ever, in cases where the fault block is nearlysubmerged (i.e., the pretilt shoreline lies updip ofthe fulcrum; Figure 4B), rotation may be accompa-nied by an enlargement of the drainage area duringthe rejuvenation of the rift topography. In general,this results in increasing erosion rates in the newlycreated uplands, possibly associated with incision

of the antecedent drainage systems or the form-ation of a consequent (reversed or juvenile)drainage. In this scenario, there may be significantproduction of sand on the updip reaches of thehanging wall. This material may form a forcedregressive shoreline or deltaic lithosome, or maybypass the shoreline areas, forming sandy sedi-ment gravity flows delivered directly to the basinalareas. The case shown in Figure 4B may result inthicker units of basal sandstone in the synrift suc-cession or in sandy units that coarsen upward frombasal mudstones.

Footwall uplift of submerged areas commonlyresults in blanketing of hemipelagic mudstone (thecoarser material being trapped in the evolving,adjacent lows), and the formation of condenseddeposits, which in areas updip of the fulcrumformed during a relative sea level fall (Ravns andBondevik, 1997; Ravns and Steel, 1997). Whencrestal areas become emergent, an evolutionarypath on the hanging wall similar to that describedis envisaged.

Hanging-wall areas adjacent to footwall scarpsexperience continuous subsidence during all of therifting, although the subsidence may vary both inrate and magnitude. In such downdip areas, thereis an initial retreat of any transverse and axial sys-tem during the early synrift period, probablyreflecting sediment supply being outpaced byincreasing fault-related subsidence, as has beenadvocated by Prosser (1993), Wignall and Pickering(1993), and Ravns and Bondevik (1997). However,this signature may apply only for sedimentary lobessourced through transfer zones (Wignall andPickering, 1993; Underhill, 1994) or from rift-margin depositional systems (Ravns and Bondevik,



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Ravns and Steel

Figure14

CorrelationpaneloftheKimmeridgeClaya

nditsintercalatedPtarmiganandM

agnussandstonemembersinthePenguinhalfgraben,

northernNorthSea.

Backsteppingturbiditesegmentsarepresentinthesedimentsdepositedduringperiodswithhightiltrates,whereasforestep-

pingturbiditesegmentsarepresentinthelowerofthetwopackagesdepositedduringperiodswithlowertiltrates.FromRavnsandSteel(1997).


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1997) and not for evolving footwall-derived, fault-scarp depositional systems. The footwall-derived,fault-scarp depositional systems, instead, may showan initial phase of progradation represented by abasal coarsening-upward motif (Surlyk, 1978;Frseth et al., 1995; Ravns and Steel, 1997).

Extension of submerged terrains often producesa footwall-derived fault-scarp succession with a

basal fine-grained layer believed to have beendeposited during the initial rotation. This basalmudstone reflects deprivation of coarse siliciclasticmaterial in the newly formed half graben prior tothe establishment of any transverse drainage. Theresponse to footwall emergence and the establish-ment of juvenile footwall drainage may be a phaseof initial progradation. Near the outlet of any preex-isting transverse drainage basal mudstone probablydoes not occur.

Rift Climax Signatures

In sediment-overfilled and sediment-balancedbasins, the rift climax stage is characterized bymaximum retreat of all of the depositional systems,resulting in the deposition of widespread mud-prone strata over most of the subbasins (Figure15B). Coarser material commonly is confined tothe subbasin margins. Local, sand-prone lobes,however, may prograde from major sediment con-duits. Any of these subbasin marginal systems mayfeed basinal sediment gravity flows.

In the sediment-underfilled and sediment-starvedsubbasins, the depositional systems are inferred tobe confined largely to the basin margins, as well(but see Surlyk, 1978); however, if these subbasinswere or iginally submerged or the prerotationalfootwall island was small and restricted to the areaabove the fault-block fulcrum (Figure 4B), this stagewould be characterized by continued tilt-climaxprogradation and offlap of sedimentary lobes downthe hanging wall (Roberts et al., 1993a; Ravns andBondevik, 1997; Ravns and Steel, 1997). Thesehanging-wall depositional systems already may havebeen established in the early rift stage. Note that theforced regressive shoreline developed above therotational fulcrum passes into normal regressiveshorelines below the fulcrum (Figure 4B), with theamount and rate of progradation determined by thesediment supply, the depositional foundation, andthe water depth of the receiving basin.

The fault-scarp successions commonly consist ofasymmetrical forestepping-backstepping, coarseningupward to fining upward, and backstepping fining-upward packages on various scales (Surlyk, 1978,1984; Surlyk et al., 1993; Wignall and Pickering,1993; Frseth et al., 1995; Theriault and Steel, 1995;Ravns and Steel, 1997), or a series of stacked

regressive-transgressive units (e.g., Dart et al., 1994).In general, any footwall sediment yield is likely to befar less than the space available in the adjacent sink.Hence, the large-scale stacking pattern (hundreds ofmeters) through the entire rift phase is likely toshow an overall fining-upward trend (Surlyk, 1978;Ravns et al., 1994). Intermediate and small-scalefining-upward motifs have been interpreted in terms

of submarine slope processes, such as filling ofscours or gullies, retrogressive flow slides, or deposi-tion from surging high- and low-density turbiditycurrents (Surlyk, 1978, 1984). Intermediate andsmaller scale coarsening-upward to fining-upwardmotifs appear to be developed in successions ofmud-rich or poorly consolidated materials (Frsethet al., 1995; Ravns and Steel, 1997), whereas inter-mediate and smaller scale fining-upward motifs aremore typical in sand-prone successions (Surlyk,1978, 1984; Turner et al., 1987; Cherry, 1993).

Frseth et al. (1995) suggested that the coarsening-upward to fining-upward character of the intermediate-scale motifs may relate to an initial phase of fan

lobe/apron progradation prior to a retreat andbackfilling phase. Successive cycles were postulat-ed to reflect rapid influx of sediment followed bymore gradual exhaustion of stored sediment fromthe adjacent footwall high; some of these eventswere thought to have been coupled to fault move-ments that caused sediment storage or instabilitythresholds to have been exceeded. Theriault andSteel (1995) argued for a strong seismic control ondecreasing pulses of sediment supply, with small-to intermediate-scale motifs reflecting temporalfluctuations in earthquake activity, and some indi-vidual conglomerate strata related to discrete, large-magnitude earthquake tremors.

Successions immediately adjacent to the faultscarp are less organized and commonly are arrangedin an overall aggradational stacking pattern (Cherry,1993; Lonergan and Schreiber, 1993). Fault-scarpsuccessions may show large variations along strike,with areas located immediately basinward of sedi-ment conduits receiving vast amounts of coarse sili-ciclastic material in contrast to interconduit areas,which are finer grained and relatively sedimentunderfilled, as illustrated from the Brae fault scarpalong the southern Viking Graben (Figure 12)(Theriault and Steel, 1995).

Late Synrift Signatures

The late synrift stage is characterized by waning orlow rates of extensional faulting (Figure 15C). If thereis sufficient sediment yield potential, sediment sup-ply eventually would equal or exceed the accommo-dation generation, allowing progradation of coarsesiliciclastic systems. In the deep-marine realm, this



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Ravns and Steel

Figure15

Schematic,generalizedblockdiagramssho

wingtheevolutionofshallowand

deepmarine,overfilled,sedimentbalanced,sediment

underfilled,andsedimentstarvedsubbasinsthrough(A)

theinitialandearlysynriftstages,(B)theriftclimaxstage,and(C)the

latesynriftandearly

(A)

(B)

(C)

(D)

Initialand

earlysynrift

Riftclimax

Latesynriftto

earlypostrift/

tectonic

quiescence

Overfilledbasins:

Sedimentba

lancedbasins:

Sedimentun

derfilledbasins:

Sedimentstarvedbasins:

Sh

allow

marine

Deep

marine

sediment

overfilled

sediment

balanced

sediment

underfilled

sediment

starved

sealevel

sealevel

shallowmarine

deepmarine

sediment

overfilled

sediment

balanced

sediment

underfilled

sediment

starved

sealevel

sealevel

shallowmarine

deepmarine

sediment

overfilled

sediment

balanced

sediment

underfilled

sediment

starved

sealevels

ealevel

shallowmarine

deepmarine

LEGEND

Alluvial

Shoreline

Slope

Basinal

Basinalg

ravityflowlobes

Marginal

(fan-apron)gravity

flowdepo

sits

Alluvialorerosion/nondeposition

Alluvial;lowerdeltaplain

Deltaicorfandeltaic

Shore

line

Drown

edshelfw/tidalridges

Slope

w/channelsandgullies

Basin

floor

Grave

lly

Sandy

Muddy

LEGEND

Sedimentgravity-flow

deposits(slope/aprons

andbasinfloor)


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stage is heralded by the income of commonly sharp-based, forestepping turbidites (Ravns and Steel,1997), which in the overfilled basin type are overlainby submarine-slope, shallow-marine, and alluvialdeposits (e.g., Leinfelder and Wilson, in press). In theshallow-marine environment, progradation of shore-line lithosomes produces a characteristic coarsening-upward profile (Figure 10).

Half grabens located at some distance from thehinterland areas commonly receive their sedimentonly from bordering footwall islands. Because theseusually are of limited size compared to the volumeof the adjacent basins, reduced uplift during thelate synrift stage coupled with high erosion ratesmay produce low-relief or peneplaned fault-blocksource areas, resulting, in turn, in reduced sedi-ment supply. Continuous high background subsi-dence rates in concert with limited or diminishingsediment supply therefore produce a characteristic,overall sourceward retreat of any local, transversedepositional system. Depending on the size of thefootwall islands, mudstone blanketing would occur

either during the late synrift, early postrift, or tec-tonic quiescence stage.

Rift-interior sources appear to be too small to pro-vide sufficient sediment to allow renewed prograda-tion of the fault-slope clastic wedge during the latesynrift stage. Late progradation of such systems, withan upward coarsening or shallowing of facies, howev-er, can occur in marginal half grabens or half grabensbordered by large footwall provenances (e.g., Surlyk,1978, 1984, 1989; Gawthorpe et al., 1990).

Correlation of Footwall, Hanging-Wall, andAxially Derived Successions

The relationships of fault-scarp signatures totheir contemporary hanging-wall signatures are stilluncertain. Synrift stratigraphic simulation suggeststhat peak hanging-wall progradation in someinstances may correlate with fault-scarp retrograda-tion. A more rapid or immediate yield of coarse sedi-ment from the footwall drainage to the new tectonicactivity would be expected due to the steeper fault-scarp slopes. Hanging-wall systems may lag behindtheir footwall counterparts, due to the gentler hang-ing-wall slopes, the tendency to drown parts of thehanging-wall catchment areas (in partly submergentbasins), or the prerift submerged state of the catch-ment areas (in partly emergent and completely sub-mergent basins). In addition to reflecting differentstyle of accommodation generation and varyingdrainage and sediment yield developments, hanging-wall and fault-scarp successions also may reflect thediachronism of tectonic movements of basin-bound-ing master faults of two adjacent half grabens. Theaxial systems are expected to have the longest

response time to the same tectonic event. Fault-relat-ed subsidence commonly results in flooding alongthe axis of the rift basins and across the rifted ter-rains. Hence, there commonly is a major time lagbetween the tectonic activity and the sedimentaryresponse from the axial drainage. We suggest thatsynchronous and in-phase development of high-order stacking patterns likely would not occur.

Tectonic Quiescence Signatures

The infill related to tectonic quiescent stages alsomay be wedge shaped, albeit with internal stratalsurfaces subparallel to the upper boundary of the

wedge. With sufficient sediment supply, this stage ischaracterized by continued progradation of the rift-margin depositional systems, resulting in completeinfilling of subbasins in a successive manner andsmoothing of the former dissected rift basinbathymetry. Degradation of source areas, resulting inreduced sediment caliber and supply, may inhibit

the filling of distal half grabens, which thus mayretain their sediment-underfilled or sediment-starvedstatus for a considerable period. Subsequent rifting

would result in renewed retreat of the rift-marginshoreline and prevent sediment supply to subbasinsdistant to the hinterlands. Such subbasins are charac-terized by dominantly fine-grained tectonic quies-cence packages of hemipelagic or pelagic origin.

Synrift Signatures Summarized

Although the three-dimensional variability ofthe sediment architecture within the subbasins islikely to be great (Figure 15), there have beenattempts to generalize about the vertical architec-tural signature likely to dominate. The three-fold(sandstone-claystone-sandstone) synrift lithosome sig-nature of Prosser (1993) (Figure 1B) is typical insediment-overfilled and sediment-balanced subbasins.The two-fold (conglomerate-sandstone-claystone) syn-rift motif (Figure 1A) of Surlyk (1978) is more charac-teristic of sediment-underfilled and sediment-starvedbasins. Extreme cases of sediment-starved basins aretypified by a thin (condensed) claystone fill, common-ly of good source rock quality. Note that in the thinclaystone fill, there is usually no pronounced wedgegeometry to the resultant synrift infill.

The vertical variation shown by the three-fold andtwo-fold synrift lithology motifs can be linked closelyto the evolving tectonic regime. In both motifs thebasal coarse clastics represent the rift initiation andearly synrift stage, whereas the mud-prone intervalsrepresent the climax of rifting. The sandstone cap-ping of the first motif is argued to represent the latesynrift stage. Waning rotational faulting, possibly



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Ravns and Steel

Figure16

Summary

ofthecontrollingfactorsonthedevelopmentofthree-fold,two-fold,orone-foldsynriftsuccessions.


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(A)

(B)

(C)

(D)

ROTATIONAL TILT PACKAGES ( high tilt rates )

TECTONIC QUIESCENE PACKAGES ( low tilt rates )

No scale implied

No scale implied

No scale implied(modified from Underhill, 1991a)

No scale implied

Figure 17Contrastingstacking patterns of aseries of synrift packagesseparated by tectonicquiescence fill asillustrated by deep-marinesynrift successions.Schematic stackingpattern of (A) half grabens

with abundant andcontinuous sedimentsupply throughout thetectonic phases,(B) half grabensdominated volumetrically

by synrotational infill,(C) half grabensdominated volumetrically

by tectonic quiescenceinfill, and (D)sediment-starvedhalf grabens (modifiedfrom Ravns and Steel,1997).


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coupled with increased sediment supply, allows forthe establishment of transverse and axial drainagesystems and the consequent progradation of coarseclastic systems, either shallow or deep marine, acrossthe basin (see also Prosser, 1993). The absence ofsandstone capping in the two-fold lithology motif issymptomatic of sediment-underfilled and sediment-starved rift basins. In these basins, low sediment sup-ply in relation to tectonic subsidence is attributedeither to high subsidence rates (i.e., in areas subject-ed to severe extension) or to limited catchment andsediment yield potential of the subbasins sourceareas (Ravns et al., 1994).

We have chosen to summarize the synrift infill sig-natures by illustrating the three main cases, as shownin Figure 16, where the initial bathymetry relative tothe position of the rotational fulcrum and the dis-tance to hinterland areas are key parameters, at leastfor the early and climax stage architecture. The syn-rift infill signatures shown in Figure 16 broadly corre-spond with sediment-balanced, sediment-underfilled,and sediment-starved cases, respectively. In addition,the timing, duration, and volume of sediment sup-plied during the specific stages of the rift cycle exertsome influence on the resultant lithology motif.Hence, a variety of subdivisions also exists for thethree-fold, two-fold, and one-fold synrift lithologymotifs, and for the sediment-overfilled, sediment-balanced, sediment-underfilled, and sediment-starvedrift basin types (Figure 16).

MULTIPLE SYNRIFT PHASES

Superposition of Synrift and QuiescencePackages

Repeated rift phases result in the vertical stackof two or more synrotational packages, intercalawith intervals showing characteristics more typ

of pre- or postrift successions; i.e., tectonic qucence packages. Figure 17 illustrates schematica range of possible multiple rotation cases baseddata from the Magnus (Figure 17A), Oseb(Figure 17B) and Inner Moray Firth (Figure 1fields. These cases show the development of eqly thick or unequally thick superimposed synand quiescence packages caused mainly by tectcally induced variability of sediment suppalthough factors of inherited bathymetry, rapiof rotation, and fault-block size are also likely toimportant. In either scenario, each motif, whetthree-fold or two-fold, shallow-marine or demarine, relates to a single rift phase.

The common evolutionary path observedmany of the northern Viking Graben half grabthat experienced recurrent rift phases is a develment from a shallow-marine to a deeper marbasin. This evolution was generally paralleled bchange from a sediment-balanced (or sedimeoverfilled) status to a sediment-underfilled staand sometimes to a sediment-starved status. Tscenario reflects the increasing amount of accmodation, decreasing sediment supply, and incring cumulative fault-block rotation as the evolved and the rifted terrain subsided. The ovesynrift signature

Marine Synrift Sedimentation Ravnas & Steel AAPG 1998

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