Marine Synrift Sedimentation Ravnas & Steel AAPG 1998

Architecture of Marine Rift-Basin Successions1 R. Ravnås2 and R. J. Steel3

ABSTRACT Marine rift basins represent a continuum ranging from mixed nonmarine/marine through shallow marine to deep marine, or from partly emergent through partly submergent to completely submergent basin types. These rift basin types have strongly variable synrift sedimentary architectures because of temporal changes in relative sea level, accommodation creation, and sediment supply throughout the rift cycle. Accommodation changes are controlled mainly by local basin-floor rotation, basinwide background subsidence, and, to a lesser degree, by eustatic changes. Sediment supply determines how much of the accommodation is filled and in what manner, and is controlled by the distance to the main hinterland areas, and the size and sedimentyield potential of any local fault-block source area. Marine siliciclastic synrift successions, whether dominantly shallow or deep marine in nature, are classified in terms of sediment supply as overfilled, balanced, underfilled, and starved. Sediment-overfilled and sediment-balanced infill types are characterized by a threefold sandstone-mudstonesandstone synrift sediment-infill motif; the sediment-underfilled type is represented by a two-fold conglomerate-sandstone-mudstone motif; and the sediment-starved type commonly is represented by a one-fold mudstone motif. The sequential development, linked depositional systems, and stratigraph©Copyright 1998. The American Association of Petroleum Geologists. All rights reserved. 1Manuscript received December 14, 1995; revised manuscript received December 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-4001 Stavanger, Norway. 3 University 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 comments by K. T. Biddle, C. K. Morley, T. H. Nilsen, and F. Surlyk on an earlier version of the manuscript sharpened the final product and are greatly appreciated. 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 of Norway 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 this paper refers to a geological feature that originates at the rift margin but may extend out into the basin.

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ic signatures of the early synrift, the rift climax, and the late synrift to early postrift stages vary significantly between these rift basin infill types, as do the tectonic significance (timing of initiation and duration) of stratal surfaces, such as footwall unconformities, nondepositional hiatuses, and marine condensed sections. The construction of the fourfold rift basin infill classification scheme provides a first basis and a strong tool for predicting the distribution and geometry of synrift reservoir and source rock types, despite the inherent variability of the marine synrift infills. INTRODUCTION Intracratonic rift basins form by the stretching and 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 settings, their sedimentary fill shows a great variability (e.g., Leeder, 1995). Ideas on synrift sediment architecture 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 recognizing the asymmetric geometry of fault-controlled sedimentary packages, it has long been realized that changing lithological signatures and stacking patterns can reflect variation in fault-related subsidence rate, assuming a variation in the balance between sediment supply and subsidence rate. Clayprone intervals (with time-equivalent conglomerates restricted to narrow belts adjacent to the fault scarps) in syntectonic successions are known to represent periods of rapid differential subsidence (Steel, 1988), whereas the intervals of coarse facies extending farthest from the fault scarp relate to periods of relative tectonic quiescence and minimal accommodation creation (Blair, 1987). Hamblin and Rust (1989) and Lambiase (1990) combined these ideas, proposing a threefold, fining- to coarsening-upward, sandstone-mudstonesandstone lithology motif for lacustrine synrift successions. The basal and capping sandstones were AAPG Bulletin, V. 82, No. 1 (January 1998), P. 110–146.

FU

CU

Early stage

Climax stage

Late stage

Postrift

Early stage: Increasing rate of fault-related subsidence. Widening of basin and rejuvenation of relief. Retreat of basin-margin attached depositional systems; establishment of locally sourced transverse sedimentary lobes. Stacking pattern dependent on sediment supply and caliber; commonly coarsening- to fining-upward due to initial resedimentation of weathered detritus. Tectonic subsidence > sediment supply.

Climax stage: High rate of fault-related subsidence. Fault-scarp and transfer zone fed fans and aprons, coarsening- or fining-upward motifs dependent on sediment supply and caliber. Isolated randomly stacked basinal lobes. Tectonic subsidence >> sediment supply

Late stage: Waning rate of fault-related subsidence. Fining-upward to massive claystones due to drowning of (most of) source areas, which cut off or reduced siliciclastic sediment supply. Potential source rocks in sediment-starved basins. Tectonic subsidence > sediment supply.

Postrift stage: Tectonic quiescence Massive claystones; deprivation of coarse siliciclastic material. Continuous sediment starvation results in deposition of condensed successions with good source rock potential, also on structural highs. Rare turbidite incursions as isolated basinal lobes.

CU

Late stage

Postrift

Early stage

FU

Climax stage

FU

Overall fining-upward to coarsening-upward rift signature.

Commonly several 100s or >1000s of meters

(B)

Early stage: Increasing rate of fault-related subsidence. Widening of basin and rejuvenation of relief. Retreat of basin-margin attached depositional systems; establishment of locally sourced transverse and axial sedimentary lobes. Overall aggradational to backstepping stacking pattern. Tectonic subsidence > sediment supply.

Climax stage: High rate of fault-related subsidence. Overall aggradational to backstepping stacking pattern. Hanging-wall progradation to backstepping around footwall islands. Isolated randomly stacked basinal lobes. Tectonic subsidence >> sediment supply.

Late stage: Waning rate of fault-related subsidence. Aggradational to progradational stacking pattern. Progradation of locally sourced transverse and axial sedimentary lobes; renewed progradation of basin-margin attached depositional systems. Tectonic subsidence < sediment supply.

Postrift stage: Tectonic quiescence. Fining upward due to degradation of source area.

Figure 1—Schematic log of an idealized, basinal synrift succession of (A) sediment-underfilled and sediment-starved rift basins, and (B) sedimentbalanced and sediment-overfilled rift basins.

Overall fining-upward rift signature.

Several 10s to 100s of meters

(A)

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Marine Rift-Basin Architecture

argued to reflect the early and late synrift stages, respectively, whereas the intervening mudstones were related to the climax of rifting. The architectural styles of f luvial and arid synrift sediment infill probably reflect the depositional system’s response to the temporal variation in the rift basins’ basin-floor and bounding fault-scarp slopes, rather then being related to any intrabasinal base level, such as local lake level. Nevertheless, a threefold sandstone-mudstone-sandstone synrift motif has also been suggested for fluvial rift basin infills (Mack and Seeger, 1990; see also Alexander and Leeder, 1987). In contrast to the architectural styles of nonmarine synrift successions, workers have made fewer attempts to synthesize data on mixed nonmarine/marine and marine synrift sediment architecture. Surlyk (1978, 1989) developed models for hanging-wall turbidite systems sourced from the adjacent footwall, and demonstrated a four-stage evolution of the turbidite systems with variable rates of sea level changes and extensional faulting. Furthermore, Surlyk’s (1978, 1989) work resulted in a two-fold lithosome model for marine synrift successions (Figure 1A). Subsequently, Prosser (1993) proposed a complementary, three-fold lithosome model for terminal marine rift basins, analogous to that suggested for lacustrine rift basins (Figure 1B). These two models are here postulated to form type members in a series of architectural patterns present in marine synrift successions. Gawthorpe et al. (1994) examined the combined effects that varying tectonic styles and subsidence rates, as well as glacioeustatic sea level fluctuations, had on synrift sequence development in an extensional back-arc setting subjected to regional uplift. Our understanding of rift basin formation and synrift sediment architecture has been aided by numerical and analog structural and stratigraphic modeling. Such studies have highlighted the structural parameters (Barr, 1987a, b, 1991; Kusznir et al., 1991; Roberts et al., 1993a) and the onlap or offlap patterns expected to characterize the resultant synrift stratigraphy (Schlische and Olsen, 1990; Schlische, 1991; Hardy, 1993; Roberts et al., 1993a; Waltham et al., 1993). The aim of this paper is to further develop our understanding of synrift sediment architecture of marine half-graben type rift basins, building on the earlier studies of Surlyk (1978, 1989), Prosser (1993), and Gawthorpe et al. (1994). Our overall contribution is a synthesis of the controls on siliciclastic sequence architecture, with an emphasis on how variable this architecture can be, in regionally extended and subsiding marine terrains formed during nonglacial periods. In contrast to previous studies, our database from the northern North Sea

and western Portugal is wide ranging and based on field studies, three-dimensional seismic coverage, and abundant well data. An additional feature we wish to emphasize is that the Middle Jurassic–Early Cretaceous synrift successions studied herein have been derived from thick, poorly consolidated sedimentary terrains. Drainage development and sedimentary response to tilting and uplift thus are likely to have been more rapid than for basement terrains, a feature at variance with assumptions in one of the most recent models (i.e., Prosser, 1993). In Prosser’s model, the sedimentary response is assumed to be delayed with respect to an extremely rapid fault rate, leaving the rift climax relatively starved of sediment. We also suggest that the role of sediment supply, especially as enacted through prerotational bathymetry and distance to main hinterland areas, has been relatively neglected in previous studies. Likewise, the importance of the shoreline position on the hanging wall with respect to location of the rotational fulcrum for the development of sediment supply has been underestimated and not differentiated (see, however, Hardy, 1993; Roberts et al., 1993a). RIFT BASIN DEVELOPMENT AND ARCHITECTURE Synrift Subsidence Synrift extension involves upper crust extension accommodated by faulting and stretching of lower crust and lithospheric mantle by plastic deformation (Kusznir et al., 1991; Kusznir and Ziegler, 1992). Synrift subsidence is a state resultant from the balance of the following items. (1) Isostatic adjustment of the crust to mechanical stretching of the lithosphere leads to the subsidence of the thinned crust (McKenzie, 1978; Cochran, 1983). This is accompanied by flexural uplift or by downwarping of the rift zone, depending on the depth of the lithospheric necking level (Kooi et al., 1992; Ziegler, 1992). (2) Upwelling of the aesthenosphere into the space created by mechanical stretching of the lithosphere, which, together with thermal upward displacement of the aesthenosphere-lithosphere boundary, causes uplift of the rift zone (Turcotte and Emmermann, 1983; Ziegler, 1992). According to Ziegler (1992), the structural style of rift basin development is influenced by the thickness and thermal state of the crust and subcrustal lithosphere at the onset of rifting; the amount of crustal extension and width over which it is distributed (β factor); the mode of extension, whether orthogonal or oblique; and the lithological composition of prerift and synrift deposits. In addition, any

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(A) SEDIMENT SUPPLY MARGINAL & NONMARINE

(BATHYMETRY + SOURCE AREA SIZE / LITHOLOGY) INCREASING SUBMERGENCE

Hinterland source areas (large)

DEEP-WATER SUBBASINS

Intrabasinal source areas (small)

SL

INCREASING DISTANCE FROM LARGE SOURCE AREAS

(B) ACCOMMODATION

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Figure 2—Controls on rift-fill architecture. (A) Sediment supply in relation to half-graben location, bathymetry, source area size, and lithology; (B) accommodation development, synrift tilting, and subsidence (Sb = background subsidence); and (C) accommodation development in time and space, and architecture of the synrift sediment fill.

(ROTATION + BACKGROUND SUBSIDENCE + FAULT SPACING + EUSTASY) rotation

rotational pivot

fulcrum

prerift surface background Sb subsidence

(zero net vertical movement)

L

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Rift climax

Early synrift

Late synrift / tectonic quiescence

Spatial variations along fault lengths

preexisting basement grain commonly exerts a strong control on the preferential development and orientation of synrift structures (see review in Morley, 1995). Changing crustal conditions, especially temperature, can cause temporal changes in rift geometry during extension (Kusznir and Park, 1987), such as from low- to high-angle fault geometries (Morley, 1989; Bosworth, 1992). General Features of Rift Basins An analysis of rifted terrains shows clearly that the fundamental morphological element is the fault block or half graben formed within the hanging wall of major basin-bounding master faults (Figures 2, 3) (Bosworth, 1985; Rosendahl et al., 1986; Morley, 1995). The rift basin may consist of a single half graben or a series of half grabens, commonly

with pronounced variation in morphology and topography along strike. The width of classic linear rift valleys ranges from 50 to 100 km. This width contrasts with regionally extended rift provinces, which may exceed 1000 km in width. Half-graben subbasins commonly have widths of 10–35 km, although widths of more than 50 km may occur. Shallow detachments commonly produce shorter wavelength fault blocks compared to fault blocks delineated by basement faults (e.g., Rattey and Hayward, 1993). The length of recent basin-bounding master faults is typically some 35–60 km, with individual fault strands at surface rarely exceeding 15–20 km (Jackson and White, 1989). The basin-bounding master faults tend to overlap and switch polarity along the rift zone, resulting in the formation of transfer and accommodation zones (Bosworth, 1985; Rosendahl, 1987; Morley et al., 1990).

Figure 3—(A) Half-graben morphology and rift-interior sediment dispersal pattern. (B) Dip section through rifted terrains showing accommodation, intrabasinal drainage, and sediment dispersal direction in (I) a solitary half graben, (II) a series of half grabens with simple basin-bounding master faults, and (III) a series of half grabens with segmented fault-complex or collapsed footwall as basin-bounding structures.

(A)

(B)

Sediment dispersal direction Drainage divide

l

F

Footwall

H

Hanging wall

Accommodation space

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Extension generates footwall uplift and adjacent hanging-wall collapse (Stein and Barrientos, 1985; Jackson et al., 1988; King et al., 1988; Stein et al., 1988). The throw on the subbasin’s boundary fault decreases away from the central part of the fault toward its tips (Figure 3A) (Walsh and Watterson, 1988; Jackson and White, 1989), producing the typical half-graben accommodation: wedge-shaped in vertical section, and commonly scoop-shaped in plan view (Figure 3). The hanging-wall depositional sinks may attain considerable depths; marine synrift strata have thicknesses of up to 3 km (Surlyk, 1978). A complicating feature of a rifted terrain with a series of half grabens is that the basin-bounding master faults, which themselves do not move uniformly along their length (Schwartz and Coppersmith, 1984; Wallace, 1985a), may or may not move in concert across the region. Although initiation, reactivation, migration, and cessation of movement on faults is likely to vary across the area, it is still unclear whether an entire rifted area experiences synchroneity or diachronism of its tectonic maxima and minima. Dating and chronostratigraphic resolution is still too poor to determine which alternative was more likely in the northern North Sea, although there are indications of a weakly diachronous rifting acme. Another feature of rifted areas is their common pulsed tectonic development. It has been argued that there are recurrent tectonic phases in many extensional terrains, particularly well illustrated in the Middle Jurassic–Early Cretaceous infill history of the northern North Sea rift basins (Underhill, 1991a, b; Rattey and Hayward, 1993; Færseth et al., 1995), but also observed in other marine (Sellwood and Netherwood, 1984; Purser et al., 1990) and nonmarine rifts (Lambiase, 1990). During Middle Jurassic–Early Cretaceous rifting in the North Sea–Norwegian Sea, the successive rift phases commonly had a duration of some 4–6 m.y., whereas the entire rift episode lasted a few tens of m.y.

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further enhanced by the tendency for accommodation creation to outpace sediment supply for much of the rift cycle, is the potential development of prominent transgressive systems tracts. A second unique aspect of rift basins, particularly marine rift basins, is the variability of the sediment supply. Supply of sediment to any subbasin is sensitive both to hinterland distance and to the size and gradient of the local fault-block source areas (which themselves vary through the rift cycle). The sediment-yield potential of areas updip of the fault block is likely to be less than the volume of adjacent depositional sinks; therefore, the risk of marine subbasins being underfilled is great. The important implication of this specific tectonostratigraphic style is that the creation of accommodation space is complex, varying in time and space, because of background subsidence and uplift, fault-block rotation and differential subsidence, linked variable but commonly limited sediment supply, and varying eustatic sea level. In most conventional sequence-stratigraphic models, in full grabens or other gently sloping basins sediment supply and basin floor tectonics are assumed to be constant or uniformly varying in space; therefore, it is hardly surprising that sequence stratigraphy is difficult to use as a predictive tool in analyzing synrift basin fills (Steel, 1993). In a synrift half-graben situation there is updip erosion and generation of sediment concomitant with, and only a short distance from, downdip accommodation creation and sediment accumulation. The upslope-downslope migration of the shoreline zone generated by changes in rotational style and rate and influenced by changes in regional subsidence and eustatic sea level, then, in addition to the hinterland-basinward movement of the rift-margin shoreline, largely determines the regressive or transgressive tendencies of the infilling sedimentary systems. CONTROLS ON RIFT-INFILL ARCHITECTURE

Unique Aspects of Rift Basins One of the unique aspects of fault blocks or half grabens that affects their role as depositional sinks is the unusually steep slope gradient attainable by the basin floor and the systematic variation of this slope in time. This steep slope gradient arises as a result of extensional rotation, which causes simultaneous subsidence and uplift of downdip and updip areas of the fault block, respectively. Rapidly attained basin-floor tilt angles of 2–3°, with longer term cumulative tilt of up to 10°, is not uncommon in the Late Jurassic subbasins of the northern North Sea (Færseth et al., 1995). An important consequence of the steep local slopes of half grabens,

In this section, we evaluate the factors that control the variability of synrift stratigraphy in marine half grabens during a single rift phase and the preceding and succeeding tectonic quiescent stages. The resultant synrift stratigraphy appears to be controlled mainly by sediment supply in relation to tectonic subsidence and uplift (or accommodation development). Controls on Accommodation Globally synchronous changes in sea level result from changes in the volume of the ocean basins and in the volume and distribution of water in

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

Mechanism Continuing lithospheric differentiation Changes in the volumetric capacity of the ocean basins caused by sediment influx or removal Changes in the volumetric capacity of the ocean basins caused by volume changes in the mid-ocean ridge system Glacio-eustasy and glacio-isostasy Desiccation (salinity) crisis

Rate 0.2–0.4 mm/1000 yr (0.0002–0.0004 mm/yr) Not expected to exceed 1.0 mm/1000 yr (0.001 mm/yr) Maximum of 7.0 mm/1000 yr (0.007 mm/yr) Fall average ~1.5 mm/yr; rise average ~10 mm/yr (peak rate of 40 mm/yr) 10 mm/yr

Magnitude (m) 150–250

<100–150 180 (200?) 12

*Data are from Allen and Allen (1990) and Plint et al. (1992) and references therein.

them. For half grabens connected to the oceans, five possible causes are suggested for eustatic changes in sea level (Table 1) (Allen and Allen, 1990). Except for glacio-eustasy and desiccation crisis, eustatic sea level changes are of relatively slow rates. Glacio-eustatically controlled accommodation development is commonly asymmetrical, resulting from lower rates of sea level fall than rise (Table 1), and produces rather short-term (10–100 k.y. time scale) variations in sea level. These shortterm variations are attributed to Milankovitch cyclicity, and may result in the development of high-frequency sequences (e.g., see Allen and Allen, 1990; Plint et al., 1992). Structurally controlled accommodation results from the effects of background subsidence and from slip on basin-bounding (and intrabasinal) faults (Figure 2). Both tilting and vertical movements accompany slip on active normal faults. Vertical movements involve hanging-wall subsidence and, commonly, footwall uplift (Barr, 1987a, b; Jackson et al., 1988; Jackson and White, 1989). These vertical movements decrease with distance from the fault, causing a tilting of the footwall and hanging-wall blocks (Figure 3). The rotational pivot denotes the line that separates the areas subjected to footwall uplift from areas experiencing hangingwall subsidence as induced by fault-block tilting alone. The superimposed background subsidence, which causes the entire fault block to subside, results in an updip movement of the line of zero net vertical movement and thus a narrowing of the zone experiencing synrotational uplift. The fulcrum is the line separating those areas of the fault block subjected to uplift from areas subjected to subsidence, taking into account the effects of both tilting and background subsidence (Figure 2). The updip-downdip movement of the shoreline zone is determined predominantly by the combined effect

of the fulcrum position and the prerotational sea level stand (Figure 4), in addition to any eustatic fluctuations. If the rift basin includes a series of master faults, the dip section is a series of fault blocks or half grabens. The space created by fault-block rotation (Figure 2) depends upon the rate of rotation, total extension, and initial master fault spacing (Barr 1987a, b; Yielding, 1990), the initial fault dip and fault shape (planar, listric, or ramp-f lat-ramp geometries) (e.g., Morley, 1989), as well as lithosphere properties (ter Voorde and Cloetingh, 1995). In continental rifts and proto-oceanic troughs, coseismic footwall uplift commonly ranges from 5 to 25% of the total displacement along the fault (Table 2) (Jackson and McKenzie, 1983; Stein et al., 1988; Jackson and White, 1989; Yielding and Roberts, 1992). During periods of seismic inactivity (interearthquake intervals), the upper crust deforms isostatically by flexuring (King et al., 1988; Stein et al., 1988; Weissel and Karner, 1989; Kusznir et al., 1991). Hanging-wall deposition and footwall erosion further influence the isostatic relaxation path following coseismic deformation. Accordingly, a total footwall uplift of some 50% of the hanging-wall subsidence may result from coseismic and postseismic deformation (Stein and Barrientos, 1985; Roberts and Yielding, 1991). Footwall uplift, however, may be suppressed if the stretching factor is large, the initial fault spacing is small (Barr, 1987a, b), or if the background subsidence is high. The amount of stretching increases toward the rift basin axis, and is commonly expressed by larger throws on the basin-bounding faults and stronger tilting of the basin floor (Roberts et al., 1993b). Fault spacing commonly shows a systematic variance as well, and decreases toward the rift axis. This variance in fault-spacing possibly reflects the

A

fulcrum

fulcrum

Transgression and storage in backbarrier and coastal plains. Mainly sediment-balanced subbasins due to potential axial system.

Expansion of sea

SL below fulcrum

fulcrum

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A'

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lll

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shoreline footwall progradation uplift

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widening of footwall island due to

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Blanket of hemipelagic mudstones and possible submarine gravitational resedimentation.

fulcrum

Starvation or submarine erosion on submerged, updip areas

Submerged

fulcrum

fulcrum

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C'

C'

C'

Figure 4—Bathymetry of (A) slightly submerged subbasins, (B) partly emerged subbasins, and (C) completely submerged subbasins. Bathymetry at the onset of rifting is shown as block diagrams (row I) and vertical cross-sections (row II). Rift climax bathymetry and sediment infill pattern are shown in row III. Note the different development of hanging-wall shorelines in the slightly submerged and partly emerged subbasins.

A

lll

A

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(A) SLIGHTLY SUBMERGED SUBBASINS


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

Rift Basin and References*

Displacement Rates (mm/yr)

Basin and Range Province Teton fault, Wyoming (1–3) Wasatch fault zone, Utah (4–6) Stillwater front, Nevada (7–9) Owens Valley, California (10) Great Basin compilation (11)

Hanging-Wall Subsidence Rates (mm/yr)

Footwall Uplift Rates (mm/yr)

0.45–1.6 0.35–3.9

Earthquake Recurrence Intervals (yr) 700–6000 500–4000 average 2000 c.r.i. 400 2800–7600 <3000

0.26–0.5 >0.4 0.01–0.5

Rio Grande rift Espanola basin, New Mexico (12)

0.11

Lusitanian Basin, Portugal Arruda subbasin (13, 14)

0.3–2.0

Gulf of Suez** Zeit-Shadwan fault–Esh Malaha fault, Gebel Zeit (15)

0.4

East African rift Lake Tanganyika (16) Lake Rukwa (16) Northern Kenya rift (Lokichar) (16)

1.0 1.07 0.54

Viking Graben, North Sea Brage fault; Oseberg fault Block (17)

Average 0.2

Gulf of Corinth, Greece† Gulf of Corinth (18–22) Gulf of Alkyonides (23, 24) Gulf of Patras (21, 25)

0.05

Average 0.18 peak 0.45

Average 0.02 peak 0.15

2.5–3.0 0.13–5.0

0.3

0.1–5.0 0.13–5.0

Argrolikos Gulf, Greece Argrolikos Gulf (26)

300–1700 400–1500

0.5–1.0

*1 = Doser and Smith (1983); 2 = Gilbert et al. (1983); 3 = Smith et al. (1993); 4 = Swan et al. (1980); 5 = Schwartz and Coppersmith (1984); 6 = Machette et al. (1991); 7 = Wallace and Whitney (1984); 8 = Okaya and Thompson (1985); 9 = Bell and Katzer (1990); 10 = Pinter (1995); 11 = Wallace (1985b); 12 = Golombek (1981); 13 = Wilson et al. (1989); 14 = Leinfelder and Wilson (in press); 15 = Jackson et al. (1988); 16 = Morley (1989); 17 = Hamborg, personal communication (1995); 18 = Schröder and Kelletat (1976); 19 = Brooks and Ferentinos (1984); 20 = Doutsos and Piper (1990); 21 = Doutsos and Poulimenos, (1992); 22 = Dart et al. (1994); 23 = Leeder et al. (1991); 24 = Perissoratis et al. (1993); 25 = Chronis et al. (1991); 26 = van Andel et al. (1993). **Has not been corrected for any possible regional uplift or subsidence effects (Jackson et al., 1988). †Has not been corrected for regional uplift; regional uplift rates are in the range of 0.3 to 1.5 mm/yr (Keraudren and Sorel, 1987; Collier, 1990; Collier et al., 1992).

progressively attenuated lithosphere or higher heat flow beneath the more extended regions (Jackson and White, 1989). Slip on normal faults related to single large earthquakes (Richter scale magnitude >6.5) (Jackson and White, 1989) ranges between 1.5 and 4 m, with the maximum reported slip on recent faults as large as 6 m (Fraser et al., 1964; Wallace, 1985a). Reported earthquake 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 lineaments and not return for thousands of years (Wallace and Whitney, 1984). Displacement rates on normal faults in extensional terrains are from 0.01 to more than 5.0 mm/yr (Table 2), with the values including possible slip on successive small- and moderate-magnitude earthquakes. Recurrence intervals (and thus slip rates) on recent faults vary laterally along fault

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Shoreline affected by eustatic sea level variations

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

zones, the earthquake frequency and magnitude being lower on fault segments located toward the ends of the lineaments (Schwartz and Coppersmith, 1984). Displacement rates vary laterally also along individual fault segments (Wallace, 1985a). Spatial and temporal variations in style and the rate of fault-block rotation, whether linear, nonlinear, or pulsating, influence the rate of accommodation creation. During a single rift phase, the displacement rates increase during the early synrift substage toward a maximum during the rift climax, and then progressively wane during the late synrift substage. Tectonic quiescent stages are prolonged periods of low seismic activity occurring between recurrent rift phases and characterized by background subsidence and minimal fault-block rotation. A comparison of published data on rates of eustatic sea level change with data on normal fault displacement rate shows clearly that, except during glacial periods, structural control on accommodation development by far exceeds eustatic control

(Tables 1, 2). Of the eustatic mechanisms, only glacio-eustasy appears to be of sufficiently high rate and magnitude to compete with accommodation changes resulting from extensional faulting. Other eustatic mechanisms show rates and sometimes also magnitudes too low to leave a basinwide imprint, and are most likely suppressed or masked by higher tectonic subsidence rates in medial and downdip reaches of the subsiding half grabens. Nevertheless, the eustatic signal resulting from the other mechanisms (in combination with the structurally induced accommodation development) is likely to have influenced the resultant relative sea level changes at (local) footwall and hanging-wall shorelines. Although it is difficult to disentangle the eustatic signal in this type of setting (it may be clouded by variations in sediment supply and suppressed by strong, local tectonics), some general points are worth noting. Eustatic sea level changes affect broader areas on gently dipping hanging-wall

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Figure 6—Nomenclature, style, and development of drainage types. (A) Prerift to earliest synrift drainage. (B) Synrift drainage.

(A) Prerift drainage rift margin Rift-marginal Antecedent axial

Rift-marginal Antecedent transverse

transfer zone Rift-margin Antecedent transverse

(B) Synrift drainage rift margin and interior

Rift-margin or interior drainage Antecedent major incision and axial river capture (backstepping) Consequent-Juvenile hanging wall

in

arg

t-m

Rif

ge ina ra ld ina ma

rg

Antecedent transverse

ge

transfer zone

Rift-interior drainage

ina

dra

Major sediment entry point at crosscutting fault

Ri

ft-

Consequent-Juvenile fault-scarp & hanging wall

Antecedent transverse (major incision)

Consequent Reversed Redirected (axially)

Large footwallsourced system (antecedent or consequent)

slopes compared to their steeper fault-scarp equivalents (Figure 5), especially if the hangingwall shorelines are located at or close to the faultblock fulcrum. In addition, eustatic sea level changes show a greater influence during periods of low rates of fault-related subsidence and basinfloor rotation; that is, during the early and late synrift substages and intervening tectonic quiescent stages. If the rift basin development were characterized by stepwise rotation or a series of rift phases, eustatic accommodation changes probably would be more inf luential during the earlier phases because these would be represented by the formation of shallower subbasins with gentler slopes, whereas later rift phases would be commonly dominated by deeper basins with steeper gradients.

Surlyk et al. (1993) argued that a basinwide sea level signal can be recognized in the large-scale cyclicity in marine synrift sediments of eastern Greenland. The sea level changes may have been of eustatic nature, but these workers emphasized that the changes more likely were a direct result of extension. Should this hold true, it would argue strongly in favor of near-synchroneity of basinwide extension. Cases of spatially variable accommodation resulting from Pleistocene glacio-eustasy in extensional settings in Greece have been discussed by Gawthorpe et al. (1994). They noted that close to the center of the normal faults, the effects of glacioeustatic falls may be canceled by high rates of hanging-wall subsidence. Away from the fault zones and toward the fault tips, however, relative sea level variations eventually would be dominated by eustasy.

Ravnås and Steel

Sediment Supply and Drainage Development In rift basins, extensional faulting exerts the prime control on morphology and topography, and hence on drainage (e.g., Jackson and Leeder, 1993; Leeder and Jackson, 1993). Additional factors that influence the temporal and spatial variation in sediment supply and drainage development or establishment include the following items. (1) Climate, which includes factors such as temperature, wind levels, and rainfall, and their seasonal variations and fluctuations, exerts a major influence on vegetation and biotype development, weathering processes, sediment supply, depositional environment, and the resultant lithofacies. For large, elongate rifts (e.g., East African rift, Red Sea–Gulf of Suez rift, Rio Grande rift), climate may vary considerably along the length of the rift basin. (2) The distance of hinterland area to successive subbasins determines the extent to which individual half grabens receive sediment from a rift-margin drainage or are dependent on supply from local (fault-block) drainage alone. (3) The size of drainage catchment, which follows Hack’s Law (Hack, 1957), depends largely on the nature of the drainage itself, whether transverse (at high angle to basin-bounding faults), axial (at low angle or paralleling the basin-bounding faults), antecedent or consequent, with consequent drainage, including reversed, captured, or juvenile drainage (Figure 6). The dimensions and nature of the local (i.e., consequent or fault block) drainage catchments depend largely on fault geometry and spacing. In general, the hanging-wall catchments are larger than those on the footwall, resulting in greater sediment-yield potential for hanging-wall successions. (4) The half-graben morphology itself, especially the gradient and relief of its slopes, the presence, spacing, segmentation, and orientation of any intrabasinal structures, and the presence of transfer zones or crosscutting faults, influences sediment supply. The steeper fault-scarp slopes result in a more rapid response and evolution of the footwall drainage systems. Spatial variations in response time between catchment generation and the development of drainage systems and their characteristic transport, dispersal, and depositional processes would be expected. Moreover, segmentation of the footwall may shift the local faultblock drainage divide away from the original basin-bounding master fault (Figure 3). These processes, either on their own or in concert, result in an increase of the footwall catchment area at the expense of the hanging-wall complement. In addition, note that fault-scarp successions form mainly as a response to fault-related subsidence along and drainage development on the footwall of the

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basin-bounding master fault. Hanging-wall successions, however, record changing conditions on the hanging wall, usually reflecting activity also on the basin-bounding master fault of the neighboring half graben. Accommodation zones or transfer faults determine the level to which individual subbasins are filled before sediment transfer to the neighbor ing subbasin can occur (Lambiase, 1990). The presence of hinterland or basinward-dipping half-graben boundary faults would be a major control on the progradational pattern and sediment architecture related to transverse, rift-margin depositional systems (5) The prerift substrate (sediment cover and rock types) clearly influence the response time, sediment-yield potential, and caliber of any local, intrabasinal sediment source. Also important is whether the rift infill contains volcanics. Sediment flux in relation to drainage development in rift basins has been discussed extensively. Rift subbasins receive their sediment from progradation of hinterland-derived, rift-margin depositional systems or from rift-interior sources. Rift-margin depositional systems usually represent antecedent drainage systems, and may be either axial or transverse (Figure 6). Axial systems commonly consist of relatively large fluvio-deltaic systems, which drain tilted subbasins upstream. Transverse systems consist either of relatively small antecedent or of consequent hanging-wall and footwall fans, fan-deltas, deltas, and shorelines (Leeder and Gawthorpe, 1987; Leeder and Jackson, 1993; Seger and Alexander, 1993; Gawthorpe et al., 1994), or large, antecedent fluvio-deltaic and shoreline systems traversing broad, submerged, platformal areas that border the rift basin proper (e.g., Hellem et al., 1986; Ravnås and Bondevik, 1997). Rift-interior sources include footwall shoals, peninsulas, and islands (Figure 6B). Depositional systems range from antecedent to consequent hanging-wall and fault-scarp talus (collapse) cones, slope aprons, fan-deltas, and deltas, with the latter two possibly feeding submarine turbidite lobes, fans, or shorelines. Only consequent drainage would develop on footwall peninsulas or islands formed during rotation of a previously submerged terrain. Upon reaching downdip (basinal) settings, sediment delivered from the transverse systems may spread out laterally (Feretinos et al., 1988) or become redirected axially down the regional paleoslope or toward the center of individual, closed subbasins (Papatheodorou and Ferentinos, 1993). Major sediment entry points form at the outlet of antecedent drainage systems (Leeder et al., 1988; Nelson et al., 1992), transfer zones (relay ramps) (Surlyk, 1977; Surlyk et al., 1981; Leeder and Gawthorpe, 1987; Morley et al., 1990; Gawthorpe and Hurst, 1993; Morley, 1995), crosscutting faults

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(Cherry, 1993) or basinward of easily erodable substrate (unconsolidated sediments, etc.) (Leeder et al., 1991) (Figure 6B). To maintain the course of transverse antecedent river systems during rotational faulting, high downcutting potential is implied with erosion rates larger than or equal to uplift rates, leading to river incision. Steep fault-scarp slopes may lack pronounced entry points; instead, sediment may derive from a linear feeder systems or a series of smaller entry points (e.g., Surlyk, 1978). During a rift phase, subbasins located nearest the hinterlands tend to trap most of the coarse material delivered by the rift-margin depositional systems. Subbasins located far from the hinterlands receive their sediment predominantly from riftinterior sources. When considered in a simplified, schematic manner, each rift phase is characterized by fault-block rotation superimposed on the background subsidence (Sb in Figure 2) (Barr 1987a, b; Jackson and White, 1989). If the prerotation elevation of the terrain is equal to Sb, the volume of the fault block that is above sea level at the end of faulting is equal to the adjacent submer ged sinks (assuming no along-strike variations). Decreasing the prerotation elevation would diminish the sediment-yield potential of the fault blocks. Consequently, stretching of a terrain that was originally close to or below sea level results in the formation of intrabasinal sediment sources that are of limited size compared to the volume of the adjacent basinal areas (sinks). This scenario results in the sediment-underfilled or clay-prone synrift signature commonly observed in marine rift basins. Notably, if the prerotation water depth is great, synrotational footwall uplift would not be sufficient to produce any subaerial emergence, and the consequent sediment-fill processes would be entirely submarine (Figure 4C). Sediment-underfilled and sediment-starved half grabens are dependent on sediment supply from other, distant sources to be completely filled, most likely the rift-margin hinterland. The subbasins become filled in a successive manner, the rate of progradation of the rift-margin shorelines determined by sediment supply, syndepositional accommodation variations, depositional foundation (Helland-Hansen and Gjelberg, 1994), and water depth to accommodation zones or transfer faults (Lambiase, 1990). Other infill patterns are present in carbonate- and evaporite-prone rift basins. A discussion of such infill types is, however, beyond the scope of this study. From these considerations, it is clear that the prerotation sea level stand and the size of and distance to the hinterland are major controls on sediment supply variations in rift basins (assuming similar climate and substrate lithologies). The distance to and sediment-yield potential of the hinterland

determines whether sediment from this source will be delivered to any subbasin, and hence commonly is the prime control on the degree to which any subbasin will be sediment overfilled, sediment balanced, sediment underfilled, or sediment starved. Prerotation bathymetry determines the size of faultblock catchment areas and influences the timing of supply from such sources. These interrelated factors are discussed further in following sections. Influence of Prerift Bathymetry and Bathymetry Development Rift basins commonly consist of a series of regionally tilted, elongate half grabens that tend to have a lower basin floor elevation toward the axis of the rift basin (Figure 2). This tilting suggests that a continuum from essentially nonmarine or partly submerged through mixed nonmarine/ marine or partly emergent to completely submerged marine subbasins both across and along the length of the rifted terrain may exist. Hence, one should expect a significant variability in sediment-infill pattern. At the landward end of marginal half grabens, marine incursions are likely in downdip areas during the rift climax only, with other substages characterized 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 different nonmarine, synrift depositional models). Subbasins with much of their basin floor close to sea level (Figure 4A) experience rapid drowning of their low-gradient downdip reaches and retreat of the rift-margin depositional systems (both axial and transverse) during the early synrift stages. Rejuvenation of the rift topography may promote progradation of any local, hanging-wall or footwall transverse system. However, this is counterbalanced and, with time, suppressed by the increasing overall subsidence, which results in further confinement of coarse siliciclastics to near the boundar y faults and to high up on the hang ing wall. Marine influence in the succession reaches a maximum during the rift climax stage. Progradation of the shoreline may occur locally at the mouths of major sediment conduits. Increasing rates of erosion in the newly created tectonic uplands along with establishment of antecedent and consequent drainage systems eventually result in renewed progradation of the depositional systems during the late climax and late synrift substages. Provided there is sufficient sediment available, individual subbasins may be completely infilled during the late synr ift and tectonic quiescent or early postrift stages.

Ravnås and Steel

If the fault-block emergence point (and hence any prerotational exposed areas) is located above the fulcrum (as in partly submerged subbasins), footwall uplift during the succeeding rift phase is likely to increase the width of any preexisting footwall peninsulas or islands, resulting in a progressive increase in the sediment-yield potential (Figure 4B). Any transverse drainage that existed during the preceding (prerift or tectonic quiescent) stage would have a rapid and enhanced sedimentary response to the new tectonic regime, expressed as progradation and off lap of sedimentary lobes down the hanging wall. Peak downdip progradation depends on the uplift rate vs. the denudation rate, and occurs either during or, more likely, sometime after the rift climax. Afterward, the hanging-wall shoreline retreats sourceward, that is, up the hanging wall itself. This response is clearly at variance with the case described where the prerift shoreline zone was located below the fulcrum. Extension of a completely submerged terrain may trigger gravitational collapse if there was some inherited relief prior to the initiation of the rift phase. If the prerotational water depth was relatively shallow (that is, if the background subsidence and the initial water depth together were less than the footwall uplift), enhanced faulting and associated footwall uplift eventually would lead to the formation of footwall peninsulas or islands, providing a rift-interior source area (Dahl and Solli, 1993; Roberts et al., 1993a; Solli, 1995; Ravnås and Bondevik, 1997). Subaerial erosion and the establishment of newly formed, consequent, hanging-wall and footwall drainage would develop some time after the updip reaches of the fault block have become emergent. If this occurs during the early synrift stage, continued footwall uplift would lead to widening of the newly created footwall peninsulas or islands during the climax and early late-rift stages (Roberts et al., 1993a), accompanied by an increase in sediment supply and sustained downdip progradation during the climax and early late synrotational stages (Ravnås and Bondevik, 1997). If, however, footwall peninsulas or islands were formed late in the rotational tilt stage, the exposed areas would be fairly narrow and, most likely, capable of producing only limited amounts of sediment. Moreover, these limited catchment areas probably would be surrounded by narrow and fairly steep shorelines. Hence, no pronounced shoreline progradation is expected to occur. The fault-scarp systems may show either a simple backstepping or a forestepping-to-backstepping stacking pattern. In the latter motif, the lower forestepping part of the scarp represents the response to the newly created footwall uplands in the early synrift stage. As increasing basinal subsidence rate outpaces the sediment supply, continued

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retreat of any fault-scarp system would occur (e.g., Surlyk, 1989). If the prerotational water depth is great (if the sum of the background subsidence and the initial water depth together is larger than the footwall uplift), footwall uplift cannot produce emergent areas (Figure 4C). Instead, submarine shoals may form in the footwall of the boundary faults. Such subbasins may be draped by a veneer of hemipelagic mudstones across intrabasinal highs, possibly intercalated with submarine resedimentation products in depositional sinks (Underhill, 1991a, b). Deposition from the sediment gravity flows is probably initiated by gravitational instability along the f lanks of adjacent, intrabasinal highs, and the deposit may in some cases reach considerable thicknesses (Færseth et al., 1995). Sediment instability may be tectonically triggered, and the resultant deposits may increase in abundance, possibly forming the bulk of the accumulated sediment during periods of high earthquake activity (the climax stage). Updip areas may show considerable sediment starvation, in particular where they are subjected to current or wave agitation. If sufficiently shallow (i.e., usually above wave base), footwall submarine erosional unconformities may form as a result of persistent wave or current erosion. FACIES TRACTS Marine synrift successions can be divided into two main types: (1) nonmarine, paralic, and shallowmarine infills (Figure 7A); and (2) deep-marine infills (Figure 7B). Deep-marine infills may include shallow-marine deposits peripheral to emergent footwall highs and on submarine shoals. The facies tracts present in the shallow and deepmarine basin types outlined in the following sections should be regarded only as idealized models, and do not take account of all possible facies associations present in rift basins. Varying sediment yield and caliber, in particular, affect the variability and architecture of the facies tracts (Orton and Reading, 1993; Reading and Richards, 1994), as do the prevalent basinal regime, whether stormwave, tidal, or fluvial influenced. Shallow-Marine Facies Tracts The shallow-marine facies tracts derived from the hanging wall, the axial system, and the fault scarp vary greatly (Figure 7A). The stacking pattern of the resultant sediment geometries depends on the interplay of accommodation development and type of accommodation in relation to sediment supply.

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(A) T id al inf lue nc e

1

3 1

Transgr essive

2

Regressive 3 3

2

1

(B) 3

1

2

4

5

6

Figure 7—Schematic, idealized presentations of synrift sediment infill models for (A) mixed nonmarine and shallow-marine rift basins and (B) deep-marine rift basins. The various styles of shallow-marine infills are based on data from the Tarbert, Hugin, and Ling formations, axial parts of the Viking Graben, and the (point 1 in A) Oseberg, Hild, and Sleipner fields, and along the (point 2 in A) Alwyn, Brent, and Statfjord fields (see also Johannessen et al., 1995), and the (point 3 in A) Heather Formation and its intercalated sandstones on the rift-margin Horda Platform, Oseberg and Brage fields. The deep-marine variant is based on data from the (point 1 in B) Brae Formation, southern Viking Graben, (point 2 in B) Draupne Formation, Visund field, (point 3 in B) Munin sandstone member, Statfjord North field, (point 4 in B) Kimmeridge Clay Formation and intercalated Ptarmigan and Magnus sandstone members, Penguin half graben, (point 5 in B) Lusitanian Basin hanging-wall succession, and (point 6 in B) Draupne Formation, Oseberg, Brage, and Veslefrikk fields.

Hanging-wall dip-slope facies tracts include hanging-wall shorelines and associated backbarrier environments, deltas, and fan deltas (Figures 7A, 8), with fan deltas commonly represented by shallowwater, Hjulstrøm types (Colella, 1988a, 1994; Massari and Colella, 1988; Gawthorpe and Colella, 1990) and their offshore correlatives. Where backbarrier environments are volumetrically well represented in the infill, characteristic features of the

transgressive segments are (1) unusually steep, seaward-dipping ravinement surfaces; (2) an unusually thick carpet of (transgressive) sand above the ravinement surface; and (3) thick wedges of backbarrier deposits, which also accumulate during the transgression, although below the ravinement surface. Onlap up the steep hanging-wall slopes tends to enhance the thickness of transgressive deposits. Figure 9 illustrates these points from

Figure 8—Correlation panel of the Viking Group in the Oseberg-Brage area of the northern North Sea (modified from Ravnås and Bondevik, 1997). (A) During the Bathonian–Kimmeridgian, three phases of tilting (to the east) and general subsidence produced three, stacked regressive-transgressive sequences. (B) The transgressive segments formed during periods with high tilt rates and are characterized by local, transverse fault-scarp and hanging-wall shoreline and fan deltas.

(A)

(B)

1

2

GR

S. Alwyn area 3/15-4

1b ive regress

1a transgressive

3

4

25m

GR

29/6-1

Hild area

Transgressive

Regressive

GR

Maximum flooding surface

Sequence boundary

Ravinement surface

Coal-bearing lagoon/delta plain

Transgressive sand

Barrier & shoreface

5

SE 30/7-8

Figure 9—Correlation panel of the Tarbert Formation in the Hild-South Alwyn area of the northern North Sea (modified from Rønning and Steel, 1987). Five sequences (numbered 1–5) developed by slight tilting episodes during the Bajocian–Bathonian. Note the seaward partitioning of sediment during regression and landward partitioning (storage) during transgression. Farther west, marine claystones are present at the maximum flooding surface levels.

NW

126 Marine Rift-Basin Architecture

Ravnås and Steel

127

N 30/9-8

30/9-10

GR Marine mudstone

outer - middle shelf

Marine sandy siltstones

middle - inner shelf

GR

S

shoreface (wave dom.) Marine sandstones

Starved outer shelf

shoreface (mouth-bar / tide dom.) mouth bar or tidal/fluvial channels 13

Lower delta plain deposits

(Tarbert Formation) (Ness Formation) 12

Erosional unconformity 10

Inner shelflower shoreface

Time lines w/palynology control point Transgressive Regressive

30/9-3 GR

11

Backbarrier and tidally influenced embayment

50

40

30 10 20

10 (m) 0

9

Lower delta plain

Figure 10—Rift-axis parallel correlation panel of the Tarbert Formation in the Oseberg area, northern North Sea. Varying rates of extension produced four landward backstepping, regressive-transgressive shoreline prisms. The regressive part formed during periods with lower tilt rates, whereas the transgressive part formed during periods with higher tilt rates. Note the development of thick transgressive packages that show a higher degree of marine influence upward in the succession, both individually and overall.

Middle Jurassic (Brent Group) fault blocks in the northern North Sea, where the destruction of shoreline barriers by steep-trajectory transgression repeatedly resulted in sediment being partitioned both landward (backbarrier) and seaward (lower shoreface). The footwall-derived fault-scarp facies tract commonly consists of isolated to coalesced fan deltas with interfan areas dominated by shoreline-paralic to shallow-marine environments (Figure 8) (e.g., Surlyk, 1978, 1989; Leeder et al., 1991). Fan-delta types may show an evolutionar y pattern from shallow-water shelf type (Hjulstrøm-type) through shoal-water mouth-bar types (Postma and Drinia, 1993) to Gilbert-type deeper water, gravitationally modified fan deltas (Figure 7A) (Colella, 1988a, b, 1994). Waning fault-related subsidence during the

late synrift stage may be accompanied by a change from an aggradational to a progradational stacking pattern (Colella, 1994), possibly associated with a change from an alluvial fan to a braid plain as the feeder system (Gawthorpe et al., 1990). Alternatively, the late synrift substage is characterized by renewed shoreline deposition. Axial facies tracts include both fluvio-deltaic and shoreline systems (Figure 10). High constructive deltas (Orton, 1988) commonly form when basinal energy (waves and tides) is low, whereas coastal plains fronted by parallel beach ridges result from higher wave energy. However, contrasting responses to the basinal marine regime, and the development of reflective or dissipative shorelines, are also dependent on the sediment supply and the caliber of the sediment delivered to the delta front (Orton,

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Figure 11—(A) Schematic basin-fill stratigraphy of the Oxfordian–Kimmeridgian synrift succession of the Lusitanian Basin, western Portugal. (B) Detailed architecture of synrift hanging-wall succession with submarine channel fill, disturbed slope sediments, and prograding Gilbert-type fan-delta sediments. Height of section is approximately 150 m. (From Ravnås et al., in press b.)

Figure 12—Strike-oriented panel of the Brae Formation subparallel to the master fault along the west side of the South Viking Graben less than 1 km out from this fault (from Theriault and Steel, 1995). The panel shows the highly variable sedimentary architecture of a footwall-derived conglomeratic and sandy submarine fan succession.


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(A) W

E

Cretaceous

(B)

?

Visund Fault Zone

CUFU 2 FU

old ith k w ic loc ass lt B Tri au & d F sic un Juras

Bas

er

Draupne Formation

Vis

FU eC

reta

ceo

us

2

?

Bas

Formation aupne e Dr

1

CUFU

CUFU 1 FU FU

150m ~5 - 6 km

CUFU Unconformity FU

Draupne clastic wedge

1

Major sequences in well

Heather Formation

Submarine unconf. or condensed interval

MINOR MAJOR Grain size

Increasing congl./sand + shale ratio

CU= coarsening upward FU= fining upward Siltst./mudst. Sandstone Conglomerate

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

1988; Orton and Reading, 1993). The fluvio-deltaic systems, whether originally transverse or axial, may feed basinal sediment gravity flows (Cockings et al., 1992), with these f lows turning axial upon reaching downdip settings (Figure 7A). Individual half grabens commonly form as scoopshaped depressions along master fault zones. Any downdip axial fluvio-deltaic system thus would prograde downslope from the sourceward end of the basin, but upslope basinward. The stacking pattern of the fluvio-deltaic system would vary as a response to such along-strike variations in the depositional foundation. The basinal marine regime may vary temporally, reflecting enhanced or subdued basin topography developed in response to varying styles and rates of extensional faulting. Ravnås and Bondevik (1997) suggested that tidal processes were enhanced by the dissection of rifted terrains into a series of elongate subbasins (i.e., structurally controlled embayments or estuaries) during the rift phases. The funnelshaped morphology of many mixed nonmarine/ marine rift basins (e.g., Surlyk and Clemmensen, 1983) probably provides ideal conditions for the amplification of originally weak tidal currents. An extreme case of tidal influence strongly enhanced by

the weak rotation of narrow fault blocks in the Middle Jurassic of the Hebrides (Mellere and Steel, 1996), resulted in thick tidal deltaic and estuarine deposits, possibly in tide-swept straits, where fault blocks were less than 8 km wide. The relative influence of tides vs. waves probably is determined by the water depth of individual subbasins (Ravnås et al., in press a). Tidal currents may be particularly strong over shallow accommodation zones, with the shallow water depth combined with a narrow seaway accelerating the tidal f low (Collier and Thompson, 1991). In contrast, larger wavefetch can result in a predominance of wave- and storm-influenced sediments. The latter are probably more common during intervening tectonic quiescent stages when footwall islands are completely submerged, and a larger, more open-marine embayment occupies the rift basin (Ravnås and Bondevik, 1997). Deep-Marine Facies Tracts Deep-marine facies tracts (Figure 7B) also vary spatially across the hanging wall, the axial, and the fault-scarp areas of half grabens.

Ravnås and Steel

When updip areas of individual fault blocks are subaerially exposed, narrow shoreline lithosomes form around the footwall islands (Surlyk, 1978, 1984, 1989; Dahl and Solli, 1993). The steeper depositional slopes of the deep-marine setting usually result in narrow facies belts, in more rapid transitions between the individual (proximal-todistal) facies tracts, and in the scarcity of either alluvial or shoreline deposits (due to cannibalization into downdip sites). Shoreline and nonmarine strata may be important only during the late rift and subsequent stages in subbasins becoming completely filled (Guèry et al., 1986). Shoreline material may be transported to basinal settings through localized channelized features, whereas the remainder of the medial hanging-wall succession is mud prone overall. Progradation of the shallow-marine and nonmarine feeder systems into deeper waters down on the hanging wall during the late synrift stage may promote the formation of a shelf ramp or, notably, Gilbert-type deltas in this setting (Figure 11). On completely submerged fault blocks, the hanging-wall facies tracts are represented by mudstone blankets and gravitationally resedimented deposits. The footwall-derived, fault-scarp facies tracts (Figures 12, 13) commonly consist of isolated collapse breccias, talus aprons, slope aprons, and gravitationally modified, deep-water cone deltas, fandeltas, delta-front slope aprons, and canyon-fan systems (Surlyk, 1978, 1984, 1989; Ferentinos et al., 1988; Dart et al., 1994). Submarine aprons and fan deltas may coalesce to form shelf-slope systems with interfan areas that are dominated by various shoreline to shallow-marine environments (Ferentinos et al., 1988; Leeder et al., 1991). One of the best documented cases is that of the Brae oil field, which consists of resedimented deposits along the footwall of the southern Viking Graben. The great lateral variability of the fan deposits along a short strike segment near the fault scarp is shown in Figure 12. The vertical organization of fault-scarp successions characterized by a high supply of coarse siliciclastic sediments shows a development from initial submarine talus through a prograding slope apron to a coalescent fan delta to submarine fan, and then a return to renewed mud-prone deposition, commonly in the form of hemipelagic claystone blankets or mud-rich submarine fans (Surlyk, 1978, 1984; Turner et al., 1987; Cherry, 1993; Surlyk et al., 1993). Provided sediment supply was sufficient and there was complete filling of the original deep-marine basin, an upward transition from slope aprons through aggradational fan-delta sequences to progradational fan-delta sequences may develop (Gawthorpe et al., 1990). The transverse depositional systems, in combination with axial deltaic systems, may feed basinal

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turbidite systems. Sediment gravity flows traveling downslope at an angle to the basin axis may be def lected, resulting in an overall along-axis transport direction. Two different types of sediment gravity flow elements may be present: backstepping and forestepping turbidite segments (Figure 14). The backstepping turbidite segment appears to be dominated by a tendency to downdip shifting and an overall sourceward-stepping or foresteppingto-backstepping stacking pattern. Sediment gravity flow lobes tend to be isolated, embedded within thick basinal mudstones. Large-scale finingupward motifs reflect an upward decrease in the volume of coarse-grained turbidites. The stacking pattern is interpreted to be typical of early to late synrift sediment gravity f low deposits (Surlyk, 1978, 1989; Surlyk et al., 1993; Ravnås and Steel, 1997). The forestepping turbidite segment is dominated by sheetlike turbidites that show an overall basinward-stepping stacking pattern. Initial turbidites infill basinal lows, whereas subsequent flows onlap successively higher up on the hanging-wall and fault-scarp slopes, in addition to reaching progressively farther basinward. This stacking pattern is interpreted to be characteristic for late synrift to tectonic quiescence or early postrift turbidites (Surlyk, 1989; Ravnås and Steel, 1997). In some cases, the forestepping turbidites are backed by an axially prograding submarineslope to shallow-marine feeder system (Leinfelder and Wilson, in press). The development of a dissected, deep-marine rift basin, with individual half grabens silled by accommodation zones, transfer faults, and footwall highs or shoals provides ideal conditions for the development of a stratified water column, possibly leading to dysaerobic and anaerobic conditions in basinal reaches. Deep-marine rift basins thus are favorable settings for the formation of high-quality marine source rocks, as exemplified by the Late Jurassic–earliest Cretaceous Kimmeridge Clay and the Draupne and Spekk formations of the North Sea–Norwegian Sea rift system (e.g., Doré, 1991). SYNRIFT STRATIGRAPHY Rift Basin Infill Types Although prerift elevation, bathymetry, and basinal regime are important variables for the stratigraphy of rift basins, as emphasized by the varied signatures and stacking patterns expected in different basinal settings, the factors of accommodation generation and sediment supply play the dominant role

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(Figure 15). Marine half grabens, whether shallow or deep marine, may be grouped or classified according to the degree of balance attained between tectonically generated accommodation and sediment supply during infilling (Figure 15). Sediment-overfilled basins are characterized by an abundance of sediment supply, complete infilling, and transfer of excess material to adjacent subbasins. Sedimentbalanced basins are typified by sediment supply (nearly) keeping pace with tectonic subsidence (when averaged over the entire synrift interval). Sediment-underfilled and sediment-starved basins are characterized by a sediment supply that is too low to keep pace with the tectonic subsidence, and hence by an inability to completely infill the accommodation space. Note that the temporal variation that an individual subbasin experiences as it subsides and evolves through a series of successive rift phases is also implicit in this classification. The subdivision of rift basin infill types partly parallels the morphological subdivision from marginal half grabens through partly submergent and partly emergent to completely submerged rift basin types. This reflects the changing sediment yield of the main source areas for the different basin types. Riftmargin, sediment-overfilled, and sedimentbalanced half grabens receive sediment from large hinterland areas. Rift-interior half grabens, in contrast, located at distance from the rift-margin hinterland, rely increasingly on supply from rift-interior sources, thereby commonly showing a sedimentunderfilled or sediment-starved synrift lithology motif. Variation in sediment supply from rift-interior and rift-margin sources is dependent on the size of the catchments, and how the catchment’s sediment yield potential varies in time by the interaction of rotational uplift/subsidence, basinwide subsidence, and eustatic sea level stand and its changes. As a result, different infill patterns, and thus different synrift lithology motifs, result from the same tectonic scenario. Importantly, fault-block rotation result in uplift and relative sea level fall in areas above the fault-block fulcrum, whereas areas downdip of the fulcrum are subject to continuous subsidence and consequent relative sea level rise. For each rift basin infill type there are specific hanging-wall updip and downdip signatures that reflect the varying accommodation development and sediment supply throughout the rift phase. Note that basinal successions ultimately record the influence of hangingwall, footwall, and axial drainage. Early Synrift Signatures On the hanging wall, the early synrift period is characterized by deposition of relatively coarse material derived either from a preexisting depositional

system (which may deliver its load simply into the newly created deeper water) or from downslope flushing of weathered detritus during the initial tilting (Figure 15A). During this period, sediment supply is generally not able to keep pace with the increasing rate of accommodation, resulting in areal reduction of the drainage system and an overall sourceward retreat of the depositional system. Shoreface retreat during this transgression is likely to contain the sediment yield from the shrinking drainage area within narrow backbarrier and coastal plain zones. This scenario likely is common in the early synrift stage, but assumes a partly submergent fault block with its pretilt shoreline somewhat downdip of the fulcrum (Figure 4A); however, in cases where the fault block is nearly submerged (i.e., the pretilt shoreline lies updip of the fulcrum; Figure 4B), rotation may be accompanied by an enlargement of the drainage area during the rejuvenation of the rift topography. In general, this results in increasing erosion rates in the newly created uplands, possibly associated with incision of the antecedent drainage systems or the formation of a consequent (reversed or juvenile) drainage. In this scenario, there may be significant production of sand on the updip reaches of the hanging wall. This material may form a “forced regressive” shoreline or deltaic lithosome, or may bypass the shoreline areas, forming sandy sediment gravity flows delivered directly to the basinal areas. The case shown in Figure 4B may result in thicker units of basal sandstone in the synrift succession or in sandy units that coarsen upward from basal mudstones. Footwall uplift of submerged areas commonly results in blanketing of hemipelagic mudstone (the coarser material being trapped in the evolving, adjacent lows), and the formation of condensed deposits, which in areas updip of the fulcrum formed during a relative sea level fall (Ravnås and Bondevik, 1997; Ravnås and Steel, 1997). When crestal areas become emergent, an evolutionary path on the hanging wall similar to that described is envisaged. Hanging-wall areas adjacent to footwall scarps experience continuous subsidence during all of the rifting, although the subsidence may vary both in rate and magnitude. In such downdip areas, there is an initial retreat of any transverse and axial system during the early synrift period, probably ref lecting sediment supply being outpaced by increasing fault-related subsidence, as has been advocated by Prosser (1993), Wignall and Pickering (1993), and Ravnås and Bondevik (1997). However, this signature may apply only for sedimentary lobes sourced through transfer zones (Wignall and Pickering, 1993; Underhill, 1994) or from riftmargin depositional systems (Ravnås and Bondevik,

Figure 14—Correlation panel of the Kimmeridge Clay and its intercalated Ptarmigan and Magnus sandstone members in the Penguin half graben, northern North Sea. Backstepping turbidite segments are present in the sediments deposited during periods with high tilt rates, whereas forestepping turbidite segments are present in the lower of the two packages deposited during periods with lower tilt rates. From Ravnås and Steel (1997).


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1997) and not for evolving footwall-derived, faultscarp depositional systems. The footwall-derived, fault-scarp depositional systems, instead, may show an initial phase of progradation represented by a basal coarsening-upward motif (Surlyk, 1978; Færseth et al., 1995; Ravnås and Steel, 1997). Extension of submerged terrains often produces a footwall-derived fault-scarp succession with a basal fine-grained layer believed to have been deposited during the initial rotation. This basal mudstone reflects deprivation of coarse siliciclastic material in the newly formed half graben prior to the establishment of any transverse drainage. The response to footwall emergence and the establishment of juvenile footwall drainage may be a phase of initial progradation. Near the outlet of any preexisting transverse drainage basal mudstone probably does not occur. Rift Climax Signatures In sediment-overfilled and sediment-balanced basins, the rift climax stage is characterized by maximum retreat of all of the depositional systems, resulting in the deposition of widespread mudprone strata over most of the subbasins (Figure 15B). Coarser material commonly is confined to the subbasin margins. Local, sand-prone lobes, however, may prograde from major sediment conduits. Any of these subbasin marginal systems may feed basinal sediment gravity flows. In the sediment-underfilled and sediment-starved subbasins, the depositional systems are inferred to be confined largely to the basin margins, as well (but see Surlyk, 1978); however, if these subbasins were originally submerged or the prerotational footwall island was small and restricted to the area above the fault-block fulcrum (Figure 4B), this stage would be characterized by continued tilt-climax progradation and offlap of sedimentary lobes down the hanging wall (Roberts et al., 1993a; Ravnås and Bondevik, 1997; Ravnås and Steel, 1997). These hanging-wall depositional systems already may have been established in the early rift stage. Note that the “forced regressive” shoreline developed above the rotational fulcrum passes into “normal regressive” shorelines below the fulcrum (Figure 4B), with the amount and rate of progradation determined by the sediment supply, the depositional foundation, and the water depth of the receiving basin. The fault-scarp successions commonly consist of asymmetrical forestepping-backstepping, coarsening upward to fining upward, and backstepping finingupward packages on various scales (Surlyk, 1978, 1984; Surlyk et al., 1993; Wignall and Pickering, 1993; Færseth et al., 1995; Theriault and Steel, 1995; Ravnås 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 be far less than the space available in the adjacent sink. Hence, the large-scale stacking pattern (hundreds of meters) through the entire rift phase is likely to show an overall fining-upward trend (Surlyk, 1978; Ravnås et al., 1994). Intermediate and small-scale fining-upward motifs have been interpreted in terms of submarine slope processes, such as filling of scours or gullies, retrogressive flow slides, or deposition from surging high- and low-density turbidity currents (Surlyk, 1978, 1984). Intermediate and smaller scale coarsening-upward to fining-upward motifs appear to be developed in successions of mud-rich or poorly consolidated materials (Færseth et al., 1995; Ravnås and Steel, 1997), whereas intermediate and smaller scale fining-upward motifs are more typical in sand-prone successions (Surlyk, 1978, 1984; Turner et al., 1987; Cherry, 1993). Færseth et al. (1995) suggested that the coarseningupward to fining-upward character of the intermediatescale motifs may relate to an initial phase of fan lobe/apron progradation prior to a retreat and backfilling phase. Successive cycles were postulated to reflect rapid influx of sediment followed by more gradual exhaustion of stored sediment from the adjacent footwall high; some of these events were thought to have been coupled to fault movements that caused sediment storage or instability thresholds to have been exceeded. Theriault and Steel (1995) argued for a strong seismic control on decreasing pulses of sediment supply, with smallto intermediate-scale motifs reflecting temporal fluctuations in earthquake activity, and some individual conglomerate strata related to discrete, largemagnitude earthquake tremors. Successions immediately adjacent to the fault scarp are less organized and commonly are arranged in an overall aggradational stacking pattern (Cherry, 1993; Lonergan and Schreiber, 1993). Fault-scarp successions may show large variations along strike, with areas located immediately basinward of sediment conduits receiving vast amounts of coarse siliciclastic material in contrast to interconduit areas, which are finer grained and relatively sediment underfilled, as illustrated from the Brae fault scarp along the southern Viking Graben (Figure 12) (Theriault and Steel, 1995). Late Synrift Signatures The late synrift stage is characterized by waning or low rates of extensional faulting (Figure 15C). If there is sufficient sediment yield potential, sediment supply eventually would equal or exceed the accommodation generation, allowing progradation of coarse siliciclastic systems. In the deep-marine realm, this

Gravelly Sandy Muddy

Shoreline

Drowned shelf w/tidal ridges

Sediment gravity-flow deposits (slope/aprons and basin floor)

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sea level

sea level

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sea level

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shallow marine

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sea level

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Alluvial; lower delta plain

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early postrift/ tectonic sediment quiescence balanced

sediment balanced

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Sediment starved basins:

Sediment underfilled basins:

Sediment balanced basins:

Overfilled basins:

Marginal (fan-apron) gravity flow deposits

Basinal gravity flow lobes

Basinal

Deep marine

Figure 15—Schematic, generalized block diagrams showing the evolution of shallow and deep marine, overfilled, sediment balanced, sediment underfilled, and sediment starved subbasins through (A) the initial and early synrift stages, (B) the rift climax stage, and (C) the late synrift and early postrift/tectonic quiescent stages. (D) Schematic cross sections of shallow- and deep-marine basin fills of sediment-overfilled, sediment-balanced, sediment-underfilled, and sediment-starved subbasins.

sediment starved


Rift climax


Initial and early synrift

(C) Late synrift to

sediment starved

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sediment starved

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stage is heralded by the income of commonly sharpbased, forestepping turbidites (Ravnås and Steel, 1997), which in the overfilled basin type are overlain by submarine-slope, shallow-marine, and alluvial deposits (e.g., Leinfelder and Wilson, in press). In the shallow-marine environment, progradation of shoreline lithosomes produces a characteristic coarseningupward profile (Figure 10). Half grabens located at some distance from the hinterland areas commonly receive their sediment only from bordering footwall islands. Because these usually are of limited size compared to the volume of the adjacent basins, reduced uplift during the late synrift stage coupled with high erosion rates may produce low-relief or peneplaned fault-block source areas, resulting, in turn, in reduced sediment supply. Continuous high background subsidence rates in concert with limited or diminishing sediment supply therefore produce a characteristic, overall sourceward retreat of any local, transverse depositional system. Depending on the size of the footwall islands, mudstone blanketing would occur either during the late synrift, early postrift, or tectonic quiescence stage. Rift-interior sources appear to be too small to provide sufficient sediment to allow renewed progradation of the fault-slope clastic wedge during the late synrift stage. Late progradation of such systems, with an upward coarsening or shallowing of facies, however, can occur in marginal half grabens or half grabens bordered by large footwall provenances (e.g., Surlyk, 1978, 1984, 1989; Gawthorpe et al., 1990).

response time to the same tectonic event. Fault-related subsidence commonly results in flooding along the axis of the rift basins and across the rifted terrains. Hence, there commonly is a major time lag between the tectonic activity and the sedimentary response from the axial drainage. We suggest that synchronous and in-phase development of highorder stacking patterns likely would not occur. Tectonic Quiescence Signatures The infill related to tectonic quiescent stages also may be wedge shaped, albeit with internal stratal surfaces subparallel to the upper boundary of the wedge. With sufficient sediment supply, this stage is characterized by continued progradation of the riftmargin depositional systems, resulting in complete infilling of subbasins in a successive manner and smoothing of the former dissected rift basin bathymetry. Degradation of source areas, resulting in reduced sediment caliber and supply, may inhibit the filling of distal half grabens, which thus may retain their sediment-underfilled or sediment-starved status for a considerable period. Subsequent rifting would result in renewed retreat of the rift-margin shoreline and prevent sediment supply to subbasins distant to the hinterlands. Such subbasins are characterized by dominantly fine-grained tectonic quiescence packages of hemipelagic or pelagic origin. Synrift Signatures Summarized

Correlation of Footwall, Hanging-Wall, and Axially Derived Successions The relationships of fault-scarp signatures to their contemporary hanging-wall signatures are still uncertain. Synrift stratigraphic simulation suggests that peak hanging-wall progradation in some instances may correlate with fault-scarp retrogradation. A more rapid or immediate yield of coarse sediment from the footwall drainage to the new tectonic activity would be expected due to the steeper faultscarp slopes. Hanging-wall systems may lag behind their footwall counterparts, due to the gentler hanging-wall slopes, the tendency to drown parts of the hanging-wall catchment areas (in partly submergent basins), or the prerift submerged state of the catchment areas (in partly emergent and completely submergent basins). In addition to reflecting different style of accommodation generation and varying drainage and sediment yield developments, hangingwall and fault-scarp successions also may reflect the diachronism of tectonic movements of basin-bounding master faults of two adjacent half grabens. The axial systems are expected to have the longest

Although the three-dimensional variability of the sediment architecture within the subbasins is likely to be great (Figure 15), there have been attempts to generalize about the vertical architectural signature likely to dominate. The three-fold (sandstone-claystone-sandstone) synrift lithosome signature of Prosser (1993) (Figure 1B) is typical in sediment-overfilled and sediment-balanced subbasins. The two-fold (conglomerate-sandstone-claystone) synrift motif (Figure 1A) of Surlyk (1978) is more characteristic of sediment-underfilled and sediment-starved basins. Extreme cases of sediment-starved basins are typified by a thin (condensed) claystone fill, commonly of good source rock quality. Note that in the thin claystone fill, there is usually no pronounced wedge geometry to the resultant synrift infill. The vertical variation shown by the three-fold and two-fold synrift lithology motifs can be linked closely to the evolving tectonic regime. In both motifs the basal coarse clastics represent the rift initiation and early synrift stage, whereas the mud-prone intervals represent the climax of rifting. The sandstone capping of the first motif is argued to represent the late synrift stage. Waning rotational faulting, possibly

Figure 16—Summary of the controlling factors on the development of three-fold, two-fold, or one-fold synrift successions.


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(A)

No scale implied

(B)

No scale implied

(C)

(modified from Underhill, 1991a)

(D)

No scale implied

No scale implied

TECTONIC QUIESCENE PACKAGES ( low tilt rates ) ROTATIONAL TILT PACKAGES ( high tilt rates )

Figure 17—Contrasting stacking patterns of a series of synrift packages separated by tectonic quiescence fill as illustrated by deep-marine synrift successions. Schematic stacking pattern of (A) half grabens with abundant and continuous sediment supply throughout the tectonic phases, (B) half grabens dominated volumetrically by synrotational infill, (C) half grabens dominated volumetrically by tectonic quiescence infill, and (D) sediment-starved half grabens (modified from Ravnås and Steel, 1997).

Ravnås and Steel

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MULTIPLE SYNRIFT PHASES Superposition of Synrift and Quiescence Packages

Divergent fanning of synrift unconformities on hanging wall

(A)

( (a) b)

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(b)

(a) Continuity of synrift unconformities 1-3 across fault-block crest (B) due to modest footwall uplift episodes (clastic wedges 1-3) with intervening periods of subsidence and high sedimentation

1

Figure 18—Schematic illustration of synrift unconformities. (A) Composite synrift unconformity in updip areas which break up into several unconformity strands downdip on the hanging wall (modified from Nøttvedt et al., 1995); (B) continuity of separate synrift unconformity strands across fault-block crest. See text for discussion.

coupled with increased sediment supply, allows for the establishment of transverse and axial drainage systems and the consequent progradation of coarse clastic systems, either shallow or deep marine, across the basin (see also Prosser, 1993). The absence of sandstone capping in the two-fold lithology motif is symptomatic of sediment-underfilled and sedimentstarved rift basins. In these basins, low sediment supply in relation to tectonic subsidence is attributed either to high subsidence rates (i.e., in areas subjected to severe extension) or to limited catchment and sediment yield potential of the subbasins’ source areas (Ravnås et al., 1994). We have chosen to summarize the synrift infill signatures by illustrating the three main cases, as shown in Figure 16, where the initial bathymetry relative to the position of the rotational fulcrum and the distance to hinterland areas are key parameters, at least for the early and climax stage architecture. The synrift infill signatures shown in Figure 16 broadly correspond with sediment-balanced, sediment-underfilled, and sediment-starved cases, respectively. In addition, the timing, duration, and volume of sediment supplied during the specific stages of the rift cycle exert some influence on the resultant lithology motif. Hence, a variety of subdivisions also exists for the three-fold, two-fold, and one-fold synrift lithology motifs, and for the sediment-overfilled, sedimentbalanced, sediment-underfilled, and sediment-starved rift basin types (Figure 16).

Repeated rift phases result in the vertical stacking of two or more synrotational packages, intercalated with intervals showing characteristics more typical of pre- or postrift successions; i.e., tectonic quiescence packages. Figure 17 illustrates schematically a range of possible multiple rotation cases based on data from the Magnus (Figure 17A), Oseberg (Figure 17B) and Inner Moray Firth (Figure 17C) fields. These cases show the development of equally thick or unequally thick superimposed synrift and quiescence packages caused mainly by tectonically induced variability of sediment supply, although factors of inherited bathymetry, rapidity of rotation, and fault-block size are also likely to be important. In either scenario, each motif, whether three-fold or two-fold, shallow-marine or deep marine, relates to a single rift phase. The common evolutionary path observed in many of the northern Viking Graben half grabens that experienced recurrent rift phases is a development from a shallow-marine to a deeper marine basin. This evolution was generally paralleled by a change from a sediment-balanced (or sedimentoverfilled) status to a sediment-underfilled status, and sometimes to a sediment-starved status. This scenario reflects the increasing amount of accommodation, decreasing sediment supply, and increasing cumulative fault-block rotation as the rift evolved and the rifted terrain subsided. The overall synrift signature of the Middle Jurassic–Early Cretaceous North Sea rift episode thus is one of large-scale upward fining. We emphasize the likelihood of multiple rift phases, seen clearly in the Late Jurassic northern North Sea rift system (Underhill, 1991a, b; Rattey and Hayward, 1993; Færseth et al., 1995), although some workers have assumed only a single major event (Prosser, 1993). This point is important for understanding, evaluating, and mapping the development of reservoir geometries and mudstones with high source rock potential within the complex synrift succession. Synrift Unconformities The synrift unconformity represents the main erosion surface formed as a response to continental rifting. This surface has local development over fault blocks and rift shoulders, and forms as a response to footwall uplift and lithosphere unloading due to extension (Kusznir et al., 1991). The unconformity consists of a single unconformity in

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fault-block updip areas, but breaks up into several unconformity strands downdip on the hanging wall (Figure 18, fault block A). In such cases, the individual unconformity strands represent the integrated effect of footwall uplift (of successive rift phases) and eustatic sea level variations. The time represented by the synrift unconformities thus may span the entire rift episode or may be of only short duration. In the case of multiple rift phases associated with only modest footwall uplift separated by intervening periods of basinal subsidence and high sedimentation rate, the successive synrift unconformity strands may be traced in the succession across the entire fault-block crest area (Figure 18, fault-block B). The synrift unconformity may vary in geometry from a surface with considerable relief (Solli, 1995; Nøttvedt et al., 1995) to a nearly peneplaned surface (Ravnås and Bondevik, 1997). In the first case, the erosive development of the synrift unconformity may continue into the succeeding postrift or tectonic quiescent stage, whereas in the latter case the erosive development of the synrift unconformity is inferred to be restricted to the synrift interval. SIGNIFICANCE OF MUD-PRONE INTERVALS Our studies of the North Sea rift basin suggest that there are a series of mud-prone intervals that may develop (Figure 17); with incomplete data, these intervals may be confused with each other or with a true maximum flooding surface. These intervals are important to recognize because they reflect different stages in the rift basin development, and their prediction as potential source rocks is of some significance. (1) In submerged rift basins, there is commonly a time lag between footwall uplift and the creation of (significant) rift-interior sources. In the meantime, increased subsidence results in retreat or cutoff of sediment supply from the rift-margin depositional systems, thereby creating a temporarily sedimentstar ved basin. Thus, initially submerged half grabens commonly experience a period of mud deposition during the early synrift stages (Figure 19B, C), and this interval may mistakenly be correlated with basinwide flooding. Instead, this interval represents initial flooding, and can be correlated with a relative sea level fall in updip areas and a relative sea level rise in downdip areas, respectively. (2) Widespread mud deposition occurs during the rift climax in marginal, sediment-overfilled, and sediment-balanced subbasins (Figure 19A). (3) The progressive drowning of intrabasinal footwall islands in the late synrift (to early postrift) stages commonly results in the most pronounced clay blanketing of sediment-underfilled half-grabens (Figure 19B). This flooding surface is analogous to

the late rift climax blanketing of Prosser (1993). Note, however, that the progressive drowning of footwall islands produces a diachronous claystone lithosome (base is younger updip), and that complete blanketing of basinal areas is dependent on complete drowning of all bordering footwall islands. Deposition of forestepping turbidites may occur in downdip settings contemporaneous with mudstone deposition on updip reaches (Ravnås and Steel, 1997). In such cases this late synrift to early postrift flooding interval also has a diachronous top. (4) In sediment-starved half grabens, sediment supply cut off in the tectonic quiescent stage can result in deposition of organic-rich claystones, which in such settings represent periods of maximum flooding (Figure 19C). (5) In contrast to sediment-starved half grabens, completely submerged half grabens commonly show pronounced mud blanketing during the entire rotational tilt stage (Underhill, 1991a). Based on the distance from the half graben to the basin margin, the tectonic quiescent stage may be characterized either by renewed sediment infilling (Figure 19D) (Underhill, 1991a, b) or continued basinal starvation (Figure 19E). Continued basinal starvation results in a clay-prone interval representing a prolonged time period. In conclusion, we would warn against assigning specific tectonic significance to the mud-prone intervals, unless the half-graben architecture and setting is known. Furthermore, the diachroneity of some of these mud-prone intervals, both within and between subbasins, warns against assigning time-stratigraphic significance to them. This, together with the increasingly acknowledged diachronism of many of the synrift systems tracts (Gawthorpe et al., 1994), throws some doubt on the practice of subdividing the synrift successions into basinwide chronostratigraphic regressivetransgressive packages (Partington et al., 1993; Rattey and Hayward, 1993; Sneider et al., 1995) in a manner more valid for postrift successions. CONCLUSIONS A review of data from Middle Jurassic–Early Cretaceous rift subbasins in and around the northern North Sea suggests that marine siliciclastic synrift architecture is more variable than current models suggest. Marine rift basins represent a continuum from partly emergent through partly submergent to completely submergent basin types. In the Middle Jurassic–lowermost Cretaceous of the northern North Sea and neighboring time-equivalent rift basins, strongly variable synrift sedimentary architectures developed because of temporal changes in relative sea level, accommodation creation,

Ravnås and Steel (A)

(B)

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(C)

Figure 19—Schematic illustration of synrift clay- and mud-prone intervals, and their tectonic significance with respect to the three-stage (early, climax, and late stage) rifting model.

and sediment supply throughout rifting. Accommodation changes were controlled mainly by local basin-floor rotation, basinwide background subsidence, and, to a lesser degree, eustatic changes. Sediment supply determined how much of the accommodation was filled and in what manner, and was controlled by the distance to the main hinterland areas and by the prerift sea level stand. The prerift sea level stand and the amount of basin floor rotation achieved during the rift period determined the regressive-transgressive movement of the faultblock shoreline zone, and the size and sedimentyield potential of any local, fault-block source area. Marine rift basins and synrift successions, whether dominantly shallow or deep marine in nature, are classified in terms of sediment supply as sediment overfilled, sediment balanced, sediment underfilled, and sediment starved. The sequential development, linked depositional systems, and stratigraphic signatures of the early synrift, the rift climax, and late synrift to early postrift stages vary significantly among these four types of marine rift

basins. Sediment-overfilled and sediment-balanced infill types are characterized by a three-fold sandstone-mudstone-sandstone synrift sediment-infill motif, the sediment-underfilled type is represented by a two-fold conglomerate-sandstone-mudstone motif, whereas the sediment-starved type commonly is represented by a one-fold mudstone motif. The tectonic significance of marine clay-prone intervals varies between the different infill types, from representing the rift climax in the sedimentoverfilled and sediment-balanced variants to the late synrift or subsequent postrift or tectonic quiescent stage in the sediment-underfilled and sediment-starved variants. In contrast, local, faultblock synrift unconformities are correlated with periods of maximum uplift rates during successive rift phases (i.e., the rift climaxes), but nevertheless represent different initiation time and duration within the varied rift basin types. Because rift basins commonly show spatial diachronism in their tectonic minima and maxima, this throws doubts on the use of unconformities and marine condensed

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ABOUT THE AUTHORS Rodmar Ravnås Rodmar Ravnås is senior geologist with Norske Conoco A.S. in Stavanger, Norway. Ravnås holds a Cand. Scient. (1991) and a Dr. Scient. (1996) from the University of Bergen, Norway. Previously he worked as an exploration geologist with Norsk Hydro.

Ron J. Steel Ron Steel is Wold Professor of energy and professor of geology/ geophysics at University of Wyoming, and was previously professor of sedimentology and reservoir geology at the University of Bergen, Norway. He graduated from the University of Glasgow, where he also obtained a Ph.D. His main interests are sequence stratigraphy and clastic depositional environments, particularly as applied to reservoir geology and basin analysis.

Marine Synrift Sedimentation Ravnas & Steel AAPG 1998

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