Styles of soft-sediment deformation on top of a growing fold systemin the Gosau Group at Muttekopf, Northern Calcareous Alps,
Austria: Slumping versus tectonic deformation
Hugo Ortner ⁎
Institut für Geologie und Paläontologie, Universität Innsbruck, Innrain 52, 6020 Innsbruck, Austria
Abstract
Styles of soft-sediment deformation in a syngrowth coarse-grained sediment gravity flow system are described from outcrops ofthe Upper Gosau Subgroup of the Northern Calcareous Alps. The deepwater sediments were deposited along the flanks and top ofgrowing folds. Beside load and fluid escape structures, which are commonly found in sediment gravity flow systems, soft-sedimentdeformation is mainly related to fluidization of coarse-grained beds and downslope gliding of meter-thick sediment packages.Common structures affecting single or several beds are: hydroplastic folding of fine-grained beds in a coarse-grained fluidizedmatrix, symmetric or asymmetric mullions at the base of coarse-grained beds, and folding of heterolithic units into coarse-grainedbeds. Folds have varying geometries and styles, and fold axes are scattered. Soft-sediment folds also affect thick sedimentpackages. These folds are uniformly verging asymmetric metric to dekametric folds, formed at different degrees of lithification.Rheology of sandstones in such folds ranges from ductile to semi-brittle and brittle. Fold axes tightly cluster about one direction.The first group of deformational structures is attributed to slump-related deformation, whereas the second group represents foldsrelated to tectonic deformation.
Keywords: Soft-sediment deformation; Tectonic deformation; Slumping; Growth strata; Turbidite deep water; Northern Calcareous Alps; Austria
1. Introduction
Dating of deformation has been an important issue instructural geology in the last few years, and importantprogress has been made in dating of deformation in meta-morphic rocks by dating single grains that grew syn-kinematically (e.g. Müller et al., 2000). Dating ofdeformation in shallower levels of the crust is not possible
by geochronologic methods, as temperatures are too lowfor recrystallisation. If sedimentation is contemporaneouswith deformation, the maximum and minimum age offault-related deformation can be defined: the maximumage of deformation is given by the age of the sediment incontact with a fault, the minimum age by the age of thesediment overlapping a fault. Direct dating of deformationis possible in sedimentary successions if larger scalegeometries such as progressive unconformities are present(Riba, 1976; Anadon et al., 1986; Medwedeff, 1989; Fordet al., 1997; Suppé et al., 1997). In these cases, the age of
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the sediment, determined by paleontological methodsgives the age of deformation. Sedimentary successionsdeposited on growing folds will be subject to (1) loadingand dewatering intrinsic to deposition, (2) gravity-drivenmass wasting processes such as slumping due to progres-sive rotation of fold limbs in addition to (3) tectonicdeformation. Because both tectonic deformation andslumping affect unlithified or poorly lithified sediment,similar small to medium scale structures will form. Thispaper presents field observations on soft-sediment de-formation from deep water sediment gravity flow depositsand discusses their origin, with the goal to define criteriafor the distinction between slump-related deformation andtectonic deformation in the field. The unusual wide rangeof grain sizes in a wide variety of gravity flow depositsfrommegabreccias tomarls in a growth strata setting makethe Gosau Group at Muttekopf an excellent naturallaboratory for characterizing and testing the ability todiscriminate gravity versus tectonically driven processesand products.
1.1. Geological setting
This paper deals with Upper Cretaceous sediments ofthe Gosau Group at Muttekopf in the western part of theNorthern Calcareous Alps of the European Eastern Alps.The Gosau Group is an Upper Cretaceous, synorogeniccarbonatic–siliziclastic sedimentary succession, whichunconformably overlies deformed Triassic to Jurassicrocks (Wagreich and Faupl, 1994; Sanders et al., 1997).Generally, deposition started in a terrestrial environment,which subsided to neritic conditions (Lower Gosau Sub-group). After a pronounced subsidence event deep ma-rine conditions prevailed (Wagreich, 1993), and the UpperGosau Subgroup was deposited. The relationship betweenthe contracting orogenic wedge and the coeval major sub-sidence is not well understood at present, and differentmodels have been put forward (e.g. Wagreich, 1993;Froitzheim et al., 1997).
The younger, deep marine part of the sedimentarysuccession (Upper Gosau Subgroup;Wagreich and Faupl,1994) was deposited during transport of the thin-skinnednappes of the Northern Calcareous Alps over tectonicallydeeper units (Fig. 1). In the Muttekopf area, internaldeformation of the moving nappe led to the formation offault-propagation folds in the subsurface of the Gosausediments and hence to formation of several (progressive)unconformitieswithin the sedimentary succession (Ortner,2001; Figs. 1–3). The Upper Gosau Subgroup is dividedinto three sequences (Ortner, 1994, 2001; Sullivan et al.,2006; Fig. 2): All three sequences of the Upper GosauSubgroup are dominated by vertically-stacked, upward-
fining, laterally continuous, unchannelized conglomeratesand sandstones that display little to no lateral variation infacies. The boundary between Sequence 2 and 3 is theRotkopf unconformity (Figs. 1–3). The boundary betweenSequence 1 and Sequence 2 is the base of the 2nd fining-upward sequence, which is significantly below the mostprominent unconformity in the area (Schlenkerkar un-conformity) (Figs. 1, 2). In Fig. 2, Sequence 2 includesboth offlapping stratal geometries, suggesting that the rateof structural growth was greater than the rate of depositionand progressive onlapping stratal geometries, suggestingthat the rate of deposition was greater than the rate ofstructural growth.Due to limited preservation of Sequence3, its stratal geometries are less clear. Sediment transportdirections are either parallel to the regional fold axial trendand are southwest or northeast-directed, or perpendicularto the structural trend and north-or south-directed (Ortner,1993, 1994). The direction of dip of the paleoslope will bereevaluated in the discussion (Section 3.2).
The style of tectonic deformation changes fromwest toeast: in the western part of the outcrop, the Gosau sed-iments are preserved in the core of a km-scale syncline(Muttekopf syncline), which is interpreted to be the lead-ing syncline of a fault-propagation fold formed above adeep-lying blind thrust (Fig. 2). Further east, the southernpart of the Gosau outcrop is characterized by dekametricfolds, which formed above minor blind thrusts which arepartly exposed (Fig. 3). Structural style changes acrossNW-striking tear faults (Fig. 1). Fold growth of the mainfault-propagation fold was syndepositional and is docu-mented by a progressive unconformity in the southernlimb of the leading Muttekopf syncline (Schlenkerkarunconformity of Figs. 1, 2). Growth of the fold north ofthe Gosau outcrop, which led to southward tilting of thenorthern limb of the Muttekopf syncline is mainly post-depositional (Ortner, 2001), only minor stratal conver-gence can be seen in Fig. 3. Fold growth in the zone ofcomplex folding at the southeastern margin of the Gosauoutcrop east of Rotkopf (Fig. 1) was also syndepositionaland led to rotational onlap at the Alpjoch unconformity(Figs. 1, 3).
1.2. Introduction into soft-sediment deformation
Soft-sediment deformation structures are common tomost sediment gravity flow systems. Gravity flows aremixtures of water and sediment and are often rapidlydeposited due to flow collapse. This rapid depositiontraps interstitial fluid in the sediment and results in high,unstable pore pressures. Most soft-sediment deforma-tion is related to fluid-escape due the squeezing out ofexcess water from the pores of a sediment during
Fig. 1. Geological sketch of the study area. Inset a: Position of the Muttekopf Gosau outcrop in the Northern Calcareous Alps and Austria. Inset b: Nappe units of the Eastern Alps in the wider vicinityof the investigated area. 1: Schlenkerkar unconformity, 2: Rotkopf unconformity, 3: Alpjoch unconformity. Modified from Ortner (2001).
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Fig. 2. View of the Galtseitejoch and Muttekopf from the Kogelseespitze in the west (see Fig. 1 for local names). The main unconformities, which arethe progressive Schlenkerkar unconformity (Ortner, 1994), and the Rotkopf unconformity (Wopfner, 1954), are shown. Sedimentary units thin to thesouth, which records progressive tilting of the southern limb of the main syncline. “1”marks two conglomerate layers which are thickest in the core ofthe syncline and thin toward the limbs.
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compaction. The style of soft-sediment deformationstructures changes with progressive burial, going alongwith a change of the main controlling factors.
In unconsolidated sediments near sediment–water in-terface pore pressures can equilibrate by either gradual orrapid fluid escape. Rapid dewatering in many cases is
Fig. 3. Panoramic view of the eastern part of the Gosau outcrop from the Platteinwiesen in the west. Note the zone of complex folding in the southwestern part of the Muttekopf outcrop, related tominor thrusts reaching into the sedimentary succession.
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triggered by rapid deposition of a dense sediment (e.g.sand) over a less dense, water-rich sediment (e.g. mud),creating a reversed density stratification. Under theseconditions pore water pressure rises above overburdenpressure, the underlying low-density sediment is fluidized(see below) and intrudes the overlying unconsolidatedsediment. Structures created by such a process are loadcasts and flame structures (ball-and-pillows and convolutebedding; Anketell et al., 1970; Lowe, 1975; Visher andCunningham, 1987). Dewatering in the shallow subsur-face triggered by seismic shocks is another reason for theformation of dewatering pipes (Owen, 1995). If sedimentdewaters gradually near the sediment–water interfaceunder its own weight, pillars and dish structures form(Lowe and LoPiccolo, 1974; Lowe, 1975).
A more general description of soft-sediment defor-mation, which can also be applied to structures related tosubsurface sediment mobilization was given by Knipe(1986). He described deformation type as a function ofthree major parameters: deformation ratio, degree oflithification and fluid pressure. Three types of deforma-tion were defined: independent particulate flow, diffu-sion mass transfer and cataclasis.
(1) independent particulate flow describes the behav-iour of non-cohesive or poorly lithified sediments. De-pending on the ratio of excess pore fluid pressure tocohesive strength due to grain weight, three deformationalprocesses are defined: hydroplastic deformation, lique-faction and fluidization.
(A) Hydroplastic deformation occurs when fluid pres-sure is lower than grain weight and preserves butmodifies primary sedimentary structures such asbedding. Many structures created during hydro-plastic deformation resemble plastic deformationin metamorphic rocks, and can be described andinterpreted using the terminology of structuralgeology.
(B) When fluid pressure is equal to grain weight, thesediment liquefies and is subject to laminar flow.Primary sedimentary features such as bedding aredestroyed.
(C) When fluid pressure is higher than grain weight, orfluid velocity is high enough to entrain grains in thefluid, the sediment fluidizes and is subject to tur-bulent flow. All primary sedimentary structures aredestroyed. Fluid pressure increases when low-permeable sediment seals a high-permeable sedi-ment below and prevents fluid escape. Fluid pres-sure then approaches lithostatic pressure (Jolly andLonergan, 2002). If fluid pressure exceeds bothtensile strength of the seal and the principal hori-
zontal stress, which is the minimum principal stressan undisturbed system, a clastic dyke forms. Incases when the principal vertical stress is smallerthan the tensile strength of the seal, and fluid pres-sure is intermediate, a sill forms.
(2) Diffusion mass transfer and (3) cataclasis werealso observed in partly lithified rocks, but do not play animportant role in this study.
2. Description of soft-sediment deformationstructures
The following section describes centimetric-to hec-tometric deformation features observed in the field andprovides a short interpretation. Deformation structuresthat formed near the sediment–water interface aredescribed first, followed by structures that formedwith greater sediment overload. Common soft-sedimentdeformation structures such as load casts, which areabundant in the study area, are not described.
2.1. Structures affecting single or few beds
2.1.1. LoadingOn top of deformed sand–mud couplets, elongate
troughs filled by coarse-grained sediment were observed(Fig. 4). When tracing the sand-beds of the sand–mudcouplets, it appears that the thin sandstone beds slightlydiverge when moving from below a trough to the crestof a rise (Fig. 4). Moving upsection in the thin-beddedunit, amplitude of the waves increases from zero to amaximum of about a meter, and locally the tops of therises have teepee-like geometry, or, in other words, thegeometry of upward pointing cusps (Fig. 4a). Thesediment fill of the troughs is horizontally laminated,and the laminae onlap the lateral margins of the trough(Fig. 4b). Toward the top, the coarse-grained unit gradesto sand, and the top of the bed is flat (Fig. 4a). The longaxis of the troughs is oriented almost perpendicular tothe sediment transport direction indicated by flute castsin neighbouring sandstone beds (see diagram in Fig. 4a).
The flat top and the sideward onlap of laminae of thecoarse-grained bed suggest that this sedimentary unitsfilled preexisting topography. Undeformed horizontallaminations in the bed preclude loading by the coarse-grained material, and suggest the troughs should haveexisted before deposition of the coarse-grained unit. Onthe other hand, the absence of a basal glide plane, abovewhich folding of the sand–mud couplets could havetaken place, and progressive upward development of thewavy folds by divergent and convergent bedding
Fig. 4. Unusual load structures. a: Wavy base of a bed graded from fine conglomerate to sandstone, filling a local topography on top of a heterolithicunit, while the top of the bed remained flat. b: Single conglomerate-filled trough with horizontal lamination, which onlaps rises in the heterolithic unitbelow. Bedding planes in the heterolithic unit diverge away from the trough axes in a and b. All orientation data on outcrop-scale in this paper wererotated to horizontal with bedding about bedding strike. All plots and operations were done using the software Tectonic VB (Ortner et al., 2002).Unless not indicated otherwise, fold axis orientations in diagrams are shown by points and a rose diagram in the same plot.
indicates differential compaction and therefore loading.A possible interpretation of the features observedsuggests initial differential loading by sand probablyorganized in sand waves, which caused the wavy surfaceobserved. Higher current velocities during a secondevent enabled the current to remove the sand and depositthe gravel onto the existing topography.
2.1.2. Fluidized coarse-grained bedsA significant part of the sedimentary succession of
the Upper Gosau Subgroup of the Muttekopf is built bymeter to several meter-thick conglomerate beds. Usual-ly, these beds exhibit little internal organization.Frequently, clasts of finer grained sediments (coarsesandstones) float in the unorganized conglomerate (Fig.5a,b). Locally, conglomerate beds show fluid escapestructures (flame structures) into overlying sandstones.Fig. 5a shows a conglomerate flame structure that hasvented through a 10 cm sandstone bed to feed a bed-parallel injection (sill). Conglomerate intrusions mayalso have a diapiric form (Fig. 5d) with vents thatcrosscut sandstone beds and indicate fluidization.Overlying sandstone beds are only weakly deformed,and horizontal lamination preserved, indicatinghydroplastic deformation of the sandstone. Where aninitial continuous grading from conglomerates tosandstones and siltstones was present in a bed, onlythe central part of the bed was fluidized. The upper andlower boundaries of fluidization are not strictly parallelto bedding. In Fig. 5b, the boundary between thefluidized unit and overlying siltstones follows planes ofa conjugate joint set, which are related to layer-parallelshortening, because the acute angle between conjugatejoints contains the bedding plane orientation.
Two important conclusions can be drawn from theseobservations: (1) Fluidization of the conglomerate bedoccurred after burial of the now fluidized bed belowyounger sediments. Increasing overburden pressure du-ring ongoing sedimentation may have contributed to riseof pore fluid pressures needed for fluidization of con-glomerates (compare Lowe, 1976). (2) As grain weightin the conglomerate beds was supported by moving porefluid during fluidization, and the sandstone bed with itsmuch smaller and lighter grains was not fluidized, theratio between grain weight and pore fluid pressure is notthe sole control on fluidization. Sandstones were prob-ably cohesive due to a higher content of clay minerals ordue to incipient cementation or both.
Fluidized coarse-grained beds can have varyingamounts of incorporated fine-grained beds. The sand-orsiltstone beds are disrupted and folded, but internalstructures like grading and lamination are well preserved(Fig. 6a). The organization of the sandstone beds withinthe fluidized layer depends on the amount of coarsematrixand on the size of the components floating in the matrixrelative to bed thickness:
(1) Where large rafts of fine-grained sediment bedsare found, shingle-like stacking of sandstone bedsis observed beside folding (Fig. 6, fluidized layer1). The poles to bedding of sandstone beds arearranged in a girdle perpendicular to the sedimenttransport direction, and parallel to the downslopedirection, in the case the folds within fluidizedlayer 1 facing south indicate the downslope direc-tion. Many folds are isoclinal, and are refolded byopen folds (Fig. 6). Fold axes (Fig. 6d) are dis-persed about a WSW trend, and most are parallel
Fig. 5. Evidence for fluidization of conglomerate beds. a: Flame structures at the interface between sandstone and conglomerate projecting into thesandstone. To the right conglomerate was injected upwards into sandstones forming a bed-parallel sill. Clasts of coarse sandstone give evidence ofmixing in an originally continuously graded bed. Length of hammer head 20 cm. b: conglomerate dyke injected into mudstone, which is a large clastwithin the conglomerate. c: fluidization in a bed continuously graded from conglomerate to siltstone. Fluidization is restricted to the central part of thebed, where silt-and sandstone clasts float in a conglomerate matrix. The capping siltstone is cut by conjugate joints. Joint surfaces delimit the fluidizedunit against the siltstone. d: conglomerate intrusion into coarse sandstone coupled with large dish structures within the sandstone. Scale is 7 cm.
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to bedding. Axial planes of folds are inclined withrespect to bedding, and most of them strike WSW,parallel to the mean axial trend (Fig. 6b).
(2) Where fine-grained material is broken up to clastsmuch smaller in size than the thickness of thefluidized layer, such as the one shown in Fig. 6,fluidized layer 2, folds are near isoclinal, and axialplanes are all near parallel to bedding (Fig. 6g).The fold axes scatter around the slope dip direction(Fig. 6f ) and can be strongly curved (sheath folds).
(3) Wherematrix dominates the fluidized bed, fold axeswithin the isolated clasts are randomly oriented.
In all cases, folded sandstone beds show parallel foldstyle, irrespective of the interlimb angle. The mecha-nism of folding was flexural folding as shown by theoccurrence of mullions in the inner bend of isoclinalfolds where more competent sandstone is in contact withmudstone, and minor normal faults at the outer bend(Fig. 6a). No similar style folds were observed in flu-idized layers.
Incorporation of sandstone beds into fluidized con-glomerates was probably a result of conglomerate intrusion
into overlying and underlying sandstone beds. Coalescingof neighbouring conglomerate sills would isolate parts ofsandstone layers and incorporate them into a fluidized bed.Intrusions across several sandstone layers or multiple in-trusion events could produce sandstone-rich liquefied lay-ers. Once incorporated, the sandstone beds could probablybe further disrupted by intrusions across layers, as de-monstrated by Fig. 5c, which shows part of a conglomeratedyke crosscutting a mudstone clast-within the conglomer-ate layer.
Fig. 6 shows the upslope end of a fluidized layer (rightend of fluidized layer 1). There, a thrust plane cutsupsection and ends below fluidized layer 1. Immediatelyabove the thrust tip within the fluidized layer, a fold in apackage of sandstone layers faces approximately south,and is covered by a conglomerate rapidly wedging outlaterally. Fluidization lasts only for very short time in-tervals (e.g. Lowe, 1976). In fluidized state, sedimentsshould move downslope. Any mass transport on a slopehas an upslope part, from which material is transportedtoward the downslope part, and the upslope part should besubject to vertical thinning and horizontal extension,whereas the downslope part should experience shortening
Fig. 6. Key outcrop of two fluidized beds near Muttekopfhütte. Fluidized layer 1 formed at the tip of a minor thrust cutting upsection from right to left. At the right end of the outcrop, fluidized layer 1 isrefolded by semi-brittle soft-sediment folds, which are depicted in detail by in Fig. 9a. a: isoclinal fold within fluidized layer 1. Grading and lamination in sandstone bed is well preserved. b: poles to beddingof deformed sandstone bed within fluidized layer 1 and calculated fold axis. c: axial planes of folds in deformed sandstones within fluidized layer 1 and calculated fold axis. d: trend and distribution of small-scale axes of folds in deformed sandstones at fluidized layer 1. e: sediment transport directions from flute casts and tool marks at the base of sandstones and conglomerates. f: axial trends and and distributionof small-scale folds in deformed siltstones within fluidized layer 2. g: deformed heterolithics and injected sandstones in fluidized layer 2. h: sketch of upramping of a unit during fluidization.
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and thickening. The observation of folding and henceshortening at the upslope end of a fluidized layer is un-usual. Fluidizationmight be a result of initial movement atthe thrust by compressing the pore water in a layer sealedat the base and the top (compare joints of Fig. 5b). Porewater pressure would then reach the critical value forfluidization. Because the wedging conglomerate layer atthe top is not folded, the fold at the upper end of thefluidized layer must have been created during movementof the thrust and during fluidization. In surficial slumps,such a relationship between downslope gliding of co-herent units, which is described in Section 2.2.4, asso-ciated with fluidization of the lower part of the slump wasobserved by Loope et al. (2001). Normal faulting in theunits on top of fluidized layer 1 might be a result ofdownslope gliding and associated stretching of the over-lying sediment package during fluidization.
The thickness of fluidized layer 1 increases in thedownslope direction (Fig. 6). Decrease of thickness in afluidized layer of another outcrop is associated with averging fold in the sandstone layer sealing the fluidizedunit at the top, giving evidence of gliding during flui-dization (Fig. 6h). Where the sandstone layer was dou-bled, the sandstone projecting into the fluidized unit wasassimilated into the fluidized unit.
2.1.3. Hydroplastic deformation at the base of coarse-grained beds
Where a sandstone beds underlies a coarse-grainedbed, the interfacewith fine-grained deposits such as sand–mud couplets is usually undeformed (Fig. 6). However,where fine-grained sediments are in direct contact withconglomerates, the contact is deformed, as is demonstrat-ed by a conglomerate layer with a boudinaged sandstonebed at the base (Fig. 7a). Common deformational struc-tures at such contacts are decimetric to metric elongateflame structures of silt-and mudstones into conglomer-ates. Flame structures can be isolated, with distances ofseveral meters from one flame to the next, or closelyspaced. In the latter case, elongate troughs are developedbetween individual flames. In many cases, these troughsand associated flames are symmetric and slightly irregular(Fig. 7a), however in other cases, both are highly asym-metric and consistently verge in one direction (Fig. 7c,e).Unlike in Fig. 5a, where flames show evidence for fluidescape from conglomerates into overlying sediments, thevergent flames developed at the base of conglomeratelayers rather resemble folds (Fig. 7f ). Centimetric alter-nations of silt-to mudstones are folded in open to isoclinalfolds into the conglomerate.
Where flames are asymmetric, a glide horizon wasobserved within the silt- to mudstones, and the sediments
above the glide horizon are deformed (Fig. 7c,e,f). Direct-ly above the glide horizon, isoclinal folds with axialplanes subparallel to bedding are abundant, and fold stylein deformed sand-and siltstones is similar. The isoclinalfolds are frequently boudinaged, and resemble intrafolialfolds as observed in metamorphic rocks. “Flames” aredeveloped, where shear zones branch off the glide planeparallel to bedding and reach into conglomerates. Foldsassociated to the latter shear zone are not boudinaged,only locally the steep limbs of verging folds are sheared.
For interpretation, the observed structures at the base ofconglomerate beds must be split up into several subtypes.Boudinage of a sandstone at the base of a conglomerateindicates stretching of the sandstone and conglomeratebeds and boudinage of the sandstone due to higher vis-cosity of the sandstone. Initial boudinage is probably aresult of downslope gliding and stretching during initialfluidization of the conglomerate layer, as indicated bylocal conglomerate intrusions (Fig. 7c). Where the sand-stone was removed due to boudinage, conglomerate cameinto contact with fine-grained sediments.
Loading of fine-grained sediments by the conglom-erate could create load casts. Loading as a consequenceof superimposing a dense plastic or fluid layer on asubstratum of lower density should create balls-and-pil-lows separated by flames arranged in a polygonal pattern(Owen, 2003). If loading was contemporaneous withdownslope translation, elongate ridges of low-densitysediment projecting into high-density sediment parallelto transport direction would be expected (e.g. Anketelland Dzulynsky, 1968; Moretti et al., 2001). Dzulynsky(1966) suggested convection-like particle movementsduring transport of sand deposited by turbulent currents,to explain ridges parallel to transport direction found atthe base of beds. For conglomerates and breccias, whichwere transported and deposited by debris flows or high-density turbidity currents, such structures were notdescribed.
As the flames and balls do not seem to be related tosedimentary processes (see above) and are perpendicularto sediment transport direction (Fig. 7a), an alternativeinterpretation is needed. Folding of the interface betweenconglomerates and fine-grained sediments would be apossible explanation. A prerequisite for this interpreta-tion is that both lithologies deform (hydro-) plastically,and hence it has to be assumed that the structures post-date liquefaction. The geometry of the folded interfacecan be described as cuspate–lobate, with the cusps de-veloped in the fine-grained material and the lobes in theconglomerate. In structural geology terms, such foldedinterfaces are called mullions (e.g. Urai et al., 2001).According to these authors, “mullions are described as
Fig. 7. Hydroplastic deformation structures at the base of coarse-grained beds (“mullions”). a: Irregular elongate ridges at a contact between conglomerateand mudstone. Near the right margin of the photograph, the end of a boudinaged sandstone bed with flute casts at the base is visible. b: Regular elongateridges at a contact between conglomerate andmudstone. c:Deformed ridges at a contact between conglomerate andmudstone. From the conglomerate bed,a sill was injected into underlying sandstones. d: location of the documented deformation structures with respect to a syncline. For location see Fig. 3. e:Deformed ridges at a contact between conglomerate andmudstone. f: Structure of an isolated “cusp”. The angle between the glide plane and the axial planeof individual small folds, increases upsection. Near the glide plane, the small structures resemble intrafolial folds. All the structures are related to a shearzone at the base of the conglomerate.
regular cuspate–lobate folds of an interface between twomaterials with a large competence contrast, the cuspspointing towards the more competent material“. At thetime of formation of the mullions, conglomerates musthave been more competent than the silt- and mudstonesbelow.
As the mullions are more or less symmetric, theyrecord pure shear shortening parallel to bedding. Where
asymmetric mullions are found (Fig. 7c,e), either forma-tion of mullions by simple shear must be assumed, orinitially formed symmetric mullions were overprintedby simple shear. Fig. 7f shows an example of an isolated“cusp” into a conglomerate bed. There, a shear zonebranches off the basal glide plane of the conglomeratebed. Isoclinal folding with axial planes parallel to thebasal shear zone and subsequent boudinage of isoclinal
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folds can be seen as an effect of progressive simpleshear. Initially, folds above the glide plane formed withSE-dipping axial planes, but were progressively rotatedinto parallelism with the shear zone and then extended.
The strain histories of individual conglomerate layerscan be complex, as seen from the descriptions and inter-pretations above. Initially, conglomerates became stretch-ed due to downslope gliding on top of a glide plane whichremained active throughout early deformation of the bed.Then bedding parallel shortening became active leadingto formation of symmetric or asymmetric mullions. Themechanism determining the type of mullion developedremains unclear at the moment. As can be seen in Fig. 7d,the position of the outcrop in the syncline is not thecontrolling factor, because different types of mullions arefound in the same bed in one limb of theAlpjoch syncline.
2.1.4. Hydroplastic deformation within coarse-grainedbeds
The main difference between units deformed by flui-dization and units affected by hydroplastic deformation isthe absence of disruption of sandstone beds and the ab-sence of pervasivemixing of conglomerates in the latter. InFig. 8, the lateral margin of a graded breccia bed affectedby slumping is depicted. Several isoclinal folds of a het-erolithic unit consisting of sand–mud couplets originallyunderlying the breccia reach into the breccia bed. The axialplanes of the folds are subparallel to bedding and the axesof the folds are strongly curved. Folds have sheath foldgeometry (e.g. fold in the center of the Fig. 8, orientationsof fold axis represented by squares in Fig. 8c). Anasto-mosing shear zones within the conglomerate bed originateat the hinge of isoclinal folds.
The isoclinal fold in the center of Fig. 8 roots in thefolded zone at the right side (diagram Fig. 8d). There-fore the sequence of deformations resulting in the ob-served structure is probably as follows: Initially, anasymmetric mullion as described in Section 2.1.3 de-veloped. When the hinge of the progressively amplify-ing verging fold reached the interface between coarsebreccia and fine breccia, which acted as a rheologicinterface, further fold growth was guided by the rheo-logic interface. Probably the mechanism of fold growthchanged at the time the hinge zone reached the rheologicinterface from folding by subhorizontal shortening andprogressive tilting of the fold limbs to gliding of the unitabove the rheologic interface and rolling the heterolithicbeds attached at the base through the hinge zone of theisoclinal folds. Taking the axis and the facing directionof the conical fold in the heterolithic unit at the right sideof Fig. 8 as an indication for direction of slumping, theslump moved to the southeast. Fold axes measured in
the isoclinal folds within the slumped bed are dispersed,but scatter about the direction of slumping (Fig. 8c).
2.2. Structures affecting thick sediment packages
In contrast to the structures described up to this point,which affected single layers of the sediment succession, thefolds which will be described in the following paragraphsreach over sediment stacks of 10 to 100 m thickness.
2.2.1. Metric foldsAt the left side of Fig. 6, a train of folds deforms
fluidized layer 1. Fig. 9 depicts one of the folds belowfluidized layer 1 in detail. All folds in the outcrop areasymmetric and face N. The axial planes of these folds areparallel and inclined with 25° against bedding (Fig. 9a).Fold axes of the folds are well aligned and trend E–W(Fig. 9b). Axial planes of folds can be traced across 3–4coarse-grained layers (Fig. 6, lower left). Sandstone bedsare frequently faulted and stacked, and on the fault planesstretched calcite fibers indicate the movement sense (Fig.9c). Some faults were refolded in the duplex in the steeplimb of the fold. Locally individual sandstone beds orlayers within thicker sandstone beds are deformedplastically (“plastic deformation” in Fig. 9). In the coreof some folds, s–c-structures developed in mud-to silt-stones. In other cases, symmetric cuspate–lobate folds arepresent in sandstones of the steep limbs of folds, whereasthe flat limbs are boudinaged (Fig. 10a,b). In the core ofthe folds, a foliation parallel to the axial plane is present(Fig. 10c). In continuation of anticlinal hinges into anoverlying conglomerate, anastomosing shear zones in theconglomerate are present (Fig. 10a, center).
Faulting and stacking of sandstone beds indicate thatshear strength of sandstone was high enough to preventplastic folding comparable toFig. 10, but not low enough toprevent multiple faulting in sandstones. Compared to foldsdescribed up to this point, lithification is more advanced,but not complete, and local plastic folding gives evidenceof inhomogeneous lithification. Compared to fluidization, ayounger relative age for the folds described here can bededuced due to refolding of fluidized layer 1 in Fig. 6.
2.2.2. Dekametric foldsOn a hectometric scale, axial planes of folds in the
northern face of Pleiskopf can be traced across 200 m ofthe sedimentary succession (Fig. 11). Where these largefolds are present, the heterolithic units between the thickconglomerate beds show chevron fold style, indicatingfolding by flexural slip. However, the hinges of foldswithin conglomerates are thickened, indicating thatmaterial flowed toward the hinges. The contacts between
Fig. 8. Hydroplastic deformation in a conglomerate bed. A heterolithic unit is repeatedly isoclinally folded into a conglomerate bed. To the right, thelateral margin of the gliding unit is characterized by conical folds, where the heterolithic unit was dragged at the slump edge.Movement of the slump istoward SE, away from the observer. a: poles to bedding of the folded heterolithic unit and the calculated fold axis. b: axial planes of folds within theheterolithic unit. c: fold axes in the heterolithic unit, represented also by rose diagram. The squares record the changing orientation of the sheath foldaxis of the fold in the center of the figure. d: poles to bedding planes and calculated fold axis in the zone of conical folds at the lateral margin of theslump.
the heterolithic unit and conglomerates at the inner bendsof the large folds show metric mullions comparable tothose of Fig. 7b, but the sandstones in the heterolithic unitdo not show any evidence of plastic deformation.
The dekametric folds described here are part of thetight fold train at the southern margin of the MuttekopfGosau outcrop. As can be seen in Fig. 3, these foldsformed above the tip lines of minor thrusts reaching into
the sedimentary succession. During folding, the sand-stones were completely lithified, but the conglomeratesstill behaved plastically.
2.2.3. Normal faultsSynsedimentary faults are a widespread feature in the
Gosau Group of the Muttekopf area and occur at allscales of observation, from the hectometric to
Fig. 9. Semi-brittle metric vergent soft-sediment fold, also depicted inFig. 6, lower left. Multiple stacking in the sandstone bed probably dueto reduced shear strength during ongoing lithification. Note plasticdeformation (p. d.) in lower sandstone bed and near the core of theupper fold. a: axial planes and b: Fold axes of all metric vergent soft-sediment folds in the outcrop (see Fig. 6). c: Calcite slickensides onfault planes stacking the sandstone bed.
112 H. Ortner / Sedimentary Geology 196 (2007) 99–118
centimetric scale. Some of the major faults can be seen inFig. 3 and in Fig. 11. Typically, offset across the majorfaults decreases upsection and downsection. Offsets ofup to 100 m were observed. Thickness changes ofsedimentary units occur across these faults, but brecciashave never been found to be associated with such faults.Fault orientations are both parallel to the regional foldaxis and perpendicular to it.
As the faults lack breccias and only thickness var-iations are found, they probably never created a majorfault scarp. Any surficial depression was immediatelyfilled by sediment. The orientation of the faults is gen-erally parallel to the interpreted dip of the slope, with thehangingwall moving down in the downslope direction.
2.2.4. Downslope gliding of sediment packagesThe only possibility to document downslope gliding
of sediment packages is to find locally developed rampsat the base of the gliding units (Fig. 12). In most areas,the units move on bedding planes in heterolithic units,
and these glide planes are difficult to recognize. In theexample shown in Fig. 12, a ramp is developed at thebase of a gliding unit, stacking a conglomerate bed. Thecoarse-grained unit on the top of the translated hetero-lithic unit fills the topography created by gliding. Thegeometry of the ramp at the base of the gliding unitsuggests NE-directed movement of the hanging wall. Inthe example shown in Fig. 12, gliding was active undera sediment stack of 6 to 7 m thickness.
3. Discussion
3.1. Slump folds versus tectonic folding
3.1.1. Kinematic model for slumpsFarrell (1984) and Farrell and Eaton (1987) have pro-
posed a dislocation model for the movement of slumpsheets. According to the model, movement and deforma-tion of a surficial slump sheet on a basal failure of finitesize is spatially organized (Fig. 13a), assuming continu-ous strain and a layered plastic or brittle material havingbeds with differing competence which get folded. Thelower part of the slump gets shortened, while the upperpart of the slump gets extended. Initial deformation in thelower half of the slump will cause bedding-parallel short-ening creating fold axes perpendicular to the downslopedirection, whereas extension in the upper part of the slumpcauses boudinage and normal faulting. Further glidingwill amplify the existing folds and rotate them progres-sively toward the downslope direction. When the basalfailure propagates downslope, new folds with axes par-allel to the tip lines of the glide plane will form. Strainoverprinting is achieved by arresting slump movement atthe toe and shifting the area of active thickening upslumpinto the area of former extension, or by arresting slumpmovement at the head and shifting the area of activeextension downslump into the area of former shortening.In both cases, complex strain histories will be restricted tothe central part of the slump, where the basal glide planehas maximum offset.
In the dislocation model the orientation of fold axeswithin the slump body is a function of themaximumoffsetacross the basal glide plane relative to the size of the glideplane. Topography of natural examples of surficial slumpsis in accordance with the dislocation model on small andlarge scale (e.g. Lajoie, 1972; van Weering et al., 1998;Bøe et al., 2000). In the example documented by Lajoie(1972), folds in the shortened part of the slump are ofdezimetric scale, whereas the maximum offset across thebasal glide plane is about 10 m and therefore large com-pared to the dimension of the folds. Most of the fold axesare parallel to the dip of slope (Fig. 13b).Whenmeasuring
Fig. 10. Metric vergent soft-sediment folds. a: Metric vergent soft-sediment fold developed at the base of the same conglomerate bed as the folds ofFig. 8. The conglomerate beds display shear zone fabric. b: Detail of folded sandstone bed in the core of the larger fold of a, which shows cuspate–lobate folds in the steep and stretching in the flat limb. Synclinal fold hinges are thickened. c: Foliation developed in the core of vergent fold. d: foldaxes of folds, e: poles to bedding and calculated common fold axis and f: axial planes of all folds in outcrop.
slumpsheet fold axis orientations in outcrop, about 50%ofthe folded area will have fold axes parallel to slope dip(area shaded grey in Fig. 13), 25% parallel to slope strike(area shaded light grey in Fig. 13), and the remaining 25%of the area would have an intermediate fold axis orien-tation. In slumps with small offsets across the basal glideplane compared to the dimension of folds, fold axesshould be parallel to slope strike.
The model outlined above describes the geometry ofthe distribution of fold axes within slumps, but does notoffer an explanation for the development of isoclinal foldswith axial planes parallel to bedding. Previous workershave described bedding-parallel folds that are often bou-dinaged and resemble intrafolial folds (e.g. Woodcock,1976; Blewett, 1991), and folds with strongly curvedhinges (e.g. Williams et al., 1969; Farrell, 1984; Tobisch,1984; Farrell and Eaton, 1987). These fold geometries arefrequently observed in shear zones and metamorphicrocks with mylonitic foliation (e.g. Ramsay and Huber,1987). Fold axes in such rocks are commonly parallel tothe transport direction of the hanging wall (e.g. Cloos,1946; Flinn, 1962; Grujic and Mancktelow, 1995). The
reason for this is the tendency of linear elements to rotatetoward parallelism with the transport direction if offsetbecomes high in relation to the size of the material line(e.g. Cobbold and Quinquis, 1980). The occurrence offolds with geometries indicative for high strain such assheath folds in distinct layers suggests a shear zone modelis useful for some of the features observed in slump sheets.
3.1.2. Folding due to tectonic forcesMedium-to small-scale folds in the brittle crust related
to tectonic deformation are either related to a fault or shearzone, such as fault-propagation folds or fault-bend folds(e.g. Suppé, 1985), or they are parasitic folds related tolarge scale folding. In the first case, folds related to asingle phase of deformation should be coaxial and vergein one direction. Applying the dislocation model outlinedabove, tectonic deformation can be described as defor-mation related to dislocation of almost infinite size. In thesecond case, folds are also coaxial, with fold vergencehaving a systematic relationship to large-scale folds.Mostfolds related to tectonic deformation affect thick sedimentpackages and are not restricted to single layers of rocks.
Fig. 11. Dekametric vergent soft-sediment folds in the north face ofPleiskopf. See Fig. 3 for location. Hinges of folds in conglomerates arethickened due to incomplete lithification. a: Poles to bedding planes ofthe folds and calculated fold axis.
Fig. 12. Gliding of sediment packages. Ramp structure at the base of agliding unit. A conglomerate was stacked, and the gliding unit folded.Younger conglomerate units fill the relief created by upramping.
114 H. Ortner / Sedimentary Geology 196 (2007) 99–118
3.1.3. Relation of lithification to deformationThe main difference between slump-related deforma-
tion and tectonic deformation is that the first takes placewhile the sedimentation process is ongoing, but cannotcontinue under a thick sedimentary cover, while the sec-ond process is persistent through time, at least in areaswith ongoing folding. Therefore, soft-sediment deforma-tion due to slumping should be restricted to the initialstages of lithification of a rock, while soft-sediment de-formation due to tectonic forces should take place duringall stages of lithification.
features in the Upper Gosau Subgroup are related to (1)fluid escape and injection driven by pore pressure in-stability and sediment loading intrinsic to sediment grav-ity flow deposition, (2) gravity driven glide processes or(3) tectonic deformation. Deformation features closelyrelated to sedimentation are the load structures describedin Section 2.2.1 and gliding of sediment packages as-sociated with folding. In both cases, a local topography
created by the deformational process was filled by sub-sequent sedimentation.
Because the size of fluidized units (Section 2.1.2)cannot be infinite, movement of hydroplastically de-forming beds which are not disrupted into very smallparts within a fluidized matrix should in some wayconform with the model for surficial slumps. However,total translation of material is limited, as the fluidizedstate can not be maintained for a long time (Lowe,1976). Free movement and chaotic mixing during thefluidization event is not possible if the fluidized layershas abundant hydroplastically deforming “rigid” bedsthat are in contact with each other. The dominantorientation of bedding is inclined downslope orupslope and fold axes are scattered around an axisperpendicular to the downslope direction (fluidizedlayer 1 of Fig. 6). This corresponds to the initial stagesof downslope movement of a surficial slump in thedislocation model. Occurrence of refolded isoclinalparallel folds might be a consequence of folding in anenvironment of extreme competence contrasts betweenthe fluidized conglomerate matrix and the cohesivesandstone beds.
Hydroplastic deformation in single or several layersis related to deformation of conglomerate beds(Sections 2.1.3 and 2.1.4). A glide plane parallel tobedding is found in a heterolithic unit directly below adeformed conglomerate bed. Upsection from the glideplane, simple shear affects the thin heterolithic layerand the base of the conglomerate (Figs. 7 and 8). Thethickness of the sediment pile transported on top ofthe glide plane is not known. In these examples, theshear zone model should be applied. Probablylocalization of gliding just below the coarse-grained
Fig. 13. Models for the movement mechanism of slumps. a: thedislocation model (redrawn from Farrell (1984). b: Fault and foldorientations in a snow slump, redrawn and modified from Lajoie(1972). Bottom right rose diagram for all fold axis orientation in theslump. Grey shading denotes area with fold axes parallel to transportdirection, light grey shading area with fold axes parallel to slope strikec: Relation between large scale features of a slump described by thedislocation model and the shear zone at the base of a slump.
bed is related to loading after deposition of theconglomerate bed. In the case of the asymmetricmullions (Section 2.1.3), finite strains within the shearzone were small, and fold axes perpendicular to
transport, whereas in the case of hydroplastic deforma-tion within conglomerate beds (Section 2.1.4), finitestrains were high and fold axes parallel to transport. Themullions can probably be seen as an initial stage leadingto the pervasive hydroplastic deformation of coarse-grained layers.
All types of deformation discussed in this section upto this point affect single or several layers, and arecharacterized by hydroplastic folding of sandstone andmudstone beds. No deformations in single or severalbeds characterized by semi-brittle or brittle deformationwere observed. In contrast, soft-sediment folds affectingmany beds show plastic deformation (Fig. 10), semi-brittle deformation (Fig. 9) and brittle deformation (Fig.11, Section 2.2.2) of sandstone. All of the foldsdescribed in Section 2.2.2 are coaxial and face N. Theverging dekametric folds described in Section 2.2.2 areconsidered to be of tectonic origin, since they are relatedto detachments that cut upsection (Fig. 3). This is furthersupported by verging metric folds with the same geo-metry that formed during all stages of lithification (seeabove).
3.2. Relation of structures to sediment transportdirection and slope dip
At the northern limb of the main syncline, sedimenttransport directions are to the WSW, which is parallel tothe main axial trend of the large scale folds in the area, orSE, down the northern limb of the main syncline (Fig.14a). Slump folds indicate SSE-directed gliding ofsediment (Figs. 6, 8, 13b), which is down the northernlimb of the main syncline. Accordingly, synsedimentarynormal faults in the area downthrow the southern block,consistent with a SSE-dipping of the slope (Figs. 3, 9).In the zone of complex folding south of the core of themain syncline, the pattern of sediment transport is com-plex and locally suggests transverse transport down thelimbs of folds, perpendicular to anticlinal hinges. Ver-gence of slump folds in the area is in some cases op-posite to sediment transport (compare flute casts of Fig.7a and vergence of folds of Fig. 7c,e,f ), which suggestssome interaction with fold growth. However, the style ofdeformation does not change along one bed across thehinge of the hectometric Alpjoch syncline (Fig. 7d).Slope dip of turbidite fans is commonly small, rangingfrom 1° to 14° (Stow et al., 1996), and in the Muttekopfsediment gravity flow system single conglomerate layersfill topography created by folding (e.g. bed marked with“1” in Figs. 2 and 7d). A change in slope dip to theopposite direction should be connected to a change in theboundary conditions of folding, e.g. the end of growth of
Fig. 14. Sediment transport directions (a) and facing of soft-sediment folds (b) in the Muttekopf Gosau outcrop. Sediment transport directions partlytaken from Ortner (1993).
116 H. Ortner / Sedimentary Geology 196 (2007) 99–118
the hectometric Alpjoch syncline to growth and north-ward tilt of the southern limb of the kilometricMuttekopf syncline shown in Fig. 3.
3.3. Distribution of deformation structures
The different types of deformational structures de-scribed are not evenly distributed. Some of the featureslike loading structures (Section 2.1.1), fluidization (Sec-tion 2.1.2), normal faulting (Section 2.2.3) and down-slope gliding of sediment packages are found allthroughout the study area. Other types of deformationare restricted to the area of complex folding in thesoutheastern part of the Muttekopf outcrop, with onlymeter-scale tectonic folds present in a few places furtherwest (Fig. 14b). This suggests that slump processes wereenhanced where wavelengths and amplitudes of grow-ing folds were smaller. Assuming constant rate ofshortening, smaller structures have higher rates of tiltingper time unit than large structures. More work is needed
to exactly define the relationship between growing foldsand slump processes.
4. Conclusions
In the Upper Gosau Subgroup deformation occurredacross a broad spectrum of timing and rheology. It ispossible to distinguish gravity-driven deformation fromtectonic deformation related to folding and thrusting;however, it is not possible to make this distinction froma single outcrop. In the example described in this paper,following arguments are used to separate tectonic andslump structures:
1) Fold axes in slump deposits and fluidized depositsscatter strongly about the transport direction of theunit or the strike of the slope. If folds are not iso-clinal, they commonly face downslope. In contrast,axes of folds related to tectonic deformation tightlycluster in a direction perpendicular to tectonic
transport. All folds are asymmetric and consistentlyface in transport direction.
2) During slumping, shortening is commonly associatedwith extensional deformation, whereas tectonic fold-ing is not.
3) Whereas slump folds form during the earliest stages oflithification, folds related to tectonic shortening formthroughout lithification. Tectonic folds therefore displaya much wider range of structures related to folding.
Two models are used to describe slump movement inthis paper: First, the dislocation model proposed byFarrell (1984) and Farrell and Eaton (1987), whichdescribes the overall distribution of deformation in aslump. If the slump folds are present at the base of amuch larger gliding unit, a shear zone model (e.g. Cob-bold and Quinquis, 1980) may be more appropriate Theshear zone model accounts for both the formation ofsheath folds, and the rotation of axial planes and foldaxes of preexisting folds into parallelism with foliationand transport direction, respectively.
Acknowledgements
Many colleagues contributed in the discussions to theshaping of the ideas presented in this paper, recently JimBorer and Don Medwedeff during a joint field trip in thearea. Part of the field work done for this paper wassupported by the Austrian Science Foundation (FWF-project P13566). Review of the paper by Jim Borer andTim Lawton is greatly acknowledged, especially thedetailed comments by the first, which helped clarifymany of the ideas presented.
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