BasinResearch Evolutionofbasinarchitectureinanincipient ......cek, 2006).The short lifespan and low subsidence rates in theMostBasin aswell as other basins of theEgerGraben ledRajchl(2006)to

Evolution of basin architecture in an incipientcontinental rift: the Cenozoic Most Basin,EgerGraben (Central Europe)Michal Rajchl,n,w David Ulicny,z Radomır Grygar‰ and Karel MachznCzech Geological Survey, Kla¤ rov 131/3, Praha 1, CzechRepublicwInstitute of Geology and Palaeontology, Charles University, Praha 2, CzechRepubliczGeophysical Institute, CzechAcademy of Sciences, Praha 4,CzechRepublic‰VS� B ^Technical University Ostrava, Institute of Geological Engineering, Ostrava ^ Poruba, CzechRepubliczSeveroc� eske¤ doly, a.s., Doly B|¤ lina, B|¤ lina, CzechRepublic

ABSTRACT

The Oligo-MioceneMost Basin is the largest preserved sedimentary basin within the Eger Graben,the easternmost part of the European Cenozoic Rift System (ECRIS).The basin is interpreted as apart of an incipient rift system that underwent two distinct phases of extension.The ¢rst phase,characterised byNNE^SSW- toN^S-oriented horizontal extension between the end of Eocene andearlyMiocene, was oblique to the rift axis and caused evolution of a fault system characterised by en-e¤ chelon-arranged E^W (ENE^WSW) faults.These faults de¢ned a number of small, shallow initialdepocentres ofvery small subsidence rates that gradually merged during the growth and linkage of thenormal fault segments.The youngest part of the basin ¢ll indicates accelerated subsidence causedprobably by the concentration of displacement at several major bounding faults.Major post-depositional faulting and forced folding were related to a change in the extension vector to anorthogonal position with respect to the rift axis and overprinting of the E^W faults by an NE^SWnormal fault system.The origin of the palaeostress ¢eld of the earlier, oblique, extensional phaseremains controversial and can be attributed either to the e¡ects of the Alpine lithospheric root or(perhaps more likely because of the dominant volcanism at the onset of Eger Graben formation) todoming due to thermal perturbation of the lithosphere.The later, orthogonal, extensional phase isexplained by stretching along the crest of a growing regional- scale anticlinal feature, which supportsthe recent hypothesis of lithospheric folding in the Alpine^Carpathian foreland.

INTRODUCTION

The relationships between the extensional stress ¢eld andan inherited basement fabric have a major in£uence on thegeometry of fault arrays within rifts, and on the resultinggeometries of sedimentary basins in rifts. In particular,the angle between the extensionvector and the axis of a riftstructure (typically a crustal- scale zone of mechanicalweakness) is very important for the resulting three-di-mensional (3D) geometry of rift-bounding faults and theresulting rift-basin geometry (e.g. Illies & Greiner, 1978;Tron & Brun, 1991; McClay & White, 1995; Morley, 1999;McClay et al., 2002; Schumacher, 2002). Long-term evolu-tion of rifted domains typically involves changes of stress¢elds through geologic time (e.g. Aldrich et al., 1986; Zieg-ler, 1990; Dore¤ et al., 1997).This is re£ected in overprintingof older fault systems by the new ones, resulting in compli-

cated structural geometries (Bonini et al., 1997; Keep &McClay, 1997) not easy to interpret particularly in fossilrifts but also in recent rifts involving highly detailed datato elucidate the stress ¢eld history (Mortimer et al., 2005).Another in£uence on the temporal evolution of rift basinsis the growth and linkage of extensional faults, resulting inchanges in the fault number and individual fault displace-ment, which in turn control temporal changes in basinsubsidence (e.g. Cartwright et al., 1995; Gupta et al., 1998;Cowie et al., 2000; Gawthorpe & Leeder, 2000; Morley,2002). Understanding these controls on extensional faultgeometries is important because of their in£uence on thepositions of depocentres and their subsidence rates, aswell as the tectonic topography governing the sedimentdispersal paths, all critical factors for the distribution ofhydrocarbon source and reservoir rocks (Scholz, 1995;Gawthorpe & Leeder, 2000;McClay et al., 2002).

The Most Basin situated within the Eger Graben ofCentral Europe (Fig. 1) o¡ers an opportunity to study theevolution of a fossil intra-continental extensional domainthat features several distinct fault systems, with so far

Correspondence: Michal Rajchl, Czech Geological Survey, Kla¤ rov 131/3, 118 21 Praha 1, Czech Republic. E-mail: [email protected]

BasinResearch (2009) 21, 269–294, doi: 10.1111/j.1365-2117.2008.00393.x

r 2009 The AuthorsJournal Compilationr Blackwell Publishing Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists 269

poorly known spatio-temporal relationships between in-dividual fault populations and the basin’s depositional his-tory. S� pic� a¤ kova¤ et al. (2000) and Ulic� nyŁ et al. (2000) havesuggested a signi¢cant role of oblique extension in theEger Graben evolution, later replaced by orthogonal ex-tension, but an understanding of the exact timing of theseextensional phases and their relationship with post-riftdeformation and uplift of this part of Alpine foreland re-quires analysis of new datasets.TheMost Basin region of-fers a range of observational scales and types of data: (i)large-scale exposures of syntectonic strata in lignitemines; (ii) exposures of fault planes allowing mesoscopicstructural observations to be made; (iii) dense regionalborehole coverage; (iv) regional geophysical maps; and (v)several 2D seismic re£ection lines combinedwith awealthof subsurface and regional geophysical data.

The Eger Graben is the easternmost part of the Eur-opean Cenozoic Rift System (ECRIS, De' zes et al., 2004),which is currently a subject of intense research and con-troversy regarding the causes andmechanisms of extensionas well as post-rift deformation. Plume-related, collisionalcompression-driven or slab-pull-driven extension in theECRIS are discussed, e.g., byMichon &Merle (2005) andDe' zes etal. (2005, and references therein).Newdata on tec-tonic evolution of the Eger Graben should also contributetowards a better understanding of the ECRIS dynamics.

GEOLOGICAL SETTING ANDSTRATIGRAPHYOF THE MOST BASIN

The NE^SW-oriented Eger Graben (Fig. 1) contains ahigher volume of volcanics than most other ECRIS rifts.The lithosphere under the Eger Graben is thinned to ca.80 km (Babus� ka & Plomerova¤ , 2006), and the trace of thegraben roughly parallels theNE^SW-trending depth con-tours of the Moho discontinuity, as shallow as ca. 30 kmunder the Erzgebirge (Krus� ne¤ Hory) Mountains anddeepening to the southeast (De' zes et al., 2004). The EgerGraben axis roughly parallels the trend of a major crustalboundary between the Saxothuringian and the Tepla¤ -Barrandian zones of the Variscan orogen (Kossmat, 1927).This major crustal inhomogeneity, interpreted as a suturecreated during a major collisional event (Matte etal., 1990),de¢ned the northwestern border of the Late Palaeozoicpost-orogenic extensional basin system in the BohemianMassif (Jindr� ich,1971;MalkovskyŁ , 1987).The post-rift his-tory of the Eger Graben is dominated by deformation anderosion at its nortwestern £ank during the Plio-Quatern-ary uplift of the Krus� ne¤ Hory (Erzgebirge) Mts., up to ca.1000m elevations (cf. Zeman, 1988; Ziegler & De' zes,2007;Fig. 2).

TheMost Basin is the largest of ¢ve sedimentary basinspreserved within the Eger Graben (Fig. 1).The area of the

(a) (b)

(c)

Fig.1. (a) Schematic map showing the EgerGraben as a part of the EuropeanCenozoic Rift System (ECRIS),modi¢ed afterDe' zes etal.(2004). (BF, Black Forest; BG, Bresse Graben; EG, Eger Graben; FP, Franconian Platform; HG, Hessian grabens; EZ, Elbe Zone; LG,Limagne Graben; LRG, Lower Rhine (RoerValley) Graben; OW, Odenwald;VG,Vosges) (b) A schematic geological map of the EgerGrabenwith the location of individual sedimentary basins and volcanic domains. (c) A schematic map of theMost Basin showing thepresent-day extent of clastic basin ¢ll and coeval volcanics, and major tectonic structures.

r 2009 The AuthorsJournal Compilationr Blackwell Publishing Ltd, European Association of Geoscientists & Engineers and International Association of Sedimentologists270

M. Rajchlet al.

basin is ca.1400 km2 and the preserved basin- ¢ll thicknessreaches over 500m (Fig. 3). The basin is bounded by theKrus� ne¤ Hory Fault Zone, the Ohr� e Fault Zone and the B|¤ -lina Fault, together with the volcanic edi¢ces Doupovske¤Hory Mts. and C� eske¤ Str� edohor� |¤ Mts. (Fig. 1). The pre-Cenozoic basement of the Most Basin is formed mainlyby metamorphics of the Krus� ne¤ Hory crystalline complex(Saxothuringian), and Upper-Proterozoic metamorphicsof theTepla¤ -Barrandian domain (Mlc� och, 1994). Youngerunits that underlie parts of the Most Basin ¢ll are UpperPalaeozoic sediments, volcanics and Cretaceous sedi-ments (MalkovskyŁ et al., 1985).

The onset of formation of theMost Basin and the entireEgerGraben is temporally associatedwith the onset of themain phase ofvolcanic activity inNW Bohemia during thelatest Eocene (KopeckyŁ , 1978; Cajz et al., 1999; Ulrych et al.,1999).The earliest part of the basin ¢ll is the volcanogenicStr� ezov Formation, followed by clastics and carbonaceousdeposits of theMost Formation (Figs 3 and 4).

Because of partial erosion of the stratigraphic record,the time of the end of syn-rift deposition in theMost Ba-sin is notwell known. It is inferred as latest earlyMiocene,based onmagnetostratigraphy (Bucha etal.,1987;Malkovs-kyŁ et al., 1989) and palaeobotanical data (Teodoridis &Kva-c� ek, 2006).The short lifespan and low subsidence rates inthe Most Basin as well as other basins of the Eger Grabenled Rajchl (2006) to interpret the Eger Graben overall as afailed, incipient rift.

DATA ANDMETHODS

Geophysical data

The map of horizontal gravity gradients was applied to in-vestigate large-scale tectonic structures of theMost Basinthat are covered by sediments andvolcanics or overprintedby younger tectonic structures in the present topography.

Fig. 2. Interpreted digital elevation model of the present surface of theMost Basin and its surroundings. Lines in the overlay marktectonic structures displayed in the present-day topography. (BF, B|¤ lina Fault; KHFZ, Krus� ne¤ Hory Fault Zone; OFZ. Ohr� e FaultZone; SF. Str� ezov Fault; CZ, Czech Republic; GER,Germany).


Basin architecture in an incipient continental rift

Fig. 3. Simpli¢ed geological cross- sections illustrating the geometry of theMost Basin ¢ll were constructed based on archiveproprietary data made available by Severoc� eske¤ doly, a.s.The map shows the locations of individual cross-sections and seismic pro¢lesfrom Fig. 6 within the fault pattern of theMost Basin.


M. Rajchlet al.

Two seismic-re£ection pro¢leswere used to showa generalpicture of the basin- ¢ll architecture and to clarify the pro-blem of syn- vs. post-depositional tectonic deformation ofthe basin ¢ll.

Borehole data

Archive borehole data were used for veri¢cation of thegeological interpretation of seismic sections, constructionof cross-sections and construction of isopach maps of thebasin ¢ll, to reconstruct the geometry of individual depo-centres. Maps were constructed for the complete pre-served basin ¢ll and for three stratigraphic intervals:deposits overlying the main lignite seam, the main ligniteseam and deposits underlying the main lignite seam.These intervals partially coincide with division of theMost Formation sensu ShrbenyŁ et al. (1994).The degree ofprecision of the maps depends on the depth and area ofsurface erosion of individual stratal units, and on the loca-tion and the number of boreholes used. The number ofboreholes changes for individual intervals, because major-ity of the boreholes commonly did not reach below the lig-nite seam. A total of 587 boreholes were used.

Analysis of digital elevationmodel (DEM)

The reconstruction of the fault patterns, obtained by themethods mentioned above, was compared with a DEM ofthe present-day surface to identify traces of the tectonicstructures in the present-day topography.This DEM wasalso used to map the youngest tectonic deformation of theMost Basin.

Sources of chronostratigraphic dating

Mostly palaeontological data were used to assess the tim-ing of the basin ¢lling (Kovar-Eder et al., 2001), togetherwith geochronological data from volcanic rocks of theC� eske¤ Str� edohor� |¤ Mts. (Bellon et al., 1998; Cajz et al.,1999). Ages based on palaeomagnetic data in Bucha et al.(1987) were used for comparison in the subsidence rate es-timates.

Subsidence rate estimate

Lithological data fromwells LIH-17 and LB-214, situatedin the deepest part of the basin, spaced 3.8 km apart, andpalaeomagnetically dated (Bucha et al., 1987) were usedfor the estimate of the subsidence rate. They were com-

Fig.4. Chart showing the regionalstratigraphy of theMost Basin modi¢edafter ShrbenyŁ et al. (1994), with thealternative stratigraphy of Doma¤ c|¤ (1977),together with the intervals of basin ¢llingand interpreted palaeostress vectors, thetemporal extent of Eger Graben volcanicphases (Cajz, 2000) and phases of theEuropean Cenozoic Rift System (ECRIS)evolution (fromDe' zes et al., 2004).



bined into a composite section, in order to represent theentire basin- ¢ll record because one of the wells was notdrilled to the basement. The section was decompactedusing the standard backstripping procedure (Sclater &Christie, 1980), and the ‘Decompact’ spreadsheet byWaltham (2001) was used. Because of the absence ofaccurate chronostratigraphic data, two curves of decom-pacted depth to the basementwere constructed, using dif-ferent dating methods: (i) magnetostratigraphy (Bucha etal., 1987) and (ii) palaeontological and radiometric data(Bellon et al., 1998; Cajz et al., 1999; Kovar-Eder et al.,2001). Because of unconformities present in the well sec-tions, the palaeomagnetic data could only be used as acrude proxy.

Structural analysis

Field-based structural analysis used mesotectonic datasuch as brittle faults, tension gashes, joints, etc., to verifyor supplement the interpretations based mostly on large-scale fault array geometries and depocentre evolution. Re-duced deviatoric palaeostress tensors were computed fromcogenetic fault populations, in some cases separated frompolyphase sets by evaluating ¢eld observation and kine-matic compatibility. Such sets were analysed with the P^T-axis method (Peresson, 1992) and, where necessary,compared with the numerical dynamic analysis (NDA)method (Sperner et al., 1993) using the software packageTectonicsFP 1.6.2 (Reiter & Acs, 2002). Both the P- and T-axes and the NDA-method give kinematic axes that, incase of coaxial deformation, are considered to coincidewith the principal stress axess1,s2 ands3.

ARCHITECTURE OF THE MOST BASIN

Fault systems

The most prominent in the present-day topography of theMost Basin is the NE^SW fault system, basically alignedwith the Eger Graben axis (Fig. 1). Rajchl & Ulic� nyŁ(2000), Ulic� nyŁ et al. (2000), and S� pic� a¤ kova¤ et al. (2000) de-monstrated that theEgerGraben basinswere strongly con-trolled by two other faults systems during the basinevolution. These are represented by E^W-oriented faultsand NW^SE-trending faults.

E^W fault system

The EFW- to ENE^WSW-striking faults are mostlyshort in length (5^10 km locally), show small displacement(50^200m) and occur abundantly in the entire basin.Although only locally well exposed, they are shown clearlyin the map of horizontal gravity gradients (Fig. 5). Thegenerally E^W-oriented gravity gradients are identi¢ed asshallow crustal faults, based on their correlationwith seis-mic pro¢les and borehole data (Figs 3 and 6).The distribu-tion of some of the E^W faults is also demonstrated in theDEM of the Most Basin basement (Mlc� och & Mart|¤ nek,2002).The analysis of the DEMof the present-day surface

shows a number of E^W-trending structures within thetopography of areas surrounding the Most Basin (Fig. 2).In the present-day topography of the Most Basin, the E^W fault system is represented mainly by E^W-orientedsegments of the Krus� ne¤ Hory Fault Zone and the B|¤ linaFault.NumerousE^W faults probably functioned as volca-nic pathways andwere sealed and covered by volcanics (cf.structural data in Cajz, 2001). E^W-trending fabrics havebeen described from the Saxothuringian basement meta-morphics of the NW periphery of the Most Basin (Kono-pa¤ sek et al., 2001). In the southeastern part of the EgerGraben, where the Cenozoic strata are underlain by Cre-taceous and Upper Palaeozoic sediments, inherited base-ment E^W structures have not been reported.

The fault segments are arranged in an en-eche¤ lon pat-tern in plan view (Fig. 5) and some tend to be curved intoparallelismwith the basin axis.The en-eche¤ lon pattern re-sults in abundant relay ramps separating individual faultsegments (e.g. Peacock & Sanderson, 1994; McClay &White, 1995).

The most accessible example of this fault system is theB|¤ lina Fault (Fig. 8), which is recognised byVa¤ ne� (1985a) asa part of an en-eche¤ lon fault array. The fault is charac-terised by overlapping segments up to 10 km long, and ac-companied by a fault-propagation fold, marked bydeformation in the prominent lignite seam. The maxi-mum vertical throw on the B|¤ lina Fault segment 1 (Fig. 9)is ca.190m. In the immediate vicinity of the large-scale B|¤ -lina Fault, several populations of small-scale normal faultsoccur. The population of roughly E^W-trending, obliquenormal faults, with a moderate dextral component, andless abundant, dextral shear trending roughly NE, is as-sumed to be syn-kinematic with the B|¤ lina Fault (Fig. 9a)because it occurs within the trace of the B|¤ lina Fault. Faultpopulations of other orientations are considered youngerand are discussed further below.

A similar set of phenomena is observed at the northernbasin margin, at an exposed example of the E^W fault seg-ments within the Krus� ne¤ Hory FZ characterised by a nor-mal to a slightly oblique normal displacement (Fig.10).

Timing of the E^W fault system activity: Large-scalefault geometry as well as mesoscopic data indicateNNE-directed extension as causing the activity of thisfault system. Direct ¢eld evidence of the syn-sedimentarye¡ects of these faults is limited, but could be well repre-sented by two examples of syn-sedimentary forced foldingabove propagating fault segments (Figs 8 and 10). In bothcases, basinward divergence of sedimentary strata re-corded tilting of depositional surface above an upward-propagating segment of a normal fault (cf. Gupta et al.,1999).

Along the northwestern edge of the preserved basin ¢ll,Oligo-Miocene coarse-grained clastics, including gneissboulders, were interpreted by Va¤ ne� (1985a) as colluvial de-posits, suggesting that the Krus� ne¤ Hory FZ already oper-ated as a syn-sedimentary tectonic margin of the MostBasin. It is, however, likely that the active faults were theE^W segments, the relicts of which are still present as


M. Rajchlet al.

parts of the Krus� ne¤ Hory FZ.The clastics most likely en-tered the basin via relay ramps present in the locations ofthe small clastic bodies. The £uvial deposits of the Hra-dis� te� locality occur on a preserved relict relay ramp, andshow palaeocurrents towards the south (Fig. 7). Also, thesand bodies, documented by Elznic (1963), Elznic et al.(1998),Va¤ ne� (1961) and Zelenka & PolickyŁ (1964) within thelignite seam and overlying lacustrine deposits along thenorthwestern edge of the preserved basin ¢ll, were prob-

ably deposited by £uvio-deltaic systems entering the basinthrough relay ramps.

The syn-sedimentary activity of this fault system is alsosupported by seismic re£ection data that show small-dis-placement normal faults of E^W orientation, commonlyterminated within the main lignite seam and resulting inits local £exure (Fig. 6). This suggests that the activity ofthese faults ceased early during the basin evolution andonly major, graben-bounding faults remained active.

Fig. 5. (a) Uninterpreted and (b)interpreted map of horizontal gravitygradients of theMost Basin and thesurrounding area.The map was producedby Geofyzika, a.s., Brno andMiligal, s.r.o.,compiled from regional gravity mapping ona1 : 25 000 scale, using reduction density2.30 g cm� 3, with illumination fromN30E.The black lines in (b) mark those horizontalgravity gradients interpreted as faultstructures.The fault pattern is representedbyE^W,NE^SWandNW^SE fault systems.The high gradients in the southwesterncorner of the map correspond to the edge ofVariscan granitoids underlying theCenozoic volcanics of the Doupovske¤ HoryMts. (BF, B|¤ lina Fault; KHFZ, Krus� ne¤Hory Fault Zone; OFZ, Ohr� e Fault Zone;SF, Str� ezov Fault).




M. Rajchlet al.



NE^SW fault systems

Normal faults of NE^SW orientation, showing typicallygreater lengths than the E^W faults, represent the mostprominent structural trend in the topography of theMostBasin, parallel to the axis of the EgerGraben (Figs1and 2).However, a marked spread of directions between ca. N35EandN60E is observed in this group of faults. Locally, someNNE-trending fault populations contain faults alignednearly N^S.Within the entire Eger Graben, the fault sys-tem is represented by parts of theKrus� ne¤ HoryFault Zoneand the Ohr� e Fault Zone at the present-day southeasterngraben margins.Within the Most Basin, few major faultsfollow this trend (Fig. 2).

A signi¢cant morphological feature of theKrus� ne¤ HoryFault Zone is the kinked trace of the faults, characterisedby a number of relatively short E^W-trending fault seg-ments, commonly linked by short segments of nearlyNNE^SSW strike (Fig. 2). It is evident that the Krus� ne¤Hory Fault Zone is a complex structure, represented insome places by normal fault segments (Fig. 3, cross-sections 2, 3) and elsewhere by monoclinal folding of thebasin- ¢ll strata in relicts of relay ramps (Fig. 3, cross-sec-tions 5^7; cf. MalkovskyŁ , 1979; Hurn|¤ k & Havlena, 1984;Marek, 1985). On the opposite margin of the Most Basin,the Ohr� e Fault Zone is represented by several relativelystraight-, parallel-trendingNE^SW faults (15^30 km long)and by a number of small fault segments, as shown in the

Fig.7. (a) Interpreted digital elevationmodel and (b) panorama of the faulted NWmargin of theMost Basin (Krus� ne¤ HoryFZ).The images showEFW-oriented faultsegments separated by relay ramps. A relictof a £uvial clastic body (the localityHradis� te� ) suggests the syn-depositionalactivity of the ramp that functioned as asouth-directed pathway of clastics into thebasin. For location, see Fig. 2.

Fig. 6. Re£ection seismic pro¢les 68/83 (a) and 21/81 (b), recently reprocessed byGeofyzika, a.s., Brno and reinterpreted (Ulic� nyŁ &Rajchl, 2002; Rajchl et al., 2003a, b; see Jihlavec &Nova¤ k, 1986 for original interpretation), showing the architecture of the basin ¢llwithin the B|¤ lina depocentre. (a) Pro¢le 68/83 shows a number of important phenomena: (i) an extensional horst-like structure (of reliefup to100m) interpreted as an accommodation zone that separated two small grabens during the initial stage of the basin evolution andbecame inactive after the main seam deposition (location between 2.3 and 3 km); (ii) wedge shape and shingle-like internal architectureof distal parts of the B|¤ linaDelta; (iii) the e¡ect of peat compaction on accommodation creation for the earliest lacustrine deposits,NWof 2.5 km.The thickness of the lowermost part of the lacustrine deposits shows an inverse relation to the thickness of underlying deltaicclastics.The lacustrine deposits are the thickest in places where the deltaic clastics are absent; (iv) onlap of lacustrine strata on thesurface of the B|¤ lina Delta sedimentary body suggests gradual drowning of the deltaic sedimentary system; (v) post-depositional fault-propagation folding of the basin ¢ll close to theNWmargin of the basin. (b) Part of the pro¢le 21/81 (segmentsH, J, K, L,M), showing alarge horst-like structure between 8.5 and11.4 km, de¢ned by normal faults and interpreted as an accommodation zone.This syn-depositional structure is characterised by a reduction in the coal seam thickness (locally down to1.5m) and underlying deposits(including volcaniclastics). A number of normal faults a¡ecting the deposits underlying the main lignite seam and partly the main ligniteseam are evident in both (a) and (b). See the map in Fig. 3 for the location of the pro¢les.


M. Rajchlet al.

data of HradeckyŁ (1977), MalkovskyŁ (1979) and MalkovskyŁet al. (1985).

The Str� ezov Fault represents the only signi¢cant intra-basinal fault of the NE^SW direction and divides thewidest part of the basin into two parts (Figs 1 and 2).Thisfault is ca. 25 km long and its SW tip terminates in the cen-tre of the Doupovske¤ Hory volcanic complex. In the base-ment, the trace of this fault generally coincides with theposition of the faulted margin of the Permo-Carbonifer-ous Kladno-Rakovn|¤ k Basin (Mlc� och & Mart|¤ nek, 2002;Fig. 3 ^ cross-sections 7, 8).

Timing of the NE^SW fault system activity: The Krus� ne¤Hory FZ (Figs 1, 2 and 5) was formed by linkage of shortsegments of the E^W fault system across their relay ramps,and as awhole, is clearly younger that the E^W faults. It is,however, not completely clear whether the linkage oc-curred only after the deposition in the basin had ceasedor whether it had already occurred during the younger de-positional phases. The presumed syn-depositional exis-tence of the Krus� ne¤ Hory FZ as one of the main syn-rift‘deep faults’ (e.g. KopeckyŁ , 1978) has been disproved longago (Va¤ ne� , 1985a), but this assumption is still found impli-citly or explicitly in some recent literature (Michon et al.,2003). NE- to NNE-trending segments of the Krus� ne¤Hory Fault Zone suggest a roughly NW-directed exten-sion; it is discussed belowwhether this represents a regio-nal palaeostress ¢eld or an evolution of a local stress ¢eldbetween propagating E^W fault tips that led to relay rampbreaching.

Post-depositional displacement along the Ohr� e FZ isdocumented by faulting of the entire basin ¢ll and by re-licts of Neogene strata in post-depositional grabenssoutheast of the Most Basin border. The large thicknessof £uvial clastics of the so-called %atec Delta (Fig. 3,cross-sections 6^8) indicates that syn-depositional ac-commodation existed in the area of the Ohr� e FZ duringthe lignite seam formation. It is unclear, however, whetherthe active faults had the same strike as the present-dayNE^SW structures ^ this also applies to the presumedfaults further southeast that governed the formation of

the lower Miocene hot-spring freshwater limestones ofTuchor� ice (Va¤ ne� , 1985a; Fejfar &Kvac� ek, 1993).The faultedmargin of the basin in the %atec Delta area apparently didnot function as the syn-depositional basin edge. Indices ofE^W-trending structures occur in the gravity gradientmaps, marking possible unmapped, E^W-trending, pre-cursors of the later Ohr� e FZ.

In addition to undoubted post-depositional displace-ment, Neogene-age syn-sedimentary activity of the in-tra-basinal Str� ezov Fault is documented by the change inthe thickness ofNeogene clastics between the footwall andthe hangingwall blocks (Fig. 3 ^ cross-section 7).The un-changed thickness of Str� ezov Formation volcanics acrossthe fault (MalkovskyŁ , 1979) and the absence of clastics un-derlying the main lignite seam on the hangingwall side ofthe fault, together with the architecture of the seam (Fig. 3^ cross-sections 7, 8), suggest that the NE^SW Str� ezovFault began to be active during the seam evolution.

Other, minor faults of SSW^NNE occur in variousplaces within the basin and generally show post-deposi-tional normal displacement in NW-directed extension,such as the Elis� ka Fault (Fig.11; Brus &Hurn|¤ k, 1987).Me-soscopic brittle structures along the B|¤ lina Fault (Fig. 9a)suggest that, in addition to the syn-depositional E^W-trending normal faults, another population of mesoscopicfaults occurs, trending NE^SW to NNE^SSW, and showsnormal displacement with a weak sinistral component.This population, occurring pervasively both in the foot-wall and in the hangingwall of the B|¤ lina Fault Segment 1(Fig. 9), indicates local NW^SE-directed extension. Post-vs. syn-depositional age of these faults with respect to theseam and clastics in theB|¤ linaFault hangingwall cannot bedetermined with certainty. In addition, localised sets offault^slip data are not su⁄cient for determining a larger-scale palaeostress ¢eld (e.g. Gapais et al., 2000).

NW^SE fault system

Both the DEM (Fig. 2) and the horizontal gravity gradi-ents (Fig. 5) indicate that theMost Basin is signi¢cantly af-

Fig. 8. Deformation of the main ligniteseam and clastics of the B|¤ lina Delta alongthe B|¤ lina Fault ^ southern margin of theB|¤ lina open cast mine (as of 2000).Theclastic wedge of the B|¤ lina Delta progradedgenerally westward, alignedwith the activesegments of the B|¤ lina Fault (Rajchl et al.,2008). Basinward divergence of deltaic andlacustrine strata suggests syn-depositionaltilting of a sedimentary surface caused byfault propagation folding. Local divergencebetween the top of lignite seam (251) andpalaeohorizontal markers within theuppermost deltaic clastics (121) is 131.Thelocation of Fig. 8 in Figs 2 and 9 shows theposition of the uppermost terrace of themine.



fected by a number of NW^SE-trending faults occurringas 20^30-km-long segments. Several faults close to theNW^SE orientation were also detected by mining (Mal-kovskyŁ , 1979; MalkovskyŁ et al., 1985; Brus & Hurn|¤ k, 1987).A number of faults of similar orientation are documentedin other basins of the rift (Kasin� ski, 2000; S� pic� a¤ kova¤ et al.,2000; Roj|¤ k, 2004). Apart from the straight, near-vertical,

clearly basement-derived faults of NW^SE strike thatcross the entire basin (Fig. 2), some intra-basinal faultstrendNW^SE only along a part of their trace. An exampleis the Victoria Fault recently exposed in the B|¤ lina mine(Fig. 9) locally curved from a nearly E^W strike to a pureNWstrike. AWNW-trending segment of this faultwas ex-posed where it joins the B|¤ lina Fault and shows normal to

(a)

(b)

Fig.9. (a) A structural map showing theelevation (in metres above the sea level) ofthe base of the main seam in the vicinity ofthe B|¤ lina Fault (location in Fig. 2).Thecontour lines illustrate the displacementassociatedwith one of the en-eche¤ lonsegments of the B|¤ lina Faults (Segment1 inthe ¢gure), including the fault-propagationfold accompanying it, and the topography ofthe relay ramp between fault segments1 and2.The interpretation of meso-scalestructural data is shown to illustrate thedetails of fault kinematics.‘Basin edge’ refersto the present-day erosional edge ofpreserved basin ¢ll. (b) Changing verticaldisplacement along segments1 and 2 isillustrated by cross sections highlighting thetwo-dimensional geometry of the mainlignite seam close to the B|¤ lina Fault.

Fig.10. Deformation of basin ¢ll associatedwith the activity of segment1of the Krus� ne¤ Hory FZ from Fig. 7 (see also Fig. 7 and 2 forlocation). (a)A closeupviewof anE^Wfault segment of theKrus� ne¤ HoryFZjuxtaposing the basin ¢ll against the crystalline basement inthe footwall.The fault is characterised by normal to slightly oblique normal displacement. (b), (c) Two examples of syn-depositionalforced folding above the propagating E^W fault segment of the Krus� ne¤ Hory FZ. In (b), the trace of base of the lignite seam showsapparent curvature due to a curved quary wall. Subtle basinward divergence of sedimentary strata recorded tilting of the depositionalsurface above the upward-propagating normal fault segment.The apparent divergence angle is decreased due to di¡erentcialcompaction above surface A ^ note the transition from mudstones to lignite in the basinward direction.


M. Rajchlet al.



slightly dextral displacement (Fig.9a), similar to the B|¤ linaFault mesoscopic data.The trace of theVictoria Fault (Fig.9) shows that the fault is a composite structure formed bylinkage of former E^W fault segments. Seismic pro¢le 68/83 documents syn-depositional vertical displacement onan E^W segment of theVictoria Fault.

Timing of the NW^SE fault system activity: The NW^SEorientation of the faults essentially coincides with thetrend of the Elbe Zone, one of the signi¢cant shear zonesformed during the Late Palaeozoic Variscan orogeny(Arthaud & Matte, 1977; Schr˛der, 1987; Scheck et al.,2002) (Fig. 1a). The existence of this pre-rift fault systemsigni¢cantly a¡ected the geometry of later fault systemsthat formed during opening of the Eger Graben. Bendingof some of the E^W faults to parallelismwith the NW^SEfaults suggests their coeval activity. Post-depositionalvertical displacement of several tens of metres is docu-mented at the NW-trending segment of the Victoria Fault(Fig.9).No mesoscopic datawere found to assess the kine-matics of these faults during the basin ¢lling; a sinistralstrike^slip regime is interpreted for their post-rift, Plio-cene phase of activity in peripheral parts of the Eger Gra-ben (Cheb Basin, e.g. S� pic� a¤ kova¤ et al., 2000; Pliocenevolcanics at the Lu&ice Fault Zone; R. Grygar, unpub-lished data). Geomorphological data suggest a possiblenarrow, sinistral pull-apart structure with a Pliocene toQuaternary in¢ll following the faults of the ChomutovkaCreek (Fig. 2), so far not proven by independent methods.The dextral component of slip in the Victoria Fault zone(Fig.9a) is probablydue to its merging with theB|¤ linaFaultsegment1 in the location of measured exposure, where thedata correspond to the kinematics of the E^W normalfaults.

Some of theNW^SE faults served as a pathway for clas-tics transported into the area of the%atecDelta close to thesoutheastern margin of the basin (Fig. 1).This clastic beltis known as the Hlavac� ov gravels and sands (e.g. Va¤ ne� ,1985a) and is characterised by a ‘panhandle’ map-view

shape (by analogy to the ‘panhandle’ of the Okavango in-land delta in Botswana, e.g.McCarthy et al., 1992).

Most Basin depocentres: geometry andspatial arrangement

Four main depocentres of the basin are distinguishedbased on preserved basin- ¢ll geometry (Rajchl, 2006;Figs 12 and 13). However, a number of small basin- ¢ll re-licts suggest that the area of deposition exceeded the pre-sent-day limit of the basin.The depocentres are elongateand show a graben or a half-graben geometry in cross-section (Figs 3 and 6). The Chomutov, B|¤ lina and TepliceDepocentres are arranged in an en-eche¤ lon pattern,similar to their bounding E^W fault segments.The %atecDepocentre axis has a NE^SWorientation (Fig. 13). Thedepocentres are separated from one another by palaeo-highs (Fig.12), characterised by reduction of the thicknessof the basin ¢ll (Fig. 3 ^ cross-section 9; Fig. 6b). Only inthe case of the%atec andChomutovDepocentres is the se-paration partially caused by the Str� ezov Fault (Fig. 3 ^cross-section 8; Fig.13).

The orientation and geometry of the palaeohighs withrespect to individual depocentres suggest their functionas accommodation zones (cf. Peacock et al., 2000; McClayet al., 2002).The orientation of the accommodation zonesroughly coincides with the transverse, NW^SE trending,basement faults (Fig. 13) that probably de¢ned their posi-tion and the o¡set of individual depocentres (for similarobservations, see, e.g., LeTurdu et al., 1999; Morley, 1999;McClay et al., 2002).

Cross-sections and isopach maps (Figs 3, 6 and12) showthat the size and shape of the depocentres changed signif-icantly during the Most Basin evolution. The data docu-ment the gradual linkage of small initial depocentres,eventually merging into the four large depocentres de-¢ned above (e.g. Fig. 6a).The initial depocentres were onlya few square kilometres across (Fig.12).

TECTONOSEDIMENTARY EVOLUTIONOF THE MOST BASIN

Intervals of basin filling

Interval1: volcanics and volcaniclastics (latest EoceneÔligocene,36^26Ma)

During the ¢rst depositional interval, the Most Basinwas ¢lled by the volcanosedimentary Str� ezov Formation(Fig. 4) comprising alkaline volcanics close to the volcaniccentres, and pyroclastics and redeposited pyroclasticmaterial in the initial depocentres (e.g. Doma¤ c|¤ , 1977;MalkovskyŁ etal.,1985).The singleNW-elongated thicknessmaximum shown in Fig.12a is probably caused by Eoceneclastic in¢lls of inherited topography that pre-date thevolcaniclastic deposition but in archive borehole docu-mentation are mostly impossible to distinguish from thevolcaniclastics. Other deposits of the Str� ezov Formationinclude intercalations of lacustrine limestones, diatomites

Fig.11. Aphotograph of theElis� kaFault, one of the minorSSW^NNE faults showing generally post-depositional normaldisplacement in NW-directed extension. For the location of thephotograph, see Fig. 2.


M. Rajchlet al.

and carbonaceous clays or coals, deposited in lacustrineand swamp environments (e.g. Doma¤ c|¤ , 1977; MalkovskyŁet al., 1985; Bellon et al., 1998). Cajz (2000) correlates theStr� ezov Formation with the lower stratigraphic unit ofthe volcanic C� eske¤ Str� edohor� |¤ Mts. (36^26Ma) that ¢llsthe original topography and the initial depocentres ofthe incipient rift structure. Curves of subsidence rateevolution (Fig. 14) show very low subsidence ratesduring this stage of the basin evolution, on a scale of afew mMa�1.

Interval 2: clastics under the main seam (latest OligoceneêarliestMiocene, 26^21Ma)

Proluvial and alluvial deposition of material derived fromthe surrounding volcanics and the Cretaceous and crystal-line basement dominated during this interval of the basinevolution, corresponding to the lower part of theMost For-mation (Fig. 4; e.g. MalkovskyŁ et al., 1985). Local lacustrineand coal-bearing environments also formedwithin the de-pressions of newly forming relief (e.g. MalkovskyŁ et al., 1985;

Elznic et al., 1998). The cross-sections and isopach mapsshow that the small local depocentreswere formedgenerallyalong the present-day Krus� ne¤ Hory Fault Zone, controlledby a number of minor intra-basinal normal faults of smalldisplacement (Fig. 6). The geometry of the depocentresand the correlation of seismic and lithological cross-sections to the mapped fault framework suggest that thetopography of theMost Basin area was generally a¡ected byan E^W-oriented fault system during this stage of the basinevolution (Fig.15a). Subsidence curves indicate a modest in-crease of the subsidence rate in the central part of the basinduring this phase (Fig.14).The thickness of the correspond-ing deposits, shown by cross-sections (Fig. 3), suggests thatthe subsidence rate could locally exceed10mMyr�1.

Interval 3: main lignite seam and corresponding clastics (earlyMiocene, 21^18Ma)

This interval is represented by the main lignite seam in themiddle part of the Most Formation (sensu ShrbenyŁ et al.,1994; Fig. 4).The seam has a nearly basin-wide extent and

Fig.12. Isopach maps of theMost Basin ¢ll. (a) Thickness of volcanics and clastics underlying the main lignite seam. (b) Thickness ofthe main lignite seam and coeval clastics.The thickness of lignite is displayedwithout decompaction,which causes a signi¢cant increaseof the thickness towards places with clastic interbeds. (c) Thickness of lacustrine deposits overlying the main lignite seam. (d) Totalthickness of the basin ¢ll. (B, B|¤ lina Depocentre; CH, Chomutov Depocentre;T,Teplice Depocentre; %, %atec Depocentre). Boreholedata were obtained fromGeofond of the Czech Republic and the archive of Severoc� eske¤ Doly, a.s.



an average thickness of ca. 30m, reaching locally up to 50min response to the underlying topography. The average lig-nite seam thickness corresponds to a peat accumulationthickness of 180m or more before loading by clastics,based on the compaction ratio of Hurn|¤ k (1978). Duringthis interval, two large coarse-grained depositional sys-tems, the %atec Delta and the B|¤ lina Delta (Dvor� a¤ k &Mach, 1999; Rajchl & Ulic� nyŁ , 1999, 2005; Rajchl et al.,2008), entered theMost Basin from the SE and E, respec-tively (Fig.15b), and caused inter¢ngering of the seamwithclastics. Some minor clastic systems also developed on theNWmargin of the basin during accumulation of the peat(Fig.15b; e.g. Zelenka & PolickyŁ , 1964; Elznic et al., 1998).

According to Kovar-Eder et al. (2001), the main ligniteseam and its clastic equivalents represent a time span ofapproximately 2Myr between 20 and 18Ma (zone MN-3of European faunal zonation). However, the di¡erentialthickness of the main lignite seam, illustrating gradualspreading of the peat swamp from individual depocentres,suggests that the life span of the original peat swamp wasprobably longer in some places of the basin.

Interval 3 is characterised by a general increase of thedepositional area. Deposits of this interval occupy the en-tire area of the basin, in contrast to the deposits of Interval2 (Fig. 15b). Clastic equivalents of the seamwere probablydeposited as far as the volcanic complex of the C� eske¤ Str� e-dohor� |¤ Mts. during this interval as suggested by isolatederosional relicts of £uvial and lacustrine clastics (Fig. 1c;Hurn|¤ k & Kvac� ek, 1999). C� adek (1966), Elznic (1970) andElznic et al. (1998) hypothetised that a hydrological outletdraining the Most Basin existed in the area of the pre-sent-day Krus� ne¤ Hory Mountains near Jirkov (Fig. 15).This speculation is based on the extent and orientation of

major £uvial channels of the %atec Delta (Fig.15b), and itslocation coincides with a relay ramp between E^W faultsegments at the northwestern end of one of the accommo-dation zones.

The aggradation of organic material in this remarkablethickness and extent suggests a relatively continuous andaccelerated subsidence in the area of the entire basin (cf.Ayers & Kaiser, 1984; Ayers, 1986), also documented bysubsidence curves (Fig. 14). This is interpreted as the in-crease in displacement at major bounding faults at the per-iphery of the basin (Fig.15b). Segments of the E^W B|¤ linaFault actively grew and produced syn-depositional defor-mation at a time corresponding to the uppermost part ofthe main lignite seam (Rajchl et al., 2008). The Str� ezovFault created accommodation for clastics of the%atecDel-ta during this time interval (Fig. 3, cross-sections 6, 7), butit is uncertainwhether at this time the fault alreadyhad thepresent-day NE^SW strike (possible E^W-trending pre-cursors are indicated by geophysical data, Fig. 5).

The termination of small-displacement intra-basinal faultsunder or within the main lignite seam, together with thenearly basin-wide extent of the seam, suggests that the smallinitial depocentres began to merge into larger ones towardsthe end of Interval 2 and especially during Interval 3 (Fig.15b).

Interval 4: post-seam clastics (earlyMiocene^middleMiocene,18^15Ma)

Predominantly lacustrine deposits (up to ca. 400m thick)of the upper part of the Most Formation (sensu ShrbenyŁet al., 1994; Fig. 4) represent the last-known interval of theMostBasin sedimentary evolution.Only carbonaceous de-posits of the LomSeam represent a local interruption of a

Fig.13. Structural model of theMost Basinshowing individual fault systems and themain depocentres, as interpreted fromcorrelation of gravity maps, isopach mapsand digital elevation model (DEM) withbrohole-based cross sections and seismiclines. Further comments in text. (B, B|¤ linaDepocentre; CH, Chomutov Depocentre;T,Teplice Depocentre; %, %atecDepocentre).


M. Rajchlet al.

lacustrine sedimentation.The time span of this interval isestimated as either 18^15Ma (using data of Kovar-Ederet al., 2001) or 21^17.7Ma based on magnetostratigraphy(Bucha et al., 1987; MalkovskyŁ et al., 1989, but note the pro-blems with the magnetostratigraphic methodology men-tioned above).

Interval 4 started with a signi¢cant increase of the sub-sidence rate (up to ca.100mMyr�1; Fig.14) associatedwithdrowning of the basin-wide swamp and coarse-grainedclastic depositional systems (the B|¤ lina and %atec deltas)by an extensive lake (Figs 3 and 6).The onlap of lacustrinedeposits on the body of the B|¤ lina Delta and the surface ofthe seam (Fig. 6), however, suggests that the process of ¢ll-ing of the accommodation space createdwas not instanta-neous. This is also documented by backstepping of the

youngest deltaic bodies of the B|¤ lina Delta (Rajchl et al.,2008). Because of post-depositional deformation a¡ectingmuch of theMost Basin, it is di⁄cult to assess the areal ex-tent of the lake at this interval. However, it probably didnot reach beyond the present-day margins of the EgerGraben in the Most Basin region (see the discussionbelow).

The carbonaceous deposits of the Lom Seam (Fig. 3 ^cross-sections 4, 9) are interpreted as a swamp forest ormire (Teodoridis & Kvac� ek, 2006).Together with underly-ing sand bodies in the central (deepest) part of the MostBasin, it indicates a temporary shallowing of the lacustrineenvironment.This can be explained by a temporary decel-eration of tectonic subsidence and ¢lling of the lake byclastics, or by an increase in the sediment supply rate.

(a) (c)

(b)

Fig.14. (a), (b)Decompacted depth-to-basement curve illustrating the subsidence history of the deepest part of theMost Basin.Valuesfor porosity and the ‘c’ factors of individual sediments were obtained from Sclater & Christie (1980), with the exception of the initialporosity of lignite (peat) of 0.88, afterMach (2003), and the ‘c’ factor (0.001) derived from the compaction ratio of peat in theMost Basin(6 : 1), as interpreted byHurn|¤ k (1972).Dating for the curve in (a) is based on magnetostratigraphic data byBucha etal. (1987).Dating forcurve in (b) is based on palaeontological and radiometric data (Bellon et al., 1998; Cajz et al., 1999; Kovar-Eder et al., 2001). Despitedi¡erences between the dating methods used and the resulting stratigraphic extent of individual intervals, both curves show verysimilar results, documenting a gradual increase of the subsidence rate during theMost Basin evolutionwith an increase towards Interval4 of basin ¢lling.The subsidence rate estimated for Intervals 1 and 2 (accumulation of the Str� ezov Formation and the lower part of theMost Formation) is very low, ranging within a few mMyr�1.The subsidence rate during Intervals 3 and 4 (accummulation of the middleand upper parts of theMost Formation) ranged from tens of mMyr�1 (during accumulation of peat) to ca.100mMyr�1(duringsedimentation of lacustrine deposits). (c)A composite section created by compilation of lithological data fromwellsLIH-17 andLB-214,situated in the deepest part of the basin. For the location of the wells, see Fig.1.



(a)

C (d)

(b)

Fig.15. Cartoons illustrating the interpreted evolution of depocentres and syn- and post-depositional fault systems and thepalaeostress interpretations. In (a), the geometry of the earliest syn-depositional fault systems, characterised byE^W- andENE^WSW-oriented fault segments, is shown together with the position of initial depocentres during Intervals 1 and 2 of the basin evolution. (b)Fault pattern and lateral extent of deposition interpreted for Interval 3, time of peat accumulation. (c) Fault pattern and lateral extent ofsedimentation during Interval 4, characterised by formation of a basin-wide lacustrine environment. An increase in the lateral extent ofindividual depocentres and the depositional area between (a) and (c) documents gradual linkage of the faults and depocentres due toincreasing extension. (d) Fault pattern interpreted for postsedimentary orthogonal extension causing destruction of the basin.Theextent of clastics shown corresponds to the preserved part of the basin ¢ll.


M. Rajchlet al.

Post-depositional deformation of the MostBasin

TheEgerGraben region underwent a polyphase deforma-tion history, the most recent part of which is associatedwith the Pliocene toQuaternary uplift of theKrus� ne¤ HoryMts. and the related deep incision of some of the rivers inthe region, including the Labe (Elbe) River (Va¤ ne� 1985a;Tyra¤ c� ek, 2001;Tyra¤ c� ek et al., 2004; Ziegler &De' zes, 2007).Below,we are concerned onlywith the part of the deforma-tion history that more or less immediately followed theEgerGraben rifting event and is related to normal faultingin theMost Basin area.

A major unconformity, spanning approximately theMid-Miocene through Early Quaternary, truncates theclastics ¢lling theMost Basin, in a manner similar to otherbasins of the Eger Graben (Fig. 4). Post-rift, Late Pliocenestrata overlying this unconformity in the Cheb Basin indi-cate that by ca. 5Ma the destruction of the syn-rift basin¢lls had already been accomplished (S� pic� a¤ kova¤ et al.,2000). According to the study of burial-derived compac-tion by Hurn|¤ k (1978), the maximum depth ofpost-rift erosion of theMostBasin ¢ll is ca. 300m.The lackof direct sedimentary evidence prevents a more accuratedating of the rift-deformation event than betweenca. 17 and 5Ma, but shifts in the drainage patternsin the BohemianMassif (MalkovskyŁ , 1979) indicate the on-set of uplift in the Eger Graben region in the MiddleMiocene.

Along the entire Eger Graben, signs of post-deposi-tional (i.e. post-early Miocene) deformation and erosionare evident. Individual basin ¢lls are relicts preserved indownthrown blocks bounded by NE-trending normalfaults, with isolated erosional remnants locally preservedoutside the downthrown blocks.This is particularly clearbetween the Sokolov Basin and the western part of theMost Basin (S� pic� a¤ kova¤ et al., 2000). There, the Krus� ne¤Hory and other, parallel fault zones show a marked linear-ity and great length of individual segments (up to 30 km).In the central and eastern parts of the Most Basin, theNE-trending faults are slightly less prominent, mainlydue to the kinked trace of the Krus� ne¤ Hory Fault, but stillvery pronounced, and their e¡ects on the basin- ¢ll geo-metry are very important. Probably the most pronouncedtectonic feature is the large-scale £exural deformation ofthe entire basin ¢ll at its NW margin de¢ned by theKrus� ne¤ Hory Fault Zone (Fig. 3, e.g. cross^sections 3^5).This was accompanied by hard linkage of some of the ear-lier E^W faults by NE- to NNE-striking faults, as sug-gested by DEM data and some of the stratigraphic cross-sections, e.g. sections 2^5 in Fig. 3.

The most detailed insight into this deformation is pro-vided by the seismic re£ection pro¢le 68/83 (Figs 3 and 6a)reaching close to the surface trace of the Krus� ne¤ HoryFault Zone. The sedimentary package (lignite seam andlacustrine clastics) above the fault zone is fractured by anarray of secondary, synthetic normal faults in the foldedzone, which splay o¡ the master fault and mostly die out

upward. Immediately above the hinge zone of the £exure,a fan-like array of faults, synthetic and antithetic to themaster normal fault, occurs.This array clearly post-datesthe lacustrine strata of Interval 4 that show no thinning to-wards the footwall.The £exure of the entire preserved ba-sin ¢ll is interpreted here as being due to forced foldingcaused by propagation of a major normal fault in the rigidcrystalline basement (see analogue models by Withjacket al., 1990; Hardy & McClay, 1999; Schlische et al., 2002,for similar examples). Contrary to the fault-related folds il-lustrated by Ford et al. (2007, e.g. their Fig. 5), which indi-cate syn-kinematic deposition throughout the faultevolution, the £exure of theMost Basin ¢ll at the northernedge of pro¢le 68/83 is post-depositional with respect toInterval 4 strata.

Post-depositional faulting along other NE-trendingfaults, both at the southeastern basin margin and, locally,within the Most Basin, is documented above, as well asthe activity of the rift-transverse, NW-trending faults,especially in the Pliocene (S� pic� a¤ kova¤ et al., 2000).

An earlier, potentially still syn-depositional, establish-ment of NW^SE-directed extension to cause the forma-tion of the NE-trending faults would be implied by theresults of Adamovic� & Coubal (1999), who infer, on the ba-sis of intrusive body geometries, a period of NW^SE ex-tension between ca. 24 and 16Ma. However, these authorsadmitted that their conclusion was based on a very smallnumber of dated intrusive bodies.

Evolution of theMost Basin in response to changing palaeostress¢elds

There is a marked contrast between the syn-depositionalrole of the E^W-trending fault arrays, typically short andarranged en-eche¤ lon along the rift axis, and the mostlypost-depositional activity of the rift-parallel, NE-trend-ing faults.Together with the interpreted extension vectorsassociated with each of the two fault populations, thisevokes a scenario of a transition from oblique extensiondominating the depositional interval to mainly post-de-positional, orthogonal extension, as demonstrated inmodels byMcClay &White (1995), Keep &McClay (1997)and Bonini et al. (1997).

Oblique extension: from initiation to fault linkage and depocentregrowth (Intervals1^3)

An oblique-extensional regime dominated the formationof accommodation and relief at E^W-trending fault arraysduring most of the basin’s recorded lifetime, at least dur-ing Intervals 1^3, and probably also Interval 4. (Fig. 15).The angle between theMost Basin axis and the E^W faults(and the axes of depocentres) is ca. 301 and most of themeso-scale structural data associated with this fault sys-tem indicate NNE (to N)-directed extension (Figs 9a and10a). This is consistent with data from other parts of theEger Graben (Peterek et al., 1997; S� pic� a¤ kova¤ et al.,2000;and, partly, Adamovic� & Coubal, 1999).This strongly



oblique-extensional setting, with a ca. 601 angle betweenthe rift axis and the extension vector, compares well withthe analogue models of Tron & Brun (1991), McClay &White (1995), Clifton et al. (2000) and McClay et al. (2002)as well as with many ¢eld examples from oblique-extensional settings (Withjack & Jamison, 1986; Morleyet al., 1992; Souriot & Brun, 1992; Brun & Tron, 1993;Bonini et al., 1997; Henry, 1998).

On several observational scales, our data document thesuccessive growth and linkage of the initial dense popula-tion of short, oblique-extensional fault segments into sev-eral major faults that bordered the depocentres betweenIntervals1and 3 of basin evolution.This is consistent withthe scenario of fault growth by gradual linkage of small in-itial faults, resulting in gradual growth of depocentres andacceleration of subsidence (Cartwright et al., 1995; Guptaet al., 1998; Gawthorpe & Leeder, 2000; McLeod et al.,2000), and contrasts with examples of fault propagationand linkage before signi¢cant basin development, shownby (Morley et al., 1999; Morley, 2002). The death of manyof the small-displacement faults from the initial stage ofbasin formation is shown in the seismic sections (Fig. 6)and the propagation of major faults during depocentredevelopment is illustrated e.g. by the syn-depositionalfault-propagation folding at the B|¤ lina Fault. The gradualbuild-up of displacement at linking faults is also re£ectedin the evolution of subsidence rates.Very slow rates of initi-al subsidence are analogous to data from the initial phasesof rift evolution in the Gulf of Suez or the Jurassic of theNorth Sea, ranging betweeno10 and ca. 30mMyr�1 andlasting severalMyr (Gupta et al., 1998;McLeod et al., 2002).

Advanced fault linkage and subsidence acceleration (Interval 4)

The transition to Interval 4 is marked by an abrupt accel-eration of subsidence, accompanied by broadening of thedepositional area beyond the preserved limits of theMostBasin ¢ll.The relative increase in the subsidence rate canbe explained by the model of Gupta et al. (1998) invokingstrain localisation on linked fault arrays and not necessitat-ing an increased strain rate or a change in the extensionvector orientation.The fact that the increase in subsidencerates occurred simultaneously in two other basins of theEger Graben, recorded by onset of the lacustrine CyprisFormation (S� pic� a¤ kova¤ et al., 2000), however, suggests thatsome overriding, regional control may have acted in thiscase. This might have been an increase in the strain rate,possibly induced by a change of the stress ¢eld.With re-gard to indices that some of the NE^SW faults may havebeen active syn-depositionally, we discuss below whetherthe increase in the subsidence rate during Interval 4 couldbe related to the beginning of a change in the palaeostress¢eld that later led to the post-depositional deformationand termination of the rift regime of the entireEgerGraben.Based on the available data, however, the preferred interpre-tation is that Interval 4 represents an advanced stage of faultlinkage that caused basin-wide deposition in depocentresdeepened along a reduced number of major faults.

Orthogonal extension: post-rift deformation and erosion of thebasin ¢ll

The post-depositional deformation described above im-plies a local NW^SE-oriented extension, orthogonal withrespect to the rift axis, and resulting in linear fault seg-ments (McClay &White, 1995; McClay et al., 2002). In ad-dition to the large-scale geometry of the NE^SW-trending faults, the orthogonal extension is supported bysome of the mesoscopic structural data and the observedlinkage of some of the E^W fault segments in the Krus� ne¤Hory Fault Zone by NE- to NNE-trending faults. Achange in the extension vector orientation, from theNNE^SSW, oblique extension, to the roughlyNW^SE di-rection of postsedimentary orthogonal extension, wasearlier interpreted byRajchl &Ulic� nyŁ (2000) andS� pic� a¤ ko-va¤ et al. (2000). Notably, the post-depositional normalfaulting occurred in an area narrower than the region ofoblique extension, generally between the Krus� ne¤ HoryFZ and the rift axis.The only exception is the%atec depo-centre downthrown post-depositionally. This part of theEger Graben evolution is referred to as post-rift becauseit involves a partial inversion and signi¢cant erosion ofthe basin ¢ll, signi¢cant reduction of volcanism and theassociated normal faulting is interpreted below as a conse-quence of regional lithospheric folding (e.g. De' zes et al.,2004).

DISCUSSION

Evidence for the transition from the oblique tothe orthogonal extensionmode

Rift-marginal faults in oblique-extensional regimes canpropagate and link to form nearly rift-parallel faults underincreasing extensional strain either due to a change in theextension direction or under unchanged stress orientation(Tron & Brun, 1991; McClay & White, 1995; McClay et al.,2002). It is therefore important to review the combinationof evidence leading to the above interpretation of achange in extension vectors in the Most Basin.A number of features in theMost Basin structural patternare analogous to those observed byBonini etal. (1997) for achange from oblique to orthogonal extension, withthe angle between the extensional vectors moderately ex-ceeding 451, and by Keep & McClay (1997) for similarsituations:

(i) numerous oblique faults displaying an en eche¤ lon geo-metry developed in the central part of the rifted do-mains during oblique extension;

(ii) almost no previous oblique faults continued to growduring the orthogonal extension, and new normalfaults developed parallel to the rift axis, as illustratedby the formation of the Ohr� e Fault Zone and NE^SW-trending segments of the Krus� ne¤ Hory FaultZone breaching former relay ramps; and

(iii) amaster fault developedduring orthogonal extension,in a zone previously weakened by small oblique faults,


M. Rajchlet al.

resulting in a wavy surface trace, accompanied bykinks, salients and embayments that mark the loca-tions of linkage ^ this is analogous to the kinked traceof the Krus� ne¤ Hory Fault, and typical of situationswith a high degree of initial extension obliquity.

Overall, the formation of prominent rift-parallel marginalfaults in an oblique-extensional regime was observed toresult from high volumes of extension (up to 50% in mod-els of McClay et al., 2002), which is not the case for theMostBasin: total stretching estimated across theMost Ba-sin is ca. 8%, based on the throw and heave distances mea-sured at pre-rift markers at the base of the basin ¢ll.Thelength and straightness of some of the NE^SW fault seg-ments (especially the Ohr� e Fault Zone) support the inter-pretation of their formation by orthogonal (NW^SE)extension.

Unlike analogue models, ¢eld examples are strongly in-£uenced by the basement structure.The di¡erences in thegeometry of the main bounding fault zones of the MostBasin (kinkedKrus� ne¤ HoryFZvs.mainly straight and par-allel Ohr� e FZ) were probably in£uenced by basement fab-rics that supported or partially overprinted the e¡ect ofextension orientation. The kinks in the Krus� ne¤ Hory FZ(of 35^501 angle between individual segments) are devel-oped in the Saxothuringian crystalline basement underly-ing this part of the rift area that is characterised by E^Wfabrics (Mlc� och, 1994). In contrast, the southeastern partof theMost Basin is underlain by anUpper Palaeozoic gra-ben of the Kladno-Rakovn|¤ k Basin (Pes� ek, 1994) de¢nedby NE^SW faults. Although this inherited fabricmight theoreticallyhave helped to direct the faults formingduring the oblique-extensional phase into parallelismwith the Eger Graben axis, the existence ofE^W faults developed in the same substratum (e.g. thefault zone followed by Ohr� e River) indicates that the twostress ¢elds led to the formation of directionallydisctinct fault populations independent of the basementstructure. We conclude that the di¡erences in the base-ment fabric between the opposite sides of the basinprobably did not lead to development of NE^SW faults inthe oblique-extensional regime, but caused a slightlydi¡erent orientation of the NE^SW fault segments oneach side.

Timing of post-depositional faulting

One of the keys to understanding the stress ¢eld evolutionduring the latest intervals of deposition in the basin is therelationship of the Krus� ne¤ Hory Fault Zone with the de-pocentres and intrabasinal faults. The composite natureof the Krus� ne¤ Hory FZ, with parts of E^W faults hard-linked across breached relay ramps or evenwith some relayramps only tiltedwithout breaching, was a source of majordisputes about the role of the Krus� ne¤ Hory FZ in the ori-gin of theMost Basin in earlier local literature (MalkovskyŁ ,1966, 1979; Hurn|¤ k & Havlena, 1984; KopeckyŁ et al., 1985;Marek, 1985;Va¤ ne� , 1985a).

With regard to the relationship between the boundaryfaults and the intra-basinal fault framework, some analo-gies may be found between theMost Basin and the RukwaRift.The latter, interpreted by Ring et al. (1992) initially asan oblique rift (see also McClay & White, 1995), later de-formed in a strike^slip regime.Although di¡erent in kine-matics and probably of a much longer history than theKrus� ne¤ Hory FZ, the Lupa Fault of the Rukwa Rift is alsoa major border fault overprinting part of the earlier depo-centres (cf. Fig.10 inMorley, 2002). Its kinked shape in themap view indicates origin by linkage of many smaller seg-ments, but, unlike the Krus� ne¤ Hory FZ, the Lupa Faulthas a signi¢cant record of syn-depositional activity. In caseof the Krus� ne¤ Hory FZ, the post-depositional activity isclear due to basin- ¢ll deformation, but it is questionabletowhich extent its formation (by coalescence of earlier ob-lique faults) may have in£uenced the basin ¢lling duringInterval 4.

Onlap of the main lignite seam on the Cretaceous sub-stratum occurs near the northern edge of seismic line 68/83, beyond 6 km, and thinning of lacustrine deposits im-mediately above the seam is observed between 5 and 6 km(Fig. 6a).This is interpreted as a record of a relay ramp evo-lution between two propagating, syn-depositional faultsegments. However, the younger part of the lacustrinestratal package of Interval 4 shows no thinning towardsthe basin margin, and is a¡ected by a clearly post-deposi-tional deformation related to fault propagation. Mostprobably, the E^W fault segments bounding this relayramp became linked before the onset of Interval 4, or theirdisplacement was transferred to another fault in a moreoutward position. The exact timing of the post-deposi-tional deformation shown in Fig. 6a, however, remains un-clear ^ it may have occurred immediately after the Interval4 termination or later, during the Miocene or even Plio-cene.

Another example of a syn-depositional relay ramp inthe vicinity of the Krus� ne¤ Hory FZ is the locality Hradis� te�(Fig. 7) where a relict of £uvial sands is preserved beyondthe trace of the Krus� ne¤ Hory FZ. A small £uvial feedersystem transporting clastics southward most probablysourced a prograding deltaic sandwedge high in theMostFormation lacustrine deposits (Va¤ ne� , 1985a) ^ during In-terval 4. Because the Hradis� te� sandstones rest on the crys-talline basement, the ramp must have been uplifted andprobably tilted before or during this depositional episode.This could have occurred during propagation of an E^Wfault segment late during Interval 3 or 4. The Hradis� te�ramp and adjacent small relay ramps to its SWwere neverbreached. However, a straight, NE^SW-trending faulttrace occurs further NW, in the basin periphery, as a con-tinuation of theKrus� ne¤ Hory FZ.This fault is most prob-ably post-depositional because its syn-depositionalactivity during Interval 4 would have caused wholesalehangingwall subsidence of the adjacent E^W segmentsand the Hradis� te� ramp.

The above lines of evidence lend further support to theinterpretation of a post-depositional change in the exten-



sion vector, fromNNE^SSW toNW^SE, that led to post-depositional faulting and related basin- ¢ll deformation inthe Most Basin region and elsewhere in the Eger Graben.Importantly, the post-depositional normal faulting wasconcentrated in a zone narrower than the syn-deposi-tional width of the Most Basin, along the Krus� ne¤ HoryFZ and,partly, theStr� ezovFault, showing a di¡erent struc-tural style than the syn-depositional fault growth thatcaused broadening of the depositional area duringInterval 4. Yet another indirect support for the post-de-positional change in the extension regime comes fromthe consideration thatwith an unchanged stress ¢eld, con-tinued linkage of normal faults would have led to a furtherincrease in the subsidence rates in the depocentres, ratherthan to their uplift and erosion (cf. Gawthorpe & Leeder,2000).

Implications for geodynamic causes of EgerGraben and ECRIS extension

With regard to the palaeostress history of the entireECRIS and the geodynamic causes of its evolution, thedata from the Most Basin in the central part of the EgerGraben are very important. All the data presented hereshow that ca.18Myr of basin- ¢lling history were governedby regional N^S to NNE^SSWextension, similar to otherparts of the Eger Graben.This rules out the idea of Bour-geois et al. (2007) that the ECRIS formed a left-lateralwrench zone fringing the Alpine front from the Mediter-ranean to the Bohemian Massif, with the Eger Grabenopening as a purely strike^slip structure.

This Oligoceneêarly Miocene local stress ¢eld of theMostBasin, however, was di¡erent from that of other partsof the ECRIS, dominated by E^W to WNWÊSE exten-sion in the western and central parts (e.g. Michon et al.,2003). It is not easy to reconcile the N^S, and later NW^SE extension, interpreted here for the Eger Graben, witha coeval continental palaeostress ¢eld characterised byN^S-oriented compression (e.g. Bergerat, 1987; De' zes etal., 2004). One possible explanation for this autonomouspalaeostress ¢eld of the Eger Graben is the hypothesis ofMichon et al. (2003) and Michon & Merle (2005) invokinga slab-pull model, with downward gravitational stressesinduced by formation of the Alpine lithospheric root,causing formation of ECRIS rifts essentially by passiverifting. In this model, the direction of foreland extensionis approximately perpendicular to the lithospheric rootand parallel to the direction of compression induced byAfricaÊurope collision.

However, the entire Eger Graben, especially its centralpart including theMostBasin, is characterised bylarge vo-lumes ofvolcanics that pre-dated the main phase of clasticsedimentation. This may support the idea of thermal-driven doming suggested byDe' zes etal. (2004) for the EgerGraben, although in general they explained the initiationof the ECRIS mainly by the build-up of syn-collisionalcompressional intraplate stresses caused by collision ofAfrica and Europe.The mantle-plume origin of European

Cenozoic rifts, commonly invoked in the 1990s (Wilson &Downes, 1992; Granet et al., 1995, among others), has beendeemphasised in the recent literature; Wilson & Downes(2006) conclude that partial melting of the mantle was in-duced by adiabatic decompression of the asthenosphere,triggered by mantle upwelling, and envisage small-scale,plume-like diapirs possibly upwelling from ca. 400 kmdepth.The evolution of theMost Basin suggests that weakextension in response to thermal doming should be con-sidered in case of the opening of the EgerGraben case, re-gardless of the origin of the melts. The relief andstratigraphic record around the Eger Graben have beenconsiderably altered by post-Miocene erosion, and thus itis di⁄cult to verify the possibility of a thermal dome evolu-tion by palaeodrainage reconstruction. However, duringInterval 2 theMostBasin ¢lling was dominated by redepo-sition of volcanic material, and in Interval 3, peat accumu-lation dominated over clastics for a long time in most ofthe basin, except the %atec Delta region, where a clasticpathway followed a transverse basement fault zone. Thisindicates that there was little signi¢cant regional clasticinput into the centre of the Eger Graben, until ca. 18Ma,which may be indirect support for regional doming.

The later, NW^SE oriented, extension related to the for-mation of the Krus� ne¤ Hory FZ and other present-daybounding faults of the Eger Graben represents a local pa-laeostress ¢eld. Its orientation contrasts with the Miocene^Pliocene build-up of NW^SE compression that dominatesCentral Europe today (Cloetingh & Kooi, 1992; Ziegler &De' zes, 2007). The formation of a local extensional domainin an overall compressional stress ¢eld, however, lends sup-port to the idea ofMiocene-age lithospheric folding as pro-posed by De' zes et al. (2004). We interpret the post-depositional normal faulting in theMostBasin as along-crestextension during the growth of a broadly SW^NE-trending,lithosphere-scale, anticlinal feature that extends from theMassif Central via the Burgundy transfer zone towards theBohemian Massif (Fig. 1). De' zes et al. (2004) and Bourgeoisetal. (2007) dated the onset of lithospheric folding in the pre-viously thermally weakened ECRIS domain as 18 or 17Ma,respectively.This time roughly corresponds to the termina-tion of the Most Basin ¢lling (according to palaeomagneticdating, Bucha et al., 1987). Similar to other ECRIS basins af-fected by the lithospheric folding, theMostBasin is tilted andpartly eroded (cf.Bourgeois etal., 2007),with the fault-propa-gation fold shown in Fig. 6a being perhaps the most graphicevidence of basin deformation in this phase ofECRIS evolu-tion.The fact that the Eger Graben is located on the south-eastern shoulder of the interpreted lithospheric foldstructure,parallel to its axis (Fig.1), but not exactly in a crestalposition, canbe explained by the rheological heterogeneity ofthe basement.

After the Eger Graben syn-rift sedimentation was ter-minated by the lithospheric folding, the Pliocene^Quaternary sedimentation and volcanism in the regionwere con¢ned mainly to NW-trending fault zones cross-cutting the graben and activated as sinistral strike^slipzones (e.g. S� pic� a¤ kova¤ et al., 2000).The interaction of these


M. Rajchlet al.

zones with individual segments of the ancient E^W faultsystem could have caused the formation of smallpull-apart structures ¢lled by alluvial clastics (Fig. 2).The interaction between the Mid-Miocene to Recentcompressional stress ¢eld, lithospheric buckling, crustalheterogeneity and the resulting patterns of uplift, anderosion/sedimentation in the entire Eger Graben regionremains poorly understood and represents an importantchallenge for future research.

SUMMARY

TheMost Basin provides a good illustration of initiation,gradual growth and linkage of normal fault arrays in astrongly oblique-extensional situation, overprinted by apost-depositional phase of orthogonal extension thatcaused signi¢cant deformation of the basin ¢ll. The faultgeometries compare well with published analogue modelsof rifts undergoing oblique to orthogonal extension,although the di¡erent basement fabrics in the crustalblocks separated by the rift axis in£uenced the local ex-pression of some fault populations.

The preserved stratigraphic record shows that the basinevolution stopped shortly after the transition from the initialrifting stage to a more mature stagewith subsidence acceler-ated along major depocentre-bounding faults.The post-de-positional faulting and formation of an adjacent upliftresulted from a localised extensional collapse along the crestof a growing lithospheric fold.The total estimated stretchingof ca. 8% is in accordancewith the very slow subsidence ratesover most of the recorded basin history.

The timing of the basin ¢lling as well as its destructioncorrelate well with events occurring elsewhere in the EgerGraben, and also with the timing of onset of rifting (ca.37Ma) and the presumed lithospheric folding (ca. 18Ma)in the entireECRIS.The indigenous stress ¢eld, abundantvolcanism and relative clastic starvation during the de-positional Intervals 1^3 suggest that a possibility of ther-mal doming as the cause of rifting initiation should notbe excluded in further studies of the Eger Graben.

ACKNOWLEDGEMENTS

This paper is based on the Ph.D. thesis of M. Rajchl, andsummarizes the results of more than 8 years of research,supported successively from several sources, mainly GAAVC� R grant A3012705 to L. S� pic� a¤ kova¤ and GAC� R (CzechScience Foundation) grant 205/01/0629 to D.Ulic� nyŁ . Ac-quisition of seismic and gravity data was made possible bythe Ministry of Environment of the Czech Republic, con-tracts No. OG 9/02 and OG 13/02. Collection and analysisof structural data in 2006^2007 were supported by GAC� Rgrant 205/06/1823. M. Rajchl was partially supported bythe MZP0002579801 research programme of the Ministryof Environment of the Czech Republic. D.Ulic� nyŁ was sup-ported by Czech Academy of Sciences research pro-

gramme V0Z30120515. K. Mach thanks the Severoc� eske¤doly, a.s., for support.The authors are grateful to P. Coufaland O. Janec� ek (Severoc� eske¤ Doly, a.s.) for their assistancein theNa¤ stupTus� imice open cast mine, and toV. Rapprich(CGS) for discussions of volcanological aspects. The re-views by J. Cartwright and O. Bourgeois, an informal re-view by G. Randy Keller, as well as notes from P.Van derBeek as journal editor, helped to improve the paper greatly.However, the responsibility for any omissions or misinter-pretations remains solely with the authors.

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