Ivana Liq Turb 2013 IAS

Can liquefied debris flows deposit clean sand over large areas ofsea floor? Field evidence from the Marnoso-arenacea Formation,Italian Apennines

PETER J. TALLING*, GIUSEPPE MALGESINI*,� and FABRIZIO FELLETTI�*National Oceanography Centre, European Way, Southampton SO14 3ZH, UK(E-mail: [email protected])�School of Ocean and Earth Science, University of Southampton, National Oceanography Centre,European Way, Southampton SO14 3ZH, UK�Dipartimento di Scienze della Terra, Universita degli Studi di Milano, Via Mangiagalli 34, 20133Milano, Italy

Associate Editor – Jaco Baas

ABSTRACT

The Marnoso-arenacea Formation in the Italian Apennines is the only ancient

rock sequence where individual submarine sediment density flow deposits

have been mapped out in detail for over 100 km. Bed correlations provide new

insight into how submarine flows deposit sand, because bed architecture and

sandstone shape provide an independent test of depositional process models.

This test is important because it can be difficult or impossible to infer

depositional process unambiguously from characteristics seen at just one

outcrop, especially for massive clean-sandstone intervals whose origin has

been controversial. Beds have three different types of geometries (facies tracts)

in downflow oriented transects. Facies tracts 1 and 2 contain clean graded and

ungraded massive sandstone deposited incrementally by turbidity currents,

and these intervals taper relatively gradually downflow. Mud-rich sand

deposited by cohesive debris flow occurs in the distal part of Facies tract 2.

Facies tract 3 contains clean sandstone with a distinctive swirly fabric formed

by patches of coarser and better-sorted grains that most likely records pervasive

liquefaction. This type of clean sandstone can extend for up to 30 km before

pinching out relatively abruptly. This abrupt pinch out suggests that this clean

sand was deposited by debris flow. In some beds there are downflow transitions

from turbidite sandstone into clean debrite sandstone, suggesting that debris

flows formed by transformation from high-density turbidity currents. However,

outsize clasts in one particular debrite are too large and dense to have been

carried by an initial turbidity current, suggesting that this debris flow ran out for

at least 15 km. Field data indicate that liquefied debris flows can sometimes

deposit clean sand over large (10 to 30 km) expanses of sea floor, and that these

clean debrite sand layers can terminate abruptly.

Keywords Debrite, liquefied, sandy debris flow, sandy debrite, submarinedebris flow, submarine fan, turbidite, turbidity current.

INTRODUCTION

Submarine gravity flows dominate sedimenttransport into many parts of the deep ocean andproduce submarine fans that include some of the

most extensive and voluminous sediment accu-mulations on Earth (Allen, 2007; Nielsen et al.,2007; Talling et al., 2007a). The scale of theseflows can be impressive because a single flowcan transport more sediment than the combined

Sedimentology (2013) 60, 720–762 doi: 10.1111/j.1365-3091.2012.01358.x

720 � 2012 The Authors. Journal compilation � 2012 International Association of Sedimentologists

annual flux from all of the world’s rivers, such asin the largest beds described here (Talling et al.,2007b,c). Ancient flows of this type have depos-ited thick rock sequences that now hold some ofthe world’s largest oil and gas reserves. Theseflows are capable of depositing relatively thick(0Æ5 to 3 m) layers of sand across large areas of thedeep sea floor, and these layers are a majorbuilding block of submarine fans (Nielsen et al.,2007). These sand layers frequently contain thickintervals of relatively clean sandstone with lowinterstitial mud content that lack sedimentarystructures, such as planar or cross-lamination;and these intervals are often referred to as the TA

division of the Bouma (1962) sequence.Submarine sediment density flows have proven

to be notoriously difficult to study through directmonitoring. The sediment concentration profileof long run-out flows that reach submarine fansbeyond the continental slope has never beenmeasured directly, at any location. Understand-ing what the flows are, and how they transportand deposit sediment, remains a major challenge.Current understanding of these flows is thereforebased heavily on the deposits that they leavebehind. The origin of massive clean (TA) sandlayers has been the subject of particularly vigo-rous debate (for example, Shanmugam & Moiola,1995, that resulted in seven separate replies;such as Lowe, 1997) and a number of differentprocesses have been proposed to account for theirorigin (e.g. Walker, 1965; Kuenen, 1966a; Mid-dleton, 1967; Lowe, 1982; Arnott & Hand, 1989;Mutti, 1992; Kneller & Branney, 1995; Shanmu-gam & Moiola, 1995; Vrolijk & Southard, 1997;Stow & Johansson, 2000; Leclair & Arnott, 2003,2005; Mutti et al., 2003; Sumner et al., 2008;Breien et al., 2010).

It can be very difficult (or impossible) to deter-mine unambiguously how massive clean-sandintervals are deposited using only informationavailable from a single outcrop, as summarized ina single vertical sedimentary log. This ambiguityis one of the reasons for continued debate over theorigin of massive clean sandstone. For instance,Kneller & Branney (1995) concluded that massivesand deposition occurred progressively from azone of liquefied sediment, but acknowledged thatit was not possible to determine whether thissustained liquefied zone is supplied by verticalsediment settling from an overlying vigorouslyturbulent flow, or by lateral motion of a liquefiedflow layer. As noted by Kneller & Branney (1995),the way in which such a liquefied zone is suppliedis of profound importance for understanding

submarine flows, because it determines the loca-tion and geometry of massive sandstone intervals,their lateral relations to other types of depositsand the distance that flows run out.

The Miocene Marnoso-arenacea Formation isvery unusual in that individual flow deposits(beds) can be correlated over a large area (up to120 · 30 km) between numerous (up to 109)individual outcrops (Fig. 1; Ricci Lucchi & Val-mori, 1980; Amy & Talling, 2006; Talling et al.,2007c). The present authors are aware of no otherancient submarine fan sequence in which indi-vidual beds have been correlated in such detailover such a large area. The correlated beds weredeposited in a relatively flat basin plain withoutchannels (Fig. 2). Bed correlations are madepossible by a series of distinctive mega-turbiditemarker beds that provide the basis for correlatingintervening beds, and by the almost complete lackof erosional bed amalgamation (Ricci Lucchi &Valmori, 1980; Amy & Talling, 2006; Tallinget al., 2007c). The marker beds have been mappedout precisely by the Emilia-Romagna, Umbria andMarche geological surveys over many years (e.g.Martelli et al., 1994). Beds are consistently sepa-rated by intervals of hemipelagic mudstone(Talling et al., 2007c). Most submarine fan depo-sits are characterized by far more frequent bedamalgamation that would prevent long-distancebed correlation, even if suitable marker beds (anddetailed geological maps of those marker beds)were to be available.

The bed correlations in the Marnoso-arenaceaFormation are important because they provide theadditional information needed to better constrainhow massive clean sand is deposited by sub-marine density flows (Amy et al., 2005; Amy &Talling, 2006). This information comprises theexternal shape of clean-sandstone intervals andthe internal bed architecture, with the lattershowing how various types of clean sand in asingle bed are arranged in downflow trending‘facies tracts’ (Mutti, 1992). This additional infor-mation provides an independent test of sanddepositional models, which are based initially onfeatures visible at the scale of a single outcrop.

Aims

Ricci Lucchi & Valmori (1980) showed that bedsin the Marnoso-arenacea Formation could becorrelated between 18 sections across an area of120 · 30 km. Subsequent work provided moredetailed bed correlations between 109 locationsfor a ca 30 m thick interval of strata immediately

Deposition of deep-water massive sandstone 721

� 2012 The Authors. Journal compilation � 2012 International Association of Sedimentologists, Sedimentology, 60, 720–762

above the most prominent mega-turbidite markerbed, called the Contessa Bed (Amy & Talling,2006; Talling et al., 2007b,c). The first aim ofthis article is to outline a detailed bed correla-tion framework for a second interval of strataimmediately below the Contessa Bed (Figs 1, 2and 3). The new bed correlations documentdownflow changes in bed architecture, or ‘faciestracts’ in the sense of Mutti (1992). Four types offacies tract occur in the new stratigraphic inter-val, and only two of these bed geometries wererecognized in previous bed correlations.

The most likely depositional process for eachtype of clean sandstone is then inferred from theinformation visible at the scale of a single outcrop(Figs 4, 5 and 6). The second aim was to test thesemodels using the external shape and lateralarrangement of the various types of sandstone,using the long-distance bed correlations. Why dosome types of clean sandstone with particulardistinctive internal textures pinch out relativelyabruptly, whilst other types of clean sandstonetaper more gradually in a downflow direction?The present authors conclude that a certain type

Fig. 1. (A) Location map showing the northern part of the outcrop of the Marnoso-arenacea Formation. The figureshows the position of measured sections and the transect panels presented in subsequent figures. The RidracoliTransect is shown in Fig. 3. Sections are numbered as follows (Amy & Talling, 2006): 109 – Cabelli River; 1 – Cabelli;77 – Mantigno; 3 – Aquadalto; 85 – Cavalmagra; 84 – Marradi 2; 83 – Marradi 1; 105 – Il bagnato; 76 – Gemelli; 78 –Fiumicello Zohotecnica; 26 – Lavacchio; 71 – Corniolo; 73 – Ridracoli 2; 74 – Ridracoli 3; 106 – Ridracoli 4; 40 –Pietrapazza; 107 – Valanello; 42 – Bagno di Romagna; 108 – Poggio Pandella; 8 – Poggio Dornata; 11 – Montefreddo;12 – Bocconi; 14 – Monte Roncole; 13 – Premilcuore; 28 – Badia; 29 – Cabelli; 30 – Isola; 44 – Castel Priore; and 33 –Galeata. (B) Palaeocurrent directions measured from flutes and grooves at the base of the beds in the interval belowthe Contessa mega-bed. See Fig. 2 (and Talling et al., 2007c) for a summary of palaeocurrent measurements from bedsabove the Contessa mega-bed.

722 P. J. Talling et al.


of massive clean sandstone with a swirly orpatchy texture was deposited by liquefied debrisflows.

The third aim was to understand how suchliquefied clean-sand debris flows could originate;are they far travelled from outside the outcrop

Fig. 2. (A) Overview of the Marnoso-arenacea foredeep basin during the Serravallian, including the sources anddirections of different flows (after Di Base & Mutti, 2002). (B) Basin plain palaeogeography during deposition of bedsabove the Contessa Bed. Note the consistent mean palaeoflow direction in the 109 logged sections. Also shown is theposition of the Verghereto High, where correlated strata thin but flows appear not to be deflected strongly. Thrustfaulting that post-dated deposition of the correlated beds split the outcrops into structural elements (thrust sheets).The relative position of thrust sheets is shown before thrust motion subsequently narrowed the basin plain. SeeTalling et al. (2007c) for a full description of basin plain palaeogeography.



Fig. 3. Cross-section along the Ridracoli thrust sheet showing the correlation of the first 20 thick beds below theContessa marker bed. Figure 1A shows the location of this transect. The correlation panel is oriented approximatelyparallel to palaeoflow. The palaeoflow direction is north-west to south-east for all the correlated beds, but in theopposing direction for the Contessa marker bed.



area, or did they form via local flow transforma-tion? It has previously been suggested that lique-fied flows of sand are not likely to travel for morethan ca 1 km, and are generally restricted to steep(>1� to 3�) gradients (Lowe, 1976). Therefore, thequestion of how liquefied layers of sand couldpotentially run out for longer distances (tens ofkilometres), across much lower gradient (<0Æ1�)basin plains is discussed here.

Two sets of previous studies have inferred thatclean sand can be deposited by dense liquefiedflow (Mutti, 1992; Mutti et al., 2003, 2009) ordebris flow (Shanmugam & Moiola, 1995; Shan-mugam, 1997, 2000, 2002), although these studieslacked detailed information on sandstone intervalshape to test such models. The fourth aim is tocompare clean-sandstone debrites in the corre-lated beds with the deposits that these authors

described, and to determine whether similar ordifferent criteria are used to infer deposition fromliquefied debris flow.

Terminology

The terminology used in this article follows thatof Talling et al. (2012b), which provides moredetailed definitions.

Turbidity currentThe term turbidity current denotes a turbulentsediment suspension from which larger grainstend to segregate and settle preferentially, depo-siting a turbidite; they tend to have lower sedimentconcentrations than debris flows, liquefied flows orfluidized flows. Fluid turbulence is the primarymechanism of sediment support within a turbidity

A

B C

D E

Fig. 4. Outcrop photographs illus-trating the different types of lami-nated turbidite sandstone seen inthe outcrops, also described in Ta-ble 1. (A) Photograph of the corre-lated interval at Section 1 (Coniale)showing the position of the Cont-essa and Fiumicello marker beds.This section is the most proximaloutcrop for flow deposits that camefrom the north-west. (B) Ripple-scale cross-laminated sandstone(Cs1) in Bed )7 at Section 8 (PoggioDornata). (C) Parallel laminatedsandstone (Cs3) in Bed )7 at Section8. (D) Stepped laminated sandstone(Cs4) in Bed )26 at Section 44(Castel Priore). (E) Convolute ripple-scale cross-laminated sandstone(Cs1) in Bed )20 at Section 8.



current. Turbidity currents are subdivided accord-ing to whether they are high density or low density.Sedimentation from low-density turbidity cur-rents is not hindered, and turbulence is notdamped significantly near the bed, often allowingthe formation of bedforms such as ripples or dunes.High-density turbidity currents are characterizedby elevated near-bed sediment concentrations thatare sufficient to damp turbulence and cause sed-iment settling to be hindered. High-density tur-bidity currents can generate relatively thin layersof sheared and tractionally reworked sediment onthe bed, which are laminar or very weakly turbu-lent. These thin near-bed layers are termed tractioncarpets (Hiscott, 1994; Talling et al., 2012b); theydiffer from laminar debris flows in that they are

driven by the momentum of the overlying turbu-lent flow, rather than by their own momentum.

Debris flowDebris flow denotes a laminar or very weaklyturbulent flow in which sediment is supportedmainly by processes other than fluid turbulence,which include excess pore fluid pressure, grainto grain interactions or buoyancy due to areduction in the relative densities of clasts andmatrix (Talling et al., 2012b). This definitionencompasses ‘debris flows’ of the type describedby Iverson (1997), Major and Iverson (1999),Iverson & Vallance (2001) and Iverson et al.(2010) in a large-scale experimental facility. Thebody of work by these authors emphasizes the

A

C

D

B

Fig. 5. Outcrop photographs illus-trating the different types of massivesandstone deposits. See Table 1 fora full description of lithofacies Ms,Sub-facies Cs1 and Cs6. (A) Bed )19at Section 30 (Isola) that comprises athin basal clean-sandstone interval(Cs6) overlain by a thick muddysandstone interval (Ms1). (B) Mud-rich sandstone with dispersed mudclasts (Ms1) in Bed )19 at Section 8(Poggio Dornata). (C) Sharp grain-size break between a massive gradedsandstone interval (Cs5) and a cross-laminated interval (Cs1) in Bed 7Æ9at Coniale 1. (D) Basal massive un-graded clean sandstone (Cs6) in Bed)32 at Section 44 (Castel Priore).

Fig. 6. Textural characteristics of: (A) Sub-facies Cs5 in Bed )6; (B) Sub-facies Cs7 in Bed )3; (C) Sub-facies Cs6 inBed )10; and (D) Sub-facies Ms1 in Bed 2Æ5. Sandstone texture is characterized by measurements of the mean grainsize, the coarsest 5% of the grains and the percentage of interstitial mud finer than 20 lm in SEM images. Larger-scaletextures are shown by large thin sections, photographed in transmitted light. Samples for Sub-facies Cs5, Cs6 and Cs7were collected at Section 77 (Mantigno). The sample presented for Sub-facies Ms1 was collected at Section 14 (MonteRoncole).





importance of excess pore fluid pressure indetermining debris flow behaviour, and howthe down-slope component of sediment weightcan sometimes be supported fully by the excesspore pressure (Iverson et al., 2010). A keyfeature of debris flows is that they are drivendown-slope by their own weight and momen-tum. This feature distinguishes debris flowsfrom ‘traction carpets’ (e.g. Hiscott, 1994) or‘sustained liquefied zones’ (Kneller & Branney,1995) that are driven primarily by overlyingturbulent flow.

LiquefiedThe term fully liquefied denotes a situation inwhich excess pore fluid pressure is supportingthe entire weight of the sediment; the term partlyliquefied is used when most of the sedimentweight is borne by excess pore fluid pressure.The term liquefied is adopted for situations inwhich it is unknown whether all, or a significantpart, of the sediment weight is borne by excesspore fluid pressure. The excess pore fluid pres-sures in liquefied flow originate through consol-idation of suspended sediment, and do notinvolve an external source of fluid (Lowe,1976). The work of Iverson (1997), Major andIverson (1999) and Iverson et al. (2010) showshow debris flows are sometimes fully liquefied(especially in their central or parts) or partlyliquefied.

FluidizedFluidized flow involves an additional externalsource of fluid, unlike liquefied flow (Lowe,1976). Pore fluid passes upwards through thefluidized sediment, generating an upward direc-ted drag force that is sufficient to keep thesediment suspended. There is therefore a criticalupward fluid velocity that will fluidize sedimentgrains of different sizes (e.g. Lowe, 1976). Thefluidization velocity necessary to support sandwould result in very rapid dissipation of excesspore fluid pressures if there is no external sourceof fluid.

Continuum between non-cohesive, poorlycohesive, and cohesive debris flowThe character of debris flows changes profoundlyas cohesive mud is added to the sediment mixture(e.g. Iverson, 1997; Marr et al., 2001; Iversonet al., 2010). Cohesion results from colloidalsurface electro-chemical bonds that form betweenfine-mud particles comprising clay minerals,such as illite, smectite, kaolinite, montmorillonite

and bentonite, that occur commonly in marinesettings (Shaw, 1992; Coussot, 1997; McAnallyet al., 2007). Changes in cohesive mud contentcan alter debris flow yield strength, viscosityand the rate at which excess pore pressuresare dissipated by several orders of magnitude(Coussot, 1997; Iverson et al., 2010), therebycontrolling run-out distances and processes ofsediment deposition (Baas et al., 2009, 2011;Sumner et al., 2009).

Non-cohesive debris flows are defined as lack-ing any cohesive mud particles and have zerocohesive strength. Cohesive debris flows aredefined as having muddy pore fluid with suffi-cient cohesive strength to support sand grainswithin the debris flow matrix indefinitely,although this strength may be insufficient tosupport much larger clasts; this results in enmasse consolidation of the matrix sand and mud,and prevents preferential settling of sand grainsthrough the muddy pore fluid. Poorly cohesivedebris flows contain some cohesive mud, but thecohesive strength of muddy pore fluid is insuffi-cient to support sand grains.

DebriteDebris flow deposits are termed debrites anddebrites can therefore be deposited by fullyliquefied or partly liquefied flows. Non-cohesive,poorly cohesive and cohesive debrites are depos-ited by non-cohesive, poorly cohesive and cohe-sive debris flows. Debrites lack well-developedtractional structures, such as planar or cross-lamination, and can therefore be clearly distin-guished from some types of turbidite sandstonesforming TB, TC and TD divisions of the Boumasequence. Debrites with abundant, chaoticallydistributed clasts and ungraded matrix withunusually high interstitial mud contents provideevidence of en masse consolidation from cohe-sive debris flows (Wood & Smith, 1959; Haughtonet al., 2003; Talling et al., 2004, 2007a; b; Haugh-ton et al., 2009).

Clean sandstone and mud-rich sandstoneThe terms clean sand and mud-rich sand are usedin this article to describe the relative amount ofinterstitial fine mud within sandstone intervals.The amount of fine mud within a sandstoneinterval is important for two reasons. Firstly, itstrongly influences the permeability and petro-leum reservoir quality of the sandstone. Secondly,mud with cohesive properties strongly influencesflow behaviour. Cohesive (colloidal) bondsbetween mud particles become increasingly



important as grain sizes decrease below ca 40 to10 lm (McAnally et al., 2007). However, the valueof this threshold grain size also depends onmineralogy, as does the strength of the bonds(Shaw, 1992; Coussot, 1997; McAnally et al.,2007). Here a threshold grain-size range of either20 lm or 30 lm is used for estimating the volumefraction of cohesive fine mud. The authorsacknowledge that these measurements do notconstrain the mineralogy of the fine-mud fraction,and that some of the fine-mud grains may benon-cohesive.

The way in which grain size is measured cancause significant variations in estimated fine-mudcontent. For instance, it is difficult to comparemeasurements of grain diameters made fromdisaggregated samples or from images of thinsections (in which grains are sliced, therebyreducing their measured diameter), or fromimages of thin sections with different resolutionsmade using optical or scanning electron micro-scopes (SEMs). The approach here is to compareestimates of fine-mud content measured using thesame technique for the same beds. Clean sand-stone denotes relatively low volume fractions offine-mud, such as those seen in sandstone inter-vals deposited by turbidity current. Clean-sanddebrites would therefore contain a similar volumefraction of fine-mud as many turbidite sand-stones. Conversely, mud-rich sandstone has ahigher interstitial mud fraction than that com-monly found in turbidite sandstone.

SortingClean sandstone according to the definitionadopted here can have large variations in sorting.Many of the clean sandstones described in thisarticle have poor sorting, with standard devia-tions of 300 long-axis measurements ranging upto 300 lm, for average grain long axes of up to400 lm (see Fig. 7B and D). As these grain-sizemeasurements are for sliced grains (Johnson,1994), and do not include the fine tail of grains<20 to 30 lm, the Folk & Ward (1957) equation forsorting of disaggregated grains is not used. Use ofstandard deviation of long-axis measurementsfollows that of Sylvester & Lowe (2004).

Depositional setting

The Marnoso-arenacea Formation is late Burde-galian to Tortonian in age (ca 17 to 7 Ma) andcrops out in the northern Italian Apennines(Figs 1 and 2). It represents one of a series offoredeep basins that progressively migrated to-

wards the north-east, in response to thrust frontmigration in the same direction (Ricci Lucchi &Valmori, 1980; Gandolfi et al., 1983; Ricci Lucchi,1986; Van Wamel & Zwart, 1990; Martelli et al.,1994; Di Base & Mutti, 2002; Mutti et al., 2002;Roveri et al., 2002; Lucente, 2004; Amy & Talling,2006; Talling et al., 2007b,c). The late OligoceneMacigno Formation and the early Miocene Cer-varola Formation represent the fill of older fore-deep basins located further to the south-west, andcrop out in the innermost belt of the Apenninechain, while the younger Messinian to Pleisto-cene basin fills are buried below the Po Plain andthe Adriatic Sea (Pieri & Groppi, 1981). North-west to south-east orientated thrust structuressubdivide the outcrop area into tectonic thrustsheets, with the major periods of deformationoccurring after deposition of the studied strati-graphic interval (Figs 1 and 2; Martelli et al.,1994; Lucente, 2004; Talling et al., 2007b,c).Flows that had different sources traversed thisbasin plain in opposite directions (Fig. 2). Mostflows entered the basin from the north-west andoriginated from the Alpine and Apennine orogens(Gandolfi et al., 1983). A number of particularlylarge flows with a distinctive carbonate-richcomposition occasionally entered the basin plainfrom the south-east, covering the entire outcroparea and producing laterally extensive ‘mega-turbidites’ that are useful as stratigraphic markerhorizons (Ricci Lucchi & Valmori, 1980; Martelliet al., 1994). The thickest mega-turbidite is calledthe Contessa Bed (Fig. 3A) and it is formed oflimestone fragments and Apennine-derived lithicfragments. It is an excellent marker horizonbecause of its unusual thickness, distinctivecomposition and palaeocurrent direction. TheContessa Bed was deposited between ca 14 Maand 14Æ5 Ma in the earliest Serravallian (VanWamel & Zwart, 1990) and it is interpreted tohave originated from the collapse of a limestoneplatform that bounded the basin at its south-eastmargin (Gandolfi et al., 1983).

This study initially presents new correlationsfor beds in the stratigraphic interval immediatelybelow the Contessa Bed (Fig. 3A). It then incor-porates insights from key beds in the previouslycorrelated interval above the Contessa Bed (Amy& Talling, 2006; Talling et al., 2007b,c). Bothintervals were deposited in a relatively flat basinplan, as inferred from the ability of the flows totransverse the area in opposite directions, theabsence of channelization and the continuoussheet-like general bed geometry (Ricci Lucchi &



Valmori, 1980; Amy & Talling, 2006; Tallinget al., 2007a).

METHODOLOGY

Sedimentary logs

New field observations are presented for thestratigraphic interval below the Contessa Bed.Nineteen sections were logged in a downflowtransect along the Ridracoli thrust sheet (Fig. 3),and a further nine sections were logged withinthe Isola and Pianetto thrust sheets (Fig. 1A).Sections were logged at a scale of 1 : 10 or 1 : 5,and grain size was estimated in the field using agrain-size comparator. Palaeocurrent directionswere measured using flutes and grooves on thebase of thick beds (Fig. 1B).

A bed-numbering scheme was adopted that issimilar to the scheme used by Ricci Lucchi &Valmori (1980) such that thick (>40 cm) beds arenumbered sequentially from )1 to )37. Sectionson the Ridracoli thrust sheet extended from theContessa Bed to Bed )20 (Bed A-20 in thescheme of Ricci Lucchi & Valmori, 1980), whilstsections logged on the Isola and Pianetto thrustsheets extended down to the Fiumicello Bed(Bed )37).

Rates of interval thinning or pinch out

Rates of thinning of sandstone intervals can bemeasured by dividing the change in thickness bythe distance between the two adjacent loggedsections. Individual outcrops comprise well-exposed strata that are often continuous for tensto several hundred metres, and well-exposed

Fig. 7. Mean grain size plotted against mud-matrix content and sorting (defined here as the standard deviation ofgrain-size measurements). Grid counting was used to measure the longest axis of 300 grains in each sample. (A) and(B) Data from scanning electron microscope (SEM) images of thin sections. The mud-matrix content was estimated asthe percentage area occupied by grains finer than 20 lm. Samples from a Cs7 interval in Bed 5 at Section 29 areindicated, where the Cs7 interval contains large sandstone clasts (Fig. 10). (C) and (D) Data from optical microscopeimages of thin sections. The mud-matrix content was estimated as the percentage area occupied by grains finer than30 lm.



outcrop can extend in some places for up to2Æ5 km. However, strata are not exposed betweenthese outcrops. This means that only the mini-mum rate of pinch out can be calculated, bydividing the thickness of the interval by thedistance to the next exposed outcrop where thesandstone interval is found to be absent.

Facies scheme

The facies scheme adopted in this articledescribes strata hierarchically, initially as litho-facies based on the dominant lithology, and thenas sub-facies based on sedimentary structures andgrading patterns (Table 1 ). These facies are usedto describe the internal architecture of each bed.

Sandstone grading, sorting and mud content

Samples were taken to provide information aboutthe textural characteristics of different sandstonefacies, including fine-mud content and verticalgrading. Thin-section analysis measures obliqueslices through grains, rather than their true long-axis length (Johnson, 1994) but this techniqueprovides a consistent method for documentingrelative changes in grain size vertically through abed (Talling et al., 2004). Such data suffice todocument relative changes in grain size (grading)and mud content. Textural data are laborious tomeasure and few other studies have presentedmeasurements for over 50 samples from ancientsubmarine sediment density flow deposits (c.f.Sylvester & Lowe, 2004).

Analysis of thin sections with opticalmicroscopyImages from an optical microscope were analysedinitially for 171 samples from the above-Contessainterval of beds. These optical images were foundto allow more rapid, if somewhat less detailed,data collection via SEM images. Mean grain sizewas determined by measuring the long axis of 100framework grains, selecting grains at grid pointswithin a thin section. The percentage of grainsat grid points with long axes of <30 lm, as aproportion of clearly detrital grains, defines thefine-mud content of the sample. The grain-sizethreshold of 30 lm was chosen because it is theminimum size of grains that can be seen inthe optical images, which is determined by thethickness to which the thin section is cut. Cleanturbidite sandstones contained <10 to 25% mudfiner than 30 lm when measured in this way.

Analysis of thin sections using scanningelectron microscope imagesA total of 70 samples were collected for SEManalysis from Beds )2 to )11 in the Mantignosection (section number 77 in Fig. 1) from belowthe Contessa section. A further 82 samples weretaken from Beds 2Æ5, 3 and 5Æ1 in the above theContessa interval. Vertical trends in grain size,sorting and matrix-mud content within thesesandstone beds were quantified using imagesfrom a SEM in backscatter mode. Three imageswith 150 times magnification were taken for eachthin section. The longest axis of 100 grainscoarser than 20 lm was measured from eachphotograph using the image analysis software(Image J, National Institutes of Mental Health,Bethesda, MD, USA). The mud-matrix contentwas measured for each image as the percentagearea occupied by grains finer than 20 lm. Thislower grain-size threshold was adopted becausethe SEM images could clearly image smallergrains much more clearly. Clean turbidite sand-stones contained <14% (and typically <10%)mud finer than 20 lm when measured in thisway.

Visual analysis of large thin sectionsTranslucent slices of rock measuring 10 cm by8 cm were made from samples of Beds )2 to )11in the Mantigno section. These ‘large thin sec-tions’ were viewed and photographed in trans-mitted light. This technique facilitated thedetailed study of sedimentary fabrics that areoften difficult to discern in outcrops. See Garton& McIlroy (2006) for a full description of thistechnique. Large thin sections were especiallyuseful to document the ‘patchy’ texture seen inCs7 sandstones, which had previously beenproblematic to capture using optical or SEMimages that produced highly irregular verticalgrain-size trends.

RESULTS

Bed correlations

Bed geometry is shown by a series of verticalcross-sections for the new stratigraphic intervalimmediately below the Contessa Bed. The corre-lations presented here document bed architecturein a downflow trending cross-section along theRidracoli thrust sheet (Figs 1 and 3). Bed )20 isunusually thick, and was used as a marker bed for



Table

1.

Facie

ssc

hem

ead

op

ted

inth

isart

icle

.









correlations along the Ridracoli thrust sheet(Fig. 3). Beds were correlated by empirical pat-tern matching of vertical bed sequences, betweensections spaced typically every 2 to 5 km (Fig. 1).Almost all (36 of 37) thick beds were correlatedbetween all of the 31 logged sections in thisinterval.

Palaeocurrent directions

Palaeocurrent measurements at the base of thicksandstone beds indicate flow from the north-westfor all beds in the interval (Fig. 1B), except for theContessa Bed that shows an opposing flow direc-tion. Palaeocurrent indicators are remarkablyconsistent and indicate that flow occurred ina direction sub-parallel to the basin margins(Fig. 1B), which is similar to the palaeocurrentdirections measured for beds in the previouslystudied interval above the Contessa Bed (Amy &Talling, 2006). Talling et al. (2007c) provided adetailed analysis of palaeocurrent data and basinplain palaeogeography for the above the Contessainterval (Fig. 2B).

Character of different types of sandstonelithofacies at single outcrops

Sandstone layers are first described based on thefeatures that can be seen at the scale of a singleoutcrop (Table 1; Figs 4 and 5). In this sequence,there is little lateral variation in the beds across asingle outcrop (typically ten to several hundredsof metres wide), such that bed character can besummarized by a single vertical graphical log. It isthis type of vertical bed structure at a singlelocation that has previously been available formost turbidite beds in ancient outcrops and cores.

Mud-rich sandstone (Ms1 and Ms2)Swirly weathered, poorly sorted, ungraded mud-rich sandstone forms lithofacies Ms. (Figs 5A, 5Band 6D). The characteristic swirly fabric, coupledwith a distinctive grey hue and the poor sortingensure that these deposits are easily recognizablein the field. Scanning electron microscope thin-section analyses show absence of vertical gradingin this deposit type (Fig. 6D) and particularlyhigh (up to 50%) fine mud-matrix content (Fig. 7).Clast-rich (Sub-facies Ms1) or clast-deficient(Sub-facies Ms2) deposits are common in theinterval studied here. Millimetre to metre-longturbidite mudstone and hemipelagic marl clastsare randomly scattered through intervals of

Sub-facies Ms1. Muddy sandstone deposits arepresent in a number of thick beds, sandwichedbetween massive or laminated sandstone inter-vals and separated from the overlying laminatedclean sandstone by a sharp grain-size break.

Cross-laminated clean sandstone (Cs1 and Cs2)Ripple scale cross-lamination (Cs1; equivalent toTC) with wavelengths of <20 cm dominates thesandstone component of many thin beds, andoccurs in the upper part of the sandstone in thickbeds (Fig. 4B and 5C). Ripple cross-laminationtends to occur in finer-grained sandstone inter-vals (125 to187 lm using a grain-size compara-tor). Dune-scale cross-lamination occurs moreinfrequently, either within thick beds or imme-diately below mud-rich debrite sandstone (Ms).

Planar-laminated clean sandstone (Cs3 andCs4)Two distinct types of planar lamination areobserved. Finely (<1 to 3 mm) laminated clean-sand intervals (Cs3; Fig. 4C) are a major buildingblock of thicker beds, and occur above massiveclean-sand intervals, broadly equivalent to the TB

division. The laminations may be planar and sub-horizontal, or in fewer cases wavy and moreirregular. More widely (3 to 10 mm) steppedlaminae (Cs4; Fig. 4D) tend to occur beneathmassive clean-sand intervals, at the base of thickbeds (Sumner et al., 2012); they occur in coarser-grained sandstone (typically 500 to 750 lm asmeasured with a grain-size comparator) than thefine planar lamination. These laminae have astepped grain-size profile and are not inverselygraded (Sumner et al., 2012), but are generallysimilar to the spaced lamination of Hiscott &Middleton (1979, 1980).

Normally graded massive clean sandstone(Cs5)Normally graded, massive clean-sandstone inter-vals (Cs5) commonly form the basal interval ofthick sandstone beds (such as Bed )6 in Fig. 6A).The mud-matrix content is relatively low andcomprises between 5% and 25% in opticalimages (Fig. 7A), and 1 to 14% of measurementsin SEM images (Fig. 7B). These differences are aresult of the much higher resolution of SEMimages, as optical images do not resolve grainsmuch finer than ca 30 lm (the thickness of theslide). Scanning electron microscope data are amore accurate record of the percentage of finergrains. Clasts occur along distinct horizons if



they are present. This lithofacies is completelystructureless and lacks a patchy texture, and dishand pillar structures. Clean massive sandstone isassigned to the Cs5 lithofacies if vertical normalgrading is recognizable using a grain-size com-parator and hand lens in the field. This tech-nique tends to record the coarsest 5% of thegrain-size distribution seen in thin-section data.For a subset of beds, more precise measurementsof grading were obtained from multiple thinsections in a vertical transect through the bed.The thin-section data in most cases confirm thepresence and patterns of grading seen in the fieldusing the grain-size comparator (Fig. 6). How-ever, limits to the precision of both methodsensure that very subtle grading patterns wouldnot be identified. The Cs5 intervals commonlygrade upwards into planar laminated (Cs4) cleansandstone without an intervening grain-sizebreak. However, in a smaller number of cases, agrain-size break separates Cs5 intervals fromoverlying planar lamination (Figs 8 and 9) orripple cross-laminated sandstone (Fig. 5C).

Ungraded massive clean sandstone (Cs6)Sub-facies (Cs6) is formed by massive andungraded clean sandstone, such as in Bed )2and Bed )10 in Figs 3 and 9. The relatively lowmud content is similar to that seen in other clean-sandstone lithofacies (Figs 6C and 7). Fieldobservations and subsequent thin-section analy-sis show an absence of normal grading, within theacknowledged measurement uncertainties. Sub-facies Cs6 deposits are usually structureless(Fig. 5D), but in few cases a subtle, decimetre-spaced parallel lamination can occur, such as inBed )2 in Section 26 (Fig. 9). A sharp grain-sizebreak always separates Sub-facies Cs6 from over-lying laminated sandstone (Figs 8A and 9). Thereare no mudstone or sandstone clasts present inCs6 intervals.

Massive clean sandstone with a swirly orpatchy texture (Cs7)A swirly weathering pattern distinguishes Sub-facies Cs7 in the field from other massive andnormally graded clean sandstones (Figs 10 to 13),which results from contorted areas of coarser-grained and better-sorted sandstone. This patchyor swirly texture can sometimes be subtle and it isnot easily recognizable except in good exposure.However, large thin slices clearly show thispatchy or swirly texture when viewed in trans-mitted light (Fig. 6B). Field observations andthin-section analysis show that the Cs7 sub-facies

has similar sorting to other clean-sandstonesub-facies (Fig. 7). The Cs7 sandstone intervalshave an irregular grading pattern due to thecoarser grain-size patches, which is difficult tocapture with traditional thin sections (Fig. 6B).The very uppermost part of the Cs7 intervalscan be weakly normally graded. The upperboundary of Cs7 intervals is always a sharpgrain-size break.

The grey hue usually associated with mud-richswirly sandstones (Ms lithofacies; Figs 5A, 5Band 6D) is absent for the Cs7 sub-facies. Rela-tively low mud content (< ca 14%; Fig. 7A) isconfirmed by thin-section analyses, and this mudcontent is similar to that in other clean-sandstonelithofacies (Fig. 7). Small (<5 cm) mudstoneclasts often occur chaotically dispersed withinSub-facies Cs7. Large (centimetres to metres)contorted sandstone clasts were seen in Sub-facies Cs7 at a single location (Cabelli River;Fig. 14). These clasts comprised unstructured orparallel laminated clean sandstone. The clastsformed most of the Cs7 interval, with the clastsize increasing upwards (Fig. 14). The Cs7 depo-sits comprise the basal interval of a number ofthick sandstone beds, although in some casesthey are underlain by normally graded sandstonethat is up to 20 cm thick (Figs 5, 6 and 15). Asharp grain-size break separates Sub-facies Cs7from overlying ripple cross-laminated sandstoneintervals (Figs 6B, 8 and 15).

The swirly or patchy texture is different fromdish and pillar structures that have been de-scribed previously (Fig. 13A; Lowe & LoPiccolo,1974), which are relatively common within mas-sive clean sandstones in other sequences, and areinferred to form via dewatering of previouslydeposited sediment in situ (Fig. 13A). The swirlyor patchy pattern observed in Cs7 intervals ismore chaotic and lacks the convex-up dishes orsystematically sub-vertically oriented pillars(Fig. 13A). Dish and pillar structures were notseen in the correlated beds, although they arewell-developed in younger (Tortonian) beds with-in the Marnoso-arenacea Formation. The swirlyor patchy texture also differs from texturesformed by convolution of lamination duringpost-depositional water escape (Fig. 13B), biotur-bation or the mixed slurry facies of Lowe &Guy (2000) that is developed in mud-rich sand(Fig. 13D). It lacks well-defined circular or ovalareas caused by slices through tubular burrows.Commonly observed trace fossils (Seilacher,2007) also differ from the chaotic patches andswirls (Fig. 13C).



Bed architecture and downflow facies tracts

After describing the lithofacies seen at individualoutcrops, information on larger-scale bed geome-try from multiple outcrops is now outlined. The

external shape and internal architecture of thefirst ten thick beds below the Contessa marker bedare illustrated in Figs 3 and 8. These diagramsillustrate the variation of thickness, facies char-acteristics and grain-size distribution for each bed

Fig. 8. Detailed bed diagrams showing the lateral variations of lithofacies and grain size along the Ridracoli thrustsheet for: (A) Bed )2; (B) Bed )4; (C) Bed )6; (D) Bed )5; (E) Bed )11; (F) Bed )7; (G) Bed )8; and (H) Bed )10.



along the Ridracoli thrust sheet (Fig. 1). Thediagrams show a downflow distance of 60 km.The beds have a rather complex internalarchitecture with a wide variety of lithofaciesthat change both in downflow and across-flowdirections.

Key information from beds containing Cs7sandstone in the interval above the ContessaBedRe-analysis of four beds in the stratigraphicinterval above the Contessa Bed revealed inter-vals of clean sandstone with a patchy or swirlytexture (Cs7; Fig. 15). These beds provide impor-tant information on the origin of Cs7 intervals, forinstance through the large exotic clasts that theseCs7 intervals sometimes contain (Fig. 15), orlateral transitions into clean turbidite sandstone.These Cs7 intervals were previously incorrectlythought to have transitional mud content betweenturbidite sandstone and mud-rich (Ms) debritesandstone (Amy & Talling, 2006, CMS2 facies ofuncertain origin). Further thin-section analysis aspart of the present study shows that their matrixcomprises clean sandstone.

Bed 2Æ5 in the interval above the Contessa Bedis also described, which contains a clast-richmud-rich debrite sandstone interval (Ms2; Fig.15D). Bed 2Æ5 is included because it has a similararchitecture of mud-rich debrite sandstone to thatof Cs7 sandstone in some beds. It is laterdiscussed whether this similarity suggests thatCs7 sandstone and mud-rich debrite sandstone

originated through similar processes. The charac-ter of these four beds (Fig. 15) is now outlinedbriefly, based on further field observations, grain-size analyses and more detailed relogging at ascale of 1 : 5 rather than 1 : 10.

Bed 0. This bed is the first thick bed above theContessa mudstone (Figs 12 and 15A). A sharpgrain-size break always separates the swirly Cs7sandstone from overlying ripple cross-laminatedsandstone. The basal part of the ripple cross-laminated sandstone has locally founded intothe underlying Cs7 sandstone, whilst the upperpart of the ripple cross-laminated interval isalways flat lying. The swirl and patchy clean-sandstone interval is underlain by a thinnerinterval of massive clean sandstone that lacks aswirly or patchy texture. This basal intervalcontains dispersed coarser granules and iscoarse-tail normally graded, and weathers out(Fig. 15A).

Bed 5Æ1. This bed contains swirly and patchyclean-sandstone intervals on both the Ridracoliand Isola thrust sheets (Fig. 15B and C). The Cs7sandstone interval is always separated fromoverlying ripple cross-laminated sandstone orturbidite mudstone by a sharp grain-size break.It is commonly underlain by a thinner interval ofcoarser sandstone lacking a swirly texture, as inBed 0. The Cs7 interval pinches out betweenadjacent outcrops that are 1 km (Fig. 15C) or3Æ6 km (Fig. 15B) apart. Close to the pinch out,

Fig. 9. Downflow transect through Bed )2 along the Ridracoli thrust sheet (Fig. 1A).



7

8

12

1451,49

15 548

551316

17

28

56Bidente v.

.

30

29,50,52-54

A

B

C

D

A’

B’

12-Bocconi1

7-Lutirano

8-P.dornata

48-Val d’Alberto 30-Isola 255-S.Agata

14-M.roncole

5-M.Arsiccio

A A’0

0,5

1

1,5

2

Thic

knes

s (m

)

53-Cabelli 3

50-La vild54-Cabelli 4

52-Cabelli230-Isola 2

28-Badia 1

D A’0

0,5

1

1,5

2

Thic

knes

s (m

)

15-Giumella1

51-Giumella3

49-Giumella2

14-Monte Roncole13-Premilcuore

0

0,5

1

1,5

2

Thic

knes

s (m

)

B C

56-Betana

16-Fiumicello

28-Badia 1

116,121121

17-Ca’ di Massimo

0

0,5

1

1,5

2

Thic

knes

s (m

)

13-Premilcuore

B B’

Ridracoli Element

Isola Element

Pianetto

Element

N

2 km

3,0 km 6,1 km 2,9 km 2,6 km 0,8 km 3,7 km 5,6 km

0,9 km 0,8 km 5,6 km 4,3 km

2,2 km 1,8 km 0,6 km 1,8 km

0,9km 0,8 km 0,3 km 0,3 km 3,7 km

Bed 5 - Isola Element

29-Cabelli1

0,3km

Flow Direction

Cs7 sandstone

Cs5,6 massive sandstone

Cs1 ripple cross-laminatedMs2 muddy sandstone

Turbidite mudstone

Contorted sand clast

Mud clast

Planar laminated

Cross-laminated

Swirly patchy texture

Grain-size break

Figure 10E,F

A

B

C

D



the base of Bed 5Æ1 can comprise ripple cross-laminated sandstone deposited by dilute turbi-dity currents.

Bed 2Æ5. This bed comprises an interval of clast-rich muddy sandstone with chaotically distrib-uted clasts (Sub-facies Ms1) along the Isola thrustsheet (Fig. 15D). The mud-rich sandstone intervalis ungraded, except for a few centimetres at itstop. It is separated from overlying ripple cross-laminated sandstone and turbidite mudstone by asharp grain-size break. There is a sharp or gradualtransition over a few centimetres into a basalinterval of massive clean sandstone, which is alsocoarser grained than the overlying mud-richsandstone. The mud-rich and clast-rich sand-stone interval pinches out abruptly (Fig. 15D).Close to the location of this pinch out, the base ofthe bed comprises much finer-grained ripplecross-laminated sandstone, which is separatedfrom the overlying massive clean sand by a grain-size break.

Bed 5. This bed contains an interval of Cs7sandstone in part of the Isola thrust sheet(Fig. 10), which is equivalent to laminated ormassive turbidite sandstone in a crossflowdirection. The Cs7 interval continues to theend of available outcrop on the Isola thrustsheet, such that it is not known whether itpinches out abruptly in a downflow direction.The Cs7 interval contains boulder-sized (up to1Æ2 m) clasts that sometimes comprise three thinbeds (Fig. 10E). This sequence of thin beds isnot seen beneath Bed 5 at any of the 109 loggedsections (Talling et al., 2007c), and cannot havefoundered from the ripple cross-laminated sand-stone interval that caps the bed at this location.It also differs from the single interval of ripplecross-laminated sandstone that is occasionallyseen at the very base of the bed (Fig. 10E). Someboulder-sized clasts therefore originated outsidethe study area, and travelled at least 7 km. Theclasts in this Cs7 interval often comprise ripplecross-laminated sandstone that is deformed andwould be relatively easy to disaggregate; thissuggests that sandstone clasts were not vigor-ously tumbled during transport. Bed 5 alsocontains mud-rich debrite sandstone with

E

F

Fig. 10. (Continued).

Fig. 10. (A) to (D) Lateral changes between turbidite clean sandstone (Cs1 to Cs6) and debrite clean sandstone (Cs7)in Bed 5 on the Isola thrust sheet. (E) Outcrop photograph of Bed 5 at Section 29 (Cabelli 1). It comprises ungradedbasal clean sandstone (Cs6), overlain by Cs7 clean sandstone with 11 to 14Æ5% matrix mud content in SEM images(Fig. 7). The Cs7 sandstone interval contains small mudstone clasts and large deformed sandstone clasts (arrowed). Aca 5 cm thick interval of ripple cross-laminated turbidite sandstone (Cs1) overlies the Cs7 sandstone. The top of thebed comprises turbidite mudstone. (F) Outcrop photograph of clast within Bed 5 at Cabelli 1 that comprisesthree individual thin beds. This sequence of thin beds differs from that seen below Bed 5 in any of the loggedsections (Fig. 16). The clast cannot have been formed by foundering of the overlying ripple cross-laminated (Cs1)interval, or erosion of the single ripple cross-laminated sandstone interval seen locally at the base of the bed(Figs 10A to D and 16).



(Ms1) or without (Ms2) clasts in other parts ofthe outcrop area, as well as a range of turbiditesandstone facies (Fig. 16; Amy & Talling, 2006;Talling et al., 2012a).

Downflow facies tractsThe various bed geometries are simplified andsynthesized into downflow facies tracts (Fig. 17).Individual beds can contain more than one faciestract, along adjacent downflow oriented transects,as found by Amy & Talling (2006).

Facies tract 1: Beds with Facies tract 1 containonly clean sandstone and turbidite mudstone indownflow transects (Fig. 17A and B). Massiveclean-sandstone intervals are typically normallygraded (Cs5) but can sometimes be visuallyungraded (Cs6). Lateral transitions from massive

to planar laminated intervals are common. Thesebeds lack massive sandstone intervals with apatchy or swirly texture (Cs7). The clean-sand-stone interval initially thickens to achieve a broadthickness maximum, before thinning distally(Fig. 18). Rates of clean sandstone thinning areup to ca 5 cm km)1, but Facies tract 1 lacks theabrupt pinch out of massive sand intervals seenin Facies tracts 2b and 3 (Figs 17 and 18).

Facies tract 1a is characterized by a gradualtransition from massive graded or ungraded (Cs5or Cs6) clean sandstone into overlying laminatedsandstone intervals (Fig. 17A). Facies tract 1blocally contains an internal grain-size break thatseparates massive sandstone from overlyingcross-laminated sandstone (Figs 9 and 17B). Mas-sive sandstone below the grain-size break iscommonly ungraded (Cs6), but can be inversely

A

B

Fig. 11. Outcrop photograph illus-trating the characteristic swirly fab-ric of sub-facies Cs7 (Table 1). (A)Bed 5.1 at section 13 (Premilcuore)that comprises clean sandstone withpatches of coarser (labelled ‘c’) andfiner (labelled ‘f’) grains. (B) Patchyareas of different grain size seen in afreshly cut surface of sub-facies Cs7from Bed 5.1 in section 13.



A

C

E F

G

B

D

Fig. 12. Outcrop photographs of Bed 0, which is located immediately above the Contessa Bed (Amy & Talling, 2006).Lateral changes in Bed 0 are summarized in Fig. 15A. (A) The most proximal outcrop of Bed 0 at Section 1 (Coniale)where the bed comprises normally graded clean turbidite sandstone. (B) Bed 0 at Section 85 (Cavalmagra) where itcomprises clean sandstone with a swirly texture (Cs7) underlain by more resistant basal clean sandstone with widelydispersed larger sand grains ‘B’ and overlain by ripple cross-laminated finer sandstone ‘UR’. The ripple cross-laminated sandstone has partly foundered into the underlying Cs7 sandstone. Bed 0 is underlain by hemipelagicmudstone ‘HP’ and then by the turbidite mud cap of the Contessa Bed ‘TM’. (C) to (F) Bed 0 at Section 83 (Marradi I).Bed 0 comprises a thick central interval of swirly weathering clean sandstone (Cs7) that is underlain by moreresistant basal clean sand with widely dispersed larger sand grains ‘B’. Contorted clasts of finer sand ‘C’ appear tohave foundered into the swirly weathering Cs7 sandstone. ‘SW’ denotes intervals of clean sandstone that have aswirly or patchy fabric (lithofacies Cs7). (G) Bed 0 at Section 68 (Poggio Pan Della) located downflow from the abruptpinch out of the clean Cs7 sandstone. Bed 0 is much thinner and comprises only turbidite mud and laminated silt.‘CTM’ is the turbidite mudstone interval of the Contessa megaturbidite.



graded (Cs5) in Bed )2 (Fig. 9; Section 77) ornormally graded (Fig. 5C; Beds 2 and 7Æ9 abovethe Contessa interval).

Facies tract 2: This facies tract comprises bedswith intervals of mud-rich sandstone (lithofaciesMs; Talling et al., 2012a). Muddy sandstoneintervals can be sandwiched between clean-sand-stone layers (Fig. 17C and D). Facies tract 2bcontains mud-rich sandstone with millimetre-scale clasts or without clasts (Ms2), underlainby gradually tapering clean-sandstone intervalsthat resemble Facies tract 1a (Talling et al.,2012a). The Ms2 sandstone tapers gradually,and there is a downflow transition into finer-grained normally graded silty mud. Facies tract2b contains mud-rich sandstone intervals withchaotically distributed larger clasts (Ms1) thatpinches out abruptly (Figs 15D and 16D; Tallinget al., 2012a). This Ms1 interval is underlain by athin layer of normally graded massive cleansandstone that pinches out in a similar place tothe overlying Ms2 sandstone (Fig. 15D; Amy &Talling, 2006; Sumner et al., 2009; Talling et al.,2012a). A thin ripple cross-laminated intervalappears at the very base of the bed close to theoverlying massive sandstone pinch out (Fig. 15D).The bed comprises only a thin bed of low-densityturbidite and turbidite mud in its distal part.

Facies tract 3: This facies tract (Fig. 17E to G)comprises beds containing intervals of cleansandstone with a swirly or patchy texture (Cs7)which are always separated from overlying thinlaminated sandstone intervals by a grain-sizebreak. The Cs7 sandstone can locally containchaotically distributed clasts, and pinches outmore abruptly than massive clean-sandstoneintervals seen in Facies tracts 1 and 2a (Fig. 18).Pinch out from Cs7 sandstone that is 50 to 160 cmthick occurs between adjacent outcrops that canbe as close as 1 km (Fig. 15C). These beds com-prise only a thin layer of laminated turbidite sandin distal sections beyond the pinch out of the Cs7interval. The turbidite mudstone cap in Faciestract 3 is relatively thin, and these beds haverelatively high sand to mud ratios (Talling et al.,2007c, fig. 4).

Facies tracts 3a, 3b and 3c are distinguished bythe extent of swirly and patchy Cs7 sandstone.The Cs7 sandstone is continuous from pinch outto the outcrop closest to source in Facies tract 3a,which includes Beds )3, )4, )8 and )9 (Figs 3, 8and 17E). Facies tract 3b is characterized by(sometimes repeated) transitions from normally

graded massive and laminated sandstone (Cs5and Cs4) into swirly and patchy Cs7 sandstone, asshown by Beds 0 and 5.1 (Figs 15B, 15C and 17F).Planar laminations have a wavy and discontinu-ous character near to such transitions, and may beconvolute (Figs 10 and 15C). In locations whereswirly and patchy clean Cs7 sandstone is absent,the upper part of the bed can contain a mud-richclast-poor sandstone interval, as seen for Bed 5Æ1in the Isola transect (Fig. 15C). The Cs7 sandstoneintervals in Facies tract 3b are underlain by a thininterval of structureless clean sandstone withcoarse-tail normal grading, which appears to beassociated with the overlying Cs7 sandstonebecause both pinch out in a similar location(Fig. 15A to C). Close to this pinch out, a thininterval of ripple cross-laminated sandstone canappear at the base of the bed, in a similar fashionto Facies tract 2b.

Facies tract 3c displays downflow transitionsfrom massive and laminated clean sandstone intoswirly or patchy clean Cs7 sandstone, as was thecase for Facies tract 3b (Fig. 17G). Beds )6 and)14 are examples of Facies tract 3c, and they lackclear evidence that Cs7 sandstone pinches outabruptly beyond its most downflow location(Section 42; Fig. 3).

DISCUSSION

Depositional processes for laminated clean-sand-stone intervals are first outlined briefly, becauseplanar laminated intervals are sometimes later-ally equivalent to massive sandstone intervals.The significance of these lateral transitions isdiscussed subsequently.

Depositional processes for laminatedsandstones

Ripple-scale and dune-scale cross-stratification(Cs1) provide unambiguous evidence for deposi-tion from low-density turbidity currents (Harms &Fahnestock, 1965; Simons et al., 1965; Allen,1982; Southard, 1991; Baas, 1994; Talling et al.,2012b). Fine scale planar lamination (Cs3) can beformed incrementally by migration of low ampli-tude bedwaves beneath dilute flows (Best &Bridge, 1992), or repeated collapse of tractioncarpets beneath high-density turbidity currents(Kuenen, 1966a; Leclair & Arnott, 2005; Sumneret al., 2008). Stepped planar laminations (Cs4) aremost likely to be produced incrementally bytraction carpets formed in coarser grain sizes



beneath high-density turbidity currents (Hiscott &Middleton, 1979, 1980; Lowe, 1982), although inthe Marnoso-arenacea Formation they lack in-verse grading.

Previous models for deposition of cleanmassive sandstone

A number of different models have been pro-posed to explain how intervals of massive sand-stone can be deposited. Here, the models aresummarized as a basis for inferring the origin ofmassive Ms, Cs5, Cs6 and Cs7 sandstone intervals(Fig. 19).

Direct suspension fallout beneath high-densityturbidity currentThe experiments of Sumner et al. (2008) andLeclair & Arnott (2003) showed how very rapidsediment fallout can produce direct vertical set-tling of sediment onto the static bed, without anintervening stage of lateral motion in tractioncarpets, as suggested by Kuenen (1966a) andLowe (1982) (Fig. 19A).

Traction carpet deposition beneathhigh-density turbidity currentSumner et al. (2008) illustrated how massiveclean-sand intervals can be deposited by repeatedcollapse of thin (<5 mm) high-concentrationnear-bed laminar layers of sheared sediment(traction carpets), that freeze from the baseupwards (Fig. 19B). As rates of sediment falloutdecrease, allowing greater lateral motion andsorting of sediment by traction, there was agradual change into deposition of clean sandwith progressively better defined planar lamina-tion equivalent to Cs4. The annular flume exper-iments of Kuenen (1966a) showed that massiveclean sandstone can form via distinct sharp-topped traction carpets, and the experiments ofBannerjee (1977), Arnott & Hand (1989) andLeclair & Arnott (2003) also showed how suffi-ciently rapid sediment fallout rates can producemassive deposits.

Sustained liquefied zone beneath high-densityturbidity currentLowe (1982) proposed that rapid sediment falloutcould form a near-bed layer of liquefied sediment,which resulted in deposition of massive sandsthrough sedimentation from the base of the lique-fied layer. Resulting massive sand deposits (the S3interval of Lowe) were inferred to commonly

Dish and pillar structures

Convolute laminations

Patchy (Cs7) texture

A

B

10 c

m

Mixed slurried faciesof Lowe & Guy (2000),

Mud-rich sandstone

D

C

Fig. 13. Illustration of the differences between: (A)dish and pillar structures (e.g. Lowe & LoPiccolo,1974); (B) convolute laminations (as in Fig. 4D); (C) theswirly or patchy texture in Cs7 lithofacies (this study);and (D) mixed slurried facies of Lowe & Guy (2000)seen in subsurface cores from the Britannia Group inthe North Sea (from original photographs in figs 11and 12).



contain dewatering structures, such as dishes andpillars (Fig. 13A; Lowe & LoPiccolo, 1974).

Kneller & Branney (1995) and Kneller (1995)further developed this type of depositional modelfor massive sandstone (Fig. 19C). These authorsinferred that rapid and prolonged sediment fall-out generated a sustained liquefied zone near the

bed that was non-turbulent and laterally sheared,and within which elevated sediment concentra-tion resulted in hindered settling. This sustainedliquefied zone differs from traction carpets in thatit lacks sharp upper or lower boundaries. There isa gradual decrease in sediment concentrationsfrom the static bed to the sustained liquefied

Fig. 14. Correlation panel illustrating the lateral changes of the four beds immediately below the Contessa markerbed in the most proximal outcrops. (A) Correlation panel between the Coniale section and the Coniale River section.(B) Photograph of the top of Bed )4, Bed )3 and the base of Bed )2 in the Coniale River section. (C) Map of theConiale area in the Santerno Valley. (D) Photograph of the base of Bed )4 at the Coniale River section and linedrawing of the rafted sandstone clasts.



Fig. 15. Downflow transects through Beds 0, 2Æ5 and 5Æ1 above the Contessa Bed (modified from Talling et al.,2012b). (A) Bed 0 on the Ridracoli thrust sheet. (B) Bed 5Æ1 on the Ridracoli thrust sheet. (C) Bed 5Æ1 on the Isolathrust sheet. (D) Bed 2Æ5 on the Isola thrust sheet.



zone, and from the sustained liquefied zone intooverlying turbulent flow. Sedimentation from thebase of the sustained liquefied zone results inprogressive bed aggradation, such that the upperboundary of the static bed gradually rises, leadingto formation of a massive sand deposit that isprone to dewatering (Fig. 18C). This type ofsustained liquefied zone has yet to be reproducedexperimentally, although this may be because it isproblematic to generate prolonged (rather thanshort duration) and rapid sediment fallout in thelaboratory.

Kneller & Branney (1995) stated that the thick-ness of the deposit typically bears no resemblanceto the thickness of the flow, implying that thesustained liquefied zone is fed primarily by falloutfrom a significantly thicker overlying energeticturbulent suspension. However, Kneller & Bran-ney (1995) acknowledged that it may be difficult todistinguish between deposits of a sustained lique-fied zone that is fed by sediment settling from anoverlying energetic turbulent suspension, or fedlaterally by liquefied flow. The former situationwould constitute high-density turbidity current

Fig. 16. Downflow changes in Bed 5along different thrust sheets, whoselocation is shown in Fig. 1A. (A)Pianetto, Civorno and Borgo Pacethrust sheets. (B) Isola thrust sheet(Fig. 10). (C) Ridracoli and Pietra-lunga thrust sheet.



according to the terminology herein, as the basallayer is driven from above. The latter situationwould comprise a liquefied debris flow accordingto these definitions, as the liquefied layer is drivenby its own momentum. Kneller & Branney (1995)mentioned that this difference in how the sus-tained liquefied layer is fed could have a profoundinfluence on the location and external shape ofmassive sand deposits, and their lateral relationswith other types of sandstone.

Massive sand deposition by debris flowsA consistent feature of debris flows is that theirdeposits pinch out relatively abruptly at their

margins (Amy et al., 2005), as observed inexperiments and in the field for subaerial debrisflows (e.g. Johnson, 1970; Iverson, 1997; Major,1997; Major & Iverson, 1999; Iverson & Vallance,2001; Revellino et al., 2004; Iverson et al., 2010)and submarine debris flows (e.g. Aksu & Hiscott,1989; Twichell et al., 1992; Gee et al., 1999;Mohrig et al., 1998; Laberg & Vorren, 2000;Bowles et al., 2003; Lastras et al., 2005; Marret al., 2001; Elverhøi et al., 2007; Tripsanaset al., 2008). Abrupt marginal pinch out canoccur in debris flows with little or no cohesivemud, due to frictional interlocking of grains atthe flow boundaries (Iverson & Vallance, 2001,

Fig. 17. Generalized facies tractsobserved for correlated beds withinthe Marnoso-arenacea Formation.



fig. 2; Iverson et al., 2010). It can also occur inmud-rich debris flows with greater cohesivestrength, if cohesive strength prevents continuedmotion in thinner marginal areas of the flow(Coussot, 1997; Laberg & Vorren, 2000; Iversonet al., 2010). Dissipation of excess pore fluidpressures, sometimes due to coarser-grainedmargins, can also help to cause abrupt freezingof flow margins for both non-cohesive andcohesive flows (Iverson et al., 2010). The steep-ness of the marginal pinch out will tend to varywith the cohesive or frictional strength devel-

oped at flow margins, and the gradient of the seafloor.

Non-cohesive liquefied debris flow (Fig. 19D):This type of debris flow lacks any cohesive finemud. It corresponds to the liquefied flowsdescribed by Lowe (1976). Pore fluid lacks anycohesive strength, such that sand particles arecontinuously settling from the base of the flow.Grain interactions cause such settling to bestrongly hindered (Richardson & Zaki, 1954).Lowe (1976) proposed that flow would stop once

Fig. 18. (A) Shape of sandstone intervals in beds displaying different facies tracts (Fig. 17). (B) Shape of sandstoneintervals with abrupt pinch out of Cs7 clean-sandstone debrite. These include Beds 0, 5Æ1 and )8 with Facies tract 3aand 3b, Bed )6 with Facies tract 3c, and Bed )2 with Facies tract 2 (Fig. 9). Rates of pinch out in Beds 0, 5Æ1 and )8are minimum values because pinch out occurs within covered areas between adjacent outcrops.



Fig. 19. Generalized summary ofdifferent process models for depo-sition of massive sandstone. Thethick dotted line denotes the profileof sediment concentration.



the rising boundary between the static bed andoverlying liquefied suspension reached the top ofthe flow (Fig. 19D). Such incremental depositionfrom the base of the non-cohesive debris flowmight be expected to produce massive sandlayers with a gradually tapering shape. However,such a conclusion neglects dissipation of excesspore pressure from the frontal and lateral marginsof the flow, which may cause the margins of theflow to freeze en masse whilst the centre of themixture continued to flow. This would mostlikely result in a clean-sand deposit with abruptmargins. Iverson & Vallance, 2001 (fig. 2) provideexamples of flows lacking any mud whosedeposit has abrupt margins due to en massefreezing.

Lowe (1976) suggested that the deposits of non-cohesive liquefied flows may be either coarse tailgraded or distribution graded, based on theexperiments of Middleton (1967) with variableconcentrations of rather low-density non-cohe-sive plastic beads. Middleton (1967) found thatsediment volume concentrations of <30% formeddistribution grading, whilst settling of more con-centrated non-cohesive dispersions generatedonly coarse tail grading. The static settling exper-iments of Amy et al. (2006) suggest there isefficient size segregation at volume concentra-tions of less than ca 20%, and that no segregationoccurs at volume concentrations exceeding ca50%. Hindered settling at intermediate volumeconcentrations will tend to generate an ungradedor very poorly graded basal interval, with gradingonly in the uppermost part of the deposit (Dorrellet al., 2011).

Cohesive debris flow (Fig. 19F): Cohesive debrisflows are defined here as having muddy pore fluidwith sufficient yield strength to support sandgrains, which remain locked within the matrix(Fig. 19E). Deposition occurs through en masseconsolidation of the mixture once flow hasstopped, generating ungraded deposits. Basalexcess pore pressures will take several orders ofmagnitude longer to dissipate than for non-cohe-sive debris flows. The yield strength of muddyfluid necessary to support sand can be very low(ca 1 Pa), corresponding to ca 14% volume ofkaolin (Amy et al., 2006; Sumner et al., 2009) orsmaller concentrations of other clay mineralswith stronger colloidal behaviour (Coussot,1997) found in the Marnoso-arenacea beds (suchas illite and smectite from XRD analyses). Debrisflows with such low yield strength can still flowas thin (1 to 2 m) flows on gradients of just 0Æ05�

(Talling et al., 2010), and do not need to hydro-plane to run out for very long distances. Suchweak cohesive flows will tend to be turbulent onsteeper slopes, and may only become laminardebris flows on low gradients (Talling et al.,2007a, 2010). Large (>50 cm) outsize mud clastscan be supported if they are buoyant and havelower densities than the surrounding flow. Thecohesive yield strength of such flows will ensurethat their deposits have abrupt margins, but theangle of these margins may be reduced by therelatively low strength of the flows (c.f. subaerialdebris flows of the type described by Iversonet al., 2010).

Poorly cohesive debris flow (Fig. 19E): Poorlycohesive debris flows contain mud, but themuddy pore fluid has insufficient strength tokeep all sand suspended indefinitely. Segregationof sand is efficient for very low mud content (forexample, <6% volume kaolin; Amy et al., 2006;Sumner et al., 2009), such that basal sedimenta-tion processes resemble that of non-cohesivedebris flow. The incrementally deposited sand islikely to be graded, but grading in the initial partof the deposit may be suppressed if sand-volumeconcentrations greatly exceed ca 20% (Amy et al.,2006). This is the type of liquefied flow that wasstudied experimentally by Breien et al. (2010)with 6Æ6% volume concentration kaolin, fromwhich sand settled out incrementally to formmassive sand.

As cohesive mud content approaches thatneeded to fully support sand, the settling beha-viour of the sand changes and becomes morecomplex (Amy et al., 2006; Sumner et al., 2009).In the static settling experiments of Amy et al.(2006, type 3 with ca 9 to 13% kaolin), grainsmoved downwards or upwards slowly(<1 mm s)1), leading to the development ofregions with variable density that displayedlarger scale plume-like movement. Elutriationpipes developed at the boundary between thestatic basal deposit and overlying convectinggrains, and this depositional boundary graduallymoved upwards at ca 0Æ16 mm s)1. Sumner et al.(2009) described how sand grains can settle fromthe later stages of a moving flow or, if deceler-ation is rapid settling, it occurs even after theflow has stopped in sand–mud mixtures with ca11 to 14% volume kaolin. Sand grains that startto settle break bonds between surrounding cohe-sive mud particles to produce weak zones,leading to the formation of subvertical pipes inwhich sand settles downwards and water is



expelled upwards. This process forms localizedblebs of sand at the base of the pipe. Theresulting deposits comprise a basal clean-sandlayer that can be either graded or ungraded, andan overlying plug that has a heterogeneoustexture and consolidated slowly. Kuenen (1965,figs 1 and 2) reported on similar convection ofmuddy sand suspensions and formation of‘trickles’ of sand.

In a laterally moving flow, such sedimentationbehaviour would incrementally deposit massivesand in locations closer to source, which mayhave sharply defined margins where excess porepressures have dissipated (as was the case fornon-cohesive flows). The convecting plug couldeventually come to an abrupt halt (freeze) furtherdown-slope on lower gradients to produce adeposit with heterogeneous internal texture thathas an abrupt margin.

Processes of formation for massive cleansandstone in the correlated beds

Depositional processes are inferred initially fromfeatures observed at the scale of a single outcrop,and then from sandstone shape.

Graded (Cs5) and ungraded (Cs6) massiveclean sandstoneBoth graded and ungraded massive sandstonecould be deposited by high-density turbiditycurrents through direct vertical settling (Fig.19A), repeated collapse of short-lived tractioncarpets (Fig. 19B), or sedimentation from near-bed sustained liquefied zones (Fig. 19C). Mostgradually aggrading high-density turbidite sand-stones are likely to be graded but steady flowcan generate ungraded or very poorly gradeddeposits. Incremental deposition in this wayexplains the commonly observed lateral andvertical transitions into laminated clean sand-stone (Cs4) that provides a clear record ofturbidity current deposition. These depositionalprocesses are consistent with the occasionaloccurrence of faint planar laminations withinCs6 intervals, which could form via tractionalreworking during sporadic periods of lowersediment fallout rates. However, the lack ofvisible grading would only occur if the turbiditycurrent was almost perfectly steady, as the flowscontained a wide range of grain sizes. A steadyturbidity current is unlikely to produce theupper grain-size break that separates Cs6 depos-its from overlying finer-grained sandstone, butthis grain-size break could form during a period

of sediment bypass as the turbidity currentstarts to wane (Kneller & McCaffrey, 2003;Sumner et al., 2008).

Graded or ungraded massive clean sandstonecould also be deposited incrementally via sedi-mentation from a non-cohesive or poorly cohe-sive debris flow. Bed correlations show that theflows contained mud, and were not entirely non-cohesive. The sharp grain-size break that over-lies ungraded or graded massive Cs5 and Cs6sandstone intervals in Facies tract 1b (Fig. 16B)could result from a sharp boundary separatingpoorly cohesive debris flow from trailing turbi-dity current, such that this is a lateral (front toback) boundary in the flow. However, the grad-ually tapering shape of Cs5 and Cs6 intervalsindicates that they were most likely depositedby high-density turbidity current, rather thandebris flow.

Massive clean sandstone with ‘patchy’ texture(Cs7) in Facies tract 3Chaotic and swirly areas of coarser-grained andbetter-sorted sandstone are common in Cs7 sand-stones (Figs 6B, 11 and 12). It was previouslyconcluded that the ‘patchy’ texture most likelyresults from pervasive liquefaction. Two possiblemodels are used to explain this liquefaction.

In situ liquefaction of turbidite sandstone:Firstly, it can be proposed that the patchy Cs7texture records pervasive in situ liquefaction ofmassive sandstone previously deposited by high-density turbidity current. This is unlikely becauseit is not clear why such post-depositional lique-faction would cause partial foundering of theoverlying ripple cross-laminated interval (as seenin Bed 0; Figs 12 and 15A), and consistently failto liquefy, or deform at all, the upper part of theripple cross-laminated interval. Moreover, thismodel for the origin of the Cs7 sandstone isless likely because the Cs7 sandstone abruptlypinches out, whilst turbidity currents that depositsediment incrementally would be expected toproduce a gradually tapering deposit shape.

Local sea floor topography can potentiallyproduce abrupt thinning and fining of turbiditesandstones, and such turbidites could thenundergo liquefaction. However, there is no evi-dence of significant topography near the pinchout of Cs7 intervals. Turbidite beds immediatelyunderlying and overlying these deposits recordno abrupt changes in thickness or grain size(Fig. 3) and there is no evidence of flow reflectionor deflection by local sea floor topography. It is



also unclear how post-depositional in situ lique-faction could produce grain-size breaks above theCs7 or Cs6 intervals (Figs 12 and 15A).

Deposition by poorly cohesive or cohesivedebris flow: In the second model, the patchytexture of Cs7 sandstone is inferred to result fromdeposition by a liquefied debris flow, which iseither cohesive such that its muddy pore fluidstrength can fully suspend the sand grains, orpoorly cohesive but from which only very slow orpartial settling of sand occurs.

Slow convection and elutriation of fine mate-rial during this liquefaction process, togetherwith blebs of sand formed by settling downvertical conduits, could form the patchy texture(Kuenen, 1965, fig. 1; Amy et al., 2006; Sumner,2007; Sumner et al., 2009). Such pervasive liq-uefaction is consistent with occasional founder-ing of the basal part of overlying ripple cross-laminated sandstone into the Cs7 sandstone(Fig. 12), and the presence of chaotically distrib-uted clasts in some locations (Figs 8, 10, 14 and16). Experiments suggest that cohesive debrisflows in which pore fluid strength traps sandtend not to develop a patchy texture, but ratherform sub-vertical pipes through which onlywater is slowly expelled (Type III deposit ofSumner et al., 2009). However, it is unknownfrom experiments whether cohesive debris flowswith higher sand to mud ratios than thosestudied by Sumner et al. (2009) might producea patchy texture. Partial segregation of sandneeded to produce patches most likely recordsdeposition from a poorly cohesive debris flow,whose pore fluid had insufficient cohesivestrength to support sand.

The upper part of the ripple cross-laminatedsandstone is flat lying (Figs 12 and 15A), sug-gesting that pore pressure dissipation from thedebris flow was completed during deposition ofthe rippled interval, which was probably depos-ited within minutes to hours of the massivesand (Baas, 1994). This would favour relativelyrapid pore pressure dissipation from a poorlycohesive debris flow rather than slower porefluid pressure dissipation from a strongly cohe-sive debris flow (Talling et al., 2012a, fig. 21).However, disaggregation of sand clasts erodedlocally from the base of the bed, such as thoseseen in the most proximal part of Bed )4(Fig. 14), may generate patches of coarser andbetter-sorted grains. Such a process could causea patchy texture to form in a fully cohesivedebris flow.

The abrupt lateral pinch out of swirly orpatchy Cs7 sandstone intervals (Figs 17 and 18)supports the conclusion that they were depos-ited by debris flow. Thin intervals of massivenormally graded massive sandstone occur belowsome Cs7 intervals (Figs 15A to C and 17F).These basal graded sands most likely formed bylate-stage settling from a poorly cohesive debrisflow, as they pinch out in the same location asthe overlying debrite sandstone. Alternatively,the basal clast-deficient massive sand layer wasdeposited by a forerunning turbidity current, andthe Cs7 sandstone was deposited by a cohesivedebris flow from which sand settled en masse. Itis equivocal as to whether the Cs7 sandstoneswere deposited by a poorly cohesive or a cohe-sive debris flow, but Cs7 sandstone is morelikely to have been deposited by poorly cohesivedebris flow from which sand could partiallysettle.

Origin of clean sand liquefied debris flowsresponsible for Cs7 sandstone

Two contrasting hypotheses can be put forwardfor the origin of the debris flows that formed Cs7clean sandstone intervals.

Long run out clean-sand debris flows thatoriginated outside basin plainThe Cs7 intervals that extend continuously tothe most proximal available outcrops (Faciestract 3a) could have been deposited by debrisflows that originated outside the outcrop area,and which subsequently ran out for 15 to 30 kmacross the basin plain (Fig. 17D). The Cs7 inter-vals can contain chaotically distributed mud-stone clasts, and sandstone clasts up to 120 cmin length (Fig. 14). The sandstone clasts in Bed)4 and Bed 5 do not come from massive basal oroverlying thinner laminated parts of the samebed. Unless clasts were low density and there-fore buoyant, it is likely that they were raftedinto place by debris flow, rather than transportedby an initial turbidity current. The turbulentnature of an initial turbidity current would alsotend to disaggregate soft sandstone clasts, againsuggesting that such clasts were rafted into placeby debris flow.

Clean-sand debris flows originated throughflow transformation from turbidity currentSome beds with Cs7 intervals have lateral andup-flow transitions from Cs7 sandstone into cleansandstone deposited by high-density turbidity



current (Figs 10 and 15). Beds 0, 5 and 5Æ1 abovethe Contessa Bed display lateral transitionsbetween Cs7 sandstone, wavy or planar-parallellaminated (Cs3) sandstone, and massive gradedsandstone (Cs5) (Figs 15A to C and 17F). Themore proximal basal part of these beds cancomprise either Cs3 or Cs5 sandstone that tran-sitions downflow into Cs7 sandstone, without asignificant change in the overall sandstone layerthickness (Fig. 15A to C). The downflow transi-tions from laminated high-density turbidite intoCs7 sandstone suggests that at least some of thepoorly cohesive debris flows formed via flowtransformation from an initial high-density tur-bidity current.

Significance of lateral transitions into cleansandstone deposited by turbidity current

Lateral transitions between liquefied debris flow(Cs7) and laminated (Cs3) or massive (Cs5)turbidite sandstone occur without majorchanges in the overall sandstone thickness(Fig. 10). This finding is significant because itsuggests that flow responsible for depositing thedifferent types of sandstone was similar, at leastin terms of the rate and duration of sanddeposition. This may suggest that the planarlaminated (Cs3) intervals were deposited fromrelatively high concentration turbidity current(Talling et al., 2012b), instead of dilute turbiditycurrent as has sometimes been inferred (Lowe,1982).

Are clean (Cs7) and mud-rich (Ms) debritesandstones formed by similar processes?

Beds within the Marnoso-arenacea basin plainoften contain intervals of mud-rich sandstone thatwere deposited by cohesive debris flow (Tallinget al., 2004, 2012a,b; Amy & Talling, 2006).A single bed can contain both clean debritesandstone (Cs7) and mud-rich debrite sandstone(Ms) in different parts of the basin plain (Fig. 15Band C), which is best illustrated by Bed 5(Fig. 16). Both mud-rich and clean-sandstonedebrites occur in a similar vertical positionwithin beds, such that they are underlain bythin massive sandstone (Cs5 or Cs6), and over-lain by ripple cross-laminated sandstone (Cs1;Figs 8, 10 and 15). The geometry of facies tractscontaining clean debrite sandstone (Cs7) canalso be similar, such as Facies tracts 2b and 3a,or Facies tracts 2a and 3c (Fig. 17). A thininterval of ripple cross-laminated sandstone can

appear at the base of the bed close to where bothtypes of debrite pinch out (Figs 15 and 17), andboth types of debrite can be absent in proximalparts of a bed (Fig. 17). These similarities sug-gest that the depositional processes for mud-richsandstone debrite and clean-sandstone debriteare broadly similar.

Thin (<2 m) mud-rich debris flows can travellong distances over very low sea floor (0Æ05�)gradients without hydroplaning, due to their lowyield strength (Talling et al., 2007a, 2010,2012a,b). The cohesive strength of their matrixis sufficient to support sand (Fig. 19F). Somemud-rich debrite sandstones in the Marnoso-arenacea Formation contain clasts that did notoriginate through erosion of the underlying sub-strate within the study area, such as the MS1interval in Bed 2Æ5 (Fig. 15; Talling et al., 2012a).These clast-rich mud-rich cohesive debris flowsoriginated from outside the basin plain, and thenran out for several tens of kilometres across thebasin plain (Talling et al., 2012a). However, mud-rich debrite sandstones that are clast-poor (MS2)may have originated through flow transformationfrom an initial turbidity current within the basinplain (Talling et al., 2012b). Such flow transfor-mation has been observed in laboratory experi-ments (Sumner et al., 2009), and results fromdevelopment of cohesive bonds between mudparticles as the flow decelerates, which suppressturbulence. Clean debrite sandstone tends tooccur in beds with higher fractional sandstonecontent in the above-Contessa interval (Tallinget al., 2007b, fig. 4F). This suggests that formationof clean debrite sandstones or mud-rich debritesandstone depends partly on the sand to mudratio within a flow.

Can liquefied debris flows of clean sand runout for long distances?

An analysis by Lowe (1976) concluded thatliquefied flows of sand would run out for dis-tances of less than ca 1Æ2 km. The factors thatcontrol the run-out distance of a liquefied sedi-ment layer are now discussed for layers thatcontain increasing amounts of cohesive mud, toexplain how liquefied debris flows could poten-tially run out for longer distances.

Run out of non-cohesive debris flow (withoutmud)The analysis by Lowe (1976) was based onsediment mixtures that lacked cohesive mudand comprised very well-sorted sand (Lowe,



1976; data from Wallis, 1969; Andersson, 1961).These liquefied flows corresponded to the non-cohesive debris flows herein (Fig. 19D). Liquefiednon-cohesive suspensions deposit sediment fromthe base up, such that a rising surface forms, thatseparates the static bed from the remainingliquefied sediment (Fig. 19D; Lowe, 1976; Sassa& Sekiguchi, 2010, 2011). The run-out distance isdetermined by the basal sedimentation rate andflow thickness, as excess pore pressure dissipatesrapidly from such mud-free sediment mixtures(Terzaghi et al., 1996).

Lowe (1976) used rates of between 0Æ08 cm s)1

and 1Æ67 cm s)1 for basal sedimentation to cal-culate the time taken for the re-sedimentationsurface to travel from the bottom to the top of theliquefied layer. It was inferred that flow wouldstop once the re-sedimentation reached the topof the liquefied layer (Fig. 19D). This inferenceneglects replenishment by sediment from therear of the flow, or changes in elevation of thetop of the flow. Run-out distance was calculatedas the product of the time taken for the re-sedimentation surface to reach the top of theflow and the lateral velocity of the flow front. Forinstance, a 1 m thick liquefied flow travelling at1 m s)1 would have flow for 60 to 1180 s, andrun out for 60 to 1180 m. Although addition ofsilt to a sand mixture can reduce permeability byorders of magnitude, thereby hindering settling(Bandini & Sathiskumar, 2009), it seems likelythat non-cohesive (mud-free) debris flows willnot run out for long distances, even on steepslopes.

Run out of cohesive debris flow (in which sandis fully supported by cohesive strength)If sand grains do not settle out and are supportedby the cohesive strength of the pore fluid, themixture forms a cohesive debris flow according tothe terminology herein (Fig. 19F). The run-outdistance of a cohesive debris flow does notdepend on rates of basal sedimentation. It mayalso not depend on the rate at which excess porepressure dissipates, as this rate of dissipation canbe very slow. Major (2000) and Iverson et al.(2010) showed how an increase from ca 1% to ca4% volume of cohesive mud in a mixture of sandand gravel can reduce its hydraulic diffusivity byca 10)4 to ca 10)7, thereby increasing the timetaken for excess pore pressure to decay by threeorders of magnitude. Dissipation of excess porepressures in 1 m thick cohesive debris flows canrange from hours to weeks (Talling et al., 2012b,fig. 18). A cohesive debris flow stops when the

cohesive strength exceeds the shearing forceexperienced at the base of the flow. Run-outdistance is determined by yield strength, flowthickness and density, and sea floor gradient.Schwab et al. (1996) and Talling et al. (2010)showed how thin (<2 m) cohesive debris flowscan travel across low (0Æ05�) sea floor gradients,even when hydroplaning (Mohrig et al., 1998)does not occur.

Run-out of poorly cohesive debris flow (inwhich sand is partly supported by cohesivestrength)Sand settles out efficiently from sediment mix-tures with relatively low mud content, such as inthe experiments of Sumner et al. (2009) contain-ing <7% of kaolin mud (Fig. 19D). Although theincreased viscosity of the muddy fluid can reducesettling velocities of sand grains, basal sedimen-tation rates are still relatively high. Breien et al.(2010) measured rates of bed aggradation of0Æ43 cm s)1 for mixtures that contained 6Æ6%volume kaolin mud. Amy et al. (2006) recordbed aggradation rates of ca 0Æ017 cm s)1 formixtures with up to 10% volume kaolin mud.A 1 m thick debris flow travelling at 1 m s)1, witha bed aggradation rate of 0Æ017 cm s)1, would runout for ca 5Æ9 km.

As the threshold mud concentration needed tofully support sand is approached, the settlingbehaviour of sand becomes more complex(Fig. 19E), such as in the experiments of Sumner(2007) and Sumner et al. (2009) that contained7 to 12% kaolin mud. Settling of sand onlyoccurred during the later stages of deceleratingflows in these experiments, or sometimes after theflow had stopped moving. Sand tended to settleslowly through pipes in which mud bonds hadbeen weakened (Fig. 19E; Sumner, 2007; Sumneret al., 2009). If sand settling occurs after a poorlycohesive debris flow has stopped moving, its run-out distance will resemble that of a cohesivedebris flow, as determined by yield strength, flowthickness and density and sea floor gradient. Ifsand settles out whilst the debris flow is moving,its run-out distance may be influenced by this rateof settling. However, sand settling will be hin-dered by the elevated mud content, and the run-out distance may approach that of non-cohesivedebris flows.

Sediment resuspension from the bedThe preceding analysis has inferred that sedimentis not re-entrained from the static bed. This isimportant because a balance between upward



re-entrainment and downward settling of sandcan lead to a situation in which there is no netloss of sand from the flow, which could thenincrease the run-out distance of the debris flowsignificantly (Fig. 19D to F). Motion of a debrisflow will lead to shear at the base of the flow,which may be vigorous enough to move sand. Ifthe debris flow is very weakly turbulent, as can bethe case for subaerial debris flows (Costa &Williams, 1984), this can lead to upward diffu-sion of sand from the bed. Such re-entrainmentcould also help to generate a patchy texture, ifclumps of sand grains are re-entrained togetherand poorly mixed within the flow.

Run out of clean-sand debris flows in theMarnoso-arenacea FormationClean-sand debrites in the Marnoso-arenaceaFormation typically contain 4 to 8% cohesivemud, and sometimes up to 15% cohesive mud,according to data from SEM images (Fig. 7A andB). This volume fraction differs from the volumefraction of cohesive mud in the pore fluid whichdetermines the cohesive strength of the flows(Kuenen, 1966b). However, these flows may havesufficient cohesive strength to support sandgrains (as a cohesive debris flow), or intermediateamounts of mud that strongly hinder settling ofsand (as a poorly cohesive debris flow). In eithercase, their run-out distance may exceed thatcalculated by Lowe (1976) for mud-free liquefiedlayers. If their mud content prevents or stronglyhinders settling of sand, then the run-out distanceof clean-sand debris flows may approach that of

the mud-rich cohesive debris flows that depositedfacies Ms1 and Ms 2.

Clean sandstone attributed previously toliquefied flow and debris flow

Dense liquefied flows of Mutti and colleaguesMutti (1992), Mutti et al. (2009), Mutti et al.(2003) and Tinterri et al. (2003) proposed thatsubmarine flows can comprise an upper turbulentsuspension underlain by a much denser layer offluidized sediment, in which sediment is sup-ported primarily by excess pore pressures(Fig. 20). The basal fluidized layer producesmassive sand layers that can be poorly graded orungraded, particularly poorly sorted (F5 facies ofMutti et al., 2003), and which can contain mud-stone clasts and dewatering structures. Deposi-tion from the basal layer occurs by rapid en masse‘freezing’. The term fluidized is used more com-monly to denote an external source for the porefluid, rather than upward flow generated by self-consolidation (Lowe, 1976). It is therefore prefer-able to use the term ‘liquefied’ rather than‘fluidized’ for this basal flow phase.

Mutti et al. (2003) proposed that beds contain-ing dense liquefied flow deposits (F5 facies)display one of two facies tracts (Fig. 20). Firstly,a downflow transition can occur from the denseliquefied flow into fully turbulent turbidity cur-rent (Fig. 20A). The massive F5 sandstone isoverlain or laterally equivalent to crudely lami-nated clean sandstone formed by turbidity cur-rent (F7 facies). Secondly, if deceleration was

Fig. 20. Submarine density flowtypes and deposits, and evolutionbetween flow types in the model ofMutti (1992) and Mutti et al. (2003).(A) Types of submarine flow andtheir deposits (F1 to F9). F5 sand-stone is deposited by laminar denseliquefied flow. (B) Downflow evolu-tion of an initially bipartite flowcomprising a basal dense liquefiedlayer and an overlying more diluteturbulent layer. Sedimentary logsindicate the resulting downflowchanges in deposit character, afterMutti et al. (2003) fig. 17.



rapid, the F5 sandstone is erosively overlain andlaterally equivalent to coarser-grained intervals ofplanar laminated or dune-scale cross-laminatedsandstone (F6 facies) formed by partial flowbypass of finer grains (Fig. 20B).

The field observations for the present studyindicate that some clean-sandstone intervals aremost likely deposited by dense liquefied flow(‘debris flow’ according to the terminology here-in). However, the distinctive swirly or patchytexture of such Cs7 deposits in the Marnoso-arenacea beds was not described previously byMutti and colleagues, who did not specify thetypes of syn-depositional and post-depositionaldewatering features (Mutti et al., 2003) that weredisplayed by dense liquefied flow deposits (F5facies). This is important because, at the scale ofsingle outcrops, it is only the swirly or patchytexture that distinguishes massive clean sand-stone deposited by dense liquefied flows in theMarnoso-arenacea beds from other types of mas-sive clean sand deposited by turbidity current,which can contain other types of dewateringfeatures, such as dish and pipe structures.

The facies tracts observed here with denseliquefied flow deposits (Cs7) also differ from thefacies tracts of Mutti et al. (2003) in key regards. Adownflow transition occurs from massive or lam-inated clean turbidite sandstone (Cs4 or Cs5) intothe liquefied debris flow deposits (Cs7) withinFacies tract 3 (Fig. 17F and G). The downflowdirection of this transition has the opposite sense(relative to palaeoflow) to that of Mutti andcolleagues (Fig. 20). The correlated beds recorddownflow transitions from more dilute to denserflow, rather than from denser into more dilute flow.

Sandy debris flows of ShanmugamShanmugam & Moiola (1995) and Shanmugam(1997, 2000, 2002) controversially proposed thatmassive layers of clean sand could be deposited enmasse by abrupt freezing of debris flows charac-terized by: (i) significant yield strength (non-New-tonian rheology); (ii) laminar flow conditions; and(iii) sediment support through matrix strength,buoyancy and grain to grain interactions. Theimportance of excess pore fluid pressure was notemphasized in these works. However, recentstudies of subaerial debris flows have shown howdebris flow behaviour is strongly dependent onvariations in excess pore fluid pressure (Iverson,1997; Iverson & Vallance, 2001; Iverson et al.,2010). The definition of debris flow used herediffers from that of Shanmugam, in that excess porepressure can be an additional important sediment

support mechanism and, unlike Shanmugam, theterm debris flow is not used for high sedimentconcentration near-bed layers driven mainly by theoverlying flow.

The criteria to identify debris flow deposits inthe present study also differ from those of Shan-mugam. The eight criteria listed by Shanmugam &Moiola (1995) and Shanmugam (1997) to identifydebris flow deposits are: (i) ungraded sand withevidence of basal shearing; (ii) concentration ofmudstone clasts at the top of the sand interval;(iii) inverse grading of clasts; (iv) dispersed(floating) larger sand grains; (v) planar clast fabric;(vi) presence of shale clasts; (vii) irregular uppersurface and lateral pinch out geometries – sug-gesting en masse freezing; and (viii) relativelyhigh detrital mud matrix sufficient to induce aplastic rheology with finite yield strength. Theclean-sand layers that the present study attributesto debris flow deposition have: (i) a swirly fabricindicative of pervasive liquefaction; (ii) a sharpgrain-size break at their upper boundary; (iii)abrupt lateral pinch outs; and (iv) sometimeschaotically distributed clasts.

This difference in how clean debrite sandstonesare identified is important because this articlediffers from that of Shanmugam in the way inwhich clean-sandstone debrites are identified.Shanmugam & Moiola (1995) proposed that theJackfork Group at the DeGray Spillway in Arkan-sas, and Shanmugam (2002) proposed that theAnnot Sandstone in the Piera Cave road section,contain clean-sand debris flow deposits. Based onthe different criteria herein and first-hand fieldobservations of the DeGray Spillway and PieraCava road sections, this article does not infer thateither of these two sections contains cleansandstones deposited en masse by debris flow,because sandstones in the two sections are oftengraded and lack a patchy or swirly texture.However, the external shape of individual sand-stone intervals is not sufficiently well-con-strained for either the DeGray Spillway or PieraCava road section to independently test thisconclusion based on features seen at the scale ofthese single outcrops.

CONCLUSIONS

The Marnoso-arenacea Formation in the northernItalian Apennines is the only place worldwidewhere individual ancient flow deposits can bemapped out in detail for up to 120 km. Theseunusually extensive bed correlations provide new



insight into how submarine sediment flowsevolve as they spread across low gradient basinplains. The shape of massive sandstone intervalsprovides an independent test of depositionalprocesses inferred initially from their internalcharacteristics at a single outcrop. Three depositgeometries (or facies tracts) occur in downflowtransects through beds in a newly studied strati-graphic interval below the Contessa marker bed(Fig. 17). Individual beds can contain more thanone of these facies tracts. Facies tracts 1 and 2contain graded or ungraded massive clean-sand-stone intervals that taper gradually downflow,and these massive clean sandstones were mostlikely deposited incrementally beneath high-density turbidity currents. Mud-rich massivesandstone deposited by mud-rich cohesive debrisflow occurs in Facies tract 2. Facies tract 3contains clean sandstone with a distinctiveswirly fabric formed by patches of coarser andbetter-sorted grains. This distinctive fabric mostlikely records pervasive liquefaction during thelater stages of flow. This type of clean sandstonepinches out abruptly, and this pinch out geome-try suggests that it was most likely deposited byliquefied debris flow. It is therefore suggested thatliquefied debris flows with elevated pore pres-sures can deposit clean sand over large expanses(up to 30 km) of sea floor. Downflow transitionsfrom turbidite sandstone to debrite sandstone(Figs 15A to C and 17F) suggest that clean-sanddebris flows are most likely to form throughtransformation from an initial turbidity current.However, in at least one case, a clean-sand debrisflow that contained chaotically distributed clasts(Figs 10 and 14) may have run out for longdistances on low gradients across this basin plain.

ACKNOWLEDGEMENTS

This work was part of a UK-TAPS project, fundedby NERC Ocean Margins LINK grant NER/T/S/2000/0106 and NER/M/S/2000/00264, and co-sponsored by ConocoPhillips, BHP Billiton andShell. The authors wish to thank Gianni Zuffaand Franco Ricci Lucchi (Bologna University)and Luca Martelli (Emilia Romagna GeologicalSurvey) for advice initially on fieldwork in north-ern Italy. Bed correlations presented in this studywere made possible by the outstanding mappingof marker beds by the Italian Geological Surveys.The assistance by Lawrence Amy (Saudi-Aramco)and Esther Sumner (University of Leeds) in thefield is much appreciated. Grain-size measure-

ments were made by Graham Blackbourn (Black-bourn Geoconsulting), Harriet Wimhurst-Brookes,Christopher Wilcox and Christopher Nutt whilstat the University of Bristol. We wish to thank JorisEggenhuisen, Jaco Baas and Suzanne Leclair fortheir insightful and constructive comments, andtwo anonymous reviewers.

REFERENCES

Aksu, A.E. and Hiscott, R.N. (1989) Slides and debris flows on

the high-latitude continental slope of Baffin-Bay. Geology,

17, 885–888.

Allen, J.R.L. (Ed.) (1982) Chapter 10. Structures and sequences

related to gravity-current surges. In: Sedimentary Struc-

tures. Their Character and Physical Basis, pp. 395–431.

Elsevier, Amsterdam.

Allen, P. (2007) Earth science: sediment en route to oblivion.

Nature, 450, 490–491.

Amy, L.A. and Talling, P.J. (2006) Anatomy of turbidites and

co-genetic debrites intervals based on long distance (120 x

30 km) bed correlation, Marnoso Arenacea Formation,

Northern Apennines, Italy. Sedimentology, 53, 161–212.

Amy, L.A., Talling, P.J., Peakall, J., Wynn, R.B. and ArzolaThynne, R.G. (2005) Bed geometry used to test recognition

criteria of turbidites and (sandy) debrites. Sed. Geol., 79,163–174.

Amy, L., Talling, P.J., Edmonds, V.O., Sumner, E.J. and

Leseuer, A. (2006) An experimental investigation of sand-

mud suspension settling behaviour: implications for bimo-

dal mud contents of submarine flow deposits. Sedimentol-

ogy, 53, 1411–1434.

Andersson, K.E.B. (1961) Sedimentation and ideal fluidiza-

tion. Sven. Kem. Tidskr., 73, 334–345.

Arnott, R.W.C. and Hand, B.M. (1989) Bedforms, primary

structures and grain fabric in the presence of suspended

sediment rain. J. Sed. Petrol., 59, 1062–1069.

Baas, J.H. (1994) A flume study on the development and

equilibrium morphology of current ripples in fine sand.

Sedimentology, 41, 185–209.

Baas, J.H., Best, J.L., Peakall, J. and Wang, M. (2009) A phase

diagram for turbulent, transitional, and laminar clay sus-

pension flows. J. Sed. Res., 79, 162–183.

Baas, J.H., Best, J.L. and Peakall, J. (2011) Depositional pro-

cesses, bedform development and hybrid flows in rapidly

decelerated cohesive (mud-sand) sediment flows. Sedi-

mentology, 58, 1953–1987.

Bandini, P. and Sathiskumar, S. (2009) Effects of silt content

and void ratio on the saturated hydraulic conductivity and

compressibility of sand-silt mixtures. J. Geotech. Geoenv.

Eng., 135, 1976–1980, doi: 10.1061/(ASCE)GT.1943-5606.

0000177.

Bannerjee, I. (1977) Experimental study on the effect of

deceleration on the vertical sequence of sedimentary

structures in silty sediments. J. Sed. Petrol., 47, 771–783.

Best, J. and Bridge, J. (1992) The morphology and dynamics of

low amplitude bedwaves upon upper stage plane beds and

the preservation of planar laminae. Sedimentology, 39, 737–

752.

Bouma, A.H. (1962) Sedimentology of some flysch deposits;

a graphic approach to facies interpretation. PhD thesis, Uni-

versity of Utrecht, Elsevier, Utrecht, The Netherlands, 168 pp.



Bowles, F.A., Faas, R.W., Vogt, P.R., Sawyer, W.B. and Ste-phens, K. (2003) Sediment properties, flow characteristics,

and depositional environment of submarine mudflows, Bear

Island Fan. Mar. Geol., 197, 63–74.

Breien, H., De Blasio, F.V., Elverhoi, A., Nystuen, J.P. and

Harbitz, C.B. (2010) Transport mechanisms of sand in deep-

marine environments-insights based on laboratory experi-

ments. J. Sed. Res., 80, 975–990.

Costa, J.E. and Williams, G.P. (1984) Debris-flow dynamics

(video). U.S.G.S. Open File Report, 84-606, 22 mins.

Coussot, P. (1997) Mudflow Rheology and Dynamics. IAHR

Monograph, Balkema, Rotterdam, Reston, Virginia, 272 pp.

Di Base, D. and Mutti, E. (2002) The ‘Proto Adriatic Basin’. In:

Revisiting Turbidites of the Marnoso-arenacea Formation

and Their Basin-Margin Equivalents: Problems With ClassicModels. 64th EAGE Conference and Exhibition Excursion

Guidebook (Eds E. Mutti, F. Ricci Lucchi and M. Roveri),

pp. I1–I4, Parma University and ENI-AGIP Division, Parma,

Italy.

Dorrell, R.M., Hogg, A.J., Sumner, E.J. and Talling, P.J. (2011)

The structure of the deposit produced by sedimentation of

polydisperse suspensions. J. Geophys. Res., 116, F01024. 12

pp., doi: 10.1029/2010JF001718.

Elverhøi, A., Norem, H., Andersen, E.S., Dowdeswell, J.A.,Fossen, I., Haflidason, H., Kenyon, N.H., Laberg, J.S., King,E.L., Sejrup, H.P., Solheim, A. and Vorren, T. (2007) On the

origin and flow behavior of submarine slides on deep-sea

fans along the Norwegian–Barents Sea continental margin.

Geo-Mar. Lett., 17, 119–125.

Folk, R.L. and Ward, W.C. (1957) Brazos River bar: a study in

the significance of grain size parameters. J. Sed. Petrol., 27,3–26.

Gandolfi, G., Paganelli, L. and Zuffa, G.G. (1983) Petrology

and dispersal directions in the Marnoso Arenacea Forma-

tion (Miocene, northern Apennines). J. Sed. Petrol., 53, 493–

507.

Garton, M. and McIlroy, D. (2006) Large thin slicing: a new

method for the study of fabrics in lithified sediments. J. Sed.

Res., 76, 1252–1256.

Gee, M.J.R., Masson, D.G., Watts, A.B. and Allen, P.A. (1999)

The Saharandebris flow: an insight into the mechanics of long

run out submarine debris flow. Sedimentology, 46, 317–335.

Harms, J.C. and Fahnestock, R.K. (1965) Stratification, bed

forms, and flow phenomenon (with examples from the Rio

Grande). In: Primary Sedimentary Structures and theirHydrodynamic Interpretation (Ed. G.V. Middleton), SEPM

Spec. Publ., 12, 84–116.

Haughton, P.D.W., Barker, S.P. and McCaffrey, W.D. (2003)

‘Linked’ debrites in sand-rich turbidite systems – origin and

significance. Sedimentology, 50, 459–482.

Haughton, P.D.W., Davis, C., McCaffrey, W. and Barker, S.P.

(2009) Hybrid sediment gravity flow deposits – classifica-

tion, origin and significance. In: Hybrid and Transitional

Submarine Flows (Eds L.A. Amy, W.B. McCaffrey and P.J.

Talling) Mar. Petrol. Geol., 26, 1900–1918.

Hiscott, R.N. (1994) Traction-carpet stratification in turbidites

– fact or fiction. J. Sed. Res., 64, 204–208.

Hiscott, R.N. and Middleton, G.V. (1979) Depositional

mechanics of the thick-bedded sandstones at the base of a

submarine slope, Tourelle Formation (Lower Ordovician),

Quebec, Canada. In: Geology of Continental Slopes (Eds L.J.

Doyle and O.H. Pilkey), SEPM Spec. Publ., 27, 307–326.

Hiscott, R.N. and Middleton, G.V. (1980) Fabric of coarse

deep-water sandstones, Tourelle Formation, Quebec, Can-

ada. J. Sed. Petrol., 50, 703–722.

Iverson, R.M. (1997) The physics of debris flows. Rev. Geo-

phys., 35, 245–296.

Iverson, R.M. and Vallance, J.W. (2001) New views of granular

mass flows. Geology, 29, 115–118.

Iverson, R.M., Logan, M., LaHusen, R.G. and Berti, M. (2010)

The perfect debris flow? Aggregated results from 28 large-

scale experiments J. Geophys. Res., 115, F03005, doi:

10.1029/2009JF001514.

Johnson, A.M. (1970) Physical Processes in Geology. Freeman,

Cooper, San Francisco, 577 pp.

Johnson, M.R. (1994) Thin-section grain-size analysis revis-

ited. Sedimentology, 41, 985–999.

Kneller, B.C. (1995) Beyond the turbidite paradigm: physical

models for deposition of turbidites and their implications

for reservoir potential. In Characterization of Deep MarineSystems (Eds A.J. Hartley and D.J. Prosser), Geol. Soc. Spec.

Publ., 94, 31–49.

Kneller, B.C. and Branney, M.J. (1995) Sustained high-density

turbidity currents and the deposition of thick massive

sands. Sedimentology, 42, 607–616.

Kneller, B.C. and McCaffrey, W.D. (2003) The interpretation of

vertical sequences in turbidite beds: the influence of longi-

tudinal flow structure. J. Sed. Res., 73, 706–713.

Kuenen, P.H. (1965) Value of experiments in geology. Geol.

Mijnbouw, 44, 22–36.

Kuenen, P.H. (1966a) Experimental turbidite lamination in a

circular flume. J. Geol., 74, 523–545.

Kuenen, P.H. (1966b) Matrix of turbidites: experimental

approach. Sedimentology, 7, 267–298.

Laberg, J.S. and Vorren, T.O. (2000) Flow behaviour of the

submarine glacigenic debris flows on the Bear Island Trough

Mouth Fan, western Barents Sea. Sedimentology, 47, 1105–

1117.

Lastras, G., DeBlasio, F.V., Canals, M. and Elverhøi, A.(2005) Conceptual and numerical modelling of the BIG’95

debris flow, western Mediterranean Sea. J. Sed. Res., 75,784–797.

Leclair, S.F. and Arnott, R.W.C. (2003) Coarse-tail graded,

structureless strata: indicators of an internal hydraulic

jump. In, Shelf Margin Deltas and Linked Down Slope

Petroleum Systems: Global Significance and Future Explo-ration Potential (Eds H.H. Roberts, N.C. Rosen, R.H. Filion

and J.B. Anderson), pp. 1–5. SEPM, Gulf Coast Section,

Houston.

Leclair, S.F. and Arnott, R.W.C. (2005) Parallel lamination

formedbyhigh-densityturbiditycurrents. J.Sed.Res.,75,1–5.

Lowe, D.R. (1976) Subaqueous liquefied and fluidized sedi-

ment flows and their deposits. Sedimentology, 23, 285–308.

Lowe, D.R. (1982) Sediment gravity flows .2. depositional

models with special reference to the deposits of high-

density turbidity currents. J. Sed. Petrol., 52, 279–298.

Lowe, D.R. (1997) Reinterpretation of depositional processes

in a classic flysch sequence (Pennsylvanian Jackfork Group,

Ouachita Mountains, Arkansas and Oklahoma: discussion.

AAPG Bull., 81, 460–465.

Lowe, D.R. and Guy, M. (2000) Slurry-flow deposits in the

Britannia Formation (Lower Cretaceous), North Sea: a new

perspective on the turbidity current and debris flow prob-

lem. Sedimentology, 47, 31–70.

Lowe, D.R. and LoPiccolo, R.D. (1974) The characteristics and

origins of dish and pillar structures. J. Sed. Res., 44, 484–

501.

Lucente, C.C. (2004) Topography and paleogeographic evolu-

tion of a middle Miocene foredeep basin plain (Northern

Apennines, Italy). Sed. Geol., 170, 107–134.



Major, J.J. (1997) Depositional processes in large-scale debris-

flow experiments. J. Geol., 105, 345–366.

Major, J.J. (2000) Gravity-driven consolidation of granular

slurries – implications for debris-flow deposition and

deposit characteristics. J. Sed. Res., 70, 64–83.

Major, J.J. and Iverson, R.M. (1999) Debris-flow deposition:

Effects of pore-fluid pressure and friction concentrated at

flow margins. Geol Soc of Am Bull, 111, 1424–1434.

Marr, J.G., Harff, P.A., Shanmugam, G. and Parker, G. (2001)

Experiments on subaqueous sandy gravity flows: the role of

clay and water content in flow dynamics and depositional

structures. Geol. Soc. Am. Bull., 113, 1377–1386.

Martelli, L., Farabegoli, E., Benini, A., DeDonatis, M., Severi,P., Pizziolo, M., Pignone, R., Cerrina Feroni, A., RicciLucchi, F., Kearton, M.C., Angelelli, A. and Forni, S. (1994)

La geologica del Foglio 265 – S. Piero in Bagno. In: La

Cartografia Geologica della Regione Emilia-Romagna, pp.

117. Servizio Cartografico e Geologico, Regione Emilia

Romagna, Bologna, SELCA, Florence.

McAnally, W.H., Friedrichs, C., Hamilton, D., Hayter, E.,Shrestha, P., Rodriguez, H., Sheremet, A. and Teeter, A.(2007) Management of fluid mud in estuaries, bays, and

lakes. I: Present state of understanding on character and

behaviour. J. Hydraul. Eng., 133, 9–22.

Middleton, G.V. (1967) Experiments on density and turbidity

currents. III. Deposition of sediment. Can. J. Earth Sci., 4,475–505.

Mohrig, D.G., Whipple, K., Ellis, C. and Parker, G. (1998)

Hydroplaning of subaqueous debris flows. Geol. Soc. Am.

Bull., 110, 387–394.

Mutti, E. (1992) Turbidite Sandstones. Istituto di Geologia

Universita di Parma & AGIP, Parma, Italy, 275 pp.

Mutti, E., Ricci Lucchi, F. and Roveri, M. (2002) Revisiting

Turbidites of the Marnoso-arenacea Formation and their

Basin-Margin Equivalents: problems with Classic Models.

64th EAGE Conference and Exhibition Excursion Guide-

Book, Parma University and ENI-AGIP Division, Parma,

Italy.

Mutti, E., Tinterri, R., Benevelli, G., di Biase, D. and Cavanna,G. (2003) Deltaic, mixed and turbidite sedimentation of

ancient foreland basins. Mar. Petrol. Geol., 20, 733–755.

Mutti, E., Benoulli, D., Ricci Lucchi, F. and Tinterri, R. (2009)

Turbidite and turbidity currents from Alpine ‘flysch’ to the

exploration of continental margins. Sedimentology, 56, 267–

318.

Nielsen, T., Shew, R.D., Steffens, G.S. and Studlick, J.R.J.(2007) Atlas of Deepwater Outcrops, A.A.P.G. Studies in

Geology 56, Shell E & P/A.A.P.G., Tulsa, Oklahoma, USA,

504 pp.

Pieri, M. and Groppi, G. (1981), Subsurface Geological

Structure of the Po Plain. Progetto Finalizzato Geodinami-

ca/Sottoprogetto ‘Modello Strutturale’’, Publ. Italian CNR,

Milano, Italy.

Revellino, P., Hungr, O., Guadagno, F.M. and Evans, S.G.(2004) Velocity and runout simulation of destructive debris

flows and debris avalanches in pyroclastic deposits,

Campania region, Italy. Eng. Geol., 45, 295–311.

Ricci Lucchi, F. (1986) The Oligocene to Recent foreland

basins of the northern Apennines. In: Forel and Basins (Eds

P.A. Allen and P. Homewood), Int. Assoc. Sedimentol.

Spec. Publ., Blackwell Scientific, 8, 105–139.

Ricci Lucchi, F. and Valmori, E. (1980) Basin-wide turbidites

in a Miocene, over- supplied deep-sea plain: a geometrical

analysis. Sedimentology, 27, 241–270.

Richardson, J.F. and Zaki, W.N. (1954) Sedimentation and

fluidisation: Part 1. Trans. Inst. Chem. Eng., 32, 35–53.

Roveri, M., F., R.L., Lucente, C.C., Manzi, V. and Mutti, E.

(2002) Stratigraphy, facies and basin fill history of the

Marnoso Arenacea Formation. In: Revisiting Turbidites of

the Marnoso-Arenacea Formation and Their Basin-Margin

Equivalents: Problems With Classic Models. 64th EAGE

Conference and Exhibition Excursion Guidebook, (Eds E.

Mutti, F. Ricci Lucchi and M. Roveri), pp. III1–III26, Parma

University and ENI-AGIP Division, Parma, Italy.

Sassa, S. and Sekiguchi, H. (2010) LIQSEDFLOW: role of two-

phase physics in subaqueous sediment gravity flows. Soils

Found., 50, 495–504.

Sassa, S. and Sekiguchi, H. (2011) Chapter 36: Dynamics of

submarine liquefied sediment flows: theory, experiments

and analysis of field behavior. In: Submarine Mass Move-

ments and Their Consequences (Eds Y. Yamada, K.

Kawamura, K. Ikehara, Y. Ogawa, R. Urgeles, D. Mosher, J.

Chaytor and M. Strasser), Adv. Nat. Technol. Hazards Res.,31, Springer, 405–416.

Schwab, W.C., Lee, H.J., Twichell, D.C., Locat, J., Nelson,C.H., McArthur, W.G and Kenyon, N.H. (1996) Sediment

mass-flow processes on a depositional lobe, outer Missis-

sippi Fan. J. Sed. Res., 66, 916–927.

Seilacher, A. (2007) Trace Fossil Analysis. Springer, Berlin,

Germany, 226 pp.

Shanmugam, G. (1997) The Bouma Sequence and the turbidite

mind set. Earth-Sci. Rev., 42, 201–229.

Shanmugam, G. (2000) 50 years of the turbidite paradigm

(1950s–1990s): deep-water processes and facies models – a

critical perspective. Mar. Petrol. Geol., 17, 285–342.

Shanmugam, G. (2002) Ten turbidite myths. Earth-Sci. Rev.,

58, 311–341.

Shanmugam, G. and Moiola, R.J. (1995) Reinterpretation of

depositional processes in a classic flysch sequence (Penn-

sylvanian Jackfork Group), Ouachita Mountains, Arkansas

and Oklahoma. AAPG Bull., 79, 672–695.

Shaw, D.J. (1992) Colloid and Surface Chemistry – Introduc-

tion. Butterworth Heinemann, Oxford, 306 pp.

Simons, D.B., Richardson, E.V. and Nordin, C.F. (1965) Sed-

imentary structures formed by flow in alluvial channels. In:

Primary Sedimentary Structures and Their Hydrodynamic

Interpretation (Ed. G.V. Middleton), SEPM Spec. Publ., 12,34–53.

Southard, J.B (1991) Experimental determination of bed-form

stability. Ann. Rev. Earth Planet. Sci., 19, 423–455.

Stow, D.A.V. and Johansson, M. (2000) Deep-water massive

sands: nature, origin and hydrocarbon implications. Mar.Petrol. Geol., 17, 145–174.

Sumner, E.J. (2007) Settling Dynamics and Deposits of

Sheared Sand-Mud Suspensions: Implications for the

Interpretation of Submarine Gravity Flow Deposits.

Unpublished PhD thesis, University of Bristol, Bristol, UK.

Sumner, E.J., Amy, L.A. and Talling, P.J. (2008) Deposit

structure and processes of sand deposition from decelerat-

ing sediment suspensions. J. Sed. Res., 78, 529–547.

Sumner, E.J., Talling, P.J. and Amy, L.A. (2009) Deposits of

flows transitional between turbidity current and debris flow.

Geology, 37, 991–994.

Sumner, E.J., Talling, P.J., Amy, L.A., Wynn, R.B., Stevenson,C.J. and Frenz, M. (2012) Facies architecture of individual

basin-plain turbidites: comparison to existing models

and implications for flow processes. Sedimentology,

doi:10.1111/j.1365-3091.2012.01329.x.



Sylvester, Z. and Lowe, D.R. (2004) Textural trends in turbi-

dites and slurry beds from the Oligocene flysch of the East

Carpathians, Romania. Sedimentology, 51, 945–972.

Talling, P.J., Amy, L.A., Wynn, R.B., Peakall, J. and Robinson,M. (2004) Beds comprising debrite sandwiched within

co-genetic turbidite: origin and widespread occurrence in

distal depositional environments. Sedimentology, 51, 163–

194.

Talling, P.J., Wynn, R.B., Masson, D.G., Frenz, M., Cronin,B.T., Schiebel, R., Akhmetzhanov, A.M., Dallmeier-Ties-sen, S., Benetti, S., Weaver, P.P.E., Georgiopoulou, A., Zu-hlsdorff, C. and Amy, L.A. (2007a) Onset of submarine

debris flow deposition far from original giant landslide.

Nature, 450, 541–544.

Talling, P.J., Amy, L.A. and Wynn, R.B. (2007b) New insights

into the evolution of large volume turbidity currents; com-

parison of turbidite shape and previous modelling results.

Sedimentology, 54, 737–769.

Talling, P.J., Amy, L.A., Wynn, R.B., Blackbourn, G. and

Gibson, O. (2007c) Turbidity current evolution deduced

from extensive thin turbidites: Marnoso Arenacea Forma-

tion (Miocene), Italian Apennines. J. Sed. Res., 77, 172–196.

Talling, P.J., Wynn, R.B., Schmidt, D.N., Rixon, R., Sumner,E.J. and Amy, L. (2010) How did thin submarine debris

flows carry boulder-sized intraclasts for remarkable dis-

tances across low gradients to the far reaches of the Mis-

sissippi Fan? J. Sed. Res., 80, 829–851.

Talling, P.J., Malgesini, G., Sumner, E.J., Amy, L.A., Felletti,F., Blackbourn, G., Nutt, C., Wilcox, C., Harding, I.C. and

Akbari, S. (2012a) Planform geometry, stacking pattern, and

extra-basinal origin of low-strength and intermediate-

strength cohesive debris flow deposits in the Marnoso-

arenacea Formation. Geosphere, in press. doi:10.1130/

GES00734.1.

Talling, P.J., Sumner, E.J., Malgesini, G. and Masson, D.G.

(2012b) Subaqueous sediment density flows: Depositional

processes and deposit types. Sedimentology, 59, 1937–2003.

doi:10.1111/j.1365-3091.2012.01353.x.

Terzaghi, K., Peck, R.B. and Mesri, G. (1996) Soil Mechanics

in Engineering Practice, 3rd edn. Wiley-Interscience,

Oxford, UK, 729 pp.

Tinterri, R., Drago, M., Consonni, A., Davoli, G. and Mutti, E.(2003) Modelling subaqueous bipartite sediment gravity

flows on the basis of outcrop constraints: first results. Mar.Petrol. Geol., 20, 911–933.

Tripsanas, E.K., Piper, D.J.W., Jenner, K.A. and Bryant, W.R.(2008) Submarine mass-transport facies: new perspectives

on flow processes from cores on the eastern North American

margin. Sedimentology, 55, 97–136.

Twichell, D.C., Schwab, W.C., Nelson, C.H., Kenyon, N.H. and

Lee, H.J. (1992) Characteristics of a sandy depositional lobe

on the outer Mississippi fan from SeaMARC 1A sidescan

sonar images. Geology, 20, 689–692.

Van Wamel, W.A. and Zwart, P.E. (1990) The structural

geology and basin development of the Romagna-Umbrian

zone (Upper Savio- and Upper Bidente Valleys,. N Italy)..

Geol. Mijnbouw, 69, 53–68.

Vrolijk, P.J. and Southard, J.B. (1997) Experiments on rapid

deposition of sand from high-velocity flows. Geosci. Can.,

24, 45–54.

Walker, R.G. (1965) The origin and significance of the internal

sedimentary structures of turbidites. Yorkshire Geol. Soc.Proc., 35, 1–32.

Wallis, G.B. (1969) One-Dimensional Two-Phase Flow.

McGraw-Hill, New York, 408 pp.

Wood, A. and Smith, A.J. (1959) The sedimentation and sed-

imentary history of the Aberystwyth Grits (Upper Llando-

verian). Q. J. Geol. Soc. London, 114, 163–195.

Manuscript received 9 October 2011; revision accepted3 July 2012



Date post:	27-Nov-2015
Category:	Documents
Upload:	duma262005
View:	13 times
Download:	0 times

Ivana Liq Turb 2013 IAS

Documents