a depositional record of deglaciation in a paleofjord...

For permission to copy, contact [email protected] 2004 Geological Society of America348

GSA Bulletin; March/April 2004; v. 116; no. 3/4; p. 348–367; doi: 10.1130/B25242.1; 29 figures.

A depositional record of deglaciation in a paleofjord(Late Carboniferous [Pennsylvanian] of San Juan Province,

Argentina): The role of catastrophic sedimentation

Ben Kneller†

School of Earth Sciences, University of Leeds, Leeds LS2 9JT, UK

Juan Pablo MilanaCONICET e Instituto de Geologıa, Universidad Nacional de San Juan, San Juan, Argentina

Clare Buckee‡

Omar al Ja’aidi§

School of Earth Sciences, University of Leeds, Leeds LS2 9JT, UK

ABSTRACT

The combination of high sediment supplyrates and ample accommodation withinfjords provides high-resolution records ofdeglaciation. Ancient fjord fills provide thepotential for three-dimensional views of theevolution of depositional environments inresponse to changing sediment supply andbase level through the deglacial process.We describe the fill (Jejenes Formation) ofa well-exposed Late Carboniferous (Penn-sylvanian) glacial valley and its tributaries;these deposits record the progressive ma-rine flooding and infilling of topography byglacially derived sediments. The geometryof the valley and its tributaries is controlledby the underlying lower Paleozoic litholo-gies: a deep valley with steep sides existswhere the bedrock is massive limestones,and a broader, shallower valley existswhere the bedrock is generally a fine-grained olistostrome. The valleys are locallyfloored with diamictites, including both insitu tillites and remobilized diamictites. Inthe trunk valley these are locally overlainby a small, shallow-water delta. The majorpart of the valley fill consists of a #150-m-thick mudstone-dominated succession

†Present address: Institute for Crustal Studies,University of California at Santa Barbara, SantaBarbara, California 93106, USA; e-mail:[email protected].

‡Present address: Office of the Deputy PrimeMinister, Whitehall, London, UK.

§Present address: Sultan Qaboos University, Mus-cat, Oman.

(probably generated by plumes of glacialoutwash) containing numerous dropstonesthat decrease in abundance down the fjord.The mudstones contain numerous thinsandstone and conglomerate turbidites thatwere supplied laterally via subaqueousgravel fans feeding in from the tributaryvalleys, each of which has a distinctive clastsuite related to the local subcrop. The en-tire succession is overlain by .300 m ofsandy turbidites, of which the upper partincludes large mass-transport complexes.

Intercalated within the succession in thetrunk valley are structureless, graded, siltymudstones lacking dropstones but withabundant large wood fragments. Close tothe steep western margin of the trunk val-ley, each of these massive mudstones is un-derlain by a slump or debris flow, locallycontaining meter-scale blocks of Ordovi-cian limestone from the valley side. We in-terpret these as a consequence of rockfallsfrom the steeper valley sides, triggering de-bris flows on the subaqueous fjord slopes.We suggest that large solitary waves weregenerated as the rockfalls entered the wa-ter, traveling along the fjord and strippingvegetation from the shoreline. Largeamounts of mud and silt were thrown intosuspension during these events and subse-quently settled from suspension to form thestructureless graded beds. These deposits,and other mass-flow deposits within thesuccession, emphasize the potential impor-tance of catastrophic sedimentation withindeglacial successions.

Keywords: deglaciation, fjord, mass trans-port, catastrophic, Carboniferous, Pennsyl-vanian, Argentina.

INTRODUCTION

Deglacial marine sedimentary sequences re-cord the succession of depositional environ-ments associated with the melting and retreatof glaciers (Powell and Domack, 1995; Bennand Evans, 1998). Fjords may provide espe-cially detailed records of environmentalchanges associated with deglaciation because(1) large amounts of accommodation space areavailable within them owing to glacial over-deepening, and (2) their very high depositionrates provide a high-resolution record (e.g.,Eyles et al., 1990; Carlson et al., 1992; Cai etal., 1997). Although some late Quaternaryglaciomarine systems are comparatively wellexposed on land because of glacio-isostatic re-bound (e.g., Nemec et al., 1999), Quaternaryfjords generally offer only a snapshot of theirpostglacial state. Their fills are generally im-aged remotely via seismic data, bathymetry,side-scan sonar, and shallow cores (e.g., Born-hold and Prior, 1990; Cowan et al., 1999;Cowan, 2001). Consequently, the facies ar-chitecture of their sediment fills and their evo-lution through the deglaciation process are in-completely known. Although ancientsuccessions lack the time resolution and/orcomplete topographic context available inQuaternary systems, they can provide three-dimensional views of the sedimentary systemand of its evolution through time. In particular,

Geological Society of America Bulletin, March/April 2004 349

DEGLACIATION IN A LATE CARBONIFEROUS (PENNSYLVANIAN) PALEOFJORD, ARGENTINA

Figure 1. Location map for the Quebrada Grande sequence and outline of local geology.Approximate area of Figure 3 and inset map outlined. Emergent areas in dark gray.Dotted line marks international frontier. Paganzo Basin includes depositional areas to theeast of the Protoprecordillera.

Figure 2. Late Carboniferous (Pennsylva-nian) paleogeography of the Paganzo andadjacent basins, modified from Buatois andMangano (1995). Study area outlined.

they may reveal the evolution of depositionalenvironments in response to changing sedi-ment supply and base level through the degla-cial process.

In this paper we present field data from asuccession that records the deglaciation of apaleofjord close to the western margin ofGondwana (Fig. 1) during the late Westphal-ian (Cesari et al., 1987). A combination of ex-ceptional exposure and topographic dissectionallows (1) a reconstruction of the local topog-raphy, and (2) establishment of the sequenceand architecture of infilling of the relict glacialvalley. The purpose of this paper is to illus-trate the evolution of depositional environ-ments through the deglaciation process, todocument the rapid lateral and vertical facieschanges, and to emphasize the role of cata-strophic sedimentation throughout much ofthe fjord’s history.

REGIONAL PALEOGEOGRAPHY

The Jejenes Formation constitutes the fill ofthe Jejenes subbasin, located at the southeast-ern rim of the Carboniferous–Permian Pagan-zo Basin of northwestern Argentina (Fernan-dez Seveso and Tankard, 1995). This basin isone of a number of pericratonic basins thatreceived glacially derived sediment during andimmediately after the Late Carboniferous(Pennsylvanian) glaciation of Gondwana

(Eyles et al., 1995). The Paganzo Basin layclose to the active western margin of the su-percontinent, at a paleolatitude of ;608S. Thetectonic setting of the Paganzo Basin is thatof a retro-arc foreland basin (Ramos, 1988);tuffs within the Upper Carboniferous succes-sion are thought to have been derived fromthe contemporaneous arc to the west. The Pa-ganzo Basin was separated from con-temporary open-marine basins to the west(Calingasta-Uspallata and Rio Blanco Basins)by one or more coastal ranges (e.g., Salfityand Gorustovich, 1983; Fig. 2) that had beenuplifted in the Devonian and that were geo-graphically more or less coincident with themodern Precordillera. Early stages of sedi-mentation were clearly controlled by the pa-leorelief generated during the glaciation, in theform of isolated valley fills. Subsequently, thegeneral subsidence caused burial of that reliefby continental sediments.

The immediately postglacial sediments ofthe Paganzo Basin have been interpreted aslargely fluvial and lacustrine in the east andfluvial and deltaic in the west (Buatois andMangano, 1995). However, in the San Juanarea, in the southwestern part of the basin rep-resented by the Jejenes Formation (Fig. 2),palynology suggests a restricted-marine envi-ronment (Cesari and Bercowski, 1997), and arestricted-marine origin has been suggestedfor the limited trace fossil suite (Peralta and

Milana, 1998). Also, to the north of the SanJuan area, Bercowski and Milana (1990), andMartinez (1991) identified a marine influencein sedimentary rocks of late Namurian–Stephanian age associated with deglaciation(Guandacol Formation). Thus, many of thesuccessions formerly interpreted as being oflacustrine origin are apparently marine, con-nected to open-marine basins via paleovalleys.Such is the case in the Jejenes Formation, thefocus of this study.

Orientation of the paleovalleys (Fig. 1) ob-tained from valley-floor reconstruction, glacialstriations, and paleocurrents within the sedi-mentary fill suggests a main flow of ice andsubsequent sediment transport toward thenorthwest (Banchig et al., 1997; Milana et al.,1985, 1987; Milana and Bercowski, 1990).Therefore, glacial events did not originate inand radiate from the coastal ranges (cf. Amosand Rolleri, 1965; Rolleri and Baldis, 1969;Lopez-Gamundı, 1983), but originated froman ice cap located on the craton, east of theJejenes subbasin (cf. Bercowski and Zambra-no, 1990). At the glacial maximum, glacierswere crossing these ranges, excavating deeppaleovalleys between the Jejenes subbasin andthe Calingasta-Uspallata Basin. A glaciatedmarine-shelf area was located on the west rimof the Precordillera (Lopez-Gamundı, 1983).The paleovalleys were flooded by marine wa-ters after glacial retreat. As the early Paleozoicstructure basically trended south-north, itseems that there was almost no tectonic con-

350 Geological Society of America Bulletin, March/April 2004

KNELLER et al.

Figure 3. Geologic map of the Quebrada Grande area. Inset shows southwest side of paleovalley at the same scale (locations shown inFig. 1).

trol of these paleovalleys on the large scale,although we have recognized a minor litho-logic control at the boundary between the low-er Paleozoic limestones and clastic rocks inthe Quebrada Grande subcrop.

The interpretation of the upper Paleozoicrocks has changed in the light of sequence-stratigraphic concepts. In the region, deglaci-ation is first indicated by a shallow-water as-semblage followed by a rapid transgression,suggesting eustatic rise due to melt-water in-put to the oceans. The widely recognized un-conformity between the Guandacol Formation(a dropstone-bearing unit found throughoutthe Paganzo Basin) and the Tupe Formation

(a postglacial coal-bearing unit) is now inter-preted as representing a forced regression dueto postglacial isostatic rebound (Bercowskiand Milana, 1990), by comparison with thelate Weichselian evolution of southern Swe-den (Milana and Lopez, 1998).

Paleobotanical evidence suggests a rapidclimatic amelioration and transgression dur-ing the late Westphalian (Raistrickia-Plicatipollenites spore subzone; Vega, 1995).Coal-bearing intervals in the Tupe Forma-tion with a Nothorhacopteris-Botrychiopsis-Ginkgophyllum plant assemblage succeededglacial intervals, suggesting cool temperateforests (Lopez-Gamundı et al., 1993).

LOCAL SETTING

The Quebrada Grande paleovalley is ex-posed on the eastern margin of the ZondaRange, in the foothills of the Eastern Precor-dillera, ;25 km south of the city of San Juan.A gentle structural dip oblique to the valleyaxis (;258 to the east-southeast), the consid-erable dissection, and the almost complete ex-posure due to the arid climate combine to pro-vide superb sections through parts of thepaleovalley system. The section lies on theforeland side of a zone of Quaternary back-thrusting (the Eastern Precordillera); structuralcomplication is slight.



Figure 4. (A) Fence diagram showing thickness variations within the paleovalley fill; line of section shown in Figure 5.

The local subcrop to the Carboniferous con-sists of Ordovician platform limestones of theSan Juan Formation (Beresi and Bordonaro,1985), stratigraphically overlain to the east byan Ordovician–Silurian olistostrome (Rincon-ada Formation; Amos, 1954; Peralta, 1990;Peralta et al., 1994) consisting largely of high-ly deformed Ordovician–Silurian siliciclasticturbidites, but including olistoliths of San JuanFormation limestone up to 1 km or more inlength (Fig. 1). Both lower Paleozoic unitsand the boundary between them dip verysteeply to the east. The Carboniferous JejenesFormation crops out as an ;1-km-wide strip,unconformably overlain to the east by Neo-gene and Quaternary sediments (Figs. 1 and3). Late Precambrian crystalline basementcrops out ;10 km to the southeast and con-sists of amphibolite schist, garnetiferousschists, and pegmatoids. More extensive inli-ers of metamorphic basement occur 40 km tothe east in the Quaternary Pie de Palo upliftof the Sierras Pampeanas.

Form of the Paleovalley

The interpretation of the paleotopography(shown schematically in Figs. 4 and 5; seealso the section on Synthesis) is constrainedby a combination of mapping of the three-dimensional form of the sub-Carboniferousunconformity (the valley floor and sides),thickness and pinch-out relationships of thevalley fill (Fig. 4), and paleocurrent directions.A broad trunk paleovalley trends northwestover the olistostrome (Rinconada Formation)and swings northward along the subverticalcontact between the more easily eroded olis-tostrome and the resistant limestones (SanJuan Formation). The northeastern side of thevalley is well constrained by the three-dimensionality of the outcrop. The unconfor-mity provides an oblique section through thevalley side and onto the trunk-valley floor.The southwestern valley side is mostly in thesubsurface because of the regional dip, but apart has been recognized in this study where

it was onlapped at high stratigraphic levels(Fig. 4). The form of the valley is inferred bythickness trends and paleocurrents. The strati-graphic relationships along the north-trendingwestern valley side are slightly complicatedby a minor thrust along the contact betweenupper and lower Paleozoic rocks; rocks of‘‘axial-valley’’ facies have been thrust up thefjord wall along the unconformity. Slumpfolds along this north-trending valley marginface east, confirming the paleoslope. A deepmodern valley cutting westward into the SanJuan Formation outcrop of the Zonda Rangerepresents an exhumed steep-sided tributarypaleovalley that drained eastward.

Unconformity

Detailed mapping in three dimensions usingtriangulation and altimetry reveals the approx-imate orientation and slope of the paleovalleyside. The local Carboniferous paleorelief isstrongly controlled by the underlying geology.


KNELLER et al.

Figure 4. (B) Graphic logs from localities 1 and 2 shown in Figure 5.

Valley configuration and valley-side orienta-tion and paleoslope are indicated by the formof the unconformity after removal of structuraldip. Restored slopes tend to be gentler (208 orless) where the Rinconada Formation turbi-dites form the subcrop, but up to 408 or steep-er over San Juan Formation limestones, in-cluding the margins of limestone olistoliths inthe Rinconada Formation. Local depressionsin the unconformity correspond to tributaryvalleys entering the trunk valley from thenorth or northeast, and the positions of thetributary valleys are clearly controlled by li-thologies within the Rinconada Formation(Fig. 3); for example, a major limestone olis-tolith forms the interfluve between two of thetributary valleys. The relief on the unconfor-mity of over 900 m, valley-side gradients inexcess of 408, and the .250 m thickness ofthe deglacial succession (suggesting over-deepening of the valley floor) confirm thefjord-like nature of the valley.

VALLEY-FILLING SUCCESSION

The deposition of the valley-filling succes-sion can be subdivided into five stages: Stage

I produced massive and crudely stratifieddiamictites with subordinate laminated sand-stones. Stage II yielded water-laid sandstonesincluding a shallow-water delta on the valleyfloor. Stage III formed gravelly fans anddropstone-bearing basin-floor mudstones withturbidite sandstones. Stage IV created sandyturbidites free of dropstones. Stage V gener-ated mass-transport complexes and turbidites.

STAGE I: FORMATION OF GLACIALDEPOSITS

Deposits formed during stage I includemassive diamictites, having a green silty mud-stone matrix, and crudely stratified matrix-supported conglomerates. Massive diamictitesshow a very wide range of clast sizes and li-thologies, and they are the facies in whichmost striated clasts were observed (Fig. 6).The deposits may form units up to 13 m thickand occur in the paleolows of the trunk-valleyaxis and tributaries (Fig. 5), including thewestern canyon. Stones in the diamictites ofthe axial valley are dominated by crystalline-basement detritus derived from the east; dia-

mictites in the tributary valleys include morelocal lower Paleozoic lithologies. Basal dia-mictite contains boulders up to 3 m. Interbed-ded with the resedimented diamictites are thin(#0.5 m) units of current-worked (water-laid)fine-grained sandstones, generally showingparallel lamination or climbing-ripple cross-lamination.

Interpretation

Massive diamictites are interpreted as lodg-ment tills. The current-worked, fine-grainedsandstones may be related to subglacial drain-age (Dowdeswell and Scourse, 1990). Crudelystratified units may be either melt-out tills orperiglacially resedimented (debrite) material.These deposits were probably laid down sub-glacially and indicate an ice-filled valley dur-ing stage I.

STAGE II: ESTABLISHMENT OF ASHALLOW-MARINE ENVIRONMENT

Deposits representing stage II consist ofwater-laid, laminated and cross-laminated



Figure 4. (C) Graphic logs from locality 3 shown in Figure 5.

sandstones overlying lower Paleozoic base-ment on the valley sides or glacially relateddiamictites of stage I on the valley floor. Inthe tributary valleys, these deposits containparallel lamination and climbing-ripple cross-lamination (Fig. 7), with some ripple formsdraped by laminae showing vertical up-building. Subhorizontal feeding traces (cf.Planolites) are locally abundant on beddingplanes.

In the axis of the trunk valley, stage II isrepresented by a shallow-water Gilbert-typedelta, with clinoforms up to 10 m high pro-grading north-northwest along the floor of thepaleovalley (Figs. 8 and 9). The topsets (Fig.10) consist of lenses of very coarse grainedsandstones within mainly fine- to medium-grained sandstones, with ripples superimposedon small dunes showing currents to the north-east. The foresets consist of intercalations ofcompletely unsorted pebbly sandstones andconglomerates and intervals of medium- to

coarse-grained sandstone with small dunes orsubcritical climbing ripples (Fig. 10), bothshowing currents to the north-northwest. Thetoesets consist of varve-like laminated silt-stones containing abundant, small, subangularlimestone dropstones, intercalated with poorlysorted thin diamictite intervals (Fig. 11).

Interpretation

The delta-top deposits in the axial valleyprobably represent a sandy braid-plain, withlenticular channel fills, and dunes formed onbar surfaces. The current directions on the del-ta top suggest that the delta was related todrainage from the western tributary fjord. Theforesets are dominated by deposits of cohe-sionless debris flows (sensu Nemec, 1990; Ne-mec et al., 1999). The cross-stratified sand-stones were produced by processes involvingmore dilute flow; sandstones are scarce in thetoesets, so if these deposits are related to sus-

pension underflows, the flows must have by-passed the proximal prodelta region. Laminat-ed siltstones on the toesets are probablyrelated to suspension fallout from plumes(Mackiewicz et al., 1984). The dropstones anddiamictite layers (undermelt diamictite; Grav-enor et al., 1984) imply floating ice.

In the tributary systems, high sediment fall-out rates are indicated by steep ripple up-building in the sandstones of the tributary val-leys (Fig. 7), whereas the Planolites traces areindicative of low sediment fall-out rates, sug-gesting episodic sedimentation, perhaps fromsuspension underflows generated directly bysubglacial outflow.

The architecture of the delta indicates a pe-riod (perhaps decades to a century; Nemec etal., 1999) during which sediment supply keptpace with creation of new accommodation bygradual rise in lake or sea level (allowingslight aggradation of the delta top). This cir-cumstance was followed by a relatively rapid


KNELLER et al.

Figure 5. Schematic local paleotopography, based on mapping of the three-dimensional form of the unconformity, thickness and pinch-out relationships of the valley fill, and paleocurrent directions. Dip and orientation of the valley floor are based on restoration of theunconformity by removal of structural dip. Orientations of tributaries 1, 2, and 3 are based on paleocurrents of tributary fans. Alsoshown (bold dashed line) is the approximate line of the fence diagram, linking the bases of strike-normal measured sections used inFigure 4A. Circled numbers show the locations of graphic logs plotted in Figures 4B and 4C. Carboniferous outcrop shaded. Darkershading indicates deposits of stages IV and V.

Figure 6. Field photograph of striated metaquartzite clast in tillite of the trunk valley;pencil for scale.

flooding associated with a significant sea-levelrise that overwhelmed the sediment supply sothat the delta top is immediately overlain bydeep-water deposits of stage III.

STAGE III: DEPOSITION OFPROGLACIAL DEEP-WATERSEDIMENTS

The eastern part of the paleovalley includesa composite conglomerate body (Fig. 3) thatcan be divided into several separate units onthe basis of differences in bedding orientation,paleocurrents, fabrics, and clast composition(Figs. 12 and 13). These form local units ofbroadly similar character that apparently issuefrom the tributary valleys and merge with abody apparently confined to the valley axis.Because they appear to have point sources anddownlap into the valley floor, we refer to thesebodies as fans. Throughout, they are dominat-ed by matrix-supported conglomerates andcobbly sandstones (Fig. 14).

Tributary Fans

The tributary fan to the west of the largelimestone olistolith (tributary 1; Figs. 5 and12) is stratigraphically oldest where seen in

outcrop. The basal part of this unit is com-posed of glacial diamictites stratigraphicallyequivalent to sediments of stage I. Overlyingcobbly sandstones and sandy conglomerates(clast-rich sandy diamictites sensu Moncrieff,1989) showing crude stratification can betraced down-fan (Fig. 15), passing downcur-rent into rhythmically bedded deposits of thin

bedded sandstone and shale on the valleyfloor. A few clast-supported horizons are pres-ent, generally one cobble or boulder thick,which have good a-axis transverse fabrics andb-axis imbrication. Paleocurrent directionsbased on a-t fabrics of large clasts in theselayers have a vector mean of 1948 (n 5 62).Depositional dip of the fan surface was cal-



Figure 7. Field photograph of laminated and cross-laminated sandstones laid down duringstage II in tributary 1, interpreted as deposits of suspension underflows.

Figure 8. Interpreted field photograph of longitudinal section through the shallow-waterdelta deposited during stage II in the valley axis, showing asymptotic foresets, abrupttoplap, and low rate of aggradation relative to progradation. The topsets are capped byan abrupt flooding surface.

Figure 9. Cartoon of stage II depositional environment.

culated by stereographic removal of structuraldip (by using the mean of bedding in the over-lying turbidites, which must have been closeto horizontal at the time of deposition); therestored depositional dip is #128 toward;1708. The fan downlaps to the south and hasan abrupt pinch-out of conglomerate bands atthe toe (Fig. 16). The clast composition isdominated by local lower Paleozoic litholo-gies, including some striated cobbles of sand-stone, though crystalline basement clasts arealso common. The grain-size distribution isbimodal as shown by the gravelly sand matrixand many grain contacts between the largerclasts. The largest clasts are limestone boul-ders locally up to 1 m in maximum dimension.

The deposits of tributary 2, lying to the eastof the large limestone olistolith (Figs. 3 and5), consist of sandy, matrix-supported con-glomerates to cobbly sandstones and gradesinto the axial sediment body. Crystalline base-ment clasts are the most abundant (includingundeformed microgranite and ultramaficrocks) but Ordovician–Silurian sandstoneclasts are also common. They display a moreor less bimodal grain-size distribution, withwell-rounded cobbles and boulders #0.5 m,and form disorganized beds up to 10 m thick.The tributary 2 deposits yielded paleocurrentdata (directions deduced from a-t fabrics oflarge clasts) with a vector mean of 2188 (n 596). The restored depositional dip of the trib-utary 2 fan is #158 toward ;2558.

Strata representing tributary 3, the eastern-most one, are poorly exposed, but appear toinclude a clast suite similar to that of the axialfan. Clast orientations in winnowed layers in-dicate currents flowing toward 2128 (n 5 73).These strata are onlapped by the turbidites ofstage IV. Tributary 0 forms a broad depressionin the unconformity ;20 m deep, floored bydiamictite, interbedded with and overlain bylaminated sands. No gravel fan is present atthe current level of exposure, but this tributaryapparently acted as a source of gravelly sedi-ment during subsequent events. There is aprobable minor tributary fan (tributary C, rec-ognized by Bercowski et al., 1991) enteringfrom the south near the western limit of Car-boniferous outcrop within this tributary fjord(restored dips on conglomerates on the southside are toward the north or north-northeast);local current indicators in the basal unit rangefrom north to east-southeast.

Axial Fan

The uppermost unit of the composite con-glomerate body, which is by far the largest,has facies, textures, and clast lithologies that


KNELLER et al.

Figure 10. Delta foreset (lower half of photograph) and topset (upper half) of axial delta,showing abrupt toplap contact between topset and foreset with maximum dip of 198;northwest is to left. Penknife at center right for scale.

Figure 11. Toesets of axial delta; note dropstones.

are generally similar to those of tributary 3(Figs. 12 and 13). Paleocurrent directionsbased on a-t fabrics of large clasts in win-nowed layers have a vector mean of 3508 (n5 43). Depositional dips were #158 toward;3408, i.e., down the axis of the trunk valley,implying drainage from the main axial glacier.The large clasts are mainly of crystallinerocks, largely metaquartzite and semipelite,but subordinate meta-igneous rock clasts anda few local Ordovician–Silurian sandstoneclasts area also found; Ordovician limestoneclasts are absent. The clast suite is similar incomposition to that of dropstones in the time-equivalent profan deposits (see below). Theaxial fan is considerably less gravelly than thetributary fans.

The main axial fan oversteps the tributaryfans and downlaps onto the deposits of stage

I, the toe of the fan from tributary 2, or time-equivalent profan deposits. It prograded alongthe valley, its aggrading toe reaching its max-imum northwesterly extent at about the levelof the T0 tuffite (Figs. 3 and 15), and subse-quently backstepped. The conglomerates passrapidly westward, at the toe of the slope, intoa profan facies of interbedded shales and sand-stones. The maximum measured thickness ofthe fan is 35 m, but because the T0 tuffite is;100 m above the valley floor in the centerof the valley, the axial fan is probably at leastthis thick in the subsurface. In the east it isonlapped by (and interdigitates with) turbi-dites deposited during stage IV (Fig. 15).

Profan

The profan is dominated by dark gray mud-stones with thin (mostly ;5 to 20 cm),

medium- to coarse-grained, locally pebblysandstones, often regularly spaced with 10 to20 cm of mudstone between them. Somecoarser beds are present, generally less than 1m in thickness, but locally up to ;6 m, rang-ing from coarse sandstone to conglomerate.These include bed GM, (with Planolites onthe upper surface) and conglomerates CGT1

and CGT2 that are sufficiently widespread tobe used as stratigraphic markers (Figs. 3 and4).

Commonly the sandstone and conglomeratebeds are normally graded (generally ratherweakly), often with sharply graded or abrupttops, sometimes with ripple marks. The sand-stones are occasionally parallel or ripple lam-inated, but more commonly the tops of thesandstones are bioturbated both with verticalburrows (cf. Skolithos) and Planolites-typesurface-feeding traces. Plant material is com-mon on the upper surfaces of the sandstoneswhere the top of the sand is sharp and frag-ments are often current aligned. The ratio ofsandstone to shale reaches a maximum atabout the level of the T0 tuff (see below andFig. 4), corresponding to the maximum pro-gradation of the axial fan. Reliable current in-dicators from clast fabrics, sole structures, rip-ples, or aligned plant fragments are rare, butlimited data suggest widely scattered dispers-al, possibly with a mode to the northwest (Fig.17). Within the limestone canyon that formsthe western tributary fjord, the profan consistsmainly of shale.

Three thick, correlatable tuffites (T0 to T2)occur within the profan unit, along with nu-merous thinner ones. They form distinctive,whitish, normally graded beds (sometimes re-peatedly normally graded) of altered, crystal-rich tuff, each with an upward-increasing pro-portion of mud. Wood and leaf fragments areespecially common immediately below thetuffs.

On the basis of their thickness distributions,sparse current indicators, and maximum par-ticle size, sources can be inferred for a few ofthe coarser beds that are thick enough to bemapped and correlated individually. MarkerCGT1 (Fig. 4) consists of poorly sorted con-glomerate locally up to 15 m thick, generallywith rounded clasts up to 55 cm but locallycontaining rafts of diamictite and angularblocks of San Juan Formation Limestone upto 3 m across (Fig. 18). Thickness distributionindicates a source in tributary 0 on the north-eastern valley side (Fig. 19A). Another thickunit is confined to a channel along the valleyaxis whose base is eroded as much as 11 minto the underlying sedimentary material; pre-ferred a-axis orientation indicates flow along



Figure 12. Logs of composite conglomerate unit from axial fan and main tributary fans from the northern valley side.


KNELLER et al.

Figure 13. Grain size (in centimeters) and clast composition data from axial and maintributary fans.

Figure 14. Field photograph of axial-fan cobbly sandstones; flow approximately from leftto right (note structural tilt).

the channel. The clasts are similar to those ofthe subaqueous outwash fans, and this unitmay have been generated largely by remobi-lization of material from fans in the paleoval-ley tributaries. Marker CGT2 (Fig. 4), whichincludes clasts up to 43 cm across, showsthickness distributions (Fig. 19B) and maxi-mum clast size indicating a similar pattern ofdispersal, with a source apparently in tributary0 and a maximum thickness within a channelalong the valley floor. The flow or flows thatdeposited the GM marker, a normally gradedpebbly to cobbly sandstone (Fig. 4), also ap-pear to have issued from tributary 0 (Fig.19C). The presence of outsize blocks suggestsa catastrophic mass-flow origin for these de-posits, further supported by the stratigraphicassociation of CGT1 with other mass-flowunits (see below).

Dropstones

Glacial dropstones are present in both thesandstones and interbedded mudstones in theeast of the area and are locally very abundantalong bedding planes (Fig. 20). They declinein abundance toward the west and are absentaltogether in the Quebrada Grande sections(Fig. 5). Dropstones are commonly very nu-merous in the lower part of the succession,and overall become less abundant upward.

Clast compositions include metamorphicrock, igneous/meta-igneous rock, Ordovicianlimestone, and Ordovician–Silurian sandstoneof local origin. Clasts are commonly up to 15cm in longest dimension (locally up to 150cm), mostly very well rounded, although somesmall limestone clasts are subrounded to su-bangular and are broadly similar to the largerclasts within the fans. Limestone clasts de-crease in abundance upward, but sandstoneclasts remain common. The clast compositionsare consistent with a dominant contributionfrom glaciers in the trunk valley and fromlarge tributaries draining the turbidites withinthe Rinconada Formation. A minor contribu-tion from the glacier in tributary 2 decreasedwith time as this (presumably local) glacierwasted earlier than larger ice masses on thelower-relief terrain to the east.

Interpretation

The composite conglomerate body consistsof a coalescence of subaqueous outwash fans(Fig. 21). The absence of ice-push structuresor any apparent ice-contact zone indicates astable or retreating ice front. The predominantpoorly sorted facies of cobbly sandstonesprobably represents the deposits of cohesion-



Figure 15. Interpreted photomosaic of the unconformity and valley fill at the site of log1, showing downlapping of the axial fan onto the irregular valley floor. Note the back-stepping of the upper part of the fan above the level of the T0 tuff.

Figure 16. Field photograph of the transition from gravel to sandstone across the toe ofthe fan from tributary 1, downlapping onto deposits of stage II (Fig. 7); flow from left toright.

less debris flows on the relatively steep slopesof these fans, resulting from resedimentationof material deposited rapidly from a sub-aqueous jet issuing from a drainage tunnel(e.g., Powell, 1990). We interpret the clast-supported conglomerates as the products ofwinnowing of the fan surface by more diluteand probably more continuous currents of gla-cial meltwater that deflated the deposits ofmore sporadic debris flows that transportedmost of the sediment to the fan surface. In-dividual outsize clasts may be the result ofdebris fall (sensu Nemec, 1990). This sedi-ment body may be partly analogous to thebank-core facies of modern morainal banksdescribed from Tarr Inlet, Alaska, by Cai etal., 1997 (see also Lønne et al., 2001).

Locally within the fan toes (especially with-in tributary 1), individual beds can be tracedfrom matrix-supported conglomerates across asharp break in slope into thin-bedded sand-stones of the profan. This relationship be-

tween the thin-bedded sandstones of the pro-fan and the debris-flow conglomeratessuggests that the debris-flow deposits and thebasin-floor sandstones were deposited duringthe same events and the coarser clastic layersof the profan were deposited from turbiditycurrents originating on the fans. Either someflow transformation occurred at the toes of thefans, and by this transformation the debrisflows were converted to more dilute gravityflows at the break of slope, or debris flowswere accompanied by (and traveled beneath)more dilute and more mobile turbidity flowsthat bypassed the slope and deposited theirloads only on the basin floor (Lønne et al.,2001). The abruptness of the transition fromthese gravel-dominated fans to the subhori-zontal basin-floor profan facies reflects thediscrete toe of slope (Fig. 16). This transitionis equivalent to the bank-core to bank-frontboundary in morainal bank systems; in TarrInlet, this coincides with a rapid decrease in

slope to ,38. The widely scattered dispersaldirections suggest that there was influx frommany tributary systems as well as the axialfan. The regularity of spacing of the sand-stones (Figs. 4B and 20) and the clustering ofdropstones near sandstone bed tops perhapssuggests a seasonal origin (cf. Cai et al.,1997). The occasional coarser beds representcatastrophic resedimentation of material fromthe fjord sides (see below).

The tuffites similarly represent large rese-dimentation events. There are no local volcan-ic rocks, which, combined with the absence ofany lapilli-grade material, suggests that thetuffs were derived from eruptions in the arc tothe west and that the tuffites represent runoffof locally accumulated ash-fall material. Theabundance of plant detritus associated withthem was probably the result of defoliation ofthe riparian forest by the ash-fall event.

During progradation of the stage III fans,an ephemeral channel was present in the axisof the fjord during at least two intervals oftime, possibly analogous to turbidite channelsobserved within the iceberg zone of modernfjords (e.g., Tarr Inlet and Queen Inlet, GlacierBay; Cai et al., 1997; Carlson et al., 1992).These channels were partly filled by cata-strophic deposits of CGT1 and CGT2. Thechannels were apparently slightly sinuous(Figs. 19A and 19B) and between 0.5 and 11m deep. Because these channels occur withinthe sandiest part of the profan succession, itappears that they were not agents of completesand bypass (cf. Carlson et al., 1992).

The abundance of plant fragments on theupper surfaces of the thin profan sandstonebeds and the presence of dropstones within theintervening mudstones imply that the mud-stones were not formed by fallout from thetails of turbidity currents that deposited thesandstones, but were deposited by more pro-tracted suspension deposition. They wereprobably deposited from buoyant plumes as-sociated with the subaqueous meltwater dis-charge. Deposition from such plumes maytake place more rapidly than typical Stokessettling as a consequence of group settling orconvective processes (Kuenen, 1968; Carey etal., 1988). Melting icebergs might also havecontributed to the fine sediments deposited infan-proximal areas (the two processes are hardto distinguish in the sediment record; Syvitskiet al., 1996; Smith and Andrews, 2000), butthe virtual absence of dropstones in westernsections containing thick profan sedimentarydeposits argues that icebergs did not move fardown the fjord.

The abundant trace fossils of the upper Cru-ziana ichnofacies (Peralta and Milana, 1998)


KNELLER et al.

Figure 17. Vector means of paleocurrent data (with number of data points) from deposits of stages II, III, and IV, excluding data fromgravel fans. Carboniferous outcrop shaded. Darker shading indicates deposits of stages IV and V.

Figure 18. Field photograph of outsize limestone block in CGT1, deposited during stageIII.

suggest that the water column was not per-manently stratified, probably because of a pe-riodic influx of cold water. However, the factthat most traces occur on the interfaces ofsandstone beds may be a consequence of win-ter breakdown of stratification (Cowan, 1992).Low ichnofaunal diversity (commonly an in-dicator of environmental stress) with generallyonly one or two forms present (Peralta andMilana, 1998) may indicate environmentalstress due to cold or brackish water.

The progradation and subsequent backstep-ping of the axial fan presumably reflect the

changing balance of sea-level rise and rate ofsediment delivery. The ratio of sandstone toshale in the profan reaches a maximum atabout the level of the T0 tuff (see below andFig. 4), corresponding to the maximum pro-gradation of the axial fan. However, the exis-tence of the fans argues for a relatively stableposition of the ice terminus to allow the fansto develop. The westward decline in abun-dance of dropstones is consistent with calvingof icebergs from floating glacier snouts to theeast (Fig. 21). The upward-decreasing abun-dance of dropstones reflects the general east-

ward retreat of floating ice as meltingprogressed.

MEGABEDS

The profan unit is punctuated by thickerbeds of apparently homogeneous mudstone orgraded silty mudstone to mudstone. Thesebeds lack glacial dropstones (implying rapiddeposition), but contain abundant dispersedplant material, ranging from small leaf frag-ments to large lignified stem fragments up to10 cm in diameter and 30 cm in length (Fig.22). They also contain scattered siderite nod-ules, at least some of which nucleated on leaffragments.

In the west of the area, along the north-south paleovalley margin against the lime-stone subcrop, at least seven of these gradedto homogeneous mudstone beds attain a thick-ness of .1 m. At least four of these beds im-mediately overlie thick (#15 m or more)mass-flow units, consisting of pebbly mud-stones and highly deformed sediment withslump folds that face eastward (down the in-ferred local paleoslope) and have highly cur-vilinear fold axes (Fig. 23). Locally the massflows can be seen to have eroded the substrate.Some also include rafts of relatively unde-formed bedded material. The contacts betweenmass-flow units and overlying graded struc-tureless muds locally appear transitional.

Two of these beds are sufficiently thick and



Figure 19. (A) Interpolated thickness distribution (in meters) of CGT1 marker. (B) Inter-polated thickness distribution (in meters) of CGT2 marker. (C) Interpolated thicknessdistribution (in meters) of GM marker. Carboniferous outcrop shaded. Darker shadingindicates deposits of stages IV and V.

distinctive to be mapped and reliably corre-lated throughout the area; they form wide-spread markers within the profan. The lowerbed (MB1) consists of a normally graded, me-dium gray silty mudstone to mudstone with agraded sandstone at its base where it is thick-est. Apart from the ubiquitous normal grading,it is completely structureless. Its thicknessranges up to at least 25 m, declining progres-sively toward the margins and also northwarddown the fjord, suggesting a source in thesouthwest (Figs. 4 and 24). In the northernpart of the area the topography of the valleyside can be constrained by the onlap of MB1

onto the basal unit of the succession. Sand-stone (maximum, coarse-sand grade) is pre-sent at the base, where the unit is thickest inthe inferred axial part of the fjord, and also inthe west, where MB1 immediately overlies acomplex mass-flow unit of pebbly mudstoneand slumped bedded strata, #15 m thick. Theslump unit immediately overlies CGT1, andcobbles from this conglomerate are locally in-corporated into the base of the overlying massflow. MB1 (#11 m thick) is also present inthe western tributary fjord with a slump unitbeneath.

The upper of the widespread thick beds(MB2) is a weakly graded, black, organic-rich,homogeneous silty mudstone to mudstone. Itattains a thickness of 7.5 m in the west, wherea graded sandstone (grain size up to medium-sand grade) is present at the base. In the cen-tral part of the area, MB2 overlies a mass-flowunit consisting of chaotic and slumped strataup to a few meters thick.

Origin of the Megabeds

The apparent absence of dropstones withinthese units, although intercalated with beds inwhich dropstones are common (at least in theeast of the area), implies that megabed depo-sition was rapid compared to the rate of ac-cumulation of dropstones. The thick, gradednature of the beds and the virtual absence oftraction structures, even within sandy parts ofthe beds, suggests catastrophic sedimentationof large amounts of sediment by settling fromsuspension, either from turbidity currents orfrom a static turbid water column. Depositionof the megabeds from turbidity currents per seseems unlikely; the volume of material in themegabeds is many times that of the massflows and thus seems too great to have beenentrained directly from the latter to form tur-bidity currents. Turbidity currents might, al-ternatively, have been generated simulta-neously with (rather than derived from) themass flows; however, the apparent absence of

any traction structures within the sandy sec-tions of the megabeds also militates against aturbidity-current origin and suggests that thesediment in the megabeds settled from a staticwater mass. Late Quaternary homogeneoussilty clay beds in Saanich Inlet, which are per-haps comparable, have been interpreted byBlais-Stevens et al. (2001) to be the result ofperiodic breakout of a glacially dammed lake.However, the stratigraphic association of thethickest and coarsest megabeds with suba-queously generated mass flows strongly sug-gests a causal link. The section of valley mar-

gin where the mass flows occur is inferred tohave been a steep part of the fjord side. Thelocal presence of angular blocks of fjord-walllimestone within the mass flows (Fig. 18) andscattered large (tens of meters) blocks of Or-dovician limestone within the lower part ofthe succession close to the fjord marginstrongly suggests that both the mass flows andthe associated megabeds were triggered bysubaerial rockfalls (Fig. 25), which were prob-ably frequent in the cold periglacial conditionsimmediately following local deglaciation(McCarroll et al., 2001).


KNELLER et al.

Figure 20. Field photograph of dropstones within the profan unit. Compass and hand lensin lower left for scale.

Figure 21. Cartoon of stage III background depositional environments.

We suggest that the megabeds were causedby giant waves, triggered either by submarineslumping on the fjord sides (Prior et al., 1982;Aarseth et al., 1989) or by rockfalls. There areseveral close analogues for such processes inmodern fjords, in which catastrophic failure ofmaterial on fjord sides, either subaerially(rock-slide or collapse of ice from a hangingglacier) or subaqueously (slumping and slid-ing of sediment), generates large solitarywaves in the water mass (e.g., Miller, 1960).These waves may be up to 30 m or more inheight and travel along the fjord with celeritiesthat depend upon the wave height and waterdepth, but that may be as much as 40 to 50m·s21. Similar but smaller waves occurred in

Ardalsfjord, Norway, in 1983, associated witha rockfall (Aarseth et al., 1989). The absenceof any associated ice-rafted debris in the Que-brada Grande megabeds suggests that glaciercollapse was not the cause (cf. Miller, 1960).The resulting large solitary waves could haveproduced transient unidirectional currents thatsuspended large amounts of sediment from thefjord floor.

We suggest that these waves were also re-sponsible for the large amount of plant debrispresent in the megabeds. Modern examples ofsuch waves have completely stripped vegeta-tion (and often overburden) from fjord sidesup to a trimline that corresponds approximate-ly to the wave height. Large quantities of the

stripped plant material may cover the fjordsurface after the passage of such waves (as hasoccurred in Lituya Bay, Alaska, several timesin the twentieth century; Miller, 1960). Pro-gressive saturation of the floating wood debrisresulted in sinking of the stems over the pe-riod of time during which mud (suspendedfrom the fjord floor by the wave) was settlingout of suspension.

The presence of woody debris indicates thatcool temperate forests grew close to the fjordshoreline and coexisted with ice, as indicatedby dropstones in the background profan sedi-mentary deposits (cf. Vega, 1995).

STAGE IV: FORMATION OFTURBIDITE FANS

At the top of the profan sedimentary de-posits a relatively abrupt transition (a few me-ters) occurs into an ;100-m-thick package ofthinly to very thickly bedded (centimeters toseveral meters), normally graded turbiditesandstones, with thin mudstone or siltstone in-tervals forming the tops of the graded units.These rocks have been more or less arbitrarilysubdivided by using distinctive horizons offiner-grained material, which define sandstoneunits that can be traced easily over the mappedarea with consistent character and thickness.The maximum grain size is dominantly me-dium to fine sand. The thinner beds displaypartial or complete Bouma sequences, andsole structures are not uncommon, althoughthe bases of many beds are flat and featureless.Many of the thicker beds are massive through-out much of their thickness and display de-layed grading sensu Lowe (1982), i.e., thebulk of the bed is ungraded and only the upperpart shows normal grading. Beds are com-monly amalgamated into composite units ofseveral meters’ thickness. Dropstones are ap-parently absent. Fragments of plant materialare common on the tops of the sandstone beds.

In general the dispersal directions are fromthe southeast, except where deflected into trib-utary valleys (e.g., in the northeast) or reflect-ed as in the western tributary fjord (Fig. 17).The micaceous composition of the medium-grained sandstones is distinct from that of thecoarse sandstones in the profan unit, suggest-ing a more distant source in the crystallinebasement. A few graded, very coarse-grainedsandstone to conglomerate beds, as thick as;0.5 m, also occur within the turbidites.These are similar to the coarser-grained bedswithin the profan unit, and although we wereunable to determine paleocurrent directionsdirectly from these beds, their composition(clasts dominantly of lithologies similar to the



Figure 22. Field photograph of woody stem in silty mudstone suspension deposit of MB1

megabed at Quebrada Grande. Compass at left for scale.

Figure 23. Field photograph of slump underlying graded part of MB1 near western limitof outcrop.

Figure 24. Interpolated thickness distribution (in meters) of the MB1 megabed.

underlying Rinconada Formation) suggests amore local derivation, i.e., from the valleysides.

The turbidites represent deposition fromboth high-density (massive beds) and low-density turbidity currents (Lowe, 1982). Thelong correlation distance and tabular nature ofindividual beds suggests deposition on a sub-horizontal fjord floor. The large volume ofsediment and the small rate of horizontal fa-cies change suggest the establishment of alarge sand supply (perhaps the distal part of aglacier-fed delta) and the generation of turbid-ity currents large enough to reach the basinmargins. This part of the succession corre-sponds to the ice-distal zone of Cai et al.(1997).

STAGE V: GENERATION OF MASS-FLOW COMPLEXES

The uppermost exposed part of the fjord fillconsists of at least 300 m of various facies ofmass flows intercalated with thinly to thicklybedded, sandy to conglomeratic turbidites(Fig. 3). The mass flows are classified intothree facies: (1) packages of folded and dis-rupted turbidites up to tens of meters thick;(2) bouldery or pebbly sandy mudstones andmuddy sandstones, some including blocks oflimestone several meters across; and (3) clast-to matrix-supported conglomerates with abun-dant limestone boulders, ranging from chaoticand massive to moderately sorted and strati-fied, in places partly confined to erosional fea-tures (possibly channels).

In some instances, the more organized andstratified conglomerates appear to be marginalto (and possibly partly confine) bodies of lessorganized conglomerate and pebbly sandymudstone. The base of the deposits of stageV consists of a very widely distributed mas-sive to stratified conglomeratic unit as thick as20 m or more, with a locally erosional contactwith the underlying turbidites of stage IV(Figs. 26 and 27). The dominant clast types inall stage V conglomeratic facies are well-rounded cobbles and boulders of metamorphicrocks and limestones. The limestones, al-though possibly of San Juan Formation, showlow-grade metamorphism and higher degreesof strain than those seen in the deposits ofstage III, suggesting a provenance farther tothe east. The mass-flow–dominated unit iscapped by laminated and ripple cross-laminated sandstones of a facies associatedwith turbid underflows (see above, and Nemecet al., 1999).

The mass-flow deposits represent a range ofprocesses from slumps (facies 1) through co-


KNELLER et al.

Figure 25. Cartoon of stage III catastrophic deposition.

Figure 26. Field photograph of erosional base of mass-flow conglomerate deposited duringstage V, showing incision into turbidites laid down during stage IV. Note lateral injectionof conglomerate into underlying sands.

Figure 27. Field photograph of mass-flow conglomerate of the basal unit of stage V.

hesive debris flows (facies 2) to cohesionlessdebris flows. Stage V appears to represent amajor rejuvenation of the sediment supplysystem involving the destabilization of slopesand transfer of coarse clastic material fromshallower environments (presumably delta-topand foreset formed during a relative sea-levelhighstand) to the basin floor.

SYNTHESIS: RELATIVE SEA-LEVELCHANGE

The end of the second phase of Gondwanaglaciation (middle Carboniferous) was asso-ciated with a rapid eustatic sea-level rise(Lopez-Gamundı, 1997). Because the totalsea-level lowering associated with the Gond-wana glaciations is estimated to have been100 to 120 m (Crowley et al., 1991), this mid-dle Westphalian sea-level rise was presumablysomewhat less than that. Because the totalthickness of the deglacial package is 500 m ormore, it follows that a large part of the accom-modation space in which the sediment accu-mulated was due to valley overdeepening (byperhaps more than 400 m).

After retreat of the valley-filling ice fromthe Quebrada Grande area (stage I), inunda-tion of the valley apparently began gradually.The stage II delta prograded into graduallydeepening water of perhaps a few tens of me-ters depth. Because the valley floor may havelain considerably below the sill, the waterbody may initially have been a fjord lake.Flooding of the delta top at the beginning ofstage III may have been related to the estab-lishment of a marine connection as the sill wasovertopped by rising sea-level. In any case,the sea-level rise was sufficiently rapid duringstage III to overwhelm completely the sedi-ment supply (Fig. 28) and create well over100 m of accommodation in which sedimentsaccumulated.

Stage III environments can be comparedwith modern fjord environments adjacent toglacier termini, e.g., in Tarr Inlet (Cai et al.,1997; Fig. 29). The fans are partly analogousto bank-core areas, and the toes of the fanscorrelate with the break of slope at the bank-core/bank-front boundary. However, in Que-brada Grande there is no clear facies differ-entiation between a bank-front and an icebergzone. Mass-flow deposits in Quebrada Grandeare not localized near the fan toes but appar-ently occur as aprons adjacent to the steeperfjord walls, associated with one particular trib-utary (T0). The extent of an iceberg zone(characterized by the presence of dropstones)varies with time, but the location of log 2 (Fig.5) remains within the iceberg zone throughout



Figure 28. Schematic plot of relative sea-level change vs. sediment supply; darker graytone denotes sediment supply exceeding rate of accommodation by sea-level rise; palergray tone—vice versa.

Figure 29. Cartoon illustrating generalized physiographic elements, depositional environments, and sedimentary processes in the Que-brada Grande paleofjord.

stage III, whereas the strata of logs 1 and 3are largely dropstone free and correspond tothe ice-distal zone. Unlike Tarr Inlet, the per-centage of sand decreases down the fjord(compare logs 2 and 3), and the ice-distal zoneis relatively sand poor. The percentage of sandat any stratigraphic level reflects the progra-dation and subsequent backstepping of the ax-ial fan, perhaps during stillstand and subse-quent retreat of the glacier terminus

The switch from a gravel- and mud-richsystem to a sand-rich system is a transitionthat occurred stratigraphically (i.e., in time),but not apparently down-fjord during stage III.We suggest that this switch occurred becauseof (1) the retreat of the source of plume mudand (2) the establishment of a glacier-fed deltathat acted as a filter for gravel and increasedthe proportion of sand reaching the fjord floor.The rather rapid transition from stage III tostage IV indicates continued sea-level riseand/or rapid glacier retreat.

Deposits laid down during stage V bearcomparison with environments immediatelybasinward of fjord-head deltas and fan deltas(Prior and Bornhold, 1988; Carlson et al.,1992), suggesting a basinward shift of facies.In sequence-stratigraphic terms, such basin-ward shifts constitute sequence boundariesand are associated with sea-level fall; in deep-water settings, they are often marked by theonset of mass transport (e.g., Weimer, 1995).


KNELLER et al.

Stage V deposition must have involved the de-stabilization of a major up-valley staging areafor coarse sediment, and liberation of this sed-iment was most likely instigated by emer-gence and base-level resetting during relativesea-level fall.

The period of time represented by stages Ito III may be rather short, given that proglacialfans may grow rapidly (perhaps on the orderof 106 m3·yr21; Powell and Molnia, 1989) anddeglacial sedimentation rates in fjords may beextremely high close to the ice margin (200–2000 cm/yr within 1 km of the grounding line;Cowan and Powell, 1991; see also Eyles et al.,1990; Cai et al., 1997; Aarseth, 1997). Makingthe more conservative assumption that eachturbidite in the profan reflects a spring melt-water event, the entire profan may represent1000 yr or less, and stages II to IV may havebeen only a few thousand years at most. Thistime frame is consistent with rapid rates ofpostglacial sea-level rise (e.g., Lambeck,1995). A possible explanation for the putativesea-level fall associated with stage V is glacio-isostatic rebound (compare Kregnes moraine,Nemec et al., 1999).

CONCLUSIONS

The evolution of the basin fill suggests thefollowing three general points: (1) The chang-es recorded in the basin fill reflect the retreatof the valley glaciers superimposed upon abackground of sea-level change. Both of thesefactors influenced the sediment supply. (2)The strata are marked by extremely rapid tran-sitions in facies that occur both laterally andvertically. The lateral facies changes resultfrom interaction with the extreme topographyand from the depositional processes involved.Gravity-driven transport mechanisms areprone to rapid transformations associated withchanges in slope (Fisher, 1983), resulting indeposition of different grain populations withconsequent changes in grain size, but alsochanges in bed thickness and texture. Rapidvertical changes come about as a result ofdownlap and onlap associated with prograda-tion and backstepping and also because of rap-id (and incompletely understood) temporalchanges in the nature of the sediment deliverysystem. (3) The sediment fill illustrates that,in addition to the two main glacial-marineprocesses (namely, sediment rainout and sed-iment gravity processes; Eyles et al., 1985),an additional component of catastrophic sed-imentation is important, involving a range ofresedimentation processes, both gravity drivenand wave driven and involving the entirespectrum of grain sizes.

ACKNOWLEDGMENTS

This study was supported by the Turbidites Re-search Consortium at the University of Leeds, fund-ed by Arco, BG, BHP, BP, Chevron, Conoco, Elf,Shell, and Texaco. Buckee was supported by aNERC (Natural Environment Research Council)studentship. Al Ja’aidi was supported by PetroleumDevelopment Oman. Bill McCaffrey contributed lo-gistically and intellectually to the work and iswarmly thanked. Helpful reviews by E. Domack, N.Eyles, and associate editor Bob Anderson contrib-uted significantly to the final form of the paper.

REFERENCES CITED

Aarseth, I., 1997, Western Norwegian fjord sediments: Age,volume, stratigraphy, and role as temporary deposi-tory during glacial cycles: Marine Geology, v. 143,p. 39–53.

Aarseth, I., Lønne, Ø., and Giskeødegaard, O., 1989, Sub-marine slides in glaciomarine sediments in some west-ern Norwegian fjords: Marine Geology, v. 88,p. 1–21.

Amos, A.J., 1954, Estructura de las formaciones paleozo-icas de La Rinconada, pie oriental de la Sierra Chicade Zonda (San Juan): Revista de la Asociacion Geo-logica Argentina, v. 9, no. 1 ,p. 5–38.

Amos, A.J., and Rolleri, E.O., 1965, El Carbonico marinoen el Valle Calingasta–Uspallata (San Juan–Mendoza): Buenos Aires, Boletın de InformacionesPetroleras, v. 1, p. 1–23.

Banchig, A.L., Milana, J.P., and Lech, R., 1997, FormacionCasa de Piedra, una nueva unidad Neopaleozoica gla-cimarina en la Sierra del Tontal, Precordillera Occi-dental, San Juan: San Juan, Argentina, II Jornadas deGeologıa de Precordillera, p. 92–97.

Benn, D.I., and Evans, D.J.A., 1998, Glaciers and glacia-tion: London, Arnold, 734 p.

Beresi, M.S., and Bordonaro, C.L., 1985, La formacion SanJuan en la Quebrada de las Lajas, Sierra Chica deZonda, Provincia de San Juan, in Actas del CongresoGeologico Argentino, Buenos Aires: Asociacion Geo-logica Argentina, v. 9, p. 95–107.

Bercowski, F., and Zambrano, J.J., 1990, Carbonıfero y Per-mico de San Juan, in Bordonaro, O.L., ed., Geologıade San Juan: San Juan, Argentina, Relatorio de geo-logıa y recursos naturales de la Provincia de San Juan,11th Congreso Geologico Argentino, p. 33–52.

Bercowski, F., and Milana, J.P., 1990, Sedimentacion gla-cimarina: Nueva interpretation para la FormacionGuandacol (Carbonıfero) en el perfil de rio Francia,Precordillera Central, San Juan: 3rd Reunion Argen-tina de Sedimentologıa: Actas, p. 31–36.

Bercowski, F., Glenisson, F., Milana, J.P., and Ruzycki, L.,1991, Esquema litofacial y paleoambiental del Car-bonıfero en Quebrada Grande, Precordillera Argenti-na, in Perez, d’A. Ernesto, ed., Congreso GeologicoChileno, Resumenes Expandidos, v. 6, p. 485–489.

Blais-Stevens, A., Bornhold, B.D., Kemp, A.E.S., Dean,J.M., and Vann, A.A., 2001, Overview of late Qua-ternary stratigraphy in Saanich Inlet, British Colum-bia: Results of the Ocean Drilling Program Leg 169S:Marine Geology, v. 174, p. 3–26.

Bornhold, B.D., and Prior, D.B., 1990, Morphology andsedimentary processes on the subaqueous Noeick Riv-er delta, British Columbia, Canada, in Colella, A., andPrior, D.B., eds., Coarse-grained deltas: InternationalAssociation of Sedimentologists Special Publication10, p. 169–181.

Buatois, L.A., and Mangano, M.G., 1995, Sedimentary dy-namics and evolutionary history of a Late Carbonif-erous lake in north-western Argentina: Sedimentology,v. 42, p. 415–436.

Cai, J., Powell, R.D., Cowan, E.A., and Carlson, P.R., 1997,Lithofacies and seismic reflection interpretation oftemperate glacimarine sedimentation in Tarr Inlet,Glacier Bay, Alaska: Marine Geology, v. 143,p. 5–37.

Carey, S.N., Sigurdsson, H., and Sparks, R.S.J., 1988, Ex-perimental studies of particle-laden plumes: Journal ofGeophysical Research, v. 93, p. 15,314–15,328.

Carlson, P.R., Powell, R.D., and Phillips, A.C., 1992, Sub-marine sedimentary features on a fjord delta front,Queen Inlet, Glacier Bay, Alaska: Canadian Journalof Earth Sciences, v. 29, p. 565–573.

Cesari, S.N., and Bercowski, F., 1997, Palinologıa de laFormacion Jejenes (Carbonıfero) en la Quebrada delas Lajas, Provincia de San Juan, Argentina, Nuevasinferencias paleoambientales: Ameghiniana, v. 34,p. 497–509.

Cesari, S., Gutıerrez, P., and Amos, A., 1987, Revisionbioestratigrafica de la Formacion Jejenes, (Carbonı-fero), Provincia de San Juan, Primerias Jornadas deGeologıa de Precordillera (San Juan, 1985): BuenosAires, Asociacion Geologica Argentina, p. 174–180.

Cowan, E.A., 1992, Meltwater and tidal currents: Controlson circulation in a small glacial fjord: Estuarine,Coastal, and Shelf Science, v. 32, p. 381–391.

Cowan, E.A., 2001, Late Pleistocene glacimarine record inSaanich Inlet, British Columbia, Canada: Marine Ge-ology, v. 174, p. 43–57.

Cowan, E.A., and Powell, R.D., 1991, Ice-proximal sedi-ment accumulation rates in a temperate glacial fjord,southeastern Alaska, in Anderson, J.B., and Ashley,G.M., eds., Glacial marine sedimentation: Paleocli-matic significance: Geological Society of AmericaSpecial Paper 261, p. 61–73.

Cowan, E.A., Seramur, K.C., Cai, J., and Powell, R.D.,1999, Cyclic sedimentation produced by fluctuationsin meltwater discharge, tides and marine productivityin an Alaskan fjord: Sedimentology, v. 46,p. 1109–1126.

Crowley, T.J., Baum, S.K., and Hyde, W.T., 1991, Climatemodel comparison of Gondwanan and Laurentide gla-ciations: Journal of Geophysical Research, v. 96,p. 9217–9226.

Dowdeswell, J.A., and Scourse, J.D., 1990, On the descrip-tion and modelling of glacimarine sediments and sed-imentation, in Dowdeswell, J.A., and Scourse, J.D.,eds., Glacimarine environments: Processes and sedi-ments: Geological Society of America Special Paper53, p. 1–13.

Eyles, C.H., Eyles, N., and Miall, A.D., 1985, Models ofglaciomarine sedimentation and their application tothe interpretation of ancient glacial sequences: Palaeo-geography, Palaeoclimatology, Palaeoecology, v. 51,p. 15–84.

Eyles, N., Mullins, H.T., and Hine, A.C., 1990, Thick andfast: Sedimentation in a Pleistocene fiord lake of Brit-ish Columbia, Canada: Geology, v. 18, p. 1153–1157.

Eyles, N., Gonzales Bonorino, G., Franca, A.B., and Eyles,C.H., 1995, Hydrocarbon-bearing late Paleozoic gla-ciated basins of southern and central South America,in Tankard, A.J., Suarez-Soruco, R., and Welsink,H.J., eds., Petroleum basins of South America: AAPGMemoir 62, p. 165–183.

Fernandez Seveso, F., and Tankard, A.J., 1995, Tectonicsand stratigraphy of the late Paleozoic Paganzo Basinof western Argentina and its regional implications, inTankard, A.J., Suarez-Soruco, R., and Welsink, H.J.,eds., Petroleum basins of South America: AAPGMemoir 62, p. 285–301.

Fisher, R.V., 1983, Flow transformations in sediment grav-ity flows: Geology, v. 11, p. 273–274.

Gravenor, C.P., von Brunn, V., and Dreimanis, A., 1984,Nature and classification of waterlain glaciogenic sed-iments, exemplified by Pleistocene, late Palaeozoicand late Precambrian deposits: Earth Science Reviews,v. 20, p. 105–166.

Kuenen, Ph. H., 1968, Settling convection and grain-sizeanalysis: Journal of Sedimentary Petrology, v. 38,p. 817–831.

Lambeck, A.M., 1995, Late Devensian and Holoceneshorelines of the British Isles and North Sea frommodels of glacio-hydro-isostatic rebound: GeologicalSociety [London] Journal, v. 152, p. 437–448.

Lønne, I., Nemec, W., Blikra, L.H., and Lauritsen, T., 2001,Sedimentary architecture and dynamic stratigraphy ofa marine ice-contact system: Journal of SedimentaryResearch, v. 71, p. 922–943.



Lopez-Gamundı, O.R., 1983, Modelo de sedimentacion gla-cimarina para la Formacion Hoyada Verde, Paleozoicosuperior de la Provincia de San Juan: Revista de laAsociacion Geologica Argentina, v. 38, p. 60–72.

Lopez-Gamundı, O.R., 1997, Glacial–post-glacial transitionin the late Paleozoic basins of southern South Amer-ica., in Martini, I.P., ed., Late glacial and postglacialenvironmental changes—Quaternary, Carboniferous–Permian, and Proterozoic: New York, Oxford Univer-sity Press, p. 147–168.

Lopez-Gamundı, O.R., Limarino, O., and Cesari, S.N.,1993, Late Paleozoic paleoclimatology of central westArgentina: Palaeogeography, Palaeoclimatology, Pa-laeoecology, v. 91, p. 305–329.

Lowe, D.R., 1982, Sediment gravity flows: II. Depositionalmodels with special reference to the deposits of high-density turbidity currents: Journal of Sedimentary Pe-trology, v. 52, p. 279–297.

Mackiewicz, N.E., Powell, R.D., Carlson, P.R., and Molnia,B.F., 1984, Interlaminated ice-proximal glacimarinesediments in Muir Inlet, Alaska: Marine Geology,v. 57, p. 113–147.

Martinez, M., 1991, Marine fauna occurrence in the Guan-dacol Formation (Carboniferous) at Agua Hedionda,San Juan, northeastern Precordillera, Argentina: Ab-stracts, 12th International Congress on Permian andCarboniferous Stratigraphy and Geology, Buenos Ai-res (addenda), p. 3.

McCarroll, D., Shakesby, R.A., and Matthews, J.A., 2001,Enhanced rockfall activity during the Little Ice Age:Further lichenometric evidence from a Norwegian ta-lus: Permafrost and Periglacial Processes, v. 12,p. 157–164.

Milana, J.P., and Bercowski, F., 1990, Facies y geometrıade depositos glaciales en un paleovalle Carbonıfero dePrecordillera Central, San Juan, Argentina: 3rd Re-union Argentina de Sedimentologıa, San Juan, Argen-tina, Actas, p. 199–204.

Milana, J.P., and Lopez, Sergio, 1998, Solar cycles recordedin Carboniferous glacimarine rhythmites (western Ar-gentina); relationships between climate and sedimen-tary environment: Paleogeography, Paleoclimatology,Paleoecology, v. 144, no. 1–2 ,p. 37–63.

Milana, J.P., Banchig, A.L., and Bercowski, F., 1985, Des-cripcion sedimentologica de la seccion basal de unasecuencia Neopaleozoica de la Precordillera Central:

1st Jornadas de Geologıa de Precordillera, San Juan,Argentina, Actas, p. 295–300.

Milana, J.P., Bercowski, F., and Lech, R.R., 1987, Analisisde una secuencia marino-continental Neopaleozoicaenla region del rıo San Juan, Precordillera Central,San Juan: 10th Congreso Geologico Argentino, Tu-cuman, Argentina, Actas III, p.113–116.

Miller, D.J., 1960, Giant waves in Lituya Bay, Alaska:U.S. Geological Survey Professional Paper 354-C,p. 51–86.

Moncrieff, A.C.M., 1989, Classification of poorly sortedsediments: Sedimentary Geology, v. 65, p. 191–194.

Nemec, W., 1990, Aspects of sediment movement on steepdeltas slopes, in Colella, A., and Prior, D.B., eds.,Coarse-grained deltas: International Association ofSedimentologists Special Publication No.10, p. 29–73.

Nemec, W., Lønne, I., and Blikra, L.H., 1999, The Kregnesmoraine in Gauldalen, west-central Norway: Anatomyof a Younger Dryas proglacial delta in a palaeofjordbasin: Boreas, v. 28, p. 454–476.

Peralta, S.H., 1990, The Silurian of the San Juan Precor-dillera, in Bordonaro, O., ed., Relatorio de geologia yrecursos naturales de la Provincia de San Juan: Bue-nos Aires, Asociacion Geologica Argentina, p. 48–64.

Peralta, S.H., and Milana, J.P., 1998, Trazas fosiles en albase de la Formacion Jejenes (Carbonıfero medio-su-perior), Precordillera Oriental de San Juan [abs.]: Ter-cera Reunion Argentina de Icnologıa, Mar del Plata,p. 23.

Peralta, S.H., Baldis, B.A., and Cecione, A., 1994, Rincon-ada Formation, eastern Precordillera of San Juan, Ar-gentina: Sedimentary melange (olistostromes) or tec-tonics?, in Campos, E., ed., 7th Congreso GeologicoChileno, Actas: Antofagasta, Chile, Universidad delNorte, v. 1, p. 129–133.

Powell, R.D., 1990, Glacimarine processes at grounding-line fans and their growth to ice-contact deltas, inDowdeswell, J.A., and Scourse, J.D., eds., Glacimar-ine environments: Processes and sediments: Geologi-cal Society of America Special Paper 53, p. 53–73.

Powell, R.D., and Domack, E., 1995, Modern glaciomarineenvironments, in Menzies, J., ed., Glacial environ-ments, Volume 1: Modern glacial environments: Pro-cesses, dynamics and sediments: Oxford, Butterworth-Heinemann, p. 445–486.

Powell, R.D., and Molnia, B.F., 1989, Glacimarine sedi-mentary processes, facies and morphology of the

south-south-east Alaska shelf and fjords: Marine Ge-ology, v. 85, p. 359–390.

Prior, D.B., and Bornhold, B.D., 1988, Submarine mor-phology and processes of fjord fan deltas and relatedhigh-gradient systems: Modern examples from BritishColumbia, in Nemec, W., and Steel, R.J., eds., Fandeltas: Sedimentology and tectonic settings: Glasgow,UK, Blackie and Son, p. 125–143.

Prior, D.B., Bornhold, B.D., Coleman, J.M., and Bryant,W.R., 1982, Morphology of a submarine slide, KitimatArm, British Columbia: Geology, v. 10, p. 588–592.

Ramos, V.A., 1988, The tectonics of central Andes: 308–338S latitude, in Clark, S.P., Jr., Burchfiel, B.C., andSuppe, J., eds., Processes in continental lithosphericdeformation: Geological Society of America SpecialPaper 218, p. 31–54.

Rolleri, E.O., and Baldis, B.A.J., 1969, Paleogeography anddistribution of Carboniferous deposits in the ArgentinePrecordillera: IUGS Symposiums: Gondwana stratig-raphy, Earth sciences: Geneva, UNESCO, v. 2,p. 1005–1024.

Salfity, J.A., and Gorustovich, S.A., 1983, Paleogeografıade la Cuenca de Paganzo (Paleozoico Superior): Re-vista de la Asociacion Geologica Argentina, v. 38,p. 437–453.

Smith, L.M., and Andrews, J.T., 2000, Sediment character-istics in iceberg-dominated fjords, Kangerlussaq re-gion, East Greenland: Sedimentary Geology, v. 130,p. 11–25.

Syvitski, J.P.M., Andrews, J.T., and Dowdeswell, J.A.,1996, Sediment deposition in an iceberg dominatedglacimarine environment, East Greenland: Basin-fillimplications: Global and Planetary Change, v. 12,p. 251–270.

Vega, J.C., 1995, La flora fosil de la Formacion Jejenes(Carbonıfero), Implicancias paleoclimaticas y paleo-biogeograficas: Ameghiniana, v. 32, p. 31–40.

Weimer, P., 1995, Sequence stratigraphy of the MississippiFan (late-Miocene–Pleistocene), northern deep Gulf ofMexico, in Pickering, K.T., Hiscott, R.N., Kenyon,N.H., Lucchi, F.R., and Smith, R.D.A., eds., Atlas ofdeep water environments: Architectural style in tur-bidite systems: London, Chapman and Hall, p. 94–99.

MANUSCRIPT RECEIVED BY THE SOCIETY 15 AUGUST 2002REVISED MANUSCRIPT RECEIVED 18 FEBRUARY 2003MANUSCRIPT ACCEPTED 14 APRIL 2003

Printed in the USA

a depositional record of deglaciation in a paleofjord...

Documents