Marine Geology 346 (2013) 31–46

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Marine Geology

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Glacial and tectonic control on fjord morphology and sedimentdeposition in the Magellan region (53°S), Chile

Sonja Breuer a,⁎, Rolf Kilian a, Dirk Schörner b, Wilhelm Weinrebe b, Jan Behrmann b, Oscar Baeza a

a University of Trier, Department of Geology, FB VI, Behringstraße, 54286 Trier, Germanyb GEOMAR Helmholtz Institute for Ocean Research, Wischhofstraße 1-3, 4148 Kiel, Germany

⁎ Corresponding author. Tel.: +49 651 2014670; fax: +E-mail address: sonja.breuer@uni-trier.de (S. Breuer).

0025-3227/$ – see front matter © 2013 Elsevier B.V. All rihttp://dx.doi.org/10.1016/j.margeo.2013.07.015

a b s t r a c t

a r t i c l e i n f o

Article history:Received 12 October 2011Received in revised form 6 July 2013Accepted 23 July 2013Available online 1 August 2013

Communicated by J.T. Wells

Keywords:reflection seismichigh-resolution bathymetrysediment sequencestectonic settingMagellan regionChile

In the Patagonian Andes erosion by temperate Pleistocene glaciers has produced a deeply incised fjord system inwhich glacial and non-glacial sedimentswere deposited since the Late Glacial glacier retreat. So far, fjord bathym-etry and structures in the sediment infill were widely unexplored. Here we report the results of an investigationof morphology and sediment characteristics of a 250 km long fjord transect across the southernmost Andes(53°S), usingmultibeam and parametric echosounder data, and sediment cores. Subaquatic morphology revealscontinuity of on-land tectonic lineaments mapped using field and remote sensing data. Our results indicate thatglacial erosion and fjord orientation are strongly controlled by three major strike-slip fault zones. Furthermore,erosion is partly controlled by older and/or reactivated fracture zones as well as by differential resistance ofthe basement units to denudation. Basement morphology is regionally superimposed by Late Glacial and Holo-cene subaquatic moraines, which are associated to known glacier advances. The moraines preferentially occuron basement highs, which constrained the glacier flows. This suggests that the extent of glacier advances wasalso controlled by basement morphology. Subaquatic mass flows, fluid vent sites as well as distinct Late Glacialand Holocene sediment infills have furthermore modified fjord bathymetry. In the western fjord system closeto the Strait of Magellan subaquatic terraces occur in 20 to 30 m water depth, providing an important tag forproglacial lake level during the Late Glacial.

© 2013 Elsevier B.V. All rights reserved.

1. Introduction

The morphology of the previously glaciated southernmost Andes(52°–54°S, Fig. 1A) is characterised by an up to 2000 m deep fjord sys-tem, mainly formed by Pleistocene glacial erosion. Orientation of thesefjords is more or less controlled by a neotectonic strike-slip fault system(Glasser and Ghiglione, 2009; Waldmann et al., 2011), related to theoblique convergence of the Antarctic and South American plates(Diraison et al., 1996; Ramos, 1999; Lodolo et al., 2003; Costa et al.,2006; Oliver and Malumian, 2008; Ghiglione et al., 2010). However,only a very crude bathymetry existed formost of the fjords, and very lit-tle information has been published regarding subaquatic moraine sys-tems and other sediments (Kilian et al., 2007a).

We have investigated the bathymetry as well as sediment types andfjord orientations. Present day glaciers originate from the 200 km2 largeGran Campo Nevado ice cap (GCN; Fig. 1C), located at the ice divide ofthe southernmost Andes at 53°S (Schneider et al., 2007). The investigat-ed fjords are aligned along a 250 km long E–W fjord transect thatcrosses the main lithological units or morphostructural provinces ofthe Andean mountain belt (trending N–S; Fig. 1B). We focus on fourareas to the west and to the east of the GCN, covering different

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morphostructural provinces: (1) fjords west of the GCN within thePatagonian Batholith, (2) the Gajardo Channel east of the GCN incisedinto a Jurassic–Cretaceous metamorphic basement, (3) the westernSeno Skyring located in the Magallanes fold-and-thrust belt and (4) theeastern Seno Skyring and Seno Otway in the Magallanes foreland basin.For the surveys we used parametric echosounding and high-resolutionmultibeam bathymetry. Additionally, twenty sediment cores were recov-ered from the fjord sediments.

2. Regional setting

Themajor orogeny of the southernmost Patagonian Andes started inthe Middle to Late Jurassic, forming an island arc and associated back-arc basin. Extensional faulting and subsidence lead to the evolution ofthe Rocas Verde marginal basin, with an infill of Jurassic volcaniclasticsand Cretaceous turbidite sequences. In the late Cretaceous and thePalaeogene crustal shortening and thickening caused the closure ofthe Rocas Verde basin, and the formation of the Magallanes fold-and-thrust belt (Fildani et al., 2003; Menichetti et al., 2008; Oliver andMalumian, 2008; Ghiglione et al., 2010; Romans et al., 2010). TheMagallanes foreland basin developed east of the Andean orogen, filledwith an up to 7 km thick sediment sequence (Ramos, 1989; Ghiglioneet al., 2009). From Neogene onwards the region was affected by NW–

SE sinistral wrench tectonics (Diraison et al., 1996). The left-lateral

Fig. 1. A) Southern South America with recent tectonic and volcanic setting (Stern and Kilian, 1996), and the Magallanes Basin with its isopachs (Ramos, 1989; Diraison et al., 1996).B) Geological and tectonic features of the Magellan region (SERNAGEOMIN, 2003; modified). C) The investigated 250 km long fjord transect across the Patagonian Andes. The boxesshow the key areas that are described in more detail in the text and other figures.

32 S. Breuer et al. / Marine Geology 346 (2013) 31–46

Magallanes–Fagnano fault zone (MFZ; Fig. 2) represents the majortransform fault system (Lodolo et al., 2003). Some minor NNE trendingright-lateral strike-slip faults can be interpreted as associated R′ Riedelshear zones (Maffione et al., 2010). Elongated pull-apart basins, horstand graben structures, pressure ridges, uplifted and faulted flower

structures and drag folds are due to releasing or restraining bends inthe fault system (Lodolo et al., 2003).

During the Last Glacial Maximum (LGM) an 80 kmwide ice field ex-tended from the GCN to the west and east (Kilian et al., 2007a). Itformed part of a large (2500 km N–S) Patagonian Icefield (Glasser and

Fig. 2. False colour Landsat ETM+ satellite image from 2000 of the Magellan fjord region between 52°S to 53.3°S with morphological lineaments representing the three main strike-slipfault directions (MFZ, SCFZ, GFZ) and their associated trendlines. The orientation of all faults and trendlines are plotted in a rose diagram. The fold axes and thrust fault derived from theofficial geological map (SERNAGEOMIN, 2003; modified). Epicentres of relevant earthquakes in this region.After USGS Earthquake Hazards Program.

33S. Breuer et al. / Marine Geology 346 (2013) 31–46

Ghiglione, 2009). Extent of the ice cover is well documented in the areaof Seno Skyring and Magellan Strait (Benn and Clapperton, 2000;McCulloch et al., 2005a; Kilian et al., 2007a). Between 18 and 15 kyrBP a major ice recession led to the formation of extended proglaciallakes, like the Seno Skyring, Seno Otway and the fjord systems of theMagellan Strait (Glasser et al., 2008; Rabassa, 2008). Proglacial lakeswere successively transformed to fjords due to the global sea level rise(Kilian et al., 2007b). The orientation ofmost fjords follows the structur-al geological lineaments (Glasser and Ghiglione, 2009;Waldmann et al.,2011). The Seno Otway, Seno Skyring and the Eastern Strait of Magellancross-cut the fold-and-thrust belt following extensional structures(Diraison et al., 2000) oriented perpendicular to the Andean orogen.The fjords and bays carved out by the glaciers (McCulloch et al.,2005a) act as depositional environments for glacial, glaciomarine, andglaciolacustrine sediments.

Five different deglacial ice retreat stages (A to E) are evident in theMagellan region (Clapperton et al., 1995; McCulloch et al., 2005b;Rabassa, 2008). In the western fjord systems the related moraines arein general subaquatic, like that of stage D in Seno Skyring and stage Ewithin the Gajardo Channel (Kilian et al., 2007a). In the westernMagellan region the glaciers retreated to present-day extents between

12.6 and 11.8 kyr BP (e.g. Kilian et al., 2007a). Pollen records from theGCN area indicate an evolved Magellanic rainforest after 11.2 kyr(Lamy et al., 2010).

Seen in a global context, the Chilean fjord system belongs to theSouthern Fjord Belt (Syvitski et al., 1987). The present-day GCN icecap reaches 1700 m (a.s.l.) and feeds temperate valley and tidewa-ter glaciers (Warren and Aniya, 1999; Benn and Clapperton, 2000;Rabassa et al., 2008; Koppes et al., 2009). The Chilean tidewater gla-ciers extend into the warmest settings, in terms of air and watertemperatures (Dowdeswell et al., 1998). In temperate fjords, likein the southernmost Chilean fjord system, the postglacial depositionis dominated by high organic input (Syvitski and Shaw, 1995; Howeet al., 2010), which can be as well measured in lake sediment coresfed by streams coming from the GCN region (Breuer et al., 2013).The sediment from the GCN-derived glaciers enters the fjords mainlyas hypopycnal flows (Fig. 3B). Several Holocene tephra layers from vol-canoes of the Austral Volcanic Zone (AVZ; Stern and Kilian, 1996) formimportant age markers in sediment cores of the investigated fjord area(Kilian et al., 2007a), with themost prominent one originating from the4.15 kyr BP Mt. Burney (Figs. 1, 2) eruption (Kilian et al., 2003; Stern,2008).

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3. Material and methods

3.1. Core data

Stratigraphies and sedimentological properties deduced from pistonand gravity cores were used to improve the interpretation of sedimentlayers seen in the seismic profiles. The sediment cores include: CG-1(Baeza, 2005; Kilian et al., 2007a); Sky-1 (Kilian et al., 2007a,b; Lamyet al., 2010). For cores BA-1, LO-1, SG-1 and MD07-3126 a chronologywas developed based on tephra chronology, 14C dating frommacroplantremnants (Poznan Radiocarbon Laboratory; Table 1) and 210Pb dating(Table 2; Appleby and Oldfield, 1992). All 14C ages were converted tocalibrated ages before present (BP 1950) using the Calib 6.0.1 softwareand SHCal04 curve for the Southern Hemisphere (Stuiver and Reimer,1993; McCormac et al., 2004; Stuiver et al., 2011). All depicted agesare means of 1-sigma values.

3.2. Acoustic data

Along the fjord transect 1600 km of echosounder profiles andaround 500 km of multibeam data were recorded, processed andinterpreted. We used the Parametric Echosounder System (SES 2000)from Innomar (Wunderlich and Müller, 2003), which provides highand variable low (4–12 kHz) frequency signals at water depths of 0.5to 800 m with a vertical resolution of less than 5 cm. The pulse lengthranges between 0.08 and 1 ms, and the beam width is 1.8 (WunderlichandMüller, 2003). The fjord bathymetrywas calculated from thehigh fre-quency signal. To obtain optimal visualization of sediments,we tested lowfrequency signals between 4 and 12 kHz to obtain an optimal trade-off ofpenetration and resolution. This allowed a penetration of up to 55 m.However, in some areas gas in the sediment hampered penetration,resulting in transparent layers (Fader, 1997). The ELAC-Nautic Seabeam1180 multibeam sonar system was used for high-resolution seafloorbathymetry in selected areas. The system is designed for water depthsup to 600 m. However, deeper basins can be imaged with a reducedswath width. Using a frequency of 180 kHz, it is able to record 126 singledepth values with a swath of max. 153°.

3.3. Data editing and processing

The echosounder data were edited using the post processing soft-ware ISE (version 2.9) from Innomar. A sound velocity correction wasapplied by using several CTD records taken in the fjord system. Acousticvelocities of water-rich sediments may vary by ±2% compared tomarine water (Hamilton, 1979). In our case, comparison of the depthof the prominent 4.15 kyr BP tephra reflector in the echosounder pro-files and related sediment cores indicated sound velocities around1500 m/s. The most important layers were digitised and surfaces ofthem were calculated by using different geostatistical interpolationmodels. The ‘Ordinary Kriging’method was used for restricted areas, be-cause it relies on the spatial correlation structure of the data to determinetheweighting values. For the extended Seno Skyringwith less coverage ofseismic lines the ‘Inverse Distance Weighted’ method (IDW) was used,which weighted the nearby data points. The multibeam data wereprocessed at the IFM-GEOMAR (Kiel, Germany). This included checkingand editing of the navigation, removing of erroneous measurementsand artefacts by automatic and manual cleaning, conversion of recordedtravel times to positions and depths by complete ray-tracing throughthe water column taking into account measured sound velocity, depthprofiles, and calculation of digital terrain models.

Fig. 3. A) Georeferenciated ortho-image (Schneider et al., 2007) combinedwith high resolutionshown. Moraines, deltas and fluvial sediments are distinguished within the Quaternary sedimecolour Landsat ETM+ satellite image from 2000 and black-and-white ortho-image (SchneidKriging). Thewhite dots show the localities of the SG-1 andMD07-3126 core. The red dashed lincan be seen clearly in the ortho-image.

3.4. Structural and lineament analysis

Lineament analysis includes terrestrial erosional incisions, fjord ori-entations, as well as distinct, straight lines in the morphology whichmay represent fractures and tectonic faults. We mainly interpretedGeoCover Landsat satellite images, recorded in 2000 by NASA, with apixel size of 14.25 m. The Landsat 7 ETM+ comprises the bands 7, 4and 2 and themosaic is projected in UTMcoordinates in theWorld Geo-detic SystemWGS 84. For the structural overviewmap (Fig. 2) only themain linear surface features were taken into account, like fjord valleys,coastlines and fracture zones. Straight continuous landforms are relatedto strike-slip faults (Glasser and Ghiglione, 2009). For the detailedmapsof the Seno Glaciar (Fig. 3A) and the Gajardo Channel (Fig. 1B) the sameprinciples were applied. Smaller faults, fractures or trendlines wereanalysed on the ice-polished bedrock. Ground-truthing of most faultswas done during field work in the Seno Glaciar area and in the Gajardoarea in 2001. We also included information, like fold axis and thrustfaults, from the geological map 1:1,000,000 (SERNAGEOMIN, 2003)and major strike-slip faults determined by Glasser and Ghiglione(2009). A total of 642 lineamentsweremapped and plotted in three dif-ferent rose diagrams (GEOrient; version 9.5.0).

4. Results

4.1. On-land morphology

Remote sensing data and field mapping show three sets of first-order fault zones (Fig. 2): (1) The sinistral strike-slip Magallanes FaultZone (MFZ) trending 136°. Northward of the Strait of Magellan thisfault fabric is less pronounced (Fig. 3); (2) A sinistral fault systemoriented NNW (165°) along the continental margin northward until atleast 50°S. In the study area it appears primarily to the east of the GCNand parallel to the Swett Channel. Thus, we name it the Swett ChannelFracture Zone (SCFZ; Fig. 3); (3) A NNE-trending fault system (44°)around the GCN, running parallel to the northern Gajardo Channel(Fig. 6). It is termed, therefore, the Gajardo Fault Zone (GFZ). These par-ticularly dextral strike-slip faults are associated to the MFZ (Maffioneet al., 2010).

4.2. Fjords to the west of the GCN (Swett Channel and adjacent bays)

The multibeam-based bathymetry indicates up to 350 m waterdepth in the Swett Channel and other parallel fjords further to theeast. Most subaquatic structures and canyons are related to the SCFZ(Fig. 3A). The geology of the Swett Channel region is characterised by al-ternating zones of granitoid rocks, mylonitic rocks and orthogneisses(Fig. 3A). The mylonites (Menichetti et al., 2008) crop out along twoNNW-trending zones parallel to the SCFZ direction, and their easy erod-ibility leads to the formation of canyons and fjords. The granitoids, resis-tant to erosion and denudation (Breuer et al., 2013), form a distinctNNW-trending ridge east of the SCFZ between the mylonitic zones.This ridge formed a morphological barrier capable of damming theGCN glaciers and hampering their westward flow. Only narrow pas-sages cross this morphological barrier. The bathymetric high at thesouthern end of the Swett Channel (Fig. 3A, B) separating it from upto 340 m deep basin (Quenca Luca) is part of the granitoid rock ridges.One of the major palaeo-glacier streams coming from the GlacierNoroeste migrated first southward through the Swett Channel, beforecrossing into the Luca Basin, andflowingwestward into the SenoGlaciarbasin (red arrows in Fig. 3B).

bathymetry. Morphologically analysed fault and trendlines, core locations and profiles arents. The rose diagram shows the orientation of faults and trendlines in this region. B) Falseer et al., 2007). The bathymetry is derived from a geostatistical interpolation (Ordinaryes indicate the former glacier flowdirections. The suspension load and transport direction

Table 114C and tephra ages for the piston cores LO-1 and BA-1.

Core depth Material Lab. ID 14C age (yr BP) Error (±yr) Calibrated age (yr BP) Sed.-rate (cm/kyr) References

LO-1760 Plant remain Poz-15817 870 40 730 40 This paper

BA-1157 Plant remain Poz-10379 1800 30 1660 50 This paper420 Mt. Burney Tephra 3860 50 4150 60 This paper

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In the south-eastern sector of the Bahia de los Glaciares (Fig. 3A)three glaciers enter this fjord area at present. Fig. 4A shows a cross sec-tion between the glacier tongues of Bahia de los Glaciares along the an-cient glacier flow until Seno Glaciar (A–A′; Fig. 3A). The profile crossesseveral well-preserved subaquatic moraine systems, faults, basementhighs and intervening small sediment basins (Fig. 4A). The glacierscould only erode two narrow and shallow passages into the above de-scribed granitoid ridge. The passages constricted the glacial streamsand thus hindered the westward flow during the Neoglacial advances,so that moraine formation occurred particularly in these narrow chan-nels (Fig. 3A;multibeam-derived bathymetrymap). In the echosounderprofiles themoraines show typical crag-and-tail structures with chaoticinternal reflectors (Fig. 4A). In small depressions on the top younger,stratified sediments with a draped channel fill pattern were deposited.A very pronounced left-lateral strike-slip fault (SCFZ) is characterisedby a deeper fjord incision.

Sedimentation in the Swett Bay, west of the GCN, is affected by theinflow of up to 500 m3/s fresh water from two large rivers. One ofthem drains Lago Muñoz Gamero with its 810 km2 catchment. Theother one is a meltwater river originating from the proglacial lake ofGlacier Noroeste (Fig. 3A). The latter river carries a high suspensionload into the fjord, causing south-westward hypopycnal flow withinthe low-salinity surface water layer over several kilometres along theSwett Channel (Fig. 3A, B) and into the adjacent bays (e.g. BahiaLobo). The wind-controlled distribution of the suspended load can beseen in several satellite images and aerial photographs (Fig. 3B), andwas also observed during our cruises at different seasons of the year.The on-land deltaic fan continues a few hundred metres into the fjord,as can be seen in the high-resolution bathymetry (Fig. 3A). Related tothe above described glacial clay plumes, relatively high sedimentationrates of 1.0 (BA-1 sediment core) to 1.2 mm/yr (SG-1 sediment core)were found in basins within the western GCN fjord system. Theechosounder data also show up to 30 m thick sediment deposits(Fig. 4B) in basins between subaquatic ridges and/or moraines, and inshallow coastal bays or platforms.

Fig. 4B shows an echosounder profile across the Swett Bay into theSwett Channel. Three seismic sequences can be differentiated. Thebasement is characterised by chaotic internal reflectors and shows apronounced morphology. The basin of Swett Bay is separated from abasin in the Swett Channel by a granitoid ridge. Both basins are filledwith N20 m of well stratified sediments, interpreted as ice distaldeposits. This indicates sedimentation under more dynamic conditions

Table 2210Pb ages for the core SG-1. The calibrated gamma lines (DL1a standard) of 210Pb, 214Pb,214Bi and 137Cs have been measured at the AWI-Bremerhaven using standard methods(Appleby and Oldfield, 1992).

Depth (cm) 210Pb age Sed-rate (cm/yr)

1.5 04.5 11 0.277.5 44 0.0910.5 81 0.0814.5 92 0.3617.5 104 0.2519.5 162 0.03

Average sedimentation rate: 0.12.

with current activity. The uppermost sequence has a draped channelfill pattern with subparallel internal reflectors documenting suspensionsedimentation.

Bahia Arévalo (Fig. 4C) is a shallow (5–6 m water depth) small baylocated 2.1 km NNW of the mouth of the river coming from GlacierNoroeste (Fig. 3A). It was investigated by echosounding and by drilling(core BA-1). The 8.8 m long piston core documented a continuousorganic matter bearing clayey sedimentation throughout the past5 kyr. The age constraints include the 4.15 kyr Mt. Burney tephra at4.2 m core depth and a calibrated 14C age of 1.7 kyr BP obtained frommacroplant remnants found at 1.6 m core depth (Table 1), indicatingrelatively constant sedimentation rates of 1 mm/yr. Below 5 m coredepth (approx. 5 kyr BP), a transition towards a coarser grained clasticfluvial sedimentation is recorded, probably documenting a less elevatedcoastline which enabled the formation of a shallow lake. This suggestscoastline uplift by at least 10 m after ~5 kyr BP (Kilian et al., 2011).The echosounder profile in Fig. 4C is next to the BA-1 core location.The reflectors of the uppermost sequence are even parallel and have adraped pattern. The lowermost lake facies with variable and partlycoarse sediment shows clinoformal reflector patterns (Fig. 4C), typicalfor prograding deltaic sediment systems (Stoker et al., 1997). At thetop of the lowermost deltaic sequence there is an erosional unconformi-ty. The upper deltaic facies II filled the depressions and levelled thesubaquatic morphology. The deltaic sequences have an Early to Mid-Holocene age. The tephra layer is conformal on the deltaic sequence,and is overlain by an organic-rich, marine sediment. At present a riverenters the bay and satellite images show a recent subaquatic delta.

Bahia Lobo is a larger fjord bay (3 × 0.8 km) 2.5 km south of theglacial river coming from the Glacier Noroeste (Fig. 3A). In the aerialimages parts of the hypopycnal flow that directly enter the bay can beseen. The whole 8.17 m long LO-1 piston core retrieved at 83 m waterdepth consists of silty clayswith low amounts of organicmaterial. A cal-ibrated radiocarbon age of 0.73 kyr BP was obtained from macroplantremnants at 761 cm core depth (Table 1) indicating very high sedimen-tation rates of 10 mm/yr during the past centuries.

4.2.1. Seno Glaciar, Seno Icy and Bahia BeaufortDuring the last Glacial Maximum there was an extended south-

westward ice stream from the GCN along the Swett Channel into SenoGlaciar, Bahia Beaufort and Seno Icy, where it merged with the largewestward-migrating ice stream of the western Strait of Magellannorth of Tamar Island (Figs. 2, 3A, B). The bathymetry of Seno Glaciardoes not showpronounced ridges. The fjord includes a N5 kmwide sed-iment basin with a water depth of 520 m, which was filled by N50 mthick sediments during the Late Glacial and Holocene. Sediment thick-ness is constrained by extrapolating seismic reflectors from the basinmargin towards its centre where a thick IRD layer hampers acousticpenetration. Two short sediment cores MD07-3126 and SG-1 couldalso not penetrate this IRD layer. The 0.48 m long QASQ box core(MD07-3126 obtained with RV Marion Dufresne in 2007) and the0.56 m long gravity core (SG-1 obtained with RV Gran Campo II) wereretrieved from the basin. Six 210Pb ages give an age frame for the last~0.55 kyr (Table 2). The base of the silty to clayey sediment sequencecontains ice-rafted debris. 210Pb ages for the past 75 years indicate aver-age sedimentation rates of 1.2 mm/yr and suggest that the massive IRD

Fig. 4.Bathymetric profile (A–A′) from theSenoGlaciar to the eastern BahiaGlaciares (line in Fig. 5) showing a high seafloor reliefwith occurringmoraine belts, basement highs, basins andfaults. The detail figure shows the echosound profile of the moraine belt with its typical reflection pattern and sediment distribution. The echosound profile (B–B′) from the Swett Bay tothe Swett Channel with a typical sediment succession of ice distal and organic-rich Holocene sediments. The echosound profile (C–C′) across the Bahia Arevalo next to the BA-1 sedimentcore showing two different deltaic facies separated by a hiatus and an organic-rich marine facies on top.

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layer was deposited during an early maximum of the Little Ice Agebetween ~0.6 and 0.5 kyr BP.

Further downstream, the ancient ice flow turned to SSW, running15 km in Bahia Beaufort and Seno Icy parallel to the GFZ (Fig. 1C). Thecrude bathymetry of Seno Icy indicates several ridges perpendicular tothe ancient ice flow, which is related to cross-cutting fractures of theMFZ (Figs. 2, 3). A detailed bathymetry was obtained across this up to600 m deep U-shaped fjord-valley (Figs. 1C, 5). In the north of thisbathymetrical map several subaquatic canyons are visible, parallel tothe MFZ and GFZ directions. They are more pronounced on the shallowfjord margins andmay have been formed by fluvial erosion after the re-gional ice retreat after 17 kyr BP (Kilian et al., 2007a), when the coast-line was much lower than at present, due to N100 m lower global sealevel. The bathymetric data also show subaquatic terraces in 20 to30 m water depth on both sides of the valley (Fig. 5). Their formation

requires relatively constant shoreline elevations during the Late Glacialand Early Holocene. This observation can be only explained, if the localglacial rebound was similar to the global sea-level rise for somemillennia, especially between 14 and 9 kyr BP. This is compatible withprevious regional coast levels estimated by Kilian et al. (2007a) andLamy et al. (2010). No clearly preserved subaquatic moraine remnantshave been detected bathymetrically along the ancient glacier streambe-tween Swett Channel and the Seno Icy area, indicating a relatively fastand continuous glacier retreat in this sector.

4.3. Gajardo Channel

The 55 km long Gajardo Channel (Figs. 1, 6) was formed by glacierstreamsoriginatingmainly fromtheGCNand frommountains onwesternRiesco Island (Kilian et al., 2007a). Along the ancient glacierflowdirection

Fig. 5. A) False colour Landsat ETM+ image from 2000 combined with high-resolution bathymetric map of the Bahia Beaufort next to the Strait of Magellan, with morphologicallyinterpreted fault directions. Submarine terraces occur in 20 to 30 m water depth which can be seen as well in the S–N profile (B) across the Bahia Beaufort.

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the NNE-trending northern section of the Gajardo Channel stretches over30 km from the Angostura de Témpanos at the GCN towards the SenoSkyring (Fig. 6).

Multibeam bathymetry and selected echosounder cross sections(Fig. 6) show a U-shaped glacial valley with a total relief of up to2000 m. The present day water depth is up to 680 m. On-landmappingin the north-eastern sector around the GCN, interpretation of aerialphotos, as well as false colour satellite images show NNE-trending ac-tive dextral strike-slip faults parallel to the GFZ direction. Faults parallelto theMFZ (NWdirection) have been also mapped in the field. The rosediagram shows a third minor E–W trending fault direction shaping theBahamondes area and Chandler Island. Parallel to this on-land faultzone several lineaments could be detected, which cross the northernGajardo Channel diagonally (Fig. 6).

The high-resolution bathymetry shows areas with irregular sub-aquatic ridges. Partly these are basement highs of less erodible rocks,which can be related to terrestrial features, and partly they are sub-aquatic moraines. The latter are seismically characterised by chaotic tocontorted reflectors (Kilian et al., 2007a). They were probably formedat around 14 kyr BP during Late Glacial glacier readvance Stage E(Fig. 6; McCulloch et al., 2005b; Kilian et al., 2007a). In front of thenorth-eastern entrance of the Gajardo Channel within the EustonChannel, an older subaquatic moraine system was likely being formedbetween 17.5 and 15 kyr BP during Stage D (Kilian et al., 2007a). Tothe north of the Angostura Témpanos and in the southward MFZ-oriented deeply incised fjord, moraines of the Little Ice Age have beenpreviously detected (Koch and Kilian, 2005; Kilian et al., 2007a). Thehigh-resolutionmultibeammap shows several curved terminalmoraines(Fig. 6), which are partly related to on-land moraine remnants.

The Gajardo Channel includes several basins (Fig. 6), which are sep-arated by subaquatic moraine systems and/or bedrock ridges. The mainHolocene sedimentation took place in these basins and in small depres-sions located on top of the morainal ridges. The deepest basin of the

fjord (680 mwater depth) is located ca. 6 kmsouthwest of the entranceof the Gajardo Channel. A 1.73 m long gravity core (CG-1) was taken inthe 250 m deep basin for which a fjord-parallel echosounder profile isshown in Fig. 6. It is characterised by parallel reflectors with a drapedchannel fill pattern. The 4.15 kyr Mt. Burney tephra layer in 5.3 mdepth in the core produces a distinct reflector within the Holocenesequence. The underlying ice distal sediments (Fig. 6) show poundedfill patterns with sub-parallel internal reflectors. The CG-1 core showsan overall clayey to silty sediment composition (Kilian et al., 2007a).The upper 14 cm of the core has been dated by the 210Pb-method.These ages indicate high sedimentation rates of 1.3 mm/yr.

A Late Holocene to recent debris fanwas detected in the bathymetricmap south of Chandler Island (Fig. 6). The associated steep (48°), hillslope on land shows clear evidence of a recent debris flow. The sub-aquatic deposits cover an area of 1.0 × 1.2 km and have a volume ofaround 80 million m3. Two major debris flow events are distinguished,indicated by the dashed lines in Fig. 6; B–B′. The most recent abrasionsurface is not visible in the Landsat TM 345 false colour image of 1986,but was observed in the field in 1999. It is possible that the mass flowis related to the M = 5.0 earthquake which occurred on 31.08.1996around 38 km to the SSW, or the 4.7 magnitude earth quake which oc-curred on 19.11.1988 around 32 km to the SWW (Fig. 1C).

4.4. Western Seno Skyring

At present, the former proglacial lake Seno Skyring (Figs. 1C, 2) rep-resents an estuarine fjord systemwith restricted connection to the Straitof Magellan through the Gajardo Channel in the west, and Seno Otway,via the Fitz Roy Channel in the east. The E–Woriented fjord has a lengthof around 80 km and is up to 15 km wide. In the west Seno Skyringpasses into the Euston Channel, where its deepest part of the basin(660 m water depth) is located. Further to the east, sediment basins

Fig. 6. Landsat ETM+ image combined high resolution bathymetric map of the Gajardo Channel. The two main fault and trendline directions are plotted in the rose diagram. Severalmoraine systems, basement highs, basins and one pronounced debris flow could be detected within the fjord bathymetry. The profiles (1–3, upper left corner) show the typical U-shapedform. The echosound profile (A–A′) shows the typical sediment succession of the isolated basins within the channel. The profile (B–B′) shows the suspected original seafloor, and thedashed lines indicating the assumed border between the two events.

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are developed in-between perpendicular oriented ridges, which belongto the Magallanes fold-and-thrust belt.

A bathymetric map of the Seno Skyring (Fig. 1C) was generated bygeostatistical interpolation from the echosounder profiles shown inFig. 7. In some areas, e.g. the Euston Channel, not enough data wereavailable for a good interpolation.

In general, the water depth decreases towards east. The 16 km2

sized Escarpada basin with amaximumwater depth of 440 m is locatedin the central sector of Seno Skyring. Echosounder profiles show amorethan 35 m thick sediment fill with parallel reflectors, which belong toHolocene and Late Glacial ice distal sequences. Escarpada Island has aU-shape (Fig. 1C), due to selective erosion of a shallowly northwardplunging fold of the Cretaceous Cerro Toro Formation (Romans et al.,

2011; KS1m in Fig. 7), which wasmore resistant to glacial erosion. Sed-iment core ES-1 was retrieved at 78 m water depth in the 2.5 kmwidebay within U-shaped coastline of Escarpada Island. The core shows the4.15 kyr BP Mt. Burney layer at 3.2 m sediment depth, indicating sedi-mentation rates of 0.78 mm/yr (Baeza, 2005).

Eastward of Escarpada Island the basin of the former Skyring glacierincised the Magellanic fold-and-thrust belt (Figs. 1, 2, 7), where UpperCretaceous and Lower Tertiary rock units of the former Magellan Basincrop out (SERNAGEOMIN, 2003). Erosionally resistant sedimentaryrock layers form N-S-trending ridges, which can be easily identified insatellite images, are subaquatically preserved. Some of them definesmall islands or peninsulas. The most important subaquatic ridgeextends from Altamirana Bay southward to Riesco Island across the

Fig. 7. False colour Landsat ETM+ image with the N-S trending Magallanes fold-and-thrust belt. The main fold axes, faults and rock units are taken from the geological map(SERNAGEOMIN, 2003; modified). The multibeam bathymetry shows the submarine continuation of a weathering resistant fold limb. The white dashed lines show the bathymetry ofthe Seno Skyring derived from interpolated echosound data. The corresponding echosound routes are indicated by the yellow dashed lines.

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whole Seno Skyring (Fig. 7). The ridge stands up to 210 mabove the gla-cially abraded basement, and subdivides Seno Skyring into two mainbasins.

Fig. 8 shows a coast-parallel echosounder E–W profile across SenoSkyring, including the above described subaquatic ridge. Four differentsediment sequences can be distinguished. A lowermost Sequence I ofice proximal deposits (5–10 m thickness) with subparallel internal re-flectors and partly chaotic reflectors documents coarse sedimentation(IRD), above the folded basement. Above this layer, a non-stratified Se-quence II occurs in isolated lenses often related to fault-like structures(Fig. 8). These may represent local subaquatic debris flow depositseither triggered by fault activity, or by reworking of IRD deposits ofthe Late Glacial (Kilian et al., 2007a). We prefer an interpretation ofthe wedge-shaped acoustically transparent bodies, lenticular in cross-section, as debris flows (Syvitski et al., 1987; Whittington and Niessen,1997; Gipp, 2003; Howe et al., 2003). The well-stratified Sequence IIIfurther up-section is interpreted as ice distal deposits (4–10 m thickness)above a basal unconformity. The uppermost sequence consists of wellstratified, organic-rich Holocene sediments (2–8 m thickness) and in-cludes the 4.15 kyr Mt. Burney tephra layer. The transition to the under-lying ice distal sediments is mostly concordant. The sediment thicknessof each sediment sequence reflects basement morphology, with highersediment thickness in basins. Faulting leads to block rotations of the base-ment (Fig. 8). The sediment cover is only partly faulted,with lowdisplace-ments. Most of these faults cause a sediment deformation and the minorfaults die out upwards. Analysis of all seismic profiles across Seno Skyringindicates that Holocene sediment thickness decreases towards the east.

The shallow (b80 mwater depth) eastern part of Seno Skyring is thearea of the ancient glacier tongue. Core Sky-1 was recovered in this

sector from 76 m water depth. Four tephra layers with well-knownages (Kilian et al., 2007a,b; Lamy et al., 2010) were identified. IRD layersin this core indicate ice retreat as early as 17 kyr BP. The sedimentationrates of 0.19 mm/yr throughout the Holocene and late Glacial are muchlower than further west.

Fig. 9 shows a representative 6.5 km long echosounder profile(A–A′) from the eastern part of the Seno Skyring. Local reflectors aredisturbed, indicating gas within the sediments that could be derivedfrom deeper sources (Fig. 9A). This layer blanking mainly affects thedeepest sequence of the ice proximal deposits. We suspect that thegas ascends along faults from the subjacent hydrocarbon-bearingCretaceous and Lower Tertiary rock units of the Magallanes Basin. Thelowermost ice proximal deposits have an uneven reflector signature;in gas-influenced areas the reflectors are disrupted. The top of the se-quence has a subdued relief, but the transition to the ice distal sedi-ments is concordant. In general, the ice distal facies is characterised byeven sub-parallel reflectors, in some areas (from 2.0 to 2.5 km profilelength) with a wavy parallel pattern. Occasional synsedimentary faultscause offsets in the range of some decimetres up to 1 m. This indicatesthat some of the faults were active during the Holocene. In a few casesan erosional unconformity separates the Holocene from the underlyingice distal sediments. One pronounced reflector within the Holocenerepresents the 4.15 kyr BP Mt. Burney tephra layer. In the SKY-E1 corethis tephra layer occurs in 1.18 m core depth with a thickness of22 cm. Within the seismic profile (Fig. 9A) the tephra layer appears ataround 1.5 m depth.

A gas-related feature appears at 5 km of the profile and extends to-wards the east from there. It is characterised by dome-shaped struc-tures, which deformed and uplifted the whole overlying sedimentary

Fig. 8. Echosound profile across theMagallanes fold-and-thrust beltwith the Late Glacial andHolocene sediment succession andmain faults cutting the sediments. Four different sedimenttypes are distinguished due to their seismic character. Within the Holocene sediments the 4.15 kyr Mt. Burney tephra occur as a distinct reflector.

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sequence (Fig. 9A). Westerly-induced currents in the upper 70 m of thewater column (Kilian et al., 2007b) eroded the uplifted sediments bybottom transport processes. These eroded sediments were depositedfurther east, above the normally deposited Holocene sediments with aclear unconformity (Fig. 9A). The reflectors of the redeposited sedi-ments have an oblique tangential pattern typical for prograding sys-tems. In contrary to this, in debris flows no internal reflectors can beseen (Fig. 9A). In the eastern part of the Seno Skyring the thickness ofthe Holocene sequence is on average 10 m in basin areas, and thins to2 m at morphological highs.

4.5. Seno Otway

SenoOtway (Fig. 9B) is an 87 km long and 30 kmwide estuarine bayand former proglacial lake, connected to the Magellan Strait in the westby the Jeronimo Channel. Most of the measured seismic data wereobtained in the north-eastern and central parts (Fig. 1).

Where the Magellanic fold-and-thrust belt intersects Seno Otway,pockmarks occur frequently. Strata below the pockmarks are stronglydisturbed, likely due to gas expulsion. Here we concentrate our descrip-tion on the easternmost sector of SenoOtway (Fig. 1) to the northeast ofthe major pockmark field. In the echosounder profile (Fig. 9) foursequences are evident. The oldest one is weakly reflective and chaotic,probably representing the bedrock or older glacial deposits (Glacial I;Fig. 9; B–B′). The transition to the ice proximal deposits is clearly visibleas well as the basin-like structure. In this sequence (till) the internal re-flectors are chaotic and contorted (DaSilva et al., 1997). Above, there is aconcordant transition intowell-stratified ice distal sediments. The inter-nal reflectors are even parallel with a draped pattern and indicatehomogeneous, fine-clastic sedimentation under fairly calm conditions.This Glacial II sequence is secondarily deformed by water-escape struc-tures, which equally affected the underlying Glacial I sediments. The

water-escape features occur in the profile as V-shaped structures(Fig. 9B), refilled with sediment. The water-escape features start withinthe ice proximal sequence (Fig. 9; B–B′), where coarser clastic sedi-ments occur. After Syvitski (1997) these events are typical for ice prox-imal areas. Water-escape events may have been triggered by periodicalearthquakes in this region (Mills, 1983), but themechanisms are not ob-vious by analysing themorphologies (Morretti and Ronchi, 2010). Rapidsedimentation can be also a triggermechanism (Postma, 1983; Syvitski,1997). It is obvious that thewater-escape structures only occur up to theunconformable transition to the uppermost sediment sequence (Fig. 9B).This erosional unconformity cuts the reflectors of the Glacial II sequence(at 5.5 km profile length). The overlying relatively thin (1.5 to 2 m thick-ness) Holocene sequence has sub-parallel reflectors and levels the surfacebetween the basin and the structural highs of the bedrock (at 3.3 kmprofile length).

5. Discussion

Cenozoic subduction and associated upper-plate deformation (Ramos,1999; Diraison et al., 2000; Polonia et al., 2007; Glasser and Ghiglione,2009) have strongly impacted on the basin and mountain topography,and the structure and erodibility of rocks in the southernmost Andes.Our investigation of fjord and on-land lineament orientation also in-dicate a predominant control of the erosion by a complex frameworkof at least three major fault zone systems, producing a complex fjordnetwork. In this framework we have attempted to identify other impor-tant mechanisms that have modified on-land morphology and bathyme-try of the fjord system. These include formation of lateral and/or terminalmoraines as well as debris flow deposits. Irregular morphological struc-tures like basement ridges, bathymetric highs, or narrow passages pro-foundly influence fjord currents and related sediment transport. In the

Fig. 9.The echosound profile (A–A′) shows the typical sediment units in the eastern Seno Skyring. Important features like gas intrusion and redeposition aremarked. Thedetailfigure givesa close-up of the internal seismic structure of the redeposition feature. The echosound profile (B–B′) from the eastern Seno Otway with the interpreted sediment facies and their typicalseismic character. Water escape structures occur between 4.0 and 4.5 km profile length. An erosional unconformity separates the ice distal deposits from the Holocene sequence.

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following chapters we discuss these aspects in the light of all the datapresented.

5.1. Tectonic control

A tectonic control on the fjord orientation has been previously pro-posed by Glasser and Ghiglione (2009), based on the investigation offjord orientation and fault zones without bathymetrical information.The detailed fjord bathymetry presented here, and the investigatedrelationship tomapped on-land tectonic lineaments, however, improvethe understanding of subaquatic erosion, sediment transport and massmovements.

The rose diagram in Fig. 2 shows the three main directions of theMFZ, SCFZ and GFZ fault systems. From the detailed maps of SwettChannel (Fig. 3A) and Gajardo Channel (Fig. 6) it becomes obviousthat the MFZ direction is present in each area, but the SCFZ directionis lacking in Gajardo Channel area and vice versa. In the Swett Channelarea zones of mylonitic rocks occur parallel to the SCFZ, probably origi-nating from strike-slip faulting in Late Cretaceous or Tertiary times(Menichetti et al., 2008). These present-day sinistral faults of SCFZ ori-entation may represent inverted or reactivated older faults.

Beside the SCFZ orientation no clear other preferred direction can beseen in the rose diagram of the Swett Channel (Fig. 3A). When taking acloser look at the satellite image (Fig. 3A) it is obvious that almost everypeninsula has its own main trendline direction. Whether this relates tofabric genesis of the different rock units (e.g. Patagonian Batholith,Tobifera Formation) separate in time and space, and a later juxtaposi-tion at terrane boundaries during the Tertiary, is not entirely clear. Animproved interpretation would require extensive radiometric datingand structural mapping. The tectonically controlled erosion features

are partly superimposed by a Quaternary detritus cover (Fig. 3A),often represented by subaerial moraines or deltas. In some areas, likearoundMt. Burney volcano (Fig. 1) the Holocene volcaniclastic depositsmask the glacially shaped fjord morphology. Such postglacial depositscomplicate the interpretation of the satellite image.

Fjord orientations are consistently linked to themain tectonic direc-tions, as described above. Rock lithology, however, has a strong local in-fluence on landscape and fjord morphology. The basement of the SwettChannel area is dominated by silicic magmatic rocks of the PatagonianBatholith and the Tobifera Formation (SERNAGEOMIN, 2003), whichare relatively resistant to glacial erosion. However, fracturing alongfault zones and tectonic lineaments greatly facilitates erosion, leadingto a fjord landscapewith steep slopes. The bathymetric map shows sev-eral subaquatic basement highs, which can be directly related to thesubaerialmorphology of continuing ridges along strike. These basementhighs and the subaquatic moraine systems subdivide the area intouneven-sized basins.

The Gajardo region is mainly affected by dextral strike-slip faults(Fig. 6) and some thrust faults, which belong to the Magallanes fold-and-thrust belt (Fig. 2). The uniform basement of the Gajardo Channelarea consists of allochthonous rocks of the Tobifera Formation, andmetamorphed sedimentary rocks of Devonian and Carboniferous age(SERNAGEOMIN, 2003). In contrast to the magmatic rocks of thePatagonian Batholith, metasedimentary rocks are generally more erod-ible (Fig. 1). Therefore the Gajardo Channel was formed as a uniformU-shape glacial valley with less subaquatic basement highs, but waslater bathymetrically modified by subaquatic moraine systems (Fig. 6).

Further to the east, the Seno Skyring is subdivided in two mainbasins by the N–S trending folds of the Magallanes fold-and-thrustbelt (Figs. 1C, 7). These steeply dipping and weathering resistant rocks

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form typical ridges in the subaerial landscape (Fig. 7), which continueinto the fjords. These structures and active faults may have played arole in creating very irregular coast lines (Figs. 8, 9A) and there is evi-dence that the Glacial and Holocene sediments in the fjord were affect-ed by very young faulting (Fig. 9A). The locations of gas expulsionstructures may also be controlled by faults and their activity.

The whole research area is a tectonically active region, documentedby recent earthquakes, with epicentres shown in Fig. 2. Along the SCFZnorthwest (Fig. 2) of our research area the epicentres of the three earth-quakes of magnitudes 5.3 to 6.6 are located. These events documentactive wrench faulting in this area. Other earthquakes of the last de-cades (Fig. 2) are also located close to the main fault zones, or withinthe foreland fold-and-thrust belt. Generally, the earthquakes in south-ernmost South America and Tierra del Fuego are characterised by shal-low hypocentres (b40 km) and by intermediate magnitudes (M =36), with higher magnitudes along MFZ and including the Lago Fagnanoarea (Lodolo et al., 2003;Waldmann et al., 2011). One of the recent earth-quakes in the GCN area may have triggered the large debris flow in theGajardo Channel.

5.2. Subaquatic moraines

In addition to the basement highs, subaquatic moraine systems areresponsible for the development of isolated basins and therefore forthe specific sediment distribution and deposition.

The cross section A–A′ through the Bahia de los Glaciares showsNeoglacial moraine ridges which were formed during the most extend-ed Holocene glacier advance at around 2.3 to 2.2 kyr (Arz et al., 2011).These moraines are preferentially located on top of basement highs(Fig. 4A), suggesting that fjord morphology is also controlled by the ex-tent of glacier advances. Post-neoglacial sedimentation is restricted tosmall depressions with a higher thickness (~11 m) or on top of themo-raines with a lower thickness (~3 m).

The stage Emoraines in theGajardoChannel (Fig. 6) are not locatedonbasement highs. Theywere formed on the nearly flat basement floor, likethe other Neoglacial moraine systems in the Gajardo Channel.Within theirregular moraine belts sedimentation is mainly restricted to small basinsand depressions in the moraine system. The Late Glacial and theNeoglacial moraines show the same sedimentation characteristics. Thedraped character of the Holocene sediments and the predominanthypopycnal sediment flow with following suspension sedimentationlead to a uniform sediment deposition. Seismic character of the morainalsystems is dominated by chaotic internal reflectors with a lower am-plitude and often characterised by strong sea bottom multiples(Carlson, 1989). Howe et al. (2003) described similar features from theKongsfjorden area, and Sexton et al. (1992) characterised moraines asstructureless seismic bodies.

Some small debris ormassflowswere detected at the steep slopes ofsubaquatic moraines. Following Ottesen et al. (2008), Carlson (1989) orCai et al. (1997) we interpret these as glacigenic debris flows. Thesemass movements affected the older ice distal sediments and partly theHolocene sediments. The older debris flows affecting the ice distal de-posits are embedded in the normal sediment succession. Slumping atmorainal ridges and at sediments deposited at basement highs andalong the steep fjordwalls is often caused by an unstable sediment con-figuration due to rapid sediment accumulation (Solheim and Pfirman,1985; Syvitski and Shaw, 1995; Stoker et al., 2010). The slumping pro-cesses modified the original shape and structure of sedimentation. Theslumps and subaqueous gravity flows (Figs. 6B, 8) may be generatedby earthquakes (Gipp, 2003; Stoker et al., 2010) or by expulsion of bio-genic gas (Fig. 9). The marine sediments are often unconsolidated, andwhen shaken by large earthquakes, they thought to move en masse(Powell and Molnia, 1989). High sedimentation rates create slopeover-steepening, and often lead to gravitational mass transfer (Powelland Molnia, 1989). In southernmost Chile the debris flows are likely

triggered by the same factors and the water-saturated soils enhancethe subaerial debris flows (Fig. 6B).

5.3. Sediment structures in fjord basins

In general, we distinguish twomain glacial units due to their seismicexpression: the ice proximal deposits and the ice distal deposits. Theheterogeneous composition of the ice proximal sediments often hasno clear internal seismic reflection pattern. Ice proximal sedimentsoccur in different thicknesses, but often the base of the units cannotbe clearly identified owing to the decreasing quality of the echosounderdata with depth. Above amainly concordant transition the ice distal de-posits succeed with a variable thickness. The ice distal sediments oftenfill depressions or small basins and smooth the submarine morphology(Figs. 6, 9B). In other cases, like the examples of the Skyring (Figs. 8, 9A),the ice distal sediments closely follow the seafloor morphology. Theirdraped character suggests sedimentation under more tranquil condi-tions, like suspension sedimentation. The seismic character of the icedistal sediments is given by continuously parallel to sub-parallel, medi-um to high-amplitude reflections (Cai et al., 1997; Stoker et al., 1997).

In most areas the transition to overlying organic-rich Holocene sed-iments is conformal, like e.g. Skyring (Figs. 8, 9A) and Gajardo Channel(Fig. 6). In the Seno Otway (Fig. 9B) a hiatus occurs between the ice dis-tal and the Holocene sediments. The hiatus contains an unknowntimespan and is associated with an erosional unconformity. Probablythe hiatusmarks the catastrophic outflow event of the former proglaciallake of the Seno Otway. The Holocene sequence in the entire researcharea is consistently characterised by even and parallel internal reflectorsand a draped pattern, which is typical for suspension sedimentation.

In the larger embayments such as Seno Skyring and Seno Otway, wefound a homogeneous and conformable sedimentation of both se-quences. The highest sediment thicknesses were found in large basin-like structures like e.g. in the 440 m deep Escarpada and in the 660 mdeep Euston basin within the Seno Skyring area (Fig. 1) and in the680 m deep basin of the north–north-eastern Gajardo Channel. Inthese depressions the reference layer of the 4.15 kyr BP Mt. Burneytephra is overlain by up to 8 m of sediments (Figs. 4B, 6, 8). The Holo-cene sequence thins out to the basin margins. Inmost of the smaller ba-sins and shallower coastal bays of the inner fjord region the 4.15 kyrMt.Burney tephra layer appears at 1–3 m sediment depth. In the narrowfjords and bays and in areas with high seafloor relief (basement highsor subaquatic moraines) the sediments are not spatially continuous.

5.4. Comparison with other fjord regions

The Chilean fjord system belongs to the world's southern fjord belt(Syvitski et al., 1987; Howe et al., 2010). The southernmost Chileanfjords described here are located in an area with a temperate climate.The annual mean temperature is about 5.7 °C at sea level (Schneideret al., 2007) and the annual precipitation rates are about 2000–8000 mm yr−1. The precipitation shows a positive correlationwith alti-tude in the GCN region and a negative correlation with the distancefrom the Pacific towards the steppe landscapes in the east. Despite therelatively mild climate, the glaciers in the GCN region reach the sealevel and form tidewater glaciers. The high precipitation rates lead tohigh glacier accumulation rates in the GCN region (Möller et al.,2007). In some cases the glacier retreat was stronger in the last years,so that the glaciers do not longer calve into the fjords and the fjordsare now under a glacio-fluvial regime. It is important to note, that thesalinity of water at the glacier terminus has no influence on the calvingrates of the glacier (e.g. Venteris, 1999).

The first order controls on glacimarine sedimentation named byPowell and Molnia (1989) are firstly tectonism, which creates accumu-lation areas for glaciers and provides sources for glacial debris. Fromtheir convergent tectonic setting the fjord regions of Alaska and Chileare correlatable. Tectonism at convergent margins causes mountain

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building and under favourable continental orientation to ocean watermasses and air masses, glaciations can be predominately accumulationcontrolled (Warren and Sugden, 1993). This setting allows glaciers inSE Alaska and in Chile at relatively low latitudes to extend down tosea level (Powell and Molnia, 1989). Furthermore, the tectonic move-ments have a strong influence on the amount of sediment production,and therefore for high sedimentation rates in the accommodation area(Hallet et al., 1996). For example, the tectonic uplift combined with lo-cally fractured and metasomatised rocks creates continuously renew-able sources for glacial erosion and debris production (Powell andMolnia, 1989). This is reflected by glacial erosion at the Bahia de losGlaciares and at the Swett Channel (Figs. 3, 4A). The investigated fjordarea is an excellent example how the tectonic setting controls the orien-tation of fjord systems and in particular changes in flow directionthrough time (Gipp, 2003). Powell and Molnia (1989) describe fromthe southeast Alaskan margin those active and dormant strike-slip faultsand their alignments separate tectonostratigraphic terranes of the mar-gin. Thus major valleys and fjords are aligned with these fault systemsand other tectonic elements such as thrust faults. Pfirman and Solheim(1989) have shown that the tidewater glaciers from Nordaustlandet,Svalbard, follow topographic depressions.

The second first order control mentioned by Powell and Molnia(1989) is the climate. It produces high rates of snow accumulationand/or high rainfall and glacier ablation, which cause high runoff andhigh sediment discharge. In southeast Alaska precipitation increaseswith altitude, like in the southernmost PatagonianAndes. High freshwa-ter discharge entering the fjord systems of southeast Alaska and south-ernmost Chile is due to a combination of high precipitation and glacialmeltwater (Powell and Molnia, 1989). Second-order controls are com-posed of the glaciers themselves and the marine-fluvial interaction(Powell and Molnia, 1989).

The fjords of southernmost Chile are dominated by siliciclastic depo-sitional systems, and thus comparable to the fjords of Alaska (PowellandMolnia, 1989). Sediment supply in the southernmost Chilean fjordsis carried by glacial meltwater, like in the fjords of Spitsbergen (Elverøiet al., 1980), of the north-western Barents Sea (Dowdeswell et al.,1998), and of southeast Alaska (Cowan, 1992). In contrast, in the fjordsof East Greenland and of the Antarctic Peninsula the high numbers oficebergs play a more significant role in transport and sedimentation(Dowdeswell et al., 1998). In temperate glacier setting like in Chileand in SE Alaska the precipitation is an important factor in glacimarinesedimentation (Cowan, 1992; Dowdeswell et al., 1998), while in moreriver-influenced fjords the circulation and sediment transport havestronger ties to the hydrological cycle (Syvitski et al., 1987).

Among the former processes fine-grained meltwater plume sedi-mentation is dominant in these temperate environments (Fig. 3B).Two types of sedimentation from meltwater plumes can be distin-guished: first, slow sedimentation from mainly fine-grained sedimentderived from the meltwater plume (Syvitski and Murray, 1981;Pfirman and Solheim, 1989); and, secondly faster deposition of thecoarse-grained bed load of the glacier streams forming ice proximaldeltas (Fig. 3A; Dowdeswell et al., 1998). The ice distal sediments inour research area are mainly characterised by fine grained input, likesilts and clay transported by glacialmeltwater, a fact that is documentedin different sediment cores taken in the research area (Kilian et al.,2007a,b). Such suspension-controlled deposits are reported from thefjord system of Svalbard during the Holocene (Elverøi et al., 1980;Dowdeswell et al., 1998). Beyond the major depocenters of tidewaterfronts and deltas, most sedimentation in the Chilean fjords is controlledby the hypopycnal flows. Particle settling from turbid plumes in upperwater layers is dependent on plume dispersal, turbulence, sedimentconcentrations and sizes (Syvitski et al., 1987). The suspension sedi-mentation is enhanced by flocculation, agglomeration and pelletisation(Syvitski and Murray, 1981; Pfirman and Solheim, 1989). These pro-cesses are dependent on the salinity of fjord water. With increasing sa-linity the settling of suspended material is faster, due to an increasing

flocculation and a decreasing velocity (Zajaczkowski, 2008; Hill et al.,2009). The high freshwater input derived from glacial meltwater andprecipitation may affect the fjord water salinity, which will affect theflocculation process and therefore the sedimentation rate (Cowan,1992). From the Svalbard archipelago Pfirman and Solheim (1989) re-ported, that the suspended material extends up to 15 km away fromthe glacier front. These values are comparable to the Seno Swett area,where the turbidity current can be traced for 12 km (Fig. 3B).

Sedimentation architecture is strongly influenced by the fjord bedmorphology (e.g. Elverøi et al., 1983; Dowdeswell et al., 1998). In the re-search area, where basins are separated by basement and morainalridges, like in the Gajardo Channel, Seno Glaciar and Seno Swett area,the main sedimentation occurs in deep troughs and basins, and only athin sedimentary cover appears on subaquatic highs (Figs. 4, 6, 8, 9;see also Syvitski and Shaw, 1995, or Sexton et al., 1992) If the depres-sions are partially filled with Holocene clayey silt, it is an indicatorthat they trap suspended sediment (Powell and Molnia, 1989). Seismicsections from the Gajardo Channel (A–A′ in Fig. 6) and from SenoSkyring (Figs. 8, 9B) document this type of sedimentation.

6. Conclusions

The on-land and subaquatic morphology of the investigated area inthe southernmost Andes is controlled by three main strike-slip faultzone directions. These are the WNW trending Magallanes Fault Zone,the Swett Channel fault zone (SCFZ), and the right-lateral GajardoChan-nel Fracture Zone, which is conjugated to the Magallanes Fault Zone(Maffione et al., 2010).

The basement highs consisting of erosion-resistant rocks partly con-trolled the extent of glacier advances, and thereby the formation of thesubaquatic moraine systems. The Late Glacial and Neoglacial morainesystems prevent a spatially continuous sedimentation, which is mainlyrestricted to isolated fjord basins. Due to the predominant suspensionsediment transport in the upper freshwater-dominated water columnwith wind-induced currents, the subaquatic geomorphology has no di-rect influence on the sediment deposition. However, fjord morphologymay have controlled locally high salinity bottom currents and theirmixing with low-salinity surface water. This could have enhanced floc-culation and settling of clay particles. Besides areas with stronger watercurrents, like in the Strait of Magellan or directly connected fjords,where sediments could be eroded, the basement highs form sedimenttrapswith increased sediment accumulation. Local sediment redistribu-tion processes like debris flows, water escape structures and gas expul-sion along faults are important and may have been triggered by seismicevents in the southernmost Patagonian Andes.

Acknowledgements

This study is funded by Grant Ki-456/11 of the German ResearchFoundation (Deutsche Forschungsgemeinschaft, DFG). Jana Friedrich isthanked for 210Pb dating at AWI Bremerhaven. We also thank MarceloArévalo for his encouragement during the different cruises since 2002,and the catchment survey. We are most grateful for sponsorship fromGran Campo Transportes Ltda. We thank the reviewers and the editorfor useful comments and suggestions, which sharpened our thinkingand helped to improve the paper considerably.

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