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Keywords:

  • Gulf of Papua;
  • submarine valley;
  • clinoform

Abstract

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[1] Much historical perspective about terrestrial and marine processes can be obtained from examination of fluvial valleys formed during lower sea level on surfaces that are now continental shelves. The continental shelf near the mouth of the Fly River (Gulf of Papua) has three incised valleys, which were not extensively modified or filled during the Holocene Transgression. Multibeam mapping of the valleys documents their morphology; seismic profiling reveals stratigraphy of sediment fill; and coring within and near the valleys suggests mechanisms of filling. Kiwai Valley is deep (20–50 m relief), narrow (∼1 km wide), steep-sided and meandering, due to river flows that caused its excavation through competent sedimentary deposits. Purutu Valley is shallow (10–20 m relief) and broad (>2 km wide). Umuda Valley is widest (∼14 km) and has multiple channels with variable incision depths, suggesting more extensive fluvial activity than the other two valleys. Valley filling has occurred in several ways, reflecting valley morphology and location relative to the present river mouth. Kiwai and Umuda valleys reveal three stages of infill: (1) hemipelagic sedimentation at distal sites, (2) gravity-driven flows spreading down the valley, and (3) subsequent clinoform progradation that completely fills the valley. Purutu Valley fill is dominated by clinoform progradation. Clinoform progradation in Umuda Valley is driven by intense sediment reworking on the surrounding topset regions, and its large width allows progradation from the sides as well as down its axis. Most shelf valleys around the world were filled long ago, and available techniques have severe limitations for documenting the details of morphology and the mechanisms of filling. The shelf valleys described in this paper provide a unique perspective to terrestrial and marine processes before, during, and after the Holocene Transgression.

1. Introduction

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[2] The MARGINS Source-to-Sink program links terrestrial and marine sedimentary environments in space and time, and an area of focused investigation is the Fly River entering the Gulf of Papua. During lowstands of sea level, rivers crossed coastal plains and formed valleys on surfaces that are now continental shelves. Examination of incised shelf valleys in the Gulf of Papua allows documentation of their morphology and provides insight to the transfer of fluvial loads from source to sink during lower stands of sea level. In addition, documentation of mechanisms filling shelf valleys helps interpret stratigraphy of infilled valleys there and elsewhere. The valleys investigated by this research provide unusual opportunities for understanding source-to-sink processes in the past and present.

[3] On continental shelves, an incised valley is an erosional feature larger than a single channel that is created by river action during a fall in base level [Zaitlin et al., 1994], such as eustatic sea level lowering or tectonic uplift. In addition, climatic change and stream capture can increase discharge and the additional flow can erode valleys. Deposits subsequently filling incised valleys can be important clastic reservoirs for hydrocarbons [Brown, 1993]. Incised valleys on continental shelves have been observed and analyzed intensively throughout the past two decades [Dalrymple et al., 1994]. Commonly, this has been done to examine the sedimentary fill as evidence of processes operating during sea level lowstand and transgression in shelf and shallow marine settings [Suter et al., 1987; Van Wagoner et al., 1990; Allen and Posamentier, 1993, 1994; Belknap et al., 1994; Thomas and Anderson, 1994; Zaitlin et al., 1994].

[4] In the study of shelf valleys, there is rarely an opportunity for detailed observation of the spatial heterogeneity associated with individual or groups of shelf valleys. Although acquisition and processing of high-resolution seismic data are improving [e.g., Lericolais et al., 1990, 1994], three-dimensional profiling is still difficult. Multibeam mapping, in contrast, produces a complete plan view image of the seafloor. Its application to the study of incised shelf valleys has been limited by the fact that fluvial and marine processes rework the coastline as sea level rises, and these processes can modify and fill incised valleys [Knebel et al., 1979; Davies et al., 1992; Davies and Austin, 1997; Nordfjord et al., 2005]. In the cases where incised valleys have been filled, seismic reflection profiling is the only option to investigate the morphology of the valleys.

[5] During three cruises spanning 2003–2004, three incised shelf valleys in the Gulf of Papua (GOP), Papua New Guinea, were imaged by multibeam mapping, examined by high-resolution seismic profiling, and sampled by coring. The maps provide a special opportunity to evaluate the morphology of valleys that have experienced little modification during and since sea level rise. In addition, profiling and coring within and near the valleys allow examination of the sedimentary processes filling them.

[6] Because the GOP valleys are in various states of burial, they provide a valuable opportunity to investigate their initial character and the evolution of the filling processes. The objectives of this study are to describe the morphology of incised valleys exposed on the continental shelf near the Fly River mouth, and to investigate the mechanisms and rates of filling.

2. Background

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

2.1. Tectonic Setting

[7] New Guinea is the largest island in the Indo-Pacific region and is part of the Australian continental plate. The tectonic history of eastern New Guinea has been dominated since the middle Miocene by the merger of its composite terrane with the Australia craton [Pigram and Davies, 1987]. The forces joining the older continental shelf strata and the northern mélange complexes caused uplift of the Finisterre, Owen Stanley, and New Guinea Highlands medial mountain ranges. As they rose, a foreland basin developed along their southern flanks and began to fill with clastic sediments eroded from the mountains [Hamilton, 1979]. This caused a transition in the GOP from carbonate substrates to those progressively dominated by siliciclastic sediments [Pigram et al., 1989; Tcherepanov et al., 2008].

[8] The foreland basin (i.e., GOP) continues to subside slowly today, with maximum rates of subsidence at the southwestern corner of the GOP reaching ∼1 mm/a [Mallarino et al., 2004]. However, subsidence rates decrease to the northeast approaching the axis of rotation, and may be as slow as 0.01 mm/a over much of the gulf [Slingerland et al., 2008a]. The slow rate of tectonic impact in the GOP allows eustatic curves to be good estimates for past sea levels during the Holocene.

[9] Unstable hillslopes and tectonic activity cause widespread landsliding in New Guinea [Dietrich et al., 1999; Densmore and Hovius, 2000]. Orographically induced rainfall, steep topography, and energetic earthquakes create conditions where large pulses of sediment can be shed to rivers [Ripper and Letz, 1993; Stevens et al., 1988; Dietrich et al., 1999], sometimes through the creation and destruction of natural dams [Costa and Schuster, 1988, 1991; Korup, 2002; Griffiths et al., 2004].

2.2. Oceanographic Setting

[10] A large drainage basin with extensive floodplains and a wet tropical environment results in a relatively constant hydrograph for the Fly River, and the maximum and minimum monthly water discharges normally differ by a factor of only 2–3 [Harris et al., 1993; Nittrouer et al., 1995; Dietrich et al., 1999]. This water enters the GOP, which is a semicircular basin with a width of ∼400 km. It is generally mesotidal (range 2–4 m, reaching a maximum of ∼5 m at the Fly delta apex during spring tides [Wolanski and Alongi, 1995]) and produces strong tidal currents on the inner shelf [Martin et al., 2008; Ogston et al., 2008]. Annual variability of physical conditions is dominated by fluctuations between monsoon (December-March) and trade wind (May-October) periods. During the monsoon period, winds blow from the northwest and waves are fetch limited, creating significant heights that are only ∼0.3 m. Waves generated by the trade winds enter the GOP from the southeast with a long fetch and significant heights ∼1.3 m [Thom and Wright, 1983]. Shear stresses produced by waves and tidal currents have been shown to remobilize sediment on the shallow topset and transport it seaward to the foreset region of the clinoform, as described in the next section 2.3 [Walsh et al., 2004; Ogston et al., 2008].

[11] The prevailing oceanic current in the GOP is the clockwise-flowing Coral Sea Current, whose axis is found beyond the shelf break [Wolanski et al., 1995]. Studies of sediment transport and accumulation [e.g., Walsh et al., 2004; Palinkas et al., 2006; Ogston et al., 2008] have indicated that, near the Fly River mouth, the dominant direction of along-shelf sediment dispersal is northeastward. This is part of a more complicated system for water circulation in the GOP [Wolanski et al., 1995; Ogston et al., 2008; Slingerland et al., 2008b], consisting of mesoscale shelf eddies forced by buoyant fluxes, winds, and seasonal variations in the Coral Sea Current.

[12] The modern continental shelf extends from the coast to the shelf break, which is located at a water depth of ∼125 m. The shelf is significantly wider in the southwest (∼150 km) than in the northeast (<25 km). Along the shelf break is an ancient shoreline formed during a lowstand of sea level [Droxler et al., 2006; Francis et al., 2008]. Today, sediment transport processes are unable to deliver much siliciclastic sediment to and beyond the shelf break in the southwest portion of the shelf [Harris et al., 1996; Carson et al., 2008; Muhammad et al., 2008]. During times of lower sea level, however, rivers crossed a broad coastal plain in the southwest region. This is demonstrated by numerous canyons that cut the surface of the outer shelf and upper slope [Francis et al., 2008], and by high fluxes of siliciclastic sediment that accumulated on the slope [Carson et al., 2008].

2.3. Sedimentary Setting

[13] The Fly River mouth is a tide-dominated delta [Galloway, 1975]. The distributary channels of the delta region have low sinuosity, and the delta plain exhibits a classic funnel-shaped morphology with three main distributaries diverging at the apex of the delta ∼100 km from the shoreline [Dalrymple et al., 2003]. Deltaic islands are shaped by tidal currents and have an elongate form, oriented parallel to the direction of tidal flow. The distributaries have a large ratio of width to depth, in contrast to river-dominated deltas in regions of low tidal range [Elliot, 1986].

[14] High relief (>4000 m) on New Guinea combines with wet tropical conditions and orographic effects to produce very high sediment yields [Milliman and Syvitski, 1992]. The Fly River is estimated to deliver ∼1.15 × 108 t/a, one third of the total sediment supply to the GOP (∼3.65 × 108 t/a [Milliman, 1995]).

[15] The large sediment supply and broad continental shelf for much of the GOP provide an environment conducive to the creation of a clinoform [Walsh et al., 2004; Slingerland et al., 2008a], which is a sigmoidal-shaped deposit that forms on many continental shelves receiving large inputs of sediment (e.g., Amazon [Nittrouer et al., 1986]; Ganges-Brahmaputra [Kuehl et al., 1989]; and Huanghe [Alexander et al., 1991]). Clinoforms consist of a gently dipping topset region close to shore, a steeper foreset region, and a nearly flat bottomset region in deeper water. Sediment accumulation is slow on the topset due to enhanced shear stresses created by waves and currents in shallow water, rapid on the foreset where shear stress decreases, and slow again on the bottomset due to reduced sediment supply. This pattern of accumulation allows the clinoform deposit to prograde into deeper water while maintaining a self-similar shape. Various sediment transport mechanisms have been suggested for creating clinoforms [e.g., Kostic et al., 2002; Schlager and Adams, 2001; Pirmez et al., 1998; Harris et al., 1993]. Gravity flows are likely an important process for maintaining the clinoform in the GOP [Walsh et al., 2004; Martin et al., 2008; Ogston et al., 2008; Slingerland et al., 2008a] and elsewhere [Kuehl et al., 1996], by transferring sediment from the topset to the foreset region.

[16] The GOP contains its modern clinoform and remnants of an older clinoform [Harris et al., 1996; Slingerland et al., 2008a]. The younger (modern) clinoform extends to ∼65 m water depth and downlaps onto the erosional surface etched into the topset of the older clinoform [Slingerland et al., 2008a]. The history of sea level change in the GOP has caused cycles of shelf erosion by subaerial rivers that created channels and valleys. Although the valleys discussed in this paper are only partially filled, other valleys in the central GOP are nearly or completely filled [Crockett et al., 2005; Slingerland et al., 2008a]. These are northeast of the Fly River mouth, and the largest filled valleys range in width from 10 to 15 km. These valleys are separated by “mesas,” and they are mostly filled by mud [Harris, 1994; Harris et al., 1996; Crockett et al., 2005; Slingerland et al., 2008a].

[17] Harris [1994] first noticed unfilled shelf valleys in close proximity to the Fly River. Because most shelf valleys in the GOP and elsewhere were filled during Holocene sea level rise, he considered explanations why the valleys remained unfilled (e.g., low sediment supply, strong currents). Subsequently, Harris et al. [2005] developed modified concepts related to coral growth and strong tidal currents, which were applied to valleys in far western areas (e.g., Torres Strait). Observations of modern processes filling the shelf valleys can provide further insight to these considerations.

3. Methods

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

3.1. Multibeam and Seismic Data

[18] A Simrad EM3000 multibeam sonar was mounted on a hydraulically stabilized pole extending through the hull of the R/V Melville, and was used to document the bathymetry of the study area (Figure 1). A single acoustic head generated 127 beams that covered a total swath of 130° at a frequency of 300 kHz. The ping rate for these surveys ranged from ∼5 to 15 kHz. Swath coverage was approximately four times the water depth. The system recorded data at all times, including transit and station work, which provided tie lines between areas of intensive mapping and information about precision and system stability. Navigation and orientation of the sonar head were provided by an Applanix Pos/mv 320 system. The data were processed (Caris software) to identify bad data points, and to correct for tidal variations using water elevations measured by instruments operating on the shelf seabed during the observation period [see Ogston et al., 2008; Martin et al., 2008]. The data were subsequently georeferenced to UTM coordinates and mapped in ARCGIS to facilitate analysis.

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Figure 1. Location of study areas with respect to Fly River Delta and Gulf of Papua. EM3000 multibeam bathymetry of three shelf valleys is overlain on bathymetric data compiled by Daniell [2008].

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[19] Subbottom profiles were collected using a ship-mounted Knudsen 320BR echo sounder operating at 3.5 kHz. Subbottom data, with penetration of 10–30 m, were collected contemporaneously with multibeam data.

3.2. Sedimentary Data

[20] To evaluate the filling of incised valleys, kasten cores (3 m long, square profile gravity cores [Kuehl et al., 1985]) were collected within the valleys and from the seabed outside the valley walls (Figure 1). Most cores are labeled by the transect number, and nominal water depth after the hyphen (for Kiwai Valley, “ch” refers to cores within its walls). Cores were X-rayed at sea to reveal internal sedimentary structure. They were subsampled every 2 cm to a depth of 50 cm in core, then every other 2 cm to the bottom of the core for sedimentary and radioisotope analyses.

3.3. Laboratory Analysis

[21] To measure grain size, the samples were wet sieved at 62 μm to separate the sand fraction from the mud (i.e., silt and clay). The sand was subsequently dried and weighed, and typically represented <5% of the sediment mass. The grain size of the silt and clay fractions was analyzed by a Sedigraph 5100, and grain size distributions were merged with the sand mass, in order to calculate percentages of the various components.

[22] The naturally occurring radioisotope 210Pb was used to calculate sediment accumulation rates [Nittrouer et al., 1979; Sommerfield et al., 2007]; 210Pb is in the 238U decay chain and has a half-life of 22.3 years. It is found dissolved in seawater, and “excess” 210Pb is permanently adsorbed onto sediment particles as they move through the water column. After deposition and during burial, the 210Pb activity decays toward a level supported by its effective parent 226Ra in the seabed. Profiles that reveal excess 210Pb activities decreasing logarithmically downward within the seabed usually reflect steady state sediment accumulation, and can be used to calculate an accumulation rate. Where nonsteady state accumulation creates a profile that does not decrease logarithmically, the presence of excess 210Pb can be used to estimate a minimum accumulation rate. Analytical techniques limit detection of excess 210Pb to approximately five half-lives (i.e., ∼100 years). Therefore any sediment containing excess 210Pb must have been deposited within the past century.

[23] The 210Pb activities were determined using a method modified from Nittrouer et al. [1979]. Sediment samples were dried and ground to expose surface areas. Approximately 5 g of sample were spiked with a known amount of 209Po to quantify the efficiency of the laboratory procedures. The 210Po, the granddaughter of 210Pb, was chemically released from the sediment by leaching in 16N HNO3 and 6N HCl. Both 209Po and 210Po were removed from the leachate onto silver planchets by spontaneous electrodeposition. Alpha decay detected from the planchets was measured over ∼24 h. From the measured activity of 210Po, the activity of 210Pb was calculated. Activities were normalized to the salt-corrected dry mass and expressed as dpm/g. All activities were decay-corrected to the time of collection.

[24] For some samples, total carbon (TC) was measured by high-temperature combustion on a Perkin Elmer 2400 Elemental Analyzer. After removal of total inorganic carbon (TIC), organic carbon (OC) was measured. TIC was then determined by the difference of TC and OC, and carbonate is presented as a weight percentage assuming all TIC was present as carbonate [Hedges and Stern, 1984].

4. Results

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

4.1. Morphology

[25] Three significant incised valleys were identified on the continental shelf near the mouth of the Fly River, and were found to be mostly unfilled. The region where the valleys are most pronounced is at the seaward edge of the modern clinoform (i.e., the base of foreset and the bottomset) extending seaward to a water depth of ∼70 m. Therefore, the shelf surface was imaged between water depths of ∼30 m and ∼70 m, and the valleys have relief below this surface of ∼10–50 m. Each valley was named for the closest deltaic island [Crockett et al., 2005; Nittrouer et al., 2003]. They are (from southwest to northeast) Kiwai, Purutu, and Umuda (see Figure 1). Subsequently, “Bramble Valley” was used as the name of a possible seaward extension of Kiwai valley [Harris et al., 2005].

4.1.1. Kiwai Valley

[26] The morphology of Kiwai Valley (Figure 2) shows little evidence of transgressive or posttransgressive modification by physical processes (e.g., waves, currents) or by sediment fill. The walls of Kiwai Valley are extremely steep (>45°) in some places and nearly vertical in others. The mapped area of Kiwai Valley encompasses four meanders numbered from north to south that join to create a valley 30 km long with a width that varies from 800 to 1100 m (Figure 2).

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Figure 2. Kiwai Valley is deep (20–50 m relief), narrow (800–1100 m wide), steep-sided, and meandering. Despite a large sediment supply from the Fly River, the valley has retained a well-defined morphology with little evidence of modification during or since sea level rise (modified from Crockett et al. [2005]).

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[27] Meander 1 is the most landward portion of Kiwai Valley and has been influenced by the prograding clinoform of the Fly River. Two tongues of sediment from the clinoform enter the valley from the west (Figure 3). In this region, the profile of Kiwai Valley appears to show a lateral channel bar [Heritage et al., 1999]; however, recent marine sediment creates the observed relief (Figure 4, profile 1). Intense modern modification of valley morphology by marine sedimentation is limited to the landward portion of the valley. Midway along meander 1 (lat 9°1.5'S), the influence of the encroaching clinoform fades, but seismic records show >20 m of sediment covering the western side of the valley with a narrow thalweg on the eastern side reaching the strong reflector that is the valley floor (Figure 5). The thalweg of the channel is very deep (almost 100 m, with ∼50 m relief) and hugs the eastern wall of the valley, where distinct ravines incise the cut bank. These ravines have an average spacing of ∼450 m and an incision depth of ∼10 m.

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Figure 3. (a) Kasten cores were collected both within Kiwai Valley and on the surrounding seabed. Evidence for subaerial modification can be seen in the ravine-riddled eastern flank of the valley wall (meander 1). A failure scarp and landslide deposit are located at ∼9°3′S (see expanded view). Marine sedimentation is filling the valley in three stages: 1, hemipelagic sedimentation; 2, gravity flows; and 3, clinoform progradation. The white line indicates the profile plotted in Figure 3b. (b) Bathymetry along the thalweg of Kiwai Valley, showing locations of core sites.

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Figure 4. Depth profiles across Kiwai Valley that show the presence of likely fluvial terraces, river bars and clinoform deposits. Maximum incision is present at the outside of each meander. The channel, where the maximum depth is reached, represents only a fraction of the total valley width.

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Figure 5. (a) Path of 3.5-kHz seismic record is shown as the black line across Kiwai Valley. (b) The U-shaped valley form is clearly visible as a strong reflector. About 20 m of acoustically transparent sediment fills the valley with a deep channel on the eastern side. The transparent fill could be a combination of fluvial sediment (deposited as a subaerial river channel), transgressive coastal deposits, and subsequent marine sediment.

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[28] Separating meander 1 from meander 2 is a region where the valley shallows by ∼20 m, from 80 m to 60 m water depth (shown in detail by the Figure 3 inset). The northern margin of this region has an arcuate scarp ∼10 m high. There are no surficial channels that cross between meanders 1 and 2 and the shoal appears to have eliminated communication between the segments.

[29] At the northwest margin of meander 2, a wide (650 m) tributary enters the main stem of the valley (Figure 2). This channel connects to a shallow (∼10 m maximum relief), narrow depression west of Kiwai Valley (Figures 2 and 5). The thalweg of meander 2 lies along the western wall of the valley, although it is not as well defined as the thalweg in meander 1, and the maximum depth reaches only ∼80 m (∼35 m relief). On the eastern side of the valley, detached elongate bars (likely lateral channel bars) can be observed midway through meander 2 (Figure 4). Some ravines have incised the western margin of the valley, but they are not as common or well defined as in meander 1.

[30] The bathymetry once again shallows between meander 2 and meander 3; however, there is a channel that allows the two regions to communicate (Figures 2 and 4). Meander 3 has a broad, clearly defined thalweg that hugs the eastern wall of the valley, and a wide, flat terrace is present on the western side of the valley. The maximum depth of the thalweg is ∼100 m, and meander 3 has relief of >45 m. The margins of meander 3 are not as linear as meanders 1 or 2 and exhibit large protrusions and recessions along both valley walls. Some marginal ravines are visible, but they are not as well defined or regular in spacing as in meander 1.

[31] Meander 4 is the least distinct of the valley regions mapped, with a maximum relief of only 20 m (Figure 2). The maximum water depth for this part of the valley is ∼70 m; the valley does not regularly deepen with increasing distance from shore. Even though the relief of meander 4 is not as distinct as other parts of Kiwai Valley, the thalweg is still apparent and is on the west side of the valley. The valley widens significantly and has a braided morphology with channels separated by topographic features resembling islands.

4.1.2. Purutu Valley

[32] Purutu Valley lies to the northeast of Kiwai Valley. It is broader (>2 km) and shallower (Figure 6) than Kiwai Valley. Purutu Valley exhibits no meanders within the region mapped, and has a relatively flat bottom with no distinct thalweg. Particularly significant is the overall relief of Purutu Valley (∼10–20 m), which is low compared to the other valleys.

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Figure 6. (a) Sedimentation within Purutu Valley is dominated by clinoform progradation, with maximum accumulation rates on the foreset of the deposit. (b) Accumulation rates are shown with bar sizes proportional to the rates.

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4.1.3. Umuda Valley

[33] Just seaward of the northernmost distributary mouth of the Fly River lies Umuda Valley, which is the broadest (∼14 km) of the three mapped near the Fly River. It exhibits complex bathymetry with elements of anastomosing channels, and a deep channel in the central part of the valley (Figure 7) with ∼20 m of relief. Seaward, areas both northeast and southwest of the main channel are characterized by bathymetric highs dissected by the anastomosing channels (some with >10 m of relief). The northeast boundary is a large shoal that separates the valley from the central portion of the GOP shelf. At the landward end, Umuda Valley is being filled on three sides by prograding sediment, and a three-dimensional clinoform morphology has developed (Figure 8, A-A'). Seaward of the clinoform, Umuda Valley displays the complex channel morphology described above, with sediment in the deep channels but not completely filling them (Figure 8, B-B') and indicating total channel relief >30 m.

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Figure 7. (a) Kasten cores collected within the thalweg of Umuda Valley have accumulation rates that maintain the sigmoidal shape of the clinoform as it progrades into the valley. Arrows indicate the directions of clinoform progradation in Umuda Valley as observed by seismic profiling (see Figure 8, A-A′). (b) Accumulation rates are shown with bar sizes proportional to the rates.

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Figure 8. Bathymetry in Umuda Valley is shown with paths of two 3.5-kHz seismic records indicated by black lines. In profile A-A', a 3.5-kHz seismic profile across Umuda Valley shows the progradation of clinoforms infilling the valley from the sides. In profile B-B', multiple river channels within Umuda Valley are partially filled with muddy sediment, which could have been deposited before, during, and after sea level rise. Modern sediment accumulation is shown in Figure 7. The ancient riverbed is a strong basal reflector in the seismic records.

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4.2. Sediment Accumulation

4.2.1. Kiwai Valley

[34] Kasten cores were collected from areas surrounding Kiwai Valley and within the thalweg (Figures 3 and 4), and these document modern sedimentation in the region. Outside the confines of the valley (cores T4-30, T4-50), 210Pb activity profiles show steady state accumulation (0.7 and 1.0 cm/a, respectively; Figure 9). Within the valley, core T4-73ch has a 210Pb activity profile that is not indicative of steady state accumulation (Figure 10). A thick sediment deposit is observed with high clay content and low 210Pb activity. The deposit contains two segments, each with sandy bases that abruptly grade into finer overlying sediment. The presence of excess 210Pb activity at the base of this core (170 cm) allows determination that the accumulation rate is ≫1.7 cm/a. Other cores farther seaward within the valley (e.g., T4-90ch, T4-70ch) exhibit steady state 210Pb accumulation rates (Figure 11) of 2.0 and 1.3 cm/a, respectively. A core collected from meander 3 (T4-95ch; Figure 12) exhibits slow accumulation (0.3 cm/a) of carbonate-rich (>50%) sediment. In contrast, carbonate content of sediment in meander 1 is <5%.

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Figure 9. The 210Pb activities for two cores collected near Kiwai Valley. Profiles show steady state sediment accumulation at these two sites, at the rates indicated.

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Figure 10. The 210Pb activity and clay percentage from core T4-73ch. The X radiograph shows two deposits, each with a sandy erosional base. The deposits visible in X radiographs correlate with a region of elevated clay content and reduced 210Pb activity. These characteristics together are indicative of high-concentration event sedimentation.

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Figure 11. The 210Pb accumulation rates along the thalweg of meander 1 in Kiwai Valley are steady state at >1 cm/a.

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Figure 12. The 210Pb profile in the distal regions of Kiwai Valley (meander 3; see Figure 5) is created by relatively slow hemipelagic sedimentation and bioturbation that causes a surface mixed layer.

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4.2.2. Purutu Valley

[35] As part of earlier work in the GOP, cores were collected near the head of Purutu Valley (Figure 6). Combining these data with the present study, we can create a longitudinal transect of sedimentation in this valley. Nearshore (i.e., clinoform topset) accumulation rates measured by Palinkas et al. [2006] are low, 0.6 cm/a at their site J21. Accumulation rates increase in deeper water, measured to be 1.7 cm/a by Walsh et al. [2004] at their site J44, and reach a maximum of 2.5 cm/a at our site T5.5-51. Accumulation rates decrease farther seaward, to 0.9 cm/a at site T5.5-69 on the bottomset of the clinoform that is prograding into Purutu Valley.

4.2.3. Umuda Valley

[36] Cores collected near and within Umuda Valley (Figure 7a) show two different 210Pb activity regimes. Cores T8-16 and T8-18 at the landward end of the valley have profiles that are representative of nonsteady state accumulation, with a thick layer (∼150 cm) of modern sediment overlying older strata (Figure 13). Core T8-20 also demonstrates a thick layer of higher activity. However, at this site, the surface layer is above sediment with excess 210Pb that decreases logarithmically, indicating steady state accumulation (Figure 14). In deeper water (Figure 7b), sediment is accumulating in a steady state manner, and the distribution of accumulation rates is consistent with typical clinoform models producing a sigmoidal shape (i.e., slow accumulation on the topset is 0.5 cm/a at T8-30, rapid accumulation on the foreset is 1.2 cm/a at T8-50, slow accumulation on the bottomset is 0.5 cm/a at T8-60). The breadth of Umuda Valley allows clinoform accumulation to fill the valley from the northeast and southwest also, as seen in 3.5-kHz seismic records (Figure 8, A-A').

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Figure 13. The 210Pb profiles from cores in shallow portions of Umuda Valley. They show a discontinuity in 210Pb activity that occurs at ∼150 cm depth. The surface sediment (down to ∼150 cm) is younger than 100 years, and the sediment below is older.

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Figure 14. The 210Pb activity at site T8-20 in Umuda Valley shows a uniform layer that was formed on top of sediment accumulating in a steady state fashion. The depth of the uniform layer (∼100 cm) and the lack of biogenic structures in the X radiograph indicate that bioturbation was unlikely to be responsible for the homogenization of the surface layer.

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5. Discussion

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

5.1. Valley Morphology and Its Burial

[37] Portions of incised valleys that remain mostly unfilled on the GOP shelf provide valuable opportunities to observe their morphology and the mechanisms for filling them. Each valley presents a different morphology: Kiwai Valley is narrow and deep, Purutu is broad and shallow, Umuda is broad and deep. These morphologies reflect diverse origins. Kiwai Valley incised deeply into the underlying substrate (likely, earlier clinoform sedimentary deposits [Slingerland et al., 2008a]), which is competent enough to maintain steep walls. However, Kiwai Valley was not active long enough to excavate a wide valley. Perhaps the tectonic conditions (earthquakes, landslides) in source areas of New Guinea caused a short-lived avulsion of the Fly River (e.g., creation and destruction of a natural dam, as described by Costa and Schuster [1988, 1991], Korup [2002], and Griffiths et al. [2004]). In contrast, Umuda Valley is much wider and it is likely the Fly River eroded channels that migrated for longer periods of time across its broad expanse. However, without further information (e.g., longer cores), the origins and early evolution of the valleys remain uncertain.

[38] The valleys almost certainly were formed some time when sea level was lower than present. However, it is difficult to say when that occurred. Harris et al. [2005] demonstrate for the GOP that during 60% of the past 120 ka, sea level has been 30 m to 70 m below its present level, which means that valleys could have formed and remained subaerial features for long periods of time (i.e., much longer than the Last Glacial Maximum).

[39] The partially filled valleys described here stand in contrast to a series of filled valleys in the central GOP (northeast of Umuda Valley) [Slingerland et al., 2008a]. Higher sediment accumulation rates [Walsh et al., 2004; Palinkas et al., 2006] are likely responsible for the filling of valleys in the central GOP. The filled valleys are observed in seismic images [Crockett et al., 2005; Slingerland et al., 2008a] but do not offer the same opportunity to document their morphology and to analyze the processes of filling.

[40] Clinoform progradation is the dominant mode of modern sedimentation on the GOP margin, and is burying the landward expression of each valley. Previous sediment budgets for the GOP [Walsh et al., 2004] have examined the clinoform, but the partially filled valleys are also sediment sinks. Their small spatial extent and their partially filled character imply that they are not major repositories for sediment. The 210Pb accumulation rates indicate that most areas of Umuda and Purutu valleys that are not being buried by the clinoform are accumulating sediment slowly (<1 cm/a). The deep portions of Kiwai Valley have higher accumulation rates, but the valley is narrow. Although not quantitatively important in shelf sediment budgets, the valleys provide conduits for modern sediment to spread seaward filling the shelf depressions. Spatially, valley filling shows a seaward progression of mechanisms. At a particular site within a valley, the same sequence is experienced: (1) slow steady state accumulation at distal valley sites, (2) rapid nonsteady state accumulation as the clinoform approaches, and (3) clinoform progradation that completely fills the valley. The following sections evaluate the diverse patterns of shelf valley evolution in the southwestern GOP.

5.2. Kiwai Valley

5.2.1. Morphologic Impacts of a Subaerial Environment

[41] At various locations in the valley, ravines incise the walls. Their presence probably indicates that Kiwai Valley existed for some time as a subaerial depression that continued to be modified by surficial processes. Modern analogies are common, such as marginal ravines along many stretches of the Colorado River that are tens of meters deep and occur with a spacing of hundreds of meters [Howard and Dolan, 1981], similar to those observed for Kiwai Valley.

[42] Across-shelf profiles of the incised valleys do not deepen regularly with distance from shore. Some of the variability in the downstream profile of Kiwai Valley (Figure 3b) is due to modification of the original feature (i.e., marine sedimentation in meander 1, landsliding between meanders 1 and 2). However, posttransgressive modification is not the only source of longitudinal variability. Incised valleys that are created during lowstand conditions may not always generate continuous across-shelf valleys [Van Wagoner, 1995; Lericolais et al., 2001]. A rapid drop in sea level that outpaces the erosional ability of the river can strand deep stretches of channel [Lericolais et al., 2001]. This may explain some of the deep channel thalwegs isolated in the Kiwai Valley profile.

[43] The common stratigraphic succession observed in the basal portions of incised shelf valleys is an erosional unconformity overlain by alluvial channel deposits [Dalrymple et al., 1994; Harris, 1994]. It is likely that a significant portion of the fill found in Kiwai Valley (Figure 5) is associated with the subaerial stage of the valley. This is especially likely for the apparent terrace and bar deposits found in seaward portions of the valley (e.g., meanders 2 and 3; see Figure 4). Long cores in Kiwai Valley could evaluate the distribution of subaerial and submarine fill.

5.2.2. Morphologic Impacts of Sea Level Rise

[44] As sea level rise causes shorelines to transgress coastal plains, incised valleys are commonly filled [see Dalrymple et al., 1994]. It is surprising that incised valleys near the mouth of the Fly River are only partially filled. The paradox of these unfilled valleys has been addressed by Harris [1994], who considered three mechanisms: strong currents preventing sediment accumulation; avulsion removing sediment supply before/during transgression; and burial followed by erosion. These are discussed in more detail below.

[45] 1. The Coral Sea current causes northeastward sediment transport near the mouth of the Fly River [Ogston et al., 2008]. During transgression, this may have led to removal of sediment to narrow northeastern regions of the shelf, where it could be lost to the continental slope [Harris et al., 1996]. In addition, numerical models predict the strongest tidal currents in the GOP when local sea level was 40–50 m below its present level [Harris et al., 2005]. These currents could have inhibited sediment accumulation during transgression.

[46] 2. Avulsion could have diverted sediment supply away from Kiwai Valley, reducing the potential for filling during transgression. The narrow morphology of the valley does not suggest that it was active for a long period, and it may have been disconnected from the sediment dispersal system during transgression.

[47] 3. Although strong tidal currents during transgression (mentioned above) were enhanced by the opening of the Torres Strait [Harris, 1994], it is unlikely that 20–50 m of sediment was deposited in and subsequently eroded from Kiwai Valley.

[48] Lack of sediment supply to Kiwai Valley, either due to strong currents or abandonment, is the likely reason it did not fill during transgression. Another consideration that might have helped reduce modification of Kiwai Valley during rise in sea level is rapid transgression. Many researchers have described periods of relatively rapid sea level rise interspersed with times of slower rise [e.g., Shepard, 1963; Fairbanks, 1990; Edwards et al., 1993; Bard et al., 1996; Harris, 1999]. It is interesting to note that the shelf depths incised by the extant valleys in the southwestern GOP (i.e., between ∼70 m and ∼40 m) correlate closely to depths that were rapidly transgressed due to meltwater pulse 1B, occurring ∼11–12.5 ka on many sea level curves [e.g., Duncan et al., 2000; Hanebuth et al., 2000; Liu et al., 2004]. These sea level curves could provide an explanation for rapid flooding and minimal modification of Kiwai and other shelf valleys.

5.2.3. Marine Sedimentation: Steady State

[49] Data collected for this project lend insight into the rates and mechanisms of sediment filling of shelf valleys under modern marine conditions. The rates generally decrease seaward in Kiwai Valley with increasing distance, from rapid nonsteady state accumulation (T4-73ch ≫1.7 cm/a) to slower steady state accumulation (T4-90ch = 2.0 cm/a, T4-70ch = 1.3 cm/a). The absence of event-dominated sedimentation downstream in Kiwai Valley probably reflects the dominance of other processes transporting and depositing sediment (e.g., hemipelagic) in distal regions (e.g., T4-90ch and T4-70ch), which is the first stage of valley filling.

[50] On the surrounding shelf, sediment accumulation rates at sites T4-30 and T4-50 (0.6 cm/a, 1.0 cm/a, respectively) are less than in the valley, and indicate a pattern of steady state accumulation that is consistent with models of clinoform progradation. However, the most active clinoforms (e.g., Amazon [Dukat and Kuehl, 1995]; Ganges-Brahmaputra [Kuehl et al., 1997]; the central GOP [Walsh et al., 2004]) are characterized by nonsteady state accumulation at higher rates (>2 cm/a).

[51] Farther seaward in Kiwai Valley, excess 210Pb is found near the surface of sediment cores, although accumulation rates are low (0.3 cm/a at site T4-95ch; Figure 12). In this region, the sediment is extremely coarse (>60% sand) and carbonate rich (>50%) compared to cores collected closer to shore within Kiwai Valley. This probably reflects the admixture of some new siliciclastic sediment from nearshore with relict carbonate sediment swept off topographic highs surrounding Kiwai Valley. With slower accumulation, benthic biological activity effectively stirs the sediment (as indicated by X radiographs) and creates a surface mixed layer ∼10 cm thick (Figure 12).

5.2.4. Marine Sedimentation: Gravity Flows

[52] The high-resolution bathymetric mapping of Kiwai Valley reveals that sediment enters from the northwest. Site T4-73ch lies on the flank of a sediment tongue, which extends into the valley from the clinoform foreset (Figure 3). The core exhibits a nonsteady state 210Pb profile (Figure 10). Sediment layers with high clay content and low 210Pb activity are deposited by high-concentration transport events [Walsh et al., 2004], and the sedimentary characteristics are attributed to limited interaction of sediment with the water column (and its dissolved 210Pb) during the event.

[53] The resulting deposit contains two segments that can be distinguished in the X radiograph (Figure 10), but which have similar grain size patterns and 210Pb activities. Each segment has an erosional base that is overlain by a coarser bed and grades abruptly into mud. These signatures of sedimentary deposits are diagnostic of rapid deposition by gravity-driven processes [Kuehl et al., 1988; Mullenbach and Nittrouer, 2000; Walsh et al., 2004]. In this case, the gravity flow had two pulses.

5.2.5. Holocene Sedimentary History

[54] On the basis of present shelf depths, it is likely that Kiwai Valley was transgressed during meltwater pulse 1B, which occurred between ∼12.5 and ∼11 ka. Therefore, Kiwai Valley has been subject to marine sedimentation in some form for at least 11 ka. A sediment accumulation rate >0.5 cm/a (∼50 m/11,000 years) would result in complete filling for the deepest parts of Kiwai Valley. Modern accumulation rates (100 year timescale) at sites within Kiwai Valley (e.g., T4-73ch, T4-90ch, T4-70ch) are greater than necessary to fill Kiwai Valley during its lifespan. Therefore, although sedimentation in parts of Kiwai Valley has been relatively constant over the last 100 years, that has not been the case for the past 11 ka.

[55] During the initial equilibration period after the rapid rise in sea level (∼5–7 ka), sedimentation would have differed from today as the long alluvial stretch of the Fly River adjusted [Lauer et al., 2008]. Sediment would not be exported to the shelf and would be trapped in an estuary near the present delta. As accommodation space filled, sediment would have started to escape and form the modern delta and shelf clinoform. Earlier research has suggested that a majority of sediment and water delivered by the Fly River enters the ocean via the southernmost distributary, proximal to Kiwai Valley [Wolanski and Eagle, 1991; Wolanski et al., 1997; Harris et al., 2004]. More recent research [Ogston et al., 2008], however, indicates that the majority of sediment is now discharged via the northernmost distributary of the Fly River delta and is transported northeastward. This sediment coalesces with sediment from the other rivers entering the central GOP, and the shelf valleys there are filled. In the southwest GOP, the valleys near the Fly River (Kiwai, Purutu, Umuda) have received sediment discharged by numerous distributaries and the relative importance of each distributary has likely changed with deltaic evolution. The combination of variable sediment discharge from distributaries and northeastward sediment transport near the mouth of the Fly has precluded complete filling of Kiwai and the other southwest valleys.

5.3. Purutu Valley

[56] Sedimentation within Purutu Valley is dominated by clinoform sedimentation at its landward end (Figure 6). Some variability can be observed in the shape of the clinoform where it enters the valley, suggesting that portions of the foreset may be prograding at different rates. Overall, the maximum accumulation rate on the clinoform near Purutu Valley (site T5.5-51, 2.5 cm/a) is high compared to accumulation on the clinoform near and within Kiwai Valley. The relief of Purutu Valley is small compared to Kiwai Valley (<20 m versus ∼50 m) and the slopes of the valley walls in Purutu Valley are gentle. Sedimentation may be more localized, with less loss down valley.

5.4. Umuda Valley

[57] Umuda Valley is the broadest of the three valleys mapped on the GOP shelf and has a complex morphology, which impacts its sedimentation. Kasten cores collected from the topset of the clinoform (i.e., sites T8-16, T8-18) do not exhibit steady state accumulation profiles (Figure 14), and they give insight into the mechanism of sediment transfer from topset to foreset in this valley region [also see Martin et al., 2008]. The 210Pb profiles at sites T8-16 and T8-18 demonstrate uniform values of excess activity above ∼150 cm, which abruptly decrease to supported activity below that level. The surface sediment is younger than 100 years and the sediment below 150 cm is an undetermined age >100 years. The uniform activity near the surface of the 210Pb profile is either due to nearly contemporaneous emplacement (on 100 year timescale) of the entire deposit, or physical homogenization of the deposit. Mixing by benthic organisms is unlikely, because biological structures such as burrows and other traces are absent and the biomass of mesobenthos and macrobenthos is extremely low [Aller and Aller, 2004].

[58] Farther seaward, core T8-20 also demonstrates a region of uniform activity at the surface, ∼100 cm thick. At this site, the homogenous deposit does not lie on top of sediment with background activities, rather it overlies a region that has logarithmic decrease of activity with depth (Figure 14), which indicates steady state accumulation at this site.

[59] Sediment reworking of 50–200 cm has been recognized in other energetic shelf locations (e.g., near the Amazon River [Kineke et al., 1996; Kuehl et al., 1996; Aller, 1998]). The effect of reworking is to homogenize the radioisotope profiles and recharge the chemical reductants. Aller et al. [2004, 2008] have identified a discontinuity in geochemical parameters of cores collected from GOP sites between the 10 m and 20 m isobaths. This discontinuity consistently appears at a depth of 50–100 cm in the cores, similar to the discontinuity for 210Pb activity observed in the topset cores collected near Umuda Valley (Figure 13). The correlation of geochemical and radioisotope data on the topset of Umuda Valley suggests that a physical process, either intense reworking by waves and tidal currents or rapid deposition has created the uniform surficial regions of radioisotope profiles.

[60] Along the length of Umuda Valley, the radioisotope profiles change from sites with excess activity lying unconformably on sediment with background activity to sites on the bottomset experiencing steady state accumulation. The transition from nonsteady state to steady state sedimentation is similar to Kiwai Valley, but the upper reaches of Umuda Valley probably reflect the greater intensity of sediment supply and reworking from the northernmost Fly distributary [see Martin et al., 2008].

6. Conclusions

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[61] The continental shelf near the Fly River mouth contains several incised valleys formed during lower stands of sea level that were not extensively modified by marine transgression and are now undergoing marine sedimentation. Multibeam mapping, seismic profiling, and core analysis provide the opportunity to contrast the valley morphology, and to examine modern sedimentary processes filling the valleys.

[62] The valleys vary in morphology from a narrow steep-sided incision (Kiwai Valley) to a broad valley with multiple channels (Umuda Valley). The proximity of the different morphologies indicates that the physical processes driving the formation of valleys can vary in a relatively small geographical area.

[63] Some filling of the valleys occurred during alluvial and transgressive phases. However, they remained only partially filled at the beginning of modern marine sedimentation. Three stages of infilling are recognized: (1) hemipelagic sedimentation at distal sites, (2) gravity-driven flows spreading down the valley, and (3) subsequent clinoform progradation that completely fills the valley. Kiwai and Umuda valleys are experiencing all three stages, and Purutu Valley is dominated by clinoform progradation.

Acknowledgments

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References

[64] We thank all U.S., PNG and Australian scientists who were involved with fieldwork, including the crew of the R/V Melville. J. P. Walsh helped with processing of 3.5-kHz profiles. We appreciate constructive comments supplied by Peter Harris, Greg Mountain, Jerry Dickens, and Bob Anderson. This paper is based on work supported by the National Science Foundation under grants OCE0203351 and OCE0504616, as part of the MARGINS Source-to-Sink Program.

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  3. 1. Introduction
  4. 2. Background
  5. 3. Methods
  6. 4. Results
  7. 5. Discussion
  8. 6. Conclusions
  9. Acknowledgments
  10. References
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