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LSU Doctoral Dissertations Graduate School
2007 Paleoceanography of the Gulf of Papua using multiple geophysical and micropaleontological proxies Lawrence A. Febo Louisiana State University and Agricultural and Mechanical College, [email protected]
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Recommended Citation Febo, Lawrence A., "Paleoceanography of the Gulf of Papua using multiple geophysical and micropaleontological proxies" (2007). LSU Doctoral Dissertations. 3357. https://digitalcommons.lsu.edu/gradschool_dissertations/3357
This Dissertation is brought to you for free and open access by the Graduate School at LSU Digital Commons. It has been accepted for inclusion in LSU Doctoral Dissertations by an authorized graduate school editor of LSU Digital Commons. For more information, please [email protected]. PALEOCEANOGRAPHY OF THE GULF OF PAPUA USING MULTIPLE GEOPHYSICAL AND MICROPALEONTOLOGICAL PROXIES
A Dissertation Submitted to the Graduate Faculty of the Louisiana State University and Agricultural and Mechanical College in partial fulfillment of the requirements for the degree of Doctor of Philosophy
in
The Department of Geology and Geophysics
by Lawrence A. Febo B.S., Ohio State University, 1999 M.S., Ohio State University, 2003 December 2007
For my late advisor, Dr. John Wrenn
ii ACKNOWLEDGEMENTS
What you are about to read, and I certainly hope enjoy, is the result of a
significant effort not just on my part, but also from so many people who I would like to
thank. First, I thank my late advisor Dr. John Wrenn. John was an excellent mentor to
me and instrumental in shaping who I have now become as a scientist and teacher. He
was also a very good friend, never ceasing to surprise me with his vast scientific knowledge, genuine willingness to help others, and his ability to brighten my days with a
casual swipe of his witty humor. Being his student was a rare privilege and it is he to
who I dedicate this volume.
After some recovery time, I was very fortunate to be adopted by Brooks Ellwood
and I have benefited greatly from his scientific expertise. In fact, many of the tools and
ideas brought together in this dissertation are a result of his guidance. I also extend many
thanks to Dr. Barun Sen Gupta for all the time he took discussing data and ideas with me.
The stable isotope data in this dissertation would not have been possible without the help
of Dr. Brian Fry, thank you for teaching me so much. I am also grateful to Sam Bentley
for inviting me to be a part of the MARGINS Source-to-Sink research initiative and also
a part of his team over the years. I came to LSU to do marine geology and
paleoceanography, and he delivered. Sam provided countless opportunities unobtainable
anywhere else, circumnavigating Australasia certainly ranking high on the list, but none
as memorable as crossing the equator, facing King Neptune’s court, and walking the
plank all in the same day!
Research results presented in this dissertation was partly supported by NSF grant
OCE #0305373 awarded to Samuel Bentley. Results presented in Chapter 4 were funded
iii by LSU and a U.S. NSF -Egypt Joint Science and Technology Board grant to Dr.
Ellwood and Dr. Aziz M. Kafafy, Grant #OTH6-008-002. I would also like to thank BP
America for providing generous financial assistance to complete my research and also
gainful future employment. I also thank the Geological Society of America, the
American Association of Stratigraphic Palynologists, and the Department of Geology and
Geophysics at Louisiana State University. The Department has supported me with
generous assistantships and scholarships, even after my term limit.
Additionally, I thank Bob Olson of Baseline Resolution for discussions about
TOC and Rock-Eval Pyrolysis data. I also received thoughtful comments from Andre
Droxler, Gerry Dickens, Chuck Nittrouer, and two anonymous reviewers on an early
version of chapter 2.
Finally, I would like to thank my parents, family, and close friends for understanding my hawkish desire to do the Ph.D. I especially thank my shipmates Jason
Francis, Luke Patterson and Zahid Muhammad, and my “paleo” officemates Rebecca
Tedford, Jeff Agnew, and Juan Chow. Thanks for listening to me squawk the last five years.
iv TABLE OF CONTENTS
ACKNOWLEDGEMENTS...... iii
ABSTRACT……………………………………………………………………………...vi
CHAPTER 1. INTRODUCTION…………………………………………………………1
CHAPTER 2. LATE PLEISTOCENE AND HOLOCENE SEDIMENTATION, ORGANIC-CARBON DELIVERY, AND PALEOCLIMATIC INFERENCES ON THE CONTINENTAL SLOPE OF THE NORTHERN PANDORA TROUGH, GULF OF PAPUA...... 9
CHAPTER 3. SURFACE SEDIMENTS AND DISPERSAL PATHWAYS ON THE OUTER SHELF-EDGE AND BASIN IN THE GULF OF PAPUA: MAGNETIC SUSCEPTIBILITY, CALCIUM CARBONATE, AND ORGANIC GEOCHEMICAL MEASUREMENTS...... 67
CHAPTER 4. DETRITAL CONTROLS ON MAGNETIC SUSCEPTIBILITY AND CYCLOSTRATIGRAPHY RECORDS …...... 100
CHAPTER 5. CONCLUSIONS...... 121
APPENDIX A: PUBLISHED AND SUBMITTED CHAPTER ABSTRACTS...... 124
APPENDIX B: SUPPLEMENTAL CHAPTER 2 DATA TABLES NOT IN TEXT….128
APPENDIX C: CORE IMAGES OF MV-51JPC………………………………………137
APPENDIX D: PERMISSION REQUEST…………………………………………….138
APPENDIX E: COMPREHENSIVE REFERENCE LIST...... 140
VITA...... 156
v ABSTRACT
Recent marine and late Pleistocene sediments examined from the Gulf of Papua
(GoP), Papua New Guinea investigate the flux and fate of detrital sediments and organic
carbon over the last glacial-interglacial cycle. Based on surface sediment magnetic
susceptibility (MS) and calcium carbonate concentrations, recent marine sediment is
exported off the narrow shelf and into deeper regions via the Kerema Canyon of the
northern Pandora Trough. Detrital clastic sediment is then dispersed deeper into the
central and southern Pandora Trough. Except for pelagic deposition, very little material
reaches the Ashmore Trough and Eastern Plateau adjacent to the Great Barrier Reef.
Rock-Eval pyrolysis data and organic petrography indicate that late Pleistocene and recent organic matter were strongly degraded prior to burial.
Late Pleistocene depositional records come from two, 12 m piston cores retrieved
from the slope of the northern Pandora Trough. MV-54 was taken at the mid-slope of the
central Pandora Trough (923 m) and MV-51 was collected from a bathymetric high in the
northeastern Pandora Trough (804 m). Core sediments were analyzed with MS, calcium
carbonate, organic geochemistry, and benthic foraminiferal assemblages. Both cores
show two periods of rapid sediment accumulation. High accumulation rates characterize
15,800-17,700 Cal. yrs B.P. in MV-54 and correspond to the early transgression when
rivers delivered sediments closer to the shelf-edge. Benthic foraminiferal assemblages in
MV-51 indicate a seasonally variable flux of organic carbon during late LGM (~18,400-
20,400 Cal. yrs B.P.), suggesting enhanced contrast between monsoon seasons.
The oldest section, >32,000 14C yrs B.P., contains the highest mass accumulation
rates and TOC fluxes, with >50% of the organic carbon derived from C3 vascular plant
vi matter. MS and benthic foraminiferal accumulation rates are orders of magnitude higher
during this interval than any younger time indicating a greater influence of detrital
minerals and labile organic carbon. Because mineralogy and detrital input are shown to
be the main controls on MS variability, the MS data in this interval suggest more direct
dispersal pathways from central and eastern PNG Rivers to the core site when sea level was lower and dispersal gradients were higher.
vii CHAPTER 1
INTRODUCTION
Continental margins are one of the ultimate storage sites for sediments shed from
continental settings. Warm, wet tropical environments such as the Amazon Basin, west
central Africa, and Indonesia supply significant quantities of sediment and terrestrial
organic carbon to marine depositional systems [Meybeck, 1988; Milliman, 1995;
Nittrouer et al., 1995; Milliman et al., 1999; Schlünz and Schneider, 2000; Syvitski et al.,
2005]. At present, a minimum of 50% of the annual sediment discharged into the global ocean originates from tropical drainage basins [Meybeck, 1988] and accounts for
approximately 60% of the total organic carbon budget [Schlünz and Schneider, 2000]. In
the tropical western Pacific, Indonesia and Australasia are estimated to be responsible for
as much as 25% of the global sediment flux [Milliman et al., 1999] and ~40% of the
terrestrial organic carbon buried in coastal and shelf deposits [Schlünz and Schneider,
2000].
In addition to the importance of global sediment and organic carbon budgets, the
tropical western Pacific region is the location of dynamic climatic and oceanographic
features, making it an area of critical paleoceanographic importance for understanding
long-term sediment storage and organic carbon cycling [Bjerknes, 1969; Meyers et al.,
1986; Beaufort et al., 2001; Clemens et al., 2003; Wang et al., 2003]. The western
Pacific warm pool, or WPWP, is a large region north of Papua New Guinea (Figure 1.1)
and contains warm water with sea surface temperatures >29°C and an atmospheric low
pressure zone (Figures 1.2, 1.3). Fluctuations in the position of the Inter Tropical
1 Convergence Zone (ITCZ) causes the position of the WPWP to move southward during
the austral summer monsoon (January – March) and deliver large quantities of rain to the
young, mountainous terrain of Papua New Guinea and other Indonesian islands (Figure
1.1). However, the position shifts north of the equator with the onset of the austral winter
(July – September) causing relatively drier conditions [An, 2000; Clemens et al., 2003].
Marine productivity, as inferred from chlorophyll a concentrations, does not appear to
change significantly between seasons with the exception of a slight decrease in the seas
between the Indonesian islands when warmer sea surface temperatures occur during the
summer monsoon (Figure 1.2).
Surprisingly, the paleoceanography of the western Pacific is just now beginning to
be investigated in greater detail despite its potential to remove atmospheric carbon over
glacial-interglacial timescales [e.g. Kawahata, 1999; Müller and Opdyke, 2000; Beaufort et al., 2001; Kilbourne and Quinn, 2004]. Because the western tropical Pacific is the location of very active depositional systems, lush tropical forests, and a prominent climatic system, it is an excellent location to examine glacial - interglacial changes in sediment flux and organic carbon burial. In particular, the Gulf of Papua (GoP) is a large
tropical mixed siliciclastic-carbonate depositional system that is located between
northeastern Australia and Papua New Guinea. The GoP is the recipient of between 300
and 400 x 106 tonnes per year (Mt y-1) in a basin one fourth the size of the Gulf of
Mexico [Milliman et al., 1999]. The contribution of sediment discharge from small, mountainous rivers for Papua New Guinea is shown in Figure 1.1 [see also Milliman et al., 1999].
2 In this dissertation, I address the issues of sediment accumulation rates, organic carbon burial, and inferred marine productivity over the past glacial-interglacial cycle
(LGM to recent). In chapter 2, I use two radiocarbon-dated piston cores to discuss changes in late Pleistocene to Holocene sediment accumulation rates, organic carbon burial, and the sources of the components preserved in the organic carbon pool. This chapter provides detailed information about how sedimentation rates and productivity responded with lower sea level and a cooler, drier climate system. Chapter 3 provides an examination of more recent sediment compositions and dispersal pathways in the Gulf of
Papua by examining surface sediments on the outer shelf-edge and basin in the Gulf of
Papua. Detrital and biogenic compositions are constrained by using magnetic susceptibility (MS), calcium carbonate concentrations, and organic geochemical measurements. Specifically, the quantity, type, and preservation of organic carbon in 68 core-top sediment samples are examined for the modern depositional system and contrasted with the late Pleistocene record.
Finally, in chapter 4, I examine the factors that control MS values and variations within marine cyclostratigraphy records. For example, MS is commonly used as a conservative tracer of terrigenous components in sediments [e.g. deMenocal et al., 1991;
Ellwood et al., 2000; Ellwood et al., 2006], but sometimes it is also argued to be a proxy for marine productivity [Meade and Tauxe, 1986; Warner and Domack, 2002]. Chapter 4 demonstrates that marine productivity alone cannot cause significant variability in MS records but instead that it is controlled by detrital fractions and mineralogy.
1.1 References
An, Z. (2000), The history and variability of the East Asian paleomonsoon climate, Quaternary Science Reviews, 19, 171-187.
3
Beaufort, L., T. de Garidel-Thoron, A.C. Mix, and N.G. Pisias (2001), ENSO-like forcing on oceanic primary production during the late Pleistocene, Science, 293, 2440-2444.
Figure 1.1. Map of Australasia and part of Indonesia showing major features of the western Pacific including seasonal weather patterns, surface circulation along the northeastern margin of Australia, and annual sediment flux directions and mass estimates for Papua New Guinea [map modified from Kershaw et al., 2003; flux estimates are in 106 tonnes per year and taken from Milliman et al., 1999]. The box denotes the Gulf of Papua study area, which is expanded in the lower right corner showing major rivers, bathymetry, and piston core sites. S.E.C. refers to the South Equatorial Current, E.A.C. is the East Australian Current, and C.S.C.C. is the Coral Sea Coastal Current. The position of the Inter Tropical Convergence Zone is shown for the austral summer (Jan. – Mar.) when it shifts below the equator. During austral winter (July - Sept.) it is positioned above the equator and out of the map area.
4
Figure 1.2. Summer Monsoon conditions. Sea surface temperatures, salinities, and chlorophyll a concentrations for part of the Indonesian Archipelago and Australasia. Temperature and salinity data are from World Ocean Atlas 2005 [Locarnini et al., 2005] and made with Ocean Data View software [Schlitzer, 2007].
5
Figure 1.3. Winter Monsoon conditions. Sea surface temperatures, salinities, and chlorophyll a concentrations for part of the Indonesian Archipelago and Australasia. Temperature and salinity data are from World Ocean Atlas 2005 [Locarnini et al., 2005] and made with Ocean Data View software [Schlitzer, 2007].
6
Bjerknes, J. (1969), Atmospheric teleconnections from the equatorial Pacific, Monthly Weather Review, 97, 163-172.
Clemens, S., P. Wang, and W. Prell (2003), Monsoons and global linkages on Milankovitch and sub-Milankovitch time scales, Marine Geology, 201, 1-3.
deMenocal, P., J. Bloemendal, and J. King (1991), A rock-magnetic record of monsoonal dust deposit ionto the Arabian Sea: Evidence for a shift in the mode of deposition at 2.4 Ma, in Proceedings ODP, Scientific Results 117, edited by W. Prell and L. Niitsuma, pp. 389-407.
Ellwood, B.B., W.L. Balsam, and H.H. Roberts (2006), Gulf of Mexico sediment sources and sediment transport trends from Magnetic Susceptibility measurements of surface samples, Marine Geology, 230, 237-248.
Ellwood, B.B., R.E. Crick, A. Hassani, S.L. Benoist, and R.H. Young (2000), Magnetosusceptibility event and cyclostratigraphy method applied to marine rocks: detrital input versus carbonate productivity, Geology, 28(12), 1135-1138.
Kawahata, H. (1999), Fluctuations in the ocean environment within the western Pacific warm pool during late Pleistocene, Paleoceanography, 14, 639-652.
Kilbourne, K.H. and T. M. Quinn (2004), A fossil coral perspective on western tropical Pacific climate ~350 ka, Paleoceanography, 19, PA1019, doi:10.1029/2003PA000944.
Locarnini, R.A., A.V. Mishonov, J.I. Antonov, T.P. Boyer, and H.E. Garcia (2005), World Ocean Atlas 2005, http://odv.awi.de/data/ocean/world-ocean-atlas-2005.html
Mead, G.A. and L. Tauxe (1986), Oligocene paleoceanography of the South Atlantic: Paleoclimatic implications of sediment accumulation rates and magnetic susceptibility measurements, Paleoceanography, 1(3), 273-284.
Meybeck, M. (1988), How to establish and use world budgets of riverine materials, in Physical and Chemical Weathering in Geochemical Cycles, edited by A. Lerman and M. Meybeck, The Hague, Kluwer Academic, pp. 242-272.
Milliman, J.D. (1995), Sediment discharge to the ocean from small mountainous rivers: the New Guinea example, Geo-Marine Letters, 15, 127-133.
Milliman, J.D. , K.L. Farnswerth, and C.S. Albertin (1999), Flux and fate of fluvial sediments leaving large islands in the East Indies, Journal of Sea Research, 41, 97-107.
Müller, A. and B.N. Opdyke (2000), Glacial-interglacial changes in nutrient utilization and paleoproductivity in the Indonesian Throughflow sensitive Timor Trough, eastern Indian Ocean, Paleoceanography, 15, 85-94.
7
Nittrouer, C.A., G.J. Brunskill, and A.G. Figueiredo (1995), Importance of tropical coastal environments, Geo-Marine Letters, 15, 121-126.
Schlitzer, R. (2007), Ocean Data View, http://odv.awi.de.
Schlünz, B. and R.R. Schneider (2000), Transport of terrestrial organic carbon to the oceans by rivers: Re-estimating flux- and burial rates, International Journal of Earth Sciences, 88, 599-606.
Syvitski, J.P.M., C.J. Vorosmarty, A.J. Kettner, and P. Green (2005), Impact of humans on the flux of terrestrial sediment to the global coastal ocean, Science, 308, 376-380.
Wang, B., S.C. Clemens, and P. Liu (2003), Contrasting the Indian and East Asian monsoons: implications on geologic timescales, Marine Geology, 201, 5-21.
Warner, N.R., and E.W. Domack,( 2002), Millennial- to decadal-scale paleoenvironmental change during the Holocene in the Palmer Deep, Antarctica, as recorded by particle size analysis, Paleoceanography, 17, Art. No. 8004, doi:10.1029/2000PA000602.
8 CHAPTER 2 LATE PLEISTOCENE AND HOLOCENE SEDIMENTATION, ORGANIC- CARBON DELIVERY, AND PALEOCLIMATIC INFERENCES ON THE CONTINENTAL SLOPE OF THE NORTHERN PANDORA TROUGH, GULF OF PAPUA*
2.1. Introduction
Rivers deliver large amounts of organic carbon (~0.4x1015 g C yr-1) from land to
continental margins, with approximately 30% of it originating from tropical rainforests alone [Schlesinger and Melack, 1981; Hedges et al., 1997]. Understanding the ultimate fate of this important carbon pool is crucial not only for understanding climatic changes, but also for accurately modeling the broader global carbon cycle [Schlünz and Schneider,
2000; Berner, 2003].
Within this context, the Gulf of Papua (GoP) (Figure 2.1) is a particularly
intriguing region to examine the flux and fate of terrestrial organic carbon over glacial to
interglacial timescales. The GoP is a marine depocenter for large and very active river
basins. In the GoP, three large rivers, the Fly, Kikori and Purari, currently deliver an
estimated 300-400 megatonnes yr-1 (Mt) of terrigenous siliciclastic material (e.g., clay, quartz, feldspars) from a young, mountainous and tropical region to the inner shelf each
year [Milliman, 1995]. This is an annual flux greater than that of the Mississippi River,
and comes from drainage basins with combined areas ~3% the size of the Mississippi
drainage basin [Milliman and Syvitski, 1992; Milliman, 1995]. With this immense siliciclastic load comes approximately 4 Mt yr-1 of terrestrial particulate organic carbon
[Bird et al., 1995]. Today, most of the fluvial discharge accumulates on the inner shelf of
______
* Submitted and accepted for publication in Journal of Geophysical Research – Earth Surface, Reproduced by permission of American Geophysical Union, see also Appendix D.
9 the GoP and in an extensive mangrove-covered deltaic system along the northern coast
(Figure 2.1); very little of the siliciclastic material and terrestrial organic carbon escapes
the shelf to deeper water [Bird et al., 1995; Walsh and Nittrouer, 2003].
The modern shelf edge of the GoP lies about 125 m below sea level (Figure 2.1).
However, during late Quaternary sea-level lowstands, water depths were ~60 to 125 m below those at present. Consequently, river mouths were at or near the shelf edge, and
large amounts of siliciclastic material and terrestrial organic carbon presumably
accumulated in deep water settings, as is known for other large river systems [e.g., Flood
and Piper, 1997, and references therein]. However, there is no information to evaluate whether massive terrestrial input indeed occurred in GoP slope and basin environments during glacial times, and, if so, what was the fate of that material with respect to burial, transformation, or remineralization.
This study aims to evaluate sediment accumulation in the northeastern GoP since
the last glacially induced sea-level lowstand. Specifically, I investigated: 1) the sediment
accumulation rates at two representative locations on the middle slope, 2) the late
Pleistocene to Holocene organic carbon contents and fluxes at these locations, and 3) the
proportions of terrestrial and marine organic carbon.
2.2. Gulf of Papua
The GoP is a tropical mixed siliciclastic-carbonate sedimentary system that is
presently affected by a warm, wet monsoon-dominated tropical climate and a young,
mountainous continental terrain. The dry season occurs during April to November when
SE trade winds dominate, setting up moderate wave fields (~1.3 m significant wave
heights) and invigorating the clockwise Coral Sea Coastal Current (Figure 2.1). In
10 contrast, the wet season occurs during December to March and is characterized by milder
NW monsoonal winds that bring significant amounts of rain and lower (~0.3 m)
significant wave heights [Thom and Wright, 1983].
Figure 2.1. General bathymetric map of the Gulf of Papua study area. Bathymetric contour interval is 100 m down to the 500-m contour line, and then every 500 m in deeper water. The 60-m contour line on the middle shelf indicates the extent of the modern shelf clinoform. The dashed arrow indicates the general direction of sediment movement from the Fly River. The bathymetric profiles A and B are plotted below the map and show the gradient from the coast to each core station - note the difference in horizontal scale. Bathymetric data set is from Daniell [2007].
11 The Papua New Guinea highlands receive 10-13 m of annual rainfall that ultimately drains into the GoP [Wolanski et al., 1984]. Coarse sediments discharged by
the Fly, Kikori, and Purari Rivers initially accumulate in deltaic sand bodies, whereas
suspended sediments are mostly advected clockwise to the northeast by geostrophic flow
[Wolanski et al., 1995; Wolanski and Alongi, 1995] and stored along the inner shelf in
waters generally shallower than 60 m [Bird et al., 1995; Brunskill et al., 1995; Harris
1996; Walsh et al., 2004; Keen et al., 2006]. Further remobilization may occur during the
wave-dominated trade-wind season [Hemer et al., 2004].
2.2.1. Pandora Trough
The Pandora Trough is a broad basin beyond the shelf edge covering >8000 km2
and is thought to receive a limited modern terrigenous sediment flux [Milliman et al.,
1999; Walsh and Nittrouer, 2003]. The Pandora Trough is rimmed by the GoP shelf that changes from broad and low-gradient (1:1000) in the northwest to a narrow and higher gradient (1:133) in the northeast (Figure 2.1). Sediments escape into the northern
Pandora Trough by nepheloid-layer transport down slope [Walsh and Nittrouer, 2003] and episodic turbidity currents in some locations [Bentley et al., 2006]. An estimated 2-
3% of the sediments entering the GoP each year may be transported off the shelf and deposited in the Pandora Trough [Walsh and Nittrouer, 2003; Muhammad et al., 2007].
2.3. Materials and Methods
2.3.1. PANASH Cruise and Sediment Cores
The Pandora and Ashmore Troughs are two large sediment sinks that ultimately
store sediments shed from the PNG highlands and coastal plain. The PANASH cruise,
which took place from March to April 2004, was part of the NSF MARGINS Source-to-
12 Sink Program to sample and quantitatively study this important sedimentary system.
During the 2004 field season aboard the R/V Melville, we collected a total of 30 multi- cores and 33 jumbo piston cores.
Piston core MV-54 was collected from the slope seaward of the broad shelf (~150 km wide) bordering the central Pandora Trough; piston core MV-51 was collected from the narrow shelf (~20 km wide) along the northern Pandora Trough (Figure 2.1). We selected cores on the basis of achieving comparable geologic setting, core length, and water depths to compare and contrast sediment accumulation in different regions of the
Pandora Trough. Cores MV-54 and MV-51 are both ~12 m in length and were obtained from 923 and 804 meters water depth, respectively, in open slope, non-channelized locations that are not likely to be inundated directly by turbidity currents (based on our multibeam surveys) [Francis et al., 2007] (Figure 2.1 and Table 2.1). MV-54 is located on the central Pandora mid-slope region (Figure 2.1) just NE of slump scars and between unfilled, relict channels as noted in the work of Francis et al., [2007, Figure 8]. MV-51 was collected on top of a NW-SE trending tectonic ridge that forms a local topographic high in the mid-slope [Francis et al., 2007, Figure 9]. The numerous tectonic ridges in the northeastern Pandora Trough slope also have troughs between them that serve to trap or divert sediments moving downslope. The ridges are probably sufficiently elevated above troughs to avoid direct inundation by gravity flows (~100 m) (Figure 1, bathymetric profile B), yet still receive suspended sediment transported off-shelf or down-slope. Consequently, MV-51 potentially holds a high-quality, high-resolution time-stratigraphic record.
13 Interpretations presented in this chapter apply specifically to the core locations
under investigation. However, the geologic settings of these cores suggest that the cores
should contain records of regional processes 1) because of the placement of MV-54 away
from surrounding channels and slump scars, 2) because MV-51 is positioned on top of a
tectonic ridge, and 3) because both are relatively close to river-influenced shelf edges.
Table 2.1. Geographic locations for sediment cores examined in this study.
Sediment Core Latitude Longitude Water Depth (m) Length (m) MV25-0403-54 8.936° S 145.389° E 923 12.16 MV 25-0403-51 8.767° S 146.018° E 804 12.28
2.3.2. Shipboard Analyses
After core retrieval, shipboard measurements were made on piston cores for bulk
density and magnetic susceptibility using a GeoTek Multi Sensor Core Logger (MSCL).
Bulk density was calculated from the attenuation of gamma rays emitted from a 10 milli-
Curie 137Cs source through sediment cores. GeoTek MSCL-derived bulk density units are g cm-3 and referred to as gamma-ray density (GRD) to separate them from traditional
wet and dry bulk density measurements [Weber et al., 1997]. Magnetic susceptibility
(MS) was measured using a Bartington MS2C Loop Sensor attached to the MSCL. MS
values were corrected with respect to volume and units are reported in 10-5 SI [Thompson
and Oldfield, 1986]. Both MS and GRD were measured on the MSCL at 1-cm intervals.
• X-radiographic Methods
Selected sections of MV-51 were imaged via X-radiography. Core subsections
were sliced into 2-cm-thick axial slabs, and imaged onboard using a Thales Flashscan 35
digital X-ray detector panel, illuminated with a Medison Acoma PX15HF X-ray
14 generator. Images were archived as 14-bit grayscale images with 127-micron pixel resolution.
2.3.3. Radiocarbon Analyses
Radiocarbon dates are based primarily on woody organic matter because of insufficient absolute concentrations of planktonic foraminifera. Nine wood samples from cores MV-54 (n=4) and MV-51 (n=5) were analyzed for their radiocarbon content at three labs (Table 2.2) using accelerator mass spectrometery (AMS). Bulk sediment samples were washed on a 63 µm sieve with distilled water and then at least 10 mg of woody organic matter were hand picked from each for analysis. Only large wood particles (>250 µm) that clearly demonstrated fibrous wood texture were selected.
Because wood and other organic matter may be reworked, particularly during periods of lower sea level, radiocarbon dates taken from wood may represent maximum ages for sediments.
Benthic foraminifera from two samples in MV-51 (Table 2.2) also were taken for radiocarbon analysis. These ages were reservoir corrected by subtracting the average age difference between planktonic and benthic foraminifera in the western Pacific Ocean
[1483 yrs; Broecker et al., 2004]. Benthic foraminifera were analyzed because commonly used planktonic foramnifera (e.g., Globigerinoides ruber, G. sacculifer) were not sufficiently abundant.
For six samples, reported radiocarbon ages were corrected for past variations in cosmogenic production using OxCal v 3.1 with the IntCal98 correction curve for dates
≤20,000 14C yrs B.P. [Stuiver et al., 1998]. Five of seven dates from MV-51 remain
“uncorrected” because internationally ratified calibration curves do not extend beyond
15 20,000 14C yrs B.P. [Stuiver et al., 1998]. Calendar ages for these samples may be as
much as 4000 to 6000 yrs older between 20,000-35,000 14C yrs B.P as a result of increased 14C production during changes in the Earth’s magnetic field intensity [Hughen et al., 2004; Fairbanks et al., 2005].
2.3.4. Magnetic Susceptibility
MS is a measure of the strength of induced magnetism experienced by mineral
grains in sediments when placed in an applied field. Low-field MS was independently
measured on piston-core samples at 5-20 cm intervals using a susceptibility bridge in a
magnetically shielded room at the LSU Rock Magnetism Laboratory. The MS Bridge
uses a balanced coil induction system and is sensitive to 1 (±0.2) x 10-9 m3 kg-1 [Ellwood
et al., 1996]. This method permits an independent cross check on data collected from the
MSCL.
MS values are excellent indicators of the total iron-containing compounds in
sediments [Nagata, 1961] and therefore MS is primarily a function of hinterland
mineralogy and climate. Terrigenous magnetic components originate from multiple
mineralogies including: 1) ferrimagnetic (such as magnetite and maghemite), 2)
paramagnetic sources such as clays (e.g., chlorite and illite), iron sulfides (e.g., pyrite), and 3) ferromagnesian silicates (such as biotite, amphibole, and pyroxene). Weakly negative diamagnetic components in sediments include calcite, quartz, and organic matter. Diamagnetic contributions may reduce the overall sediment MS because of a dilution effect, although this is generally relatively small, given the susceptibility of even small amounts of detrital components [Ellwood et al., 2000].
16 2.3.5. Calcium Carbonate
Calcium-carbonate content (Figures 2.3, 2.4) was analyzed in MV-51 and 54 using the vacuum-gasometric technique [Jones and Kaiteris, 1983] at the LSU Rock
Magnetism Laboratory. Dried sediment samples were ground to a fine powder and then a
0.25 g subsample was reacted under a vacuum in a reaction vessel with phosphoric acid for one hour. After complete reaction, the pressure difference was then read from a pressure gauge and used to calculate the sample carbonate concentration using a regression equation. Triplicate carbonate measurements on several samples resulted in a precision of +/- 0.10%. Tests of reagent grade (99%) carbonate and a mixed sample standard (95% feldspar 5% carbonate) gave accurate results to +/- 1%.
2.3.6. Sedimentary Organic Matter
The quantity and type of sedimentary organic carbon offer major evidence for
reconstructing paleoclimatic and paleoceanographic change. Organic matter in core
sediments was measured using three independent analyses. Total organic carbon (TOC)
was measured on core sediments using a LECO Carbon analyzer at 20 cm sample
intervals for both MV-51 and MV-54. Rock-Eval Pyrolysis (REP) was measured on
sediments from MV-51, because this second core contained the longest time-stratigraphic
record. TOC and REP measurements were made at Baseline Resolution Analytical Labs,
Houston Texas, with analytical errors of ± 0.05 wt. % and ± 0.02 mgHC/gTOC, respectively. TOC data quantify weight percent of organic carbon [Jarvie, 1991] and may serve as an initial proxy for productivity [Berger and Herguera, 1992; Rühlemann et
al., 1999]. Mass accumulation rates (MAR) for TOC were then calculated using the
accumulation rates in both piston cores:
17
= × × LSRDBDTOCMARTOC g g cm g (1) MARTOC =××= g cm3 kyr 2 ⋅kyrcm
where: C org g org C .%)( wtTOC .%)( ×= 100 sed g g sed DBD= Dry Bulk (gDensity 3- )cm LSR = Linear Sedimentat Rateion (cm 1- ).kyr
REP helps to determine the principle sources of the TOC by examining the amount of free (S1) and pyrolytic (S2) hydrocarbons preserved in sediments. REP was developed as a screening method to determine the hydrocarbon potential of petroleum source rocks [Tissot and Welte, 1984; Rullkötter, 2000], but it is becoming widely used in paleoceanography [Lallier-Vergès et al., 1993; Meyers, 1997; Bouloubassi et al., 1999;
Rullkötter, 2000; Wagner, 2000]. Two REP parameters are used in this study, the hydrogen index (HI) and the thermal maximum (Tmax). The hydrogen index is the mass of hydrocarbons per gram of TOC and is determined from heating kerogen in a sample until the maximum amounts of hydrocarbons are evolved [Espitalié and Bordenave,
1993]. Equation 2 was used to calculate the hydrogen index. Tmax is the temperature during pyrolysis at which the most hydrocarbons are evolved.
× S2100 mg 100g mgHC HI = =×= (2) TOC g g gTOC
where : HI = Hydrogen Index ⎛ mgHC ⎞ S2 = ⎜ ⎟ ⎝ g sediment ⎠
18 Hydrogen index values are generally low (≤150 mgHC gTOC-1) for organic
matter composed of vascular-land-plant carbon in recent deep-sea sediments, and are
characterized as Type 3 and Type 4 kerogen [Tissot and Welte, 1984; Meyers, 1997].
Algal-rich marine organic matter values range from 200 to 400 mgHC gTOC-1 for recent
marine sediments and are characterized as Type 2a kerogen [Rullkötter, 2000]. Tmax
values of modern “immature” organic carbon are below 430°C for kerogen types 3 and 4,
whereas “overmature” organic carbon from reworked fossil organic matter are >430°C
[Delvaux et al., 1990].
More precise determination of sedimentary organic-carbon types was achieved by
analyzing the C/N and δ13C ratios on selected samples from MV-51. Isotopic analyses
were performed on a Thermo Finnigan Delta Plus XP Isotope Ratio Mass Spectrometer
with an online elemental analyzer at the Coastal Ecology Institute, LSU. Dry sediment
samples were first acidified with 10% HCl until the reaction with calcium carbonate
ceased, or approximately 10 minutes. Sediments were then acidified with 50 ml
concentrated HF for eight hours to remove minerogenic matter and isolate the kerogen.
Residues were centrifuged and rinsed with distilled water several times until completely neutralized, then dried at 50°C and powdered for analysis. We also made slides from this isolated kerogen for later optical palynofacies analysis. Standard deviations of duplicate measurements show analytical precision of δ13C to be ±~0.15‰ and C/N ratios to ±~0.20
molar percent.
The terrigenous component of TOC was determined from δ13C values using a
two-source mixing model in which the marine (-20.5‰ ±0.5‰) and terrigenous (-26.5‰
±0.5‰) isotope values are used as end members [Bird et al., 1995; Aller and Blair,
19 2004]. Equations 3-5 are adapted from Amo and Minagawa [2003] and Bird et al.,
[1995].
13 13 (δ Cmeasured −δ Cmarine) f sterrigenou = (3) 13 13 (δ C lterrestria −δ Cmarine)
13 13 ⎡ δ Cmeasured −δ Cmarine ⎤ % = fTOClTerrestria TOC =× ( )× × .%100 TOCwt (4) sterrigenou ⎢ 13 13 ⎥ ⎣⎢()δ C lterrestria −δ Cmarine ⎦⎥
2.3.7. Benthic Foraminifera
Benthic foraminiferal assemblages are very useful paleoproductivity proxies because their abundances and assemblage compositions are linked to the flux of organic carbon arriving at the seabed. Several studies have demonstrated that foraminiferal densities are positively correlated with increases in nutrient fluxes, surface-water productivity, and the subsequent flux of that material to the seabed [e.g. Gooday, 1988;
Jorissen et al., 1992; Sjoerdsma and van der Zwann, 1992; Loubere, 1996]. Certain assemblages of benthic foraminifera are adapted to particular types of organic-carbon flux, such as highly variable or “seasonal” flux that is charactersistic of upwelling zones compared to relatively constant flux [Loubere and Farridudin, 1999; Smart and Gooday,
1997; Licari and Mackensen, 2005]. We use benthic foraminifera as an independent metric of organic-carbon flux, as well as the stability of flux through time.
In this study, eight to ten grams of 36 freeze-dried samples from core MV-51
were gently washed with distilled water over a 63 µm sieve and dried at 50°C. The >63
µm fraction was then weighed and further split into 150-250 µm and >250 µm fractions.
At least 300 benthic foraminifers were then picked from each fraction >150 µm, so that
20 ~600 specimens were counted from each sample. Specimens were placed on assemblage
slides, sorted, and then enumerated. Assemblage data are routinely collected from the
>150 µm fraction in ecological and paleoceanographic studies [e.g. Pérez et al., 2001]
because ecologically important taxa may be missed by examining only the >250 µm
fraction [Sen Gupta et al., 1987]. Examination of these fractions also permits comparison
with other studies. The >63 µm fraction was inspected for all sieved samples in MV-54
to determine presence or absence of foraminifera.
Benthic foraminiferal densities (BF# = total shells or specimens g-1) and benthic
foraminiferal accumulation rates (BFAR; # cm-2 kyr-1) were calculated from absolute foraminiferal densities [Herguera, 1992; Herguera and Berger, 1994]. Groupings of diagnostic species were examined with R-mode cluster analysis (Minitab v. 14.20) using a 1-Pearson correlation coefficient distance metric and an average linkage method.
Cluster analysis is based on the absolute abundance data (BF#, total specimens g-1) in 36 samples for the most abundant, recurrent, and diagnostic taxa, totaling 12 species. Q- mode cluster analysis was used to examine the stratigraphic clustering of samples based on BFAR data for the same 12 taxa. Taken together, these quantitative data are used to infer relative paleoproductivity during the past ~33,000 years in the northern Pandora
Trough.
2.3.8. Thermomagnetic Measurements
Thermomagnetic susceptibility measurements were performed on a number of
samples from MV-51 using the KLY 3S Kappa Bridge manufactured by AGICO, Czech
Republic. Results are given in SI and are dimensionless. This measurement involves
21 heating the sample from room temperature to 700oC while measuring the MS as temperatures rise.
Paramagnetic minerals show a parabolic-shaped MS decay during this process,
because the MS in these samples is inversely proportional to temperature of measurement
[Hrouda, 1994]. Based on measured standards, ferrimagnetic minerals usually show an
increase in MS up to a point where the MS decays toward the Curie temperature of the
minerals responsible for the MS. For magnetite, this temperature is ~580o C and for
maghemite it is ~610o C for most samples. Above 640oC, maghemite converts to
hematite with a Curie point of ~680oC. During measurement, some minerals that contain
iron and exhibit paramagnetic behavior at low temperatures, are unstable at higher
temperatures and chemically change during heating. This often produces diagnostic peaks
that result from the formation of a ferrimagnetic phase during this change [Ellwood et al.,
2007].
2.4. Results
2.4.1. Core Sediments
Sediments in MV-51 and MV-54 are composed of homogenous marine mud, and
contain few discernible sedimentary structures. MV-51 contains four ~5-cm thick
intervals at 1.6, 2.4, and 4.8 meters below sea floor (mbsf) that contain well preserved
biogenic sedimentary structures associated with diffuse ash layers. Another discrete,
partially bioturbated ash layer occurs at 6.9 mbsf. X-radiographs of MV-51 over the 3-
4.5 mbsf interval revealed homogeneous sediments, with no obvious bedding. MV-54
contains two coarse-grained, fining-upward, organic rich intervals at 1.5-1.9 mbsf and
22 6.1-6.3 mbsf. Sediments in the interval 6.1-6.3 mbsf are inclined by ~45°, suggesting
deposition by mass flow.
2.4.2. Geochronology and Accumulation Rates
Eleven radiocarbon dates illustrate a late transgression to Holocene section <2 m
thick, very rapid linear sedimentation rates (LSR), and high MAR for both cores during times older than ~15,000 Cal. yrs B.P. (Figure 2.2). Though we do not have radiocarbon dates or 210Pb data in the uppermost sediment of MV-51 and 54, we assume a recent age
for the core tops. Radiocarbon dates for MV-54 are generally all within the range of
error for each other, particularly when corrected to calendar years (Figures 2.2 and 2.3;
Table 2.2). The date at 1.9 mbsf in MV-54 may be an exception. We considered the dates of 17,000 Cal. yrs B.P. at 4.75 mbsf and 8.3 mbsf to be indistinguishable.
Figure 2.2. Time-Depth plot of radiocarbon ages. Arrows point to radiocarbon ages determined from mixed in situ benthic foraminifera. Numbers along profiles are sediment accumulation rates in m kyr-1. Ages and error estimates are listed in Table 2.2.
23 Two sections in MV-51, 1.40-3.50 and 10-12 mbsf, show high LSR and MAR
(Figure 2.3). The first section (~18,400-20,400 Cal. yrs B.P.) exhibits LSR of ~1.09 m
kyr-1 and ~81.2 g cm-2 kyr-1, and the LSR in the second section (>32,000 14C yrs B.P.) is
twice as much at ~2.70 m kyr-1 and ~244 g cm-2 kyr-1. Two sets of radiocarbon ages are
within the range of error for each other, 1.20 and 1.34 mbsf, and 10 and 12 mbsf (Figure
2.2, Table 2.2).
Table 2.2. Radiocarbon ages reported in this paper. Lab #: 1 Keck CCAMS Facility, UC Irvine, 2 Beta Analytic, 3 National Ocean Sciences AMS at WHOI. *Radiocarbon ages are calibrated to calendar ages (2 sigma 95% probability) using OxCal v. 3.10 [Bronk Ramsey, 1995] and the IntCal 98 calibration curve [Stuiver et al., 1998]. ** Corrected by subtracting age difference between planktic and benthic foraminifera in the Western Pacific [Average age difference: 1483 yrs; Broecker, 2004]. Depth (m) Lab Material 14C Age Error (yrs) Median Calendar Age, Range (± yrs) B.P.* MV-51 1.20 1 wood 15 565 ± 50 18 600 650 MV-51 1.34 2 wood 15 390 ± 100 18 405 505 MV-51 3.51 1 wood 20 400 ± 90 20 405 185 MV-51 7.40 3 Benthic Foraminifera 29 500 ± 220 28 017** 220 MV-51 7.94 1 wood 29 940 ± 180 out of range 180 3 Benthic 34 100 ± 250 250 MV-51 10.00 32 617** Foraminifera MV-51 12.00 1 wood 33 360 ± 320 out of range 320 MV-54 1.90 1 wood 13 320 ± 25 15 850 700 MV-54 4.75 1 wood 14 160 ± 35 17 000 500 MV-54 8.30 1 wood 14 175 ± 40 17 000 500 MV-54 9.94 2 wood 14 810 ± 80 17 740 440
The time intervals outside of high accumulation rate sections in MV-51 show 1) a
very thin late transgression to Holocene section (~1 m) and 2) a thick or expanded section
(~4 m) between ~20,400 and 29,000 14C yrs B.P., during the earliest to middle part of the
Marine-Isotope-Stage-2 glaciation that includes the LGM time of 23,000–19,000 Cal. yrs
B.P.. Sediments were deposited with very low LSR (0.07 m kyr-1) and MAR (~5 g cm-2 kyr-1) since late lowstand time (<~18,400 Cal. yrs B.P.). Intermediate LSR of 0.51 m
24 kyr-1 and moderate MAR of 94 g cm-2 kyr-1 occurs between 20,400 and 32,000 14C yrs. A short duration decrease in LSR (0.28 m kyr-1) and MAR (24 g cm-2 kyr-1) occurs between
28,000-29,900 14C yrs (Figures 2.2 and 2.3).
Two distinct intervals of LSR and MAR occur in MV-54 (Figure 2.4). The first
interval is from 0-1.9 mbsf (0-15,850 Cal. yrs B.P.) and is characterized by LSR of 0.11
m kyr-1 and MAR of 9.7 g cm-2 kyr-1. The second interval is from 1.9-9.4 mbsf and
contains the highest LSR of 3.96 m kyr-1 and MAR of 428 g cm-2 kyr-1.
2.4.3. Magnetic Susceptibility
Trends in data collected from the MS bridge at Louisiana State University are generally in good agreement with MSCL Bartington Loop measurements (R=0.81, p<<0.005; n=213; Figures 2.3 and 2.4). Values from the MSCL are relatively low for both cores (<40x10-5 SI) while MS bridge values are higher and show averages between
1.20 x 10-7 m3 kg-1 (MV-51) and 1.98 x 10-7 m3 kg-1 (MV-54). High frequency MS
fluctuations are evident in MV-51, particularly between 5.5-7.5 mbsf that are not present
in the Bartington Loop data alone (Figure 2.3). This is due to the diminished sensitivity
of the Bartington Loop Sensor at low MS values. Results presented below are limited to
detailing the more sensitive MS bridge data.
In MV-51, sediments less than 10 mbsf, or younger than 32,000 14C yrs B.P., have
low MS values, with some notable exceptions. MS maxima occur at 2.40, 4.80, and 6.94
mbsf (mean 3.56 x 10-7 m3 kg-1, n=3; Figure 2.3) and are nearly four times higher than
background levels over the interval from 0-10 mbsf (mean 9.35 x 10-7 m3 kg-1, n=51).
Higher MS values are present in MV-51 between 1.15 and 1.55 mbsf, depths
corresponding to approximately 18,400 Cal. yrs B.P.
25 A major shift to higher MS values by two to three orders of magnitude in MV-51
occurs below 10 m (>~32,000 14C yrs B.P.) and indicates enhanced input of detrital
minerals during that time. The magnetic detrital fraction in these sediments is restricted
to mud-sized particles because the sand-sized fraction (>63 µm) comprises <1.5% of sediments below 10 mbsf (Figure 2.3), and those coarse fractions are composed entirely
of foraminifera.
MS values are much more variable in MV-54 (Figure 2.4). Sediments from 0 to
1.9 and 9 to 12 mbsf have high MS values (mean 2.7 x 10-7 m3 kg-1, n=31), or
approximately <15,800 and 17,700 Cal. yrs B.P., respectively. A more variable interval
occurs from 2 to 9 mbsf (mean 1.33 x 10-7 m3 kg-1, n=35) and contains two maxima at
3.40 and 6.20 mbsf with values more similar to the upper and lower parts of the core.
The high values from ~1.5 – 1.9 mbsf are associated with dark, muddy sand and at 6.2
mbsf with dark, sandy, laminated and contorted sediment. The sand-sized fractions (>63
µm) comprise as much as 85% of the sediment from 1.5 – 1.9 mbsf and 25% at 6.20 mbsf
(Figure 2.4). Sediments are generally coarser in MV-54 (mean 15% >63 µm) than in
MV-51 (mean 2%).
2.4.4. Calcium Carbonate and Volcanic Ash Layers
Siliciclastic particles dominate sediments from cores MV-51 and MV-54.
Samples in MV-51 have very low calcium carbonate percentages (mean 3.8%, n=68)
whereas values are slightly lower in MV-54 (mean 3.0%, n=37). Calcium carbonate in
MV-54 occurs only in the upper 3.6 m and is completely absent below 4.0 mbsf (Figure
2.4).
26
27
28 Several volcanic ash layers are present in MV-51. MS maxima at 2.4, 4.8, and
6.9 mbsf are a result of elevated levels of paramagnetic grains such as biotite,
hornblende, and amphibole that were identified from sieved samples (>63 µm) using a
binocular microscope. These grains are accompanied by abundant pumice and glass
shards and result in sharp increases in the sand-sized fraction (>63 µm, Figures 2.3, 2.7),
suggesting a sudden input by volcanic origin. TOC and calcium-carbonate values decline
at the same levels, probably due to particle dilution, but they may also reflect a decline in biological productivity during deposition of ashfall. Another pumice-rich layer is present at 1.0 mbsf, but it does not correspond to an MS maximum because it lacks abundant biotite and other iron-rich paramagnetic minerals. Additionally, the pumice-rich layer at
1.0 mbsf does not show a significant decrease in either TOC or calcium carbonate contents.
2.4.4. Sedimentary Organic Matter
Organic carbon content in MV-51 and MV-54 is relatively low overall and at site
MV-51, appears largely to have a source from land-plant debris. TOC in core MV-51 ranges from 0.41 to 0.97 wt. % (mean 0.75), generally lower than values in core MV-54, which range from 0.63 to 2.29 wt. % (mean 0.92). In MV-51, TOC is highest from 1.4 to
1.8 mbsf and >7.4 mbsf, or during 18,400-20,400 Cal. yrs B.P. and >28,000 14C yrs B.P.,
respectively. MAR of TOC is also highest during these time intervals (Figure 2.3).
MV-54 shows nearly uniform values in TOC over most of the core length. Four
prominent subsurface maxima in excess of 1.3 wt. % occur at 1.8, 3.2, 6.2, and 7.0 mbsf
and correspond to maxima in both MS and the sand-sized fraction (Figure 2.4).
Sediments from 1.5-1.9 mbsf and at 6.2 mbsf are composed of a dark, sandy material,
29 with abundant and well preserved pteropods, foraminifera, and woody material. Organic
particles are common in all sieved samples from MV-54, many of which are unpyritized
and can be identified as plant cuticle and woody cortex under binocular microscopy. One
sample from 7.0 mbsf contained tree resin (amber) and a complete leaf (~2mm long) with
well preserved venation.
Hydrogen index values of sedimentary organic matter from MV-51 are very low
(7 to 152.6 mgHC gTOC-1; average 74) and contain very low S2 pyrolytic yields (Figure
2.5a). These data reflect hydrogen-poor organic carbon and therefore may indicate either
type 3 organic matter or a mixture of type 3 and oxidized type 2 marine organic matter.
Figures 5a and 5b show the data plotted in two different diagrams, both generally plotting
below the Type 3 terrestrial organic carbon thresholds. However, low S2 values, below 1
mgHC g sediment-1, may be a consequence of mineral matrix dilution of hydrogen-poor
hydrocarbons which makes it difficult to accurately determine the organic matter type
[Espitalié et al., 1980; Katz, 1983; Langford and Blanc-Valleron, 1990]. No correlation
exists between hydrogen index and TOC values, as expected given the very low hydrogen
index values and narrow range in TOC values [Bouloubassi et al., 1999].
Tmax measurements generally indicate that MV-51 sediments contain immature organic carbon, as expected for recent marine sediments. However, six samples from 3.0,
3.2, 4.6, 5.2, 8.8, 9.8 mbsf (Figure 2.5b) indicate overmature organic carbon, which probably had a source from reworked sedimentary rocks inland (e.g., coal). There is no stratigraphic significance to the timing of reworked organic matter in MV-51.
30
Figure 2.5. Rock-Eval 2 data from MV-51 plotted in an S2 versus TOC diagram (5a, after Langford and Blanc-Valleron, 1990) and a Van Krevelen diagram (5b, after Wagner, 2000). Boundaries in the Van Krevelen diagram between kerogen type 2, 3, and 4 are plotted with dashed lines and derived from reference values reported in Delvaux et al. [1990]. Hydrogen index values for MV-51 plot in the vascular land plant type 3 to type 4 kerogen window, with six samples representing reworked fossil organic matter (3.0, 3.2, 4.6, 5.2, 8.8, and 9.8 meters below sea floor, labeled on plot).
31
32 The percentages of terrestrial components calculated from stable isotopic data
exhibit nearly parallel changes with TOC, with maximum values (>40%) between 1.4-1.8
mbsf and deeper than 7.4 mbsf (Figure 2.6). Three terrestrial input maxima occur in
MV-51, one at 1.80 mbsf (62%), another at 8.08 mbsf (55%), and the third at 10 mbsf
(57%). A strong positive correlation exists between TOC concentrations and percent terrigenous TOC (R=0.85, p<0.001; n=24), indicating that increases in TOC concentrations result from enhanced input of continental sources and not from marine primary productivity.
Figure 2.7. Organic geochemical data from MV-51 are plotted in an elemental (atomic C/N) and stable-isotopic data cross plot. Data (empty circles) generally plot along a theoretical mixing line (dashed line) between two sources, modern marine organic matter and C3 land plants. OM refers to organic matter and SOM refers to soil organic matter. Plot after Meyers [1997] and sources mapped from values reported in Meyers [1994] and Goni et al., [2006].
Elemental and stable-isotope data provide further constraint on organic-carbon sources. A plot of elemental (atomic C/N) and stable isotopic (δ13C, ‰) values from
33 MV-51 illustrates two groupings of data (Figure 2.7). The first group (0.4-6.8 mbsf)
plots closer to modern marine organic matter values than does the second group (core-
top, and >7.4 mbsf). This second group (core-top, and >7.4 mbsf) clusters nearest to the
region of C3 land plants. Atomic C/N and stable-isotopic data generally fall on a mixing
line between the two sources, except for one point from an ash layer at 6.9 mbsf that rests
well below the mixing line (Figure 2.7). The ash layer forms an outlier, possible as a
result of a different organic matter source such as woody debris transported in pumice
rafts.
2.4.5. Benthic Foraminifera
Benthic foraminifera are abundant in MV-51 sediments. Absolute densities and
BFAR (Figure 2.8) generally give the same trends but with some differences in the lower
seven meters. Lowest densities of foraminifera (mean ~98 specimens g-1, n=22) occur
shallower than 6.8 mbsf, whereas densities between 6.8 and 12 mbsf are nearly two times
higher (mean ~174 specimens g-1, n=14; Figure 2.8). A sample at 6.8 mbsf contains the
highest density of foraminifera in MV-51 (~283 specimens g-1). In contrast, lower BFAR
(mean ~5800 cm-2 kyr-1, n=31) characterize core depths less than 10 mbsf, except for the
maximum at 6.8 mbsf (12,600 cm-2 kyr-1). Deeper than 10 mbsf, benthic foraminiferal densities are six times higher than in samples less than 10 mbsf (mean ~34,800 cm-2 kyr-1, n=5; Figure 2.8). The significance of benthic foraminiferal densities, BFAR, and organic carbon flux at site MV-51 (Figure 2.8) was examined with multiple linear regression analysis.
34
35 A statistically significant association exists between total benthic foraminiferal
densities >150 µm and TOC concentrations (R=0.46; p=0.003; n=35). However, the low
correlation coefficient suggests that while TOC plays a role in the densities of
foraminifera, other factors are important as well. If the mass accumulation rates are
incorporated into the raw data so that TOC MAR and BFAR are calculated, a much
stronger correlation of analytical data exists (R=0.95; p<0.001; n=35). These statistics
emphasize the importance of incorporating sediment accumulation rates with raw
foraminiferal data [Herguera, 1992]. Together, they show that the mass accumulation of
organic-carbon and foraminifera are intrinsically linked, and therefore that BFAR is a
reflection of organic-carbon flux.
Two periods of high BFAR and high TOC MAR occur in MV-51 (Figure 2.8).
The first interval is 1.4-3.4 mbsf and corresponds to ~18,400-20,400 Cal. yrs B.P., or middle to late lowstand time. The second interval occurs below 10 mbsf in sediments greater than ~32,000 14C yrs B.P., or approximately the late Stage 3 interstadial. The
latter period contains the highest BFAR and TOC MAR in the entire core. In addition to
total BFAR, important paleoecological information is contained in the accumulation rates
for groups of benthic foraminiferal species.
R-mode cluster analysis divided benthic foraminiferal species into two distinct
groups (Figure 2.9). The first cluster is characterized by low-abundance taxa including
Oridorsalis tener and Melonis barleeanus. Species of this cluster rarely exceed 1-5
specimens per gram or 150 specimens cm-2 kyr-1. The second cluster is composed of ten
taxa (Figure 2.9) with three distinct subclusters being the most abundant and
paleoceanographically significant. The first subcluster (A) is composed of Uvigerina
36 peregrina/hispida, Cibicidoides pachyderma, and Bolivina robusta (Figure 2.9). The
second subcluster (B) is composed of Bulimina aculeata, Sphaeroidina bulloides, and
Bolivinita quadrilatera as a companion taxon. The more distant third subcluster (C) contains Gavelinopsis translucens and Globocassidulina subglobosa. Species distributions in Figure 2.9 are arranged in order of the clusters described above.
Figure 2.9. Cluster analysis of absolute densities (specimens g-1) for the twelve most abundant species in MV-51 sediments. Analysis used is a 1-Pearson distance metric and average-linkage clustering, n=36 samples. Major clusters are numbered and smaller subclusters within cluster 2 are identified by alphabetic letters.
Three distinct species BFAR intervals are present in MV-51 (Figure 2.8, shaded
areas). The first interval (10-12 mbsf; >32,000-33,000 14C yrs B.P.) is characterized by
elevated abundances of nearly all foraminiferal species with the exception of M.
barleeanus (Figure 2.8). The most abundant taxa are, in order of decreasing abundance,
B. robusta, U. peregrina/hispida, S. bulloides, Bul. aculeata, C. pachyderma, and O.
37 tener. The second interval is composed of a single sample at 6.8 mbsf (<28,000 14C yrs
B.P.) and is composed chiefly of B. robusta, U. peregrina/hispida, and C. pachyderma.
The third abundance interval occurs from 1.4-3.4 mbsf (~18,400-20,400 Cal. yrs B.P.) and is composed mostly of B. robusta, G. translucens, Gc. subglobosa, O. tener, and M.
barleeanus. Overall, the most abundant species throughout MV-51 is B. robusta.
Large sections of MV-51 contain fewer benthic foraminifera and stand in contrast to previously mentioned high BFAR and abundance intervals. In particular, two intervals
0-1.4 and 3.4-6.8 mbsf contain the lowest BFAR values (Figure 2.8). The interval 0-2.4 mbsf (<18,400 Cal. yrs B.P.) is characterized by B. robusta, U. peregrina/hispida, Bul. aculeata, and Gc. subglobosa, in order of decreasing average BFAR values. In contrast, the second interval 3.4-6.8 mbsf (>20,400 Cal. yrs B.P. to <28,00014C yrs B.P.) is
composed mostly of B. robusta, Gc. subglobosa, Bul. aculeata, and G. translucens. The
interval 7-10 mbsf (>28,000 to <32,000 14C yrs B.P.) shows intermediate BFAR values
and is composed of the same dominant species as for 10-12 mbsf.
The Q-mode dendrogram shows the clustering of samples based on the BFAR
values for 12 species at each depth interval (Figure 2.10). In particular, the stratigraphic
record is grouped according to BFAR and the species assemblages discussed above.
2.4.6. Thermomagnetic Results
Samples measured from MV-51 over the entire range of the core (Figure 2.11, n=8)
exhibit a ferrimagnetic behavior at low temperatures but alter at higher temperatures to a
secondary, stronger ferrimagnetic phase. Such phases are typical of iron-containing
detrital minerals. The character of all of the samples measured is very similar, suggesting
that no major change in magnetic mineralogy is present in the core.
38
Figure 2.10. Q-mode cluster dendrogram of benthic foraminiferal samples from MV-51. Analysis used is a 1-Pearson distance metric and average-linkage clustering, n=36 samples.
We interpret this result to indicate that high MS values observed in the base of MV-51
(Figure 2.3) are the result of an increase in the detrital component at the site, with a corresponding decrease in percent calcium carbonate.
2.5. Discussion
2.5.1. Magnetic Susceptibility and Climate
Variations in MS signals primarily reflect the processes governing terrestrial erosion and transport into the marine environment, largely the result of climate change and eustacy
[Tite and Linington, 1975; Verosub et al., 1993; Banerjee, 1996; Ellwood et al., 1996;
39 Ellwood et al., 2000] and biological productivity [Ellwood and Ledbetter, 1977; Mead et al., 1986]. Fluvial and airborne delivery mechanisms account for most detrital input and are themselves primarily functions of precipitation and aridity, respectively. Volcanic ashfall may provide ferri- and paramagnetic compounds to marine sediments. Also affecting the MS, albeit to a lesser extent, are fluxes of diamagnetic calcium carbonate, biogenic silica, and organic matter to the seabed, which are functions of primary productivity in the surface water and therefore of nutrient supply.
Figure 2.11. Thermomagnetic results for eight samples from core MV-51JPC. Numbers adjacent to each curve represent the core depth (cm).
The effects of climate change and provenance on MS records are well documented for terrestrial and marine sediments [e.g. de Menocal et al., 1991; Heller and
40 Evans, 1995; Ellwood et al., 1999; Weedon et al., 1999; Balsam et al., 2005].
Precipitation is one of the principal drivers of chemical and physical weathering,
particularly in the monsoon-dominated tropics. Periods of increased precipitation are
generally associated with enhanced rates of runoff and erosion of weathered detrital
particles. Likewise, drier conditions are less favorable for erosion and transport of
weathered material, particularly during glacial ice advances, because they remove water
vapor available for erosion. During eustatic sea-level lowstands, river mouths may
prograde across exposed shelves to create coastlines that are much closer to shelf breaks,
and therefore may deliver greater amounts of detrital sediment to marine basins.
MS is not strongly influenced by diamagnetic minerals and therefore reflects a
dominantly detrital signal. Linear regression analysis was used to test MS values against
other variables such as TOC and calcium carbonate in cores MV-51 and 54. No association exists between TOC and MS in either MV-51 (R=0.04, p=0.29, n=70) or MV-
54 (r2=0.06, p=0.26, n=62). There is no correlation between calcium carbonate and MS
in either MV-51 (R=0.16, p=0.195, n=68) or MV-54 (R=0.17, p=0.311, n=37).
Additionally, no correlation exists between TOC and calcium carbonate content (MV-51:
R=.14, p=0.255, n=68; MV-54: R=.1, p=0.548, n=37).
Eroded detrital minerals dominate the MS signal, and their production, fate and
transport are controlled by mineral provenance and climate. Changes in river source(s)
and/or climate may explain MS patterns in MV-51. The most notable pattern in the MS
data for MV-51 is the major change at 10 mbsf (Figure 2.6). We propose some
possibilities to explain the higher MS values based on the surrounding PNG geology.
41 • PNG Geology
Papua New Guinea is composed of a complex, heterogeneous geological terrain
consisting of predominantly folded sedimentary rocks in western PNG and volcanogenic
rocks in the central-eastern PNG region [Rickwood, 1968; Davies and Smith, 1971]. The
Fly-Strickland River drainage basin in western PNG erodes predominantly limestone,
siltstone, and sandstone lithologies [Rickwood, 1968, Brunskill, 2004] and as a result discharges relatively mature sediments with high quartz/feldspar ratios (Figure 2.12, fluvial source 1) [Slingerland et al., 2007]. In contrast, central and eastern PNG are dominated by metamorphic and volcanic rocks that result in immature sediment lithologies draining from the other PNG rivers such as the Turama-Kikori-Purari Rivers
(Figure 2.12, fluvial source 2) [Davies and Smith, 1971; Slingerland et al., 2007].
In eastern PNG, the Owen Stanley Range is a large metamorphic region that
forms the prominent mountainous spine of the PNG highlands (Figure 2.12).
Metasedimentary rocks from the Owen Stanley Range consist of greenschist, slate, and
phyllite, and are interrupted by numerous igneous intrusions [Davies and Smith, 1971;
Blake, 1976]. Basaltic, volcanic rocks intrude the Owen Stanley Range and compose a
1000-km belt of Cretaceous to Pliocene rocks that are potassium rich and silica
undersaturated [Mackenzie, 1976; Smith, 1982]. Although the precise mineralogies of these metamorphic and igneous rocks are not well documented, a number of studies suggest that these rocks are mineralogically distinct from the dominantly sedimentary and intrusive andesitic volcanic rocks in western PNG [Rickwood, 1968; Mckenzie, 1976].
Therefore, we hypothesize that in the past, river sediments from central and eastern PNG
42 (Figure 2.12, fluvial sources 2-3) may have been more important than the Fly-Strickland as a source of iron-bearing clastics.
During the Last Glacial Maximum (LGM), lowered sea-level in the GoP enabled
river dispersal systems to incise across the exposed shelf to the shelf-edge [Harris et al.,
1996]. Given the proximity of MV-51 to central and eastern PNG and the higher
gradient, we suggest that the Turama-Kikori-Purari river system may have incised a more
direct pathway to MV-51 so that high MS values below 10 mbsf result from a more direct
input of volcaniclastic material from central and eastern PNG (Figure 2.12, fluvial sources 2-3). The Turama-Kikori-Purari river system would have had far less distance to travel to the northeastern GoP than the Fly River. This hypothesis is testable in the future by collecting quantitative mineralogical data from MV-51 and comparing them to different PNG river sediments. Another test could include age dating amphiboles from above and below 10 m. For example, Pleistocene age dates would indicate that amphiboles came from younger volcaniclastic sources of western PNG while older ages
(e.g. Cretaceous-Pliocene) would indicate a source from the Papuan Peninsula.
• Volcanic ash layers
The source(s) of MV-51 ash layers remain uncertain because few ash layers have
been reported in and around PNG that date to this time interval. Some ash deposits (e.g.
Tomba Tephra) have been studied around Mt. Hagen in western PNG and dated to
~28,000-50,000 yrs B.P. [Pain and Blong, 1976; Chartres and Pain, 1984]. Other ash
deposits, such as the Sagamasi and Natanga Tephras, have been examined around Mt.
Lamington in eastern PNG and dated to 15,000 ± 500-20,100 ±600 Cal. yrs B.P. [Ruxton,
1966].
43 Our radiocarbon dates provide some constraint on the ash layers present in MV-
51. The ash layer at 2.4 mbsf is well constrained between ~18,400 and ~20,400 Cal. yrs
B.P. The ashes at 4.8 and 6.9 mbsf are not as well constrained but were deposited
sometime between ~20,400 Cal. yrs B.P. and ~28,000 14C yrs B.P. Mt. Lamington may
have been a source for the ash layer at 2.4 mbsf but there are not enough dated ash layers
of the correct time frame on PNG to hypothesize the source for the older ash layers.
Therefore, they may originate from any of the active late Quaternary volcanic peaks
shown in Figure 2.12. One exception is Madilogo north of Port Moresby because it is
thought to have formed during the past ~1000 yrs [Blake, 1976]. These ash layers may
serve as correlation points in future GoP studies over this time interval.
2.5.2. Organic Geochemical Data
Organic geochemical proxies suggest that there were significant variations over
time in the paleoflux of organic carbon and the types of sedimentary organic matter
buried in the northeastern Pandora Trough. These data illustrate two important points.
First, TOC and hydrogen index values of sediments are consistently low. Secondly,
stable isotopic and elemental data indicate a complex mixture of marine and C3 vascular
plant matter in the organic carbon, with a potential contribution of other terrestrial plant
matter.
Terrestrial plant matter is common in the modern GoP and therefore it is not surprising that hydrogen index values are indicative of refractory types 3 and 4 organic matter. Small, mountainous rivers are important to the GoP [Milliman, 1995] and on average, 65% of river-borne terrestrial organic carbon is refractory [Ittekkot, 1988]. We frequently observed floating megadetritus (branches to whole tree trunks) in the
44 northeastern GoP during the cruise. In addition, Robertson and Alongi [1995] estimated
that as much as 9 Mt C yr-1 of floating macro and megadetritus occur in the surface water
of the Fly Delta, and that significant quantities are also found on the shelf and in deep-sea
troughs of the GoP. In addition, sediments with hydrogen poor organic matter (i.e.
terrestrial plant matter) are susceptible to much lower hydrogen indices because the
hydrocarbons are retained on mineral matrix during pyrolysis [Katz, 1983; Langford and
Blanc-Valleron, 1990].
Low hydrogen index values in MV-51 are consistent with recent studies of
fluvial-marine dispersal systems [Aller et al., 2004; Goni et al., 2006]. These systems
can promote efficient remineralization of labile organic compounds, particularly during
sediment reworking on the inner shelf, and leave only refractory organic matter.
Therefore, it is possible that a greater amount of labile organic carbon was originally
delivered to MV-51, but that it was later removed from the burial record by oxidation.
We note that trends in pyrolytic hydrogen index data do not correlate consistently
with TOC and stable-isotope-derived percentage of terrestrial TOC (Figure 2.7). One
explanation is that the hydrogen-index data have been biased by hydrogen-rich (algal) labile organic matter in sediments and their oxidation to lower hydrogen index values to reflect type 3 organic matter [Meyers, 1997]. Palynomorph observations indicate common pollen, spores, leaf and grass cuticles, but rare marine palynomorphs such as resistant dinoflagellate cysts. However, total kerogen observations also indicated an abundance of amorphous organic matter, which is derived from degraded marine organic matter and supports the presence of oxidized labile compounds [Lewan, 1986].
45 Stable-carbon isotopic data indicate that TOC contains >50% marine organic carbon throughout much of MV-51. However, the mixing model only takes into consideration two sources, marine and C3 plants. There is also the possibility of an influence from vascular C4 plant matter, soil organic matter upon which C3 and C4 plants grew, or also freshwater algae from rivers and estuaries (Figure 2.7). In addition, it is also possible that the values for the marine end member used in the mixing equation
(-20.5‰ ±0.5‰) changed during the recent geologic past to a more enriched value (e.g. -
19.5‰), which would cause the percentage to change toward more (>50%) terrestrial fractions. At this time, we are unsure about a dominance of marine organic matter as suggested by the stable isotope data and recognize that the estimates of terrestrial sedimentary organic carbon may be minimums.
While keeping in mind the assumptions made in the simple, two-end-member
mixing model, the stable-isotopic and elemental data illustrate a mixture of marine and
predominantly C3 vascular plant matter (Figure 2.7). C3 plants are arboreal/herbaceous
angiosperms and gymnosperms that occur in a wide range of settings where precipitation
is abundant, such as warm and wet tropical rainforests [e.g. mangroves; Tyson, 1995].
C4 plants are fast growing, herbaceous angiosperms (e.g. grasses) that prefer warm and
arid tropical environments such as savannah grasslands [Tyson, 1995; Wagner, 2000;
Wagner et al., 2003]. C3 plants are geochemically distinct from C4 plants in that they
are ~10-15‰ depleted in δ13C (-27‰) compared to C4 plants (-14‰) [Tyson, 1995;
Meyers, 1997; Wagner et al., 2003]. Because of these geochemical distinctions, we suggest that a large proportion, though not all, of the vascular plant matter is from C3 plants.
46 Many global climate records indicate that LGM was both a cooler and drier time
than present; even in the warm tropics [Thompson et al., 1995; Broecker, 1996; Lea et al.,
2000]. Therefore, we had hypothesized that during LGM, a grassland savannah may
have covered much of the PNG coastal plain as it did in northeastern Queensland
[Kershaw, 1978; 1988]. However, none of the samples measured from MV-51 exhibit an
obvious C4 contribution during LGM. Instead, values show a C3 signature that seems to
indicate that wet conditions and C3 vegetation persisted in PNG for the past ~33,000 14C
yrs B.P.
Terrestrial climate records from the GoP are mainly from the Papuan Highlands
[Bowler et al., 1976; Hope, 1976; Hope and Tulip, 1994], although some records exist from the coastal plains [Garrett-Jones, 1979]. Pollen records from the PNG highlands indicate an atmospheric cooling of 6-10°C from 40,000 - 30,000 14C yrs B.P. and during
LGM from 25,000 - 15,000 Cal. yrs B.P. [Bowler et al., 1976; Webster and Streten,
1978]. During the past ~10,000 Cal. yrs B.P. the coastal plain in northern PNG, adjacent to the Solomon Sea, was similar to today in that it was covered by a Lower Montane
Rainforest with some savannah grasses [Garrett-Jones, 1979]. Very little is known about
the coastal plain surrounding the GoP during the past ~10,000 Cal. yrs or preceding time
intervals. Pollen records from Lynch’s Crater, northeastern Queensland (~1000 km SW)
indicate dry conditions and Eucalyptus woodlands dominated during 26,000 – 8,000 Cal.
yrs B.P. [Kershaw, 1978; 1988].
We tentatively suggest that the C3 signal preserved in MV-51 sediments is
indicative of wet tropical vegetation covering the PNG coastal plain and mangroves
rimming the coast similar to today [Harrison and Dodson, 1993; Robertson and Alongi,
47 1995]. Nevertheless, we are unable to rule out the possibility of C4 grasslands in the
vicinity of the GoP without pollen data. Therefore, we recommend a future examination
of pollen spectra from GoP sediment cores to further test this interpretation.
TOC values are slightly higher in MV-54 than they are in MV-51. One
explanation for higher TOC values may be that the high accumulation rates found in MV-
54 would have limited exposure time to organic carbon oxidation more than at site MV-
51 [Hartnett et al., 1998; Hedges et al., 1999]. However, wood-rich organic matter is
abundant throughout MV-54 in the >63µm grain-size fraction and that probably best
explains the overall higher TOC values compared to MV-51. MV-54 also contains a
higher sand-sized fraction of terrigenous siliciclastic particles (mean 15%) compared to
the sand-poor sediments of MV-51 (mean 2%) (Figures 2.3, 2.4). Even without carbon
isotope data, it is apparent from microscopic observations of sand-sized particles, and the limited biogenic calcium carbonate, that MV-54 is dominated by vascular plant matter.
2.5.3. Benthic Foraminifera and Organic Carbon Fluxes
Accumulations of fossil benthic foraminifers are byproducts of multiple
competing biological and taphonomic factors, but overall their abundance and taxonomic composition can indicate relative abundance of labile organic carbon at the time of
deposition. In general, benthic foraminifera live in the upper few cm of sediment, and
their microhabitat selection and maintenance primarily are controlled by food and oxygen, and secondarily by competition for those resources [Jorissen et al., 1995; Van
der Zwaan et al., 1999]. Shell accumulation results from growth and reproduction within
their microhabitat and increases as nutrient supplies increase [Nees and Struck, 1999; Van
der Zwaan et al., 1999]. Postmortem shell abandonment is followed by mixing via
48 bioturbation, transport, and possible removal by dissolution. Shells that eventually pass
into the historical record are a reflection of all of those processes [Loubere, 1989;
Loubere et al., 1993].
In core MV-51, BFAR and assemblages during >32,000 14C yrs B.P. and 18,400-
20,400 Cal. yrs B.P. indicate substantially different oceanographic conditions compared
to modern. The older interval >32,000 14C yrs B.P. contains the highest BFAR of B.
robusta, U. peregrina/hispida, S. bulloides, Bul. aculeata, C. pachyderma, and O. tener.
Species such as U. peregrina/hispida, C. pachyderma, and S. bulloides are characteristic
of high productivity surface waters and higher flux rates of organic matter under low
seasonality [Altenbach and Sarnthein, 1989; Wahyudi and Minagawa, 1997; Loubere and
Farridudin, 1999]. Cibicidoides pachyderma is an epifaunal taxon (0-1 cm) and adapted
to uptake of labile organic matter [Rathburn et al., 1996; Jorissen et al., 1998].
Collectively, these taxa are epifauna to shallow infauna (upper 2 cm), meaning they are
adapted to abundant food and oxygen present in the near surface [Jorissen et al., 1995;
Jorissen et al., 1998].
It is surprising that the lower part of MV-51, with highest MAR and terrestrial
organic-matter content also contains the highest densities of benthic foraminifera.
Normally, high sediment accumulation rates dilute benthic foraminiferal concentrations, and productivity taxa are not adapted to refractory organic carbon. However, the higher accumulation rates and burial of refractory, terrigenous dominated TOC suggests that
there was a much greater river influence at the time. We account for these discrepancies
by suggesting that enhanced river discharge may have supplied nutrients to increase productivity. This provided labile organic carbon for foraminifera, but the labile carbon
49 was not preserved. Highest BFAR before ~32,000 14C yrs B.P. suggests that this was the
time of highest marine productivity compared to any other time interval in MV-51.
The benthic foraminiferal assemblage during ~18,400-20,400 Cal. yrs B.P. indicates a much higher seasonally variable flux of organic carbon to the seabed compared with assemblages from >32,000 14C yrs B.P. Bolivina robusta, G. translucens,
Gc. subglobosa, O. tener, and M. barleeanus comprise the assemblage during this time
interval. Globocassidulina subglobosa is a shallow infaunal (0-2 cm) and opportunistic
species that is able to respond rapidly to phytodetrital pulses from the photic zone [Linke
and Lutze, 1993; Rathburn et al., 1996; Smart and Gooday, 1997]. Gavelinopsis
translucens is typically a shallow infaunal taxon and is found associated with fresh, labile
organic carbon [Jorissen et al., 1998]. In the eastern South Atlantic, an association
between G. translucens and E. exigua is correlated to a higher seasonality of export
production resulting from seasons of maximal upwelling [Licari and Mackensen, 2005].
Melonis barleeanus is an intermediate infaunal taxon (1-4 cm) adapted to degraded
organic matter in areas of high upwelling seasonality [Corliss, 1991; De Stigter et al.,
1998; Jorissen et al., 1998; Loubere and Fariduddin, 1999].
Oceanographic settings with high seasonality in the flux of organic carbon to the
seafloor can be found in many locations from the high latitudes to the tropics. In the
northeast Atlantic Rockall Trough, primary productivity and the flux of phytodetrital
material to the seafloor is restricted to warmer, ice-free summer months and as a result,
specific benthic foraminifera rapidly increase in numbers with the arrival of labile
material at the seafloor [Gooday and Hughes, 2002]. In the Gulf of Guinea, coastal
upwelling is perennial, but it also has strong seasonal upwelling events during May-
50 September, and December which are in part also related to maximal discharge from the
Congo River [Licari and Mackensen, 2005]. Unlike the Gulf of Guinea, the GoP lacks a
western boundary current to promote upwelling, therefore we cannot invoke an upwelling
seasonality in the past. However, changes in the monsoons are seasonal and have
significant effects on the ocean. The enhanced seasonality of organic carbon flux
inferred from benthic foraminifera may be a reflection of past monsoons.
Enhanced seasonality in the GoP may have occurred in the past by increasing the contrast between the present day monsoonal climate so that the southeast trades were stronger and the northwest trades were weaker. Marine productivity, which is dependent on nutrient availability in the surface waters, would have increased when ocean currents were more invigorated from stronger southeast blowing trades, and then reduced during calmer surface waters from lower energy northwest trades. This type of seasonal contrast has been demonstrated in the present day Sulu Sea where primary productivity is highest when wind speeds and upper ocean nutrient mixing are maximal during the East Asian winter monsoon (January – March) [de Garidel-Thoron et al., 2001].
A third interval at 6.8 mbsf is inhabited by B. robusta, U. peregrina/hispida, and
C. pachyderma (Figure 2.8). This assemblage is indicative of high productivity, but it is
a very short-lived event and interrupts an interval of otherwise low BFAR. We suggest
that this sample may be a recovery fauna after deposition of ash at 6.9 mbsf. The
presence of epifaunal species such as C. pachyderma is consistent with post-ashfall
recovery faunas [Hess et al., 2001].
Intervening lower BFAR intervals experienced low organic carbon flux. Bolivina
robusta is consistently abundant in MV-51 and probably is a generalist species, though
51 little is known about its ecology (Figure 2.8, unshaded areas). Dissolution may be one
cause for the low numbers of foraminifera in MV-51; however, the pattern is similar to
TOC MAR. Additionally, there is no correlation between the percent calcium-carbonate
content and benthic foraminiferal densities (R=-0.14; p=0.59; n=32) as one would expect
if dissolution controlled the carbonate content of sediments.
2.5.4. Mass Fluxes, Sea Level, and Sediment Sources
Two periods of high sediment and TOC accumulation are recorded in MV-54 and
MV-51 sediments. The first period (32,000-33,000 14C yrs B.P., late regression)
corresponds to a time when sea-level fall slowed and an extensive mid-shelf clinoform
was aggrading on the mid-shelf [Slingerland et al., 2007]. High accumulation rates are
also observed at other locations in the GoP during this time and may have contributed to
mass transport deposits in the Pandora and Moresby Troughs [Francis et al., 2007]. For
example, sediment cores from the northern Ashmore Trough and southern Pandora
Trough are characterized by turbidites during late regression [Jorry et al., 2007] and a similar “regressive package,” characterized by high terrigenous sediment flux, but deposited during MIS stage 4, is also observed on the slopes of the Ashmore Trough
[Francis et al., 2006; Dickens et al., 2006]. The source for the Ashmore Trough regressive unit may have been sediment resuspended by waves and tides from the prograding mid-shelf clinoform [Dickens et al., 2006] when sea level was ~60 m lower
[Chappell and Shackleton, 1986; Lambeck and Chappell, 2001]. However, MV-51 rests on a bathymetric high (not strongly influenced by turbidity currents), and so sediment must have been delivered through the water column via nepheloid-layer advection or similar means.
52 The MV-51 MS and thermomagnetic results suggest a source shift from sediments with high concentrations (early) to lower concentrations (later) of iron-rich paramagnetic minerals around 32,000 14C yrs B.P. Additionally, benthic foraminiferal
species indicate oxygenated bottom water and greater organic-carbon flux before that
time than after. Organic geochemical proxies suggest greater terrestrial input before
32,000 14C yrs B.P. than after. These observations implicate greater fluvial discharge for
these sediments before 32,000 14C yrs B.P, possibly originating from the Papuan
Peninsula (Figure 2.12, fluvial source 3).
Sediment delivery was very high during ~15,000 – 18,000 yrs B.P. at MV-54 and
likely corresponds to a stillstand time after the first major transgression (10-15 m) associated with a meltwater pulse at 19,000 yrs B.P. [Clark et al., 2004]. During this
time, sea level was ~110 m lower and likely permitted the Fly, Turama, Kikori, and
Purari rivers to empty much closer to the shelf edge [Figure 2.12; Lambeck and
Chappell, 2001; Clark et al., 2004; Milliman et al., 2004]. For example, the chronology of MV-54 indicates extremely rapid emplacement of sediments, on the order of six meters in <1000 yrs (Figures 2.2, 2.4). Mass failure from upslope may be one explanation, because some coarse, inclined beds were observed in that interval. Fluxes at
MV-51 during 18,400-20,400 yrs B.P. were also high and likely resulted from direct delivery of river sediment when sea level was lower (Figure 2.12, fluvial source 2).
MV-54 and MV-51 show significantly lowered accumulation rates from the late
transgression to early Holocene time (<15,000 yrs B.P.). This partly resulted from inland
sediment storage as sea level flooded the shelf and increased accommodation space
[Harris et al., 1996].
53
Figure 2.12. Map of the GOP and hinterland topography and geology showing major sources of river input. The geological map was constructed based on information from various sources including Davies and Smith [1971], Mackenzie [1976], Smith [1982] and Steinshouer et al., [1999]. Fluvial source 1 is the Fly-Strickland-Bamu river system, 2 is the Turama-Kikori-Purari rivers, and 3 is a proposed input from the Papuan Peninsula to explain data in MV-51. Two bathymetric lines at 60 and 120 m are shown to reconstruct probable position of the stage 3 (>30,000 yrs B.P.) and the Last Glacial Maximum (23- 19,000 yrs B.P.) coastlines, respectively. Bathymetric profiles A and B illustrate decreased shelf space during LGM relative to recent. The coastlines shown were subsampled from data provided by Daniell [2007].
A coralgal ridge developed at the shelf-edge during sea-level rise from 14,500-12,000 yrs
B.P. suggesting less suspended sediment in the water column; this ridge may also have
54 served as a barricade to sediment escape in the central and northeastern Pandora Trough
[Droxler et al., 2006]. These reefs started growing on the shelf-edge up dip of MV-51 at
~19,000 yrs B.P. with additional growth during the meltwater pulse 1a (14,500–12,500
yrs B.P.) [Droxler et al., 2006].
2.6. Conclusions
Sediment cores examined in this study span the time interval from ~15,000 to
33,000 yrs B.P. and provide a unique window into depositional events at sites MV-54 and
MV-51 in the northeastern GoP. Geochemical and paleontological evidence indicate
different oceanographic conditions in the GoP >32,000 14C yrs B.P. compared to any younger time First, higher TOC MAR and total BFAR prior to 32,000 14C yrs B.P suggest that greater amounts of organic carbon were being delivered to the seabed.
Increases in the densities of productivity taxa (U.peregrina/hispida-C. pachyderma-S. bulloides) suggest some fraction of this organic carbon was in labile form. However, low hydrogen index values may suggest that if any labile compounds were originally present, they were mostly oxidized and not buried. During this time, increased percentages of terrestrial organic carbon over marine organic carbon suggest an abundant input of C3 and possibly other vascular plant matter. A major clastic sediment source change at
~32,000 14C yrs B.P. is suggested by MS, but thermomagnetic results indicate similar
magnetic mineralogy throughout the core. Higher MS values during this interval are
probably the result of increased clastic flux from the central and eastern PNG rivers
rather than significant changes in mineralogy.
The time from 15,000-20,400 yrs B.P. was a period of high sediment
accumulation rates. MV-54 experienced the greatest accumulation rates during late
55 transgression (~15,000-18,000 yrs B.P.) when rivers were situated on the exposed shelf
and delivered sediments directly to the slopes, although we do not know what
accumulation rates were prior to 18,000 yrs B.P. at this site. Stable-carbon isotopes from
MV-51 point to enhanced terrestrial organic carbon delivery, while the G. translucens,
Gc. subglobosa, and M. barleeanus benthic foraminiferal assemblage indicates higher
seasonality of organic carbon flux during this time. The significantly lower accumulation
rates at MV-54 and MV-51 during late transgression to Holocene time are evidence of
inland and coastal sediment storage and possible isolation of sediment supply from the
GoP slope by a shelf-edge coralgal complex upslope from MV-51.
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66 CHAPTER 3 SURFACE SEDIMENTS AND DISPERSAL PATHWAYS ON THE OUTER SHELF-EDGE AND BASIN IN THE GULF OF PAPUA: MAGNETIC SUSCEPTIBILITY, CALCIUM CARBONATE, AND ORGANIC GEOCHEMICAL MEASUREMENTS
3.1. Introduction
Deep marine sedimentary basins beyond the shelf-edge are one of the ultimate storage sites for siliciclastic and biogenic particles over long-term geological timescales.
During times of sea-level lowstand, siliciclastic and organic particles are transported directly into deep-water environments and form mass transport deposits usually in the form of submarine fans [Posamentier and Vail, 1988; Flood and Piper; 1997; Maslin and
Mikkelsen, 1997]. However, detrital and organic particles are stored in clinoform shelf deposits during times of higher sea-level (e.g. highstands) and are rarely transported farther offshore. For example, in mixed siliciclastic-carbonate depositional systems such as the Queensland Margin [Francis et al., 2007] and the Gulf of Papua (GoP), modern siliciclastic sediments discharged from rivers are stored in near-shore inner-shelf sediment wedges or clinoforms [Walsh et al., 2004; Palinkas et al., 2006; Slingerland et al., 2007]. Nevertheless, certain amounts of recent sediment are episodically re- deposited onto the slopes and troughs through gravity flows and dense nepheloid layer plumes. Therefore, understanding the composition and distribution of sediments blanketing the seafloor in deep-marine environments can potentially help to discern the overall net transport pathways of sediments [Nittrouer and Wright, 1994; Ellwood et al.,
2006]. It is also important to accurately model sediment dispersal systems from a source- to-sink perspective as well as to improve our understanding of global biogeochemical cycles.
67 In this study, we examine sediments blanketing the upper slope and basin of the
GoP, which is one of the National Science Foundation’s MARGINS Source-to-Sink
(S2S) focus areas. Our main goal is to determine the nature of recent deep-sea sediments
and the origin of their formation. In particular, we examine the detrital and biogenic
composition of outer shelf-edge and basinal sediments using magnetic susceptibility
(MS), calcium carbonate content, organic geochemical measurements, and organic
petrography on surficial sediments, and then map the distribution of sediment types. On
the basis of these maps, the overall transport pathways are then determined.
3.2. Gulf of Papua (GoP)
3.2.1. Recent Marine Sedimentation
The GoP (Figure 3.1) is a tropical mixed siliciclastic-carbonate depositional
setting between the northeastern Great Barrier Reef (GBR) platform and Papua, New
Guinea, that receives a very large supply of terrigenous sediment from numerous river
systems. Between 300 and 400 megatonnes (Mt) yr-1 of terrigenous clastic sediment is
supplied to the GoP by the Fly-Kikori-Purari rivers, a sediment load in size exceeding
that from the Mississippi River [Milliman and Syvitski, 1992; Milliman, 1995; Wolanski
and Alongi, 1995; Milliman et al., 1999]. However, this load is deposited in a basin just
one fourth the size of the Mississippi River basin. Recent studies indicate that a majority
(up to 75%) of the sediment exiting these rivers is deposited within river floodplains and
mangrove intertidal zones [Walsh et al., 2004]. The remainder of this sediment is deposited in water depths shallower than 60 m, resulting in an inner shelf clinoform that is rapidly prograding across older, late Holocene deposits [Harris et al., 1996; Walsh and
Nittrouer, 2003; Walsh et al., 2004].
68 Inner-shelf sediment deposits account for approximately 20% of the modern Fly-
Kikori-Purari river sediment load [Walsh et al., 2004]. The maximum short-term accumulation of sediment (up to 4 cm yr-1) is found on clinoform foresets in water depths between 20 and 45 m, whereas very little sediment accumulates seaward of the 60m isobath [Walsh et al., 2004; Palinkas et al., 2006]. Sediment deposited on the inner shelf clinoform forms a prograding clastic wedge that is slowly moving toward the northeast as a result of seasonal changes in currents driven by spring tides, wind, and thermohaline currents [Wolanski and Alongi, 1995; Walsh and Nittrouer, 2003; Walsh et al., 2004;
Slingerland et al., 2007]. In contrast, siliciclastic sediment accumulation rates are minimal on the middle and outer shelf as evidenced by the presence of relict carbonate sands and Halimeda banks [Alongi and Robertson, 1995; Brunskill et al., 1995; Harris et al., 1996].
Approximately 3.3% of inner-shelf sediment, or ~1x107 tons yr-1, is transported to the outer shelf-edge where it may episodically spill onto the upper slope as the shelf accommodation space narrows significantly in the vicinity of the Kerema Canyon
[Figure 3.1; Brunskill, 2004; Walsh et al., 2004]. A combination of hemipelagic depositional and nepheloid layer transport processes are thought to be the dominant sediment transport mechanisms [Walsh and Nittrouer, 2003; Muhammad et al., 2007].
Siliciclastic sediment that is liberated from the shelf-edge is mixed with calcium carbonate rich sediment that has been exported off carbonate platforms. Added to this is a calcareous neritic marine plankton sedimentary component that has settled into deep- water environments [Francis et al., 2007]. Unlike the relatively well-studied shallow
69 water depositional environments in the GoP, very few studies have examined the
composition of sediment deposited in deep-water or mapped its distribution.
Figure 3.1. General Map of the Gulf of Papua (GoP) with labeled major physiographic features, bathymetry, and sample locations. The green shaded region highlights the area of the inner shelf clinoform (shallower than 60m) where most recent sediment is deposited. The contour interval is 20 m to -100 m, and then 500 deeper than 1000 m. Contour lines between the -100 and -1000 m isobaths are the 200 and 500 m isobath, respectively. Bathymetry from Daniell, [2007].
3.2.2. Seafloor Physiography
The GoP (Figure 3.1) is a complex basin composed of a variable gradient shelf and a series of troughs and carbonate pinnacles [see Francis et al., 2007]. In general, the shelf changes from a broad and low-gradient shelf (1:1000) seaward of the Turama River
70 to a narrower and higher gradient shelf (1:133) up dip of the Kerema Canyon. The
Pandora Trough is a ~2000 m deep basin exhibiting a NE-SW trend beyond the central and northeastern shelf edge of the GoP. Unfilled paleo channels extend from the shelf- edge and onto the slope as evidenced by irregular inlets in the 100 m isobath [Harris et al., 1996; Francis et al., 2007; Slingerland et al., 2007]. To the west and adjacent to the
GBR is the Ashmore Trough, a ~1000 m deep basin separated from the Pandora Trough by the Ashmore Reef. The Portlock Reef extends outward from the modern shelf-edge and is perched between the Ashmore and Pandora Trough. The GoP deepens to >2000 m in the Moresby Trough that ultimately leads into the deeper Coral Sea Basin.
3.2.3 Regional Geology
Sediment lithologies discharged from the Fly-Kikori-Purari rivers generally reflect the hinterland geology of their drainage basins. Mature siliciclastic sediments with a high illite:smectite ratio and detrital carbonate content (~60-70%) characterize the sediment load of the Fly and Kikori rivers [Brunskill, 2004; Slingerland et al., In Press].
However, much less of the Purari River sediment load (~20%) is composed of carbonate, whereas 30-35% of sediments from the Kikori and Purari rivers are volcaniclastic. These differences reflect the carbonate-dominated folded sedimentary rocks in the Fly drainage basin and an increasing amount of extrusive igneous rocks in parts of the Kikori and
Purari river basin [see Chapter 2; Rickwood, 1968; Davies and Smith, 1971].
3.3. Materials and Methods
3.3.1. PANASH and PECTEN Cruises
Data presented in this study are based on 68 measurements of surficial sediments taken from selected cores collected during the PANASH and PECTEN cruises of the R/V
71 Melville and R/V Marion Dufresne (Table 3.1). The PANASH cruise (Pandora-
Ashmore) took place in the GoP from March to April, 2004, during which time three
gravity cores, two box cores, 30 multi-cores, and 33 jumbo piston cores were collected
primarily in outer shelf edge and trough settings. The PECTEN cruise (Past Equatorial
Climate: Tracking El Niño) took place during June to July, 2005 on board the R/V
Marion Dufresne and collected 16 jumbo CALYPSO cores, six CASQ cores, and one
gravity core within the GoP study area.
3.3.2. Surficial Sediments
A set of 68 undisturbed core-tops (0-2cm) were selected for analyses from the
PANASH and PECTEN core collections. Multi-core, gravity core, box core, and CASQ
core-tops contain the least disturbance in the upper two cm, whereas jumbo piston core
tops may potentially have several centimeters (cm) of loss due to over-penetration. We
excluded piston core tops where they overlapped with multi-core locations, and when
possible, trigger core-tops were used. Samples collected from piston core tops were
carefully inspected for sediment loss by ensuring that they contained characteristics of
surface sediments. The selection criteria we used were 1) the sediment should be water
charged, and 2) sediment color should be lighter than downcore segments to indicate
surface oxidation.
3.3.3. Magnetic Susceptibility (MS) Measurements
MS measurements are a simple, quick, and nondestructive way to determine the
relative abundance of iron-bearing detrital minerals within a sediment or rock sample.
MS values record the strength of induced magnetism experienced by mineral grains when placed in an applied magnetic field. We measured the bulk low-field MS on dried
72 sediment samples using a susceptibility bridge in a magnetically shielded room at the
LSU Rock Magnetism Laboratory [Ellwood et al., 2006]. The LSU Bridge uses a balanced coil induction system that is sensitive to <1 (±0.2) x 10-9 m3 kg-1 [Ellwood et al.,
1996] and calibrated using values for standard salts reported by Swartzendruber [1992].
We report the mean of three measurements for each sample (Table 3.1).
3.3.4. Calcium Carbonate
Calcium carbonate content was analyzed in surficial sediment samples using both wet chemical HCl acid digestion and the gasometric technique at the LSU Rock
Magnetism Laboratory. Multi-core samples were sampled for 0.18 g dry weight and then treated with 50 ml of HCl. The weight loss after 1 hr of digestion was used to calculate total percent carbonate.
Other sediment samples were treated with the gasometric technique [Jones and
Kaiteris, 1983]. This method uses 5 ml of phosphoric acid and 0.25 g of dry, powdered sediment. Triplicate carbonate measurements on several samples resulted in a precision of +/- 0.10%. Tests of reagent grade (99%) carbonate and a mixed sample standard (95% feldspar 5% carbonate) gave accurate results to +/- 1%.
3.3.5. Sedimentary Organic Matter
Organic matter in core sediments was measured using two independent analyses.
Total organic carbon (TOC) was measured on sediment samples using a LECO Carbon analyzer, while Rock-Eval Pyrolysis (REP) was measured on samples using the Rock-
Eval II system. Both measurements were made at Baseline Resolution Analytical Labs, in Houston Texas. Analytical errors for TOC are ±0.05 wt. % and for REP measurements they are ± 0.02 mg of HC per gram of TOC.
73 TOC is a basic parameter measured to quantify the amount of organically-bound
carbon in a sample, whether it is of marine or terrestrial plant origin [Tissot and Welte,
1984; Jarzie, 1991]. The origin of sedimentary organic matter is examined through two
REP parameters, the amount of free (S1) and pyrolytic (S2) hydrocarbon (HC) preserved in sediment samples [Meyers, 1997; Rullkötter, 2000; Wagner, 2000]. In particular, the
hydrogen index (HI in Equation 1) is calculated using TOC and S2 numbers and is given as:
HI = 100xS2/TOC (1)
HI is the mass of S2 hydrocarbons per gram of TOC and is determined by heating
kerogen in a sample until the maximum amount of nonvolatile and volatile hydrocarbons
is evolved [Espitalié and Bordenave, 1993].
Hydrogen index values range from near 0 to >400 and may help to distinguish if
sediments are dominated by vascular land plant matter (≤150 mg of HC per gram of
TOC), or dominated by algal-rich marine organic matter (200 to 400 mg of HC per gram
of TOC). Sediment with hydrogen index values ≤150 mg of HC per gram of TOC are
characterized as Types 3 and 4 kerogen [Tissot and Welte, 1984; Meyers, 1997]. Organic
matter rich in marine plankton are Type 2a kerogen and are in the range of 200-400 mg
of HC per gram of TOC [Rullkötter, 2000].
Tmax is the temperature above which a sample must be heated before most of the
hydrocarbons are evolved. Tmax values of recent or “immature” organic matter are
<430°C, whereas organic matter is “overmature” (>430°C) if derived from reworked
fossil organic matter such as coal [Delvaux et al., 1990].
74 3.3.6. Organic Matter Petrography
Acid insoluble organic matter was optically examined to determine the major
types of kerogen particles contained in selected surface sediment samples and to constrain geochemical measurements [Teichmüller, 1986; Tyson, 1995; Batten, 1996;
Meyers, 1997; Omura and Hoyanagi, 2004; Sebag et al., 2006]. Standard total kerogen processing techniques were utilized in sample preparation [e.g. Tyson, 1995; Oboh-
Ikuenobe and Villiers, 2003]. Approximately 5 g of dry sediment from ten samples were acidified in 10% HCl until reaction ceased, and then acidified in concentrated HF overnight. After complete acidification, samples were then neutralized with distilled
water and reacidified with concentrated HCl to remove any calcium fluoride crystals that
formed during reaction with HF. The remaining residue was diluted with distilled water,
centrifuged, and this process repeated until a neutral pH was achieved. Three permanent
kerogen slides were made for each sample using glycerine jelly. Kerogen residues were
not sieved or oxidized prior to slide making because those steps can remove most of the amorphous organic matter and bias the total kerogen assemblage [Oboh-Ikuenobe and de
Villiers, 2003].
Organic particles were counted using a Leitz Dialux-20 micropscope with a 40X
objective. The identification criteria for counting follows Hart [1986] and total kerogen
categories such as % amorphous organic matter, % phytoclasts, and % marine
palynomorphs are reported [e.g. Tyson, 1989; Tyson, 1993; Table 3.2]. A minimum of
300 particles were counted over several traverses and two slides per sample were
examined to minimize within-sample heterogeneity.
75 Kerogen preservation state was examined using fluorescence microscopy [Van
Gizel, 1979; Tyson, 1995]. After point counts were completed, total kerogen slides were
subjected to incident blue-light excitation and examined using a Leitz Ortholux-II
fluorescing microscope and a 40x NPL Fluotar objective.
3.3.6.1. Phytoclast Group
Phytoclasts are organic particles that are of terrestrial plant origin such as leaf cuticles and tree cortex. Seven subcategories are differentiated in this main category and
include fungal hyphae, well and poorly preserved phytoclasts, amorphous structured and
nonstructured phytoclasts, and opaque particles (Figure 3.2, 1a-h). Amorphous
structured and nonstructured phytoclasts correspond to reddish-brown, highly degraded
or “gelified” plant particles [Hardy, 2000; Sebag et al., 2006]. In reality, amorphous
structured and nonstructured are better described as degraded vs. destroyed, respectively
[Hardy, 2000]. Opaque phytoclasts are dark brown to black and represent extensively
oxidized wood or coal particles [Tyson, 1995; Hardy, 2000; Oboh-Ikuenobe and de
Villiers, 2003; Sebag et al., 2006].
3.3.6.2. Palynomorph Group
Palynomorphs are complete organic-walled microfossils derived from pollen and
spores (sporomorphs), phytoplankton such as dinoflagellate cysts, and are the organic
resistant remains of zooplankton, foraminifera, and annelid worms. Zooclasts are the
recognizable remains of planktonic and benthic marine metazoans (Figures 3.2, 2a-f).
Zooclasts identified in this study include foraminiferal organic linings, annelid worm jaw
parts (scolecodonts), and zooplankton egg-cases [see also Van Waveren, 1992]. Because
76 zooclasts are strictly of marine origin and dinoflagellate resting cysts are exceedingly rare
to absent, zooclasts are combined in the marine “paly” category for data analysis.
Figure 3.2. Acid insoluble organic particles identified in kerogen slides. 1, Phytoclast group: (1a) fungal hyphae surrounded by amorphous organic matter (AOM) (1b) well- preserved phytoclast (1c) poorly-preserved phytoclast with partial gelification (1d) amorphous structured phytoclast (1e-g) nonstructured amorphous phytoclast (1h) opaque phytoclast. 2, Palynomorph group: (2a) pollen grain (2b) trilete spore (2c) copepod egg case (2d) foraminiferal organic lining (2e) fecal pellet containing and surrounded by AOM (2f) scolecodont. 3, AOM surrounding a foraminiferal organic lining. All images to same scale, scale bars = 125µm.
77 3.3.6.3. Amorphous Organic Matter Group
Amorphous organic matter (AOM; Figure 3.2, 3) is a structureless organic constituent that is usually yellow-brown to grey in color and has a thin, flaky to microparticulate and globular morphology. AOM, or “sapropel”, is the end product of accumulating sinking particulate organic matter or “marine snow” that has undergone extensive bacterial degradation [Alldredge, 1979; Alldredge and Silver, 1988; Tyson,
1995]. Carbon isotope analysis of AOM indicates that it is derived from autochthonous marine sources [Lewan, 1986].
3.3.6.4. Absolute Densities
Absolute kerogen abundance was determined by relating the percentage of
kerogen types to the TOC wt. % of each sample [e.g. Hart, 1986; Hardy, 2000].
Equation 2 was used to convert the percent of AOM, phytoclasts, and marine palynomorphs (including zooclasts) to grams of TOC for each category, where
×TOCX X oc = (2) 100
and Xoc = the wt. % of organic kerogen Type X and X = relative percentage of organic
kerogen type.
3.3.7. Map Bathymetry
Several maps are presented in this paper to illustrate the concentrations and
distributions of MS, calcium carbonate, and TOC concentrations. All maps were
constructed using ArcGIS 9.1 and then sediment properties were hand-contoured on base
maps using Corel Draw 9.0. The bathymetry contours for the GoP are from a gridded
78 dataset that utilizes data from numerous sources including multibean sonar coverage collected during the PANASH and PECTEN cruises [see Daniell, 2007].
3.4. Results
3.4.1. Magnetic Susceptibility Measurements from Surface Sediments
Surficial sedimentary MS values are shown in Figure 3.3. In general, MS values
exhibit a gradient from highest values (6-10 x 10-7 m3kg-1) on the upper slope of the
central Pandora Trough to lowest values in the Ashmore Trough and south of Eastern
Field’s Reef. A zone of moderately high values rims the upper slope from the central to
the northern Pandora Trough and also one station in the Moresby Trough (29JPC).
Deeper and seaward in the GoP, moderate values (1-4 x 10-7 m3kg-1) extend from and the
northern to the southern Pandora Trough, into the Moresby Trough and around to the
Eastern Plateau.
3.4.2. Calcium Carbonate Concentrations in Surface Sediments
Calcium carbonate concentrations display a seaward increasing gradient
from low to high concentrations (Figure 3.4) and are approximately the inverse
distribution of the MS values. Lowest concentrations (<30%) are found on the upper
slopes of the Pandora Trough, which also wrap around the margin of the Papuan
Peninsula and into the eastern portion of the Moresby Trough. A transition or inter-
fingering zone characterizes the central and southern Pandora Trough from moderately
low (30-49%) to moderately high calcium carbonate concentrations. Highest
concentrations are found in the basin to upper slopes of the Ashmore Trough.
79
Figure 3.3. Distribution of magnetic susceptibility (MS) values in surface sediment from the GoP.
Figure 3.4. Distribution of calcium carbonate concentrations in surface sediment from the GoP.
80 Two exceptions to this general pattern are found on the shelf-edge (<100 mwd) of
the eastern margin of the Papuan Peninsula (Figure 3.4). Higher calcium carbonate concentrations are found at 43MC (54%) and slightly higher at 44MC (33%) compared to other adjacent deeper cores that have <30% calcite.
3.4.3. TOC Concentrations from Surface Sediment
TOC concentrations are low (≤1 wt. %) throughout the GoP, but they still display
a high (landward) to low (seaward) distribution similar to MS and inverse of calcium
carbonate. Highest concentrations (>0.8 wt. %) are observed in core-tops taken from the
upper slope of the northern Pandora Trough and form a bi-lobed distribution extending
south and southwestward to deeper water depths (Figure 3.5).
Figure 3.5. Distribution of total organic carbon (TOC) concentrations in surface sediment from the GoP.
81 Surrounding this and deeper into the Pandora and Moresby Troughs is an area of
low to moderate TOC concentrations (0.5-0.79 wt. %). A region of low organic carbon
content rims the upper slope of the central Pandora Trough and then extends into the
southern Pandora Trough and throughout the Ashmore Trough. Lowest values are also
found in a single core-top from the Moresby Trough (29JPC) and also the outer shelf-
edge of the eastern Papuan margin (43MC).
3.4.4. RE Pyrolysis Measurements from Surface Samples
HI values of sedimentary organic matter in core-tops are generally low (28 to 346
mg of HC per gram of TOC; mean 125, n=68) and contain low S2 pyrolytic yields
(Figure 3.6a). HI and S2 values indicate hydrogen-poor organic carbon that is typical of terrestrial organic carbon such as Type 3 kerogen or a mixture of Type 3 and oxidized
Type 2 marine organic matter (Figure 3.6b). Low S2 yields are typical of all kerogen regardless of the TOC content (Figure 3.6a); however, highest HI values are
characterized by moderate to low TOC concentrations (Figure 3.6b).
The elevated HI values correspond to TOC depleted sediments found in the
carbonate-rich southern Pandora and Ashmore Troughs (Figure 3.5, Table 3.1). This is consistent with a reduced detrital input and greater marine influence, and to higher sedimentary organic matter levels in this region (i.e. low MS, Figure 3.3). In addition, recent marine HI values are generally shifted higher compared to late Pleistocene values observed in MV-51JPC (Figure 3.6b, see also Chapter 2). However, it is difficult to accurately determine the organic matter type when S2 yields are < 1 mg of HC per gram of TOC, because they may be the result of mineral matrix dilution of hydrogen-poor hydrocarbons [Espitalié et al., 1980; Katz, 1983; Langford and Blanc-Valleron, 1990].
82 Tmax measurements for core-top sediments contain immature organic carbon, as
expected for recent marine sediment. All core-tops are <430°C and range from 296 to
428°C (mean = 390, n=68, Table 3.1.). In contrast, late Pleistocene sediment contains
six samples with higher Tmax values, indicative of older, more mature carbon exported to the basin during lower sea level stands (Figure 3.6b, see also Chapter 2).
Figure 3.6. Rock-Eval 2 data from core-tops, plotted in an S2 versus TOC diagram (6a, after Langford and Blanc-Valleron, 1990) and a Van Krevelen diagram (6b, after Wagner, 2000). Boundaries in the Van Krevelen diagram between kerogen Types 2, 3, and 4 are indicated with dashed lines and are derived from reference values reported in Delvaux et al. [1990]. Hydrogen index (HI) values for MV-51 (Chapter 2) are plotted in Figure 6b for comparison of recent marine sediments with late Pleistocene samples.
83
Table 3.1. Data for all stations presented in this paper. Geographic coordinates are given in decimal degrees and water depths are listed in meters below sea level. Magnetic susceptibility 3 (MS) data are reported in m /kg and both CaCO3 and TOC concentrations are given as weight percentages of dry sediment. Rock-Eval II pyrolysis data are presented for the hydrogen index (HI, mgHC per g TOC), thermal maximum (Tmax, °C), and hydrocarbon sources S1 and S2 (mg HC per g sediment) and S3 (mg CO2 per g sediment). BC=box core, MC=multi-core, JPC=jumbo piston core, TC=trigger core, C2=Casq core (R/V Marion Dufresne).
Coretop Long Lat Depth CaCO3 MS TOC S1 S2 S3 Tmax HI 1 BC 144.00 -10.31 342 94.457 5.97E-09 0.18 0.05 0.39 1.09 418 216.67 2 BC 144.16 -10.31 661 88.919 4.09E-08 0.31 0.09 0.54 1.58 408 174.19 10 MC 144.48 -9.87 584 80.554 2.92E-08 0.37 0.15 0.84 1.81 418 227.03 12 MC 144.46 -9.81 372 90.5 4.65E-08 0.44 0.16 1.27 1.60 417 288.64 14 MC 144.66 -9.91 760 80.032 6.72E-08 0.42 0.15 0.83 2.21 405 197.62 16 MC 144.74 -9.74 686 74.657 1.36E-07 0.40 0.09 0.51 2.12 414 127.50 18 MC 145.51 -8.51 267 39.209 5.00E-07 0.42 0.08 0.33 1.72 366 78.57 19 MC 145.51 -8.52 229 24.115 5.42E-07 0.21 0.05 0.10 0.97 314 47.62 20 MC 145.02 -9.29 661 57.667 1.99E-07 0.35 0.08 0.32 1.92 379 91.43 21 MC 145.06 -9.36 1001 65.242 1.47E-07 0.43 0.13 0.52 2.41 379 120.93 24 MC 146.25 -9.79 2102 41.131 1.99E-07 0.51 0.14 0.43 2.75 354 84.31 26 MC 146.38 -10.01 2232 38.748 2.24E-07 0.51 0.15 0.40 3.07 372 78.43 28 MC 147.13 -10.17 2426 34.498 2.47E-07 0.55 0.15 0.46 3.06 349 83.64 30 MC 146.17 -9.67 2023 44.279 2.23E-07 0.55 0.16 0.56 2.83 374 101.82 32 MC 145.35 -9.55 1620 56.431 1.76E-07 0.51 0.16 0.73 2.68 378 143.14 38 MC 145.35 -9.01 961 47.859 2.23E-07 0.56 0.16 0.58 3.01 372 103.57 39 MC 145.61 -8.61 690 42.519 2.06E-07 0.74 0.19 0.79 3.15 384 106.76 42 MC 145.84 -8.37 91 27.75 4.77E-07 1.14 0.20 1.11 2.94 402 97.37 43 MC 146.40 -8.91 62 54.586 2.77E-07 0.46 0.09 0.36 1.82 368 78.26 44 MC 146.19 -8.61 99 33.701 5.35E-07 0.61 0.12 0.50 2.21 382 81.97 47 MC 145.83 -8.43 399 31.408 2.19E-07 1.01 0.18 1.01 2.90 415 100.00 50 MC 146.05 -8.59 795 27.634 2.76E-07 1.00 0.24 1.09 4.18 372 109.00 53 MC 145.94 -8.73 1111 31.504 2.61E-07 0.80 0.21 0.85 3.38 355 106.25 56 MC 145.23 -8.94 450 19.779 1.15E-06 0.25 0.02 0.07 1.13 353 28.00 59 MC 144.92 -9.39 577 40.765 6.82E-07 0.27 0.03 0.15 1.41 383 55.56 60 MC 145.81 -9.24 1671 47.006 2.17E-07 0.52 0.15 0.62 2.65 373 119.23 65 MC 145.35 -10.47 1012 76.033 6.52E-08 0.24 0.07 0.27 1.94 296 112.50 67 MC 144.90 -9.58 1147 71.981 1.55E-07 0.38 0.11 0.51 2.36 384 134.21 69 MC 145.87 -9.98 1638 69.746 1.10E-07 0.50 0.15 0.81 2.51 386 162.00 70 MC 145.15 -9.78 1758 61.711 1.47E-07 0.46 0.13 0.56 2.53 371 121.74 72 MC 144.69 -10.25 1320 80.936 6.91E-08 0.33 0.09 0.46 2.00 406 139.39 76 MC 144.86 -10.44 1551 67.287 1.14E-07 0.34 0.10 0.46 2.30 375 135.29 7 JPC 144.21 -10.44 773 82.31 3.58E-08 0.38 0.29 0.75 2.31 424 197.37 9 JPC 144.48 -9.87 574 82.26 3.29E-08 0.36 0.25 0.66 2.25 424 183.33
84
Table 3.1. Continued Coretop Long Lat Depth CaCO3 MS TOC S1 S2 S3 Tmax HI 13 JPC 144.46 -9.81 376 85.84 4.46E-08 0.35 0.21 0.74 1.86 423 211.43 22 JPC 146.47 -9.77 2063 20.96 2.51E-07 0.65 0.31 0.39 3.78 369 60.00 23 JPC 146.00 -9.88 2073 32.97 2.28E-07 0.51 0.31 0.37 3.70 386 72.55 25 JPC 146.56 -9.93 2199 21.95 1.76E-07 0.57 0.34 0.41 3.63 378 71.93 27 JPC 146.68 -10.17 2077 27.17 1.40E-07 0.51 0.32 0.47 3.63 386 92.16 29 JPC 146.63 -10.03 2295 14.91 4.58E-07 0.38 0.20 0.28 1.82 387 73.68 31 JPC 146.07 -9.56 1964 21.77 2.01E-07 0.54 0.29 0.35 3.16 368 64.81 33 JPC 145.60 -9.87 1795 44.13 1.39E-07 0.49 0.31 0.44 3.19 394 89.80 37 JPC 145.19 -9.24 1011 45.62 2.01E-07 0.51 0.37 0.61 2.96 412 119.61 48 JPC 145.55 -8.56 494 24.61 5.72E-07 0.47 0.23 0.46 2.12 395 97.87 49 TC 145.83 -8.58 859 20.49 1.96E-07 0.70 0.37 0.61 2.85 391 87.14 51 JPC 146.02 -8.77 804 5.1437 7.93E-08 0.73 0.42 0.56 2.51 393 76.71 51TC 146.02 -8.77 804 17.047 2.21E-07 0.85 0.52 0.82 3.41 405 96.02 52 JPC 145.94 -8.73 1110 15.98 1.83E-07 0.88 0.42 0.70 3.47 390 79.55 54 JPC 145.39 -8.94 923 20.94 1.73E-07 0.64 0.33 0.47 3.32 389 73.44 54TC 145.39 -8.94 923 31.149 1.71E-07 0.60 0.34 0.46 3.55 398 76.79 55 JPC 145.24 -8.94 405 8.53 8.04E-07 0.52 0.25 0.35 2.05 395 67.31 57 JPC 145.02 -9.29 654 49.17 2.51E-07 0.31 0.24 0.39 2.19 392 125.81 58 JPC 144.91 -9.41 415 29.73 6.14E-07 0.22 0.12 0.12 1.43 410 54.55 61 JPC 145.97 -9.00 1616 19.76 2.47E-07 0.66 0.38 0.62 3.34 399 93.94 63 TC 146.25 -10.62 1668 1.87E-07 0.57 0.67 1.00 3.13 410 175.44 64 JPC 145.51 -10.37 994 69.61 8.31E-08 0.25 0.24 0.29 2.27 390 116.00 66 JPC 145.15 -10.00 1793 50.74 1.12E-07 0.43 0.31 0.44 2.62 372 102.33 68 JPC 144.91 -9.58 1142 63.27 7.51E-08 0.45 0.22 0.40 2.51 398 88.89 71 JPC 145.02 -9.69 1593 52.66 9.11E-08 0.49 0.30 0.53 3.12 388 108.16 73 JPC 144.42 -9.78 137 87.85 3.62E-08 0.06 0.33 1.04 2.05 418 236.36 74 JPC 144.08 -10.73 684 87.59 1.40E-08 0.44 0.03 0.02 0.88 402 33.33 75 JPC 144.02 -10.71 530 88.59 1.58E-08 0.41 0.50 1.42 2.08 420 346.34 6 GC 144.21 -10.44 775 81.88 2.96E-08 0.33 0.32 0.80 2.07 427 242.42 8 GC 144.48 -9.87 594 83.66 3.07E-08 0.34 0.39 0.89 2.08 420 261.76 11 GC 144.46 -9.81 377 88.09 4.53E-08 0.47 0.46 1.41 2.09 428 300.00 2932 C2 146.01 -8.74 837 18.484 3.04E-07 0.82 0.63 1.05 3.19 400 128.05 2933 C2 145.88 -8.33 71 5.7973 3.81E-07 0.78 0.56 0.92 2.84 412 117.95 2951 C2 144.66 -9.90 765 70.224 9.13E-08 0.53 0.43 1.10 2.84 413 207.55
Statistical tests of TOC, calcium carbonate, and HI values demonstrate strong and significant correlations among variables. TOC and calcium carbonate concentrations are inversely correlated (R=-0.58, n=68, P<0.001) whereas HI values and calcium carbonate are positively correlated (R=0.74, n=68, P<0.001; Figure 3.7).
85
Figure 3.7. Cross-plot of core-top TOC vs. calcium carbonate (R=0.58, n=68, significant at P<0.001) and hydrogen index (HI) vs. calcium carbonate R=0.74, n=68, significant at P<0.001).
3.4.5. Organic Petrography
All sediment samples examined are dominated by AOM kerogen and contain
fewer amounts of phytoclasts and palynomorphs (Tables 3.2, 3.3). AOM ranged from
62-87% (mean 76%) regardless of water depth or location. A strong and significant
correlation exists for AOM abundance and TOC (Figure 3.8a, Table 3.4). Phytoclasts
ranged from 10-34% (mean 21%) in kerogen residues and are also significantly
correlated with TOC content. Phytoplankton were the least abundant kerogen
components ranging from 0.6 to 8%, the majority of which are zooclasts. Dinocysts and
other identifiable algal remains are essentially absent from surface sediments.
Terrestrialy-derived palynomorphs, such as pollen and spores, comprise <3% of organic
components.
Phytoclasts are dominated by degraded amorphous nonstructured constituents and
recalcitrant opaque phytoclasts. Fewer than 3% were well to poorly preserved
86 phytoclasts, which is indicative of limited terrigenous plant input. On a ternary diagram, all kerogen data plot near the % AOM apex (Figure 3.8b). Data plotting at the AOM apex, or in region 4, are indicative of basinal palynofacies dominated by marine organic matter under oxic to dysoxic conditions [Tyson, 1993].
Table 3.2. Raw data for kerogen counts in core-top samples. Kerogen categories are explained in section 3.6 and illustrated in Figure 3.2.
Palynomorph Group Phytoclast Group AOM Core-top Phytoplankton Zoomorphs Spores & pollen fungal hyphae Well- preserved Poorly- preserved Amorphous structured Amorphous Nonstructured Opaque Phytoclasts AOM Total Non AOM Total kerogen
43 MC 13 1 4 1 0 0 5 26 30 226 80 306 33 C2 4 4 10 2 0 1 4 15 38 229 78 307 42 MC 3 2 2 1 0 1 1 27 35 251 72 323 42 MC 2 2 2 0 1 3 4 18 53 218 85 303 44 MC 2 1 2 3 0 0 2 7 22 261 39 300 47 MC 8 7 11 5 0 0 6 47 94 300 178 478 39 MC 2 8 3 0 0 1 3 21 17 253 55 308 51 C2 19 7 2 0 0 2 2 17 26 233 75 308 51 C2 13 6 4 3 1 3 8 50 123 550 211 761 51 JPC 1 2 13 3 0 0 7 60 69 302 155 457 32 C2 6 5 2 1 0 2 0 3 24 257 43 300 32 C2 4 2 5 2 0 0 1 17 32 239 63 302 63 TC 1 1 2 1 0 0 1 17 40 242 63 305
Table 3.3. Percentage data and counting error for kerogen groups. The counting error is 1 standard deviation of duplicate counts on two separate slides.
Core Marine Pollen Structured top Paly error +/- AOM error +/- Terrestrial error +/- &Spores Phytos Opaques 43 MC 4.58 73.86 21.57 1.31 10.13 9.80 33 C2 2.61 74.59 22.80 3.26 6.51 12.38 42 MC 1.55 0.16 77.71 4.07 20.74 4.24 0.62 8.98 10.84 42 MC 1.32 71.95 26.73 0.66 8.58 17.49 44 MC 1.00 87.00 12.00 0.67 3.00 7.33 47 MC 3.14 62.76 34.10 2.30 11.09 19.67 39 MC 3.25 82.14 14.61 0.97 8.12 5.52 51 C2 8.44 75.65 15.91 0.65 6.82 8.44 51 C2 2.50 4.20 72.27 2.39 25.23 6.59 0.53 8.15 16.16 51 JPC 0.66 66.08 33.26 2.84 14.66 15.10 32 C2 3.67 85.67 10.67 0.67 1.67 8.00 32 C2 1.99 1.19 79.14 4.62 18.87 5.80 1.66 5.96 10.60 63 TC 0.66 79.34 20.00 0.66 5.90 13.11
87
Table 3.4. Kerogen absolute abundance data. Percent data from Table 3.3 and TOC from Table 3.1 were used in Equation 2. Errors were calculated using the 1 standard deviation from Table 3.3.
Core-top g Marine Paly error +/- g AOM error +/- g Phytoclasts error +/- 43 MC 0.02 0.34 0.10 33 C2 0.02 0.58 0.18 42 MC 0.02 0.00 0.89 0.03 0.24 0.03 42 MC 0.02 0.82 0.30 44 MC 0.01 0.53 0.07 47 MC 0.03 0.63 0.34 39 MC 0.02 0.61 0.11 51 C2 0.04 0.02 0.40 0.01 0.08 0.02 51 C2 0.01 0.38 0.13 51 JPC 0.00 0.48 0.24 32 C2 0.03 0.01 0.70 0.03 0.09 0.03 32 C2 0.02 0.65 0.15 63 TC 0.00 0.45 0.11
Figure 3.8. Kerogen abundance data. (a) Absolute kerogen data vs. sample TOC. Vertical error bars are based on two sample count replicates and the horizontal error bar is the analytical error for TOC measurements (Table 3.4). The AOM abundance is positively correlated with whole sample TOC wt. % (R=0.94, n=10, P<0.001). The phytoclast kerogen component is also correlated with whole sample TOC wt. % (R=0.77, n=10, P=0.02). (b) Ternary diagram of major kerogen categories [adapted from Tyson, 1993]. Sub-fields on the palynofacies diagram correspond to: (1) source proximal environment such as river delta or inner-shelf; (2) proximal oxic shelf; (3) distal oxic shelf; (4) distal basinal facies.
88 AOM particles are microparticulate and vary from thin, discrete flakes to large
grumose pieces (Figure 3.2). Surprisingly, almost all of the kerogen is nonfluorescent
with the exception of rare autofluorescing particles within AOM, which would not be
expected if marine material dominated the organic matter [Batten, 1996; Omura and
Hoyanagi, 2004]. However, strong oxidation of labile organic compounds in kerogen
can destroy the fluorescent characteristics of marine-derived constituents, such as AOM
[Tyson, 1995]. The fecal pellet (Figure 3.2, 2f) is microparticulate and similar in color to other typical AOM, suggesting that some AOM particles may be components liberated
from fecal excreta. Unfortunately, it is difficult to determine whether the provenance of
the fecal pellets is from zooplankton or from detritus-feeding benthic invertebrates.
3.5. Discussion
Based on the distribution of MS, calcium carbonate, and TOC, there is a clear
particle gradient from the shelf-edge and upper slope near the Kerema Canyon to the
Ashmore Trough. MS values are a function of the mineralogy and/or abundance of
detrital siliciclastic components [deMenocal et al., 1991; Ellwood et al., 2000; Ellwood et
al., 2007; see also Chapter 4]. The MS distribution pattern clearly demonstrates a change
from high to low values on the source proximal shelf-edge to the Ashmore Trough. In
contrast, calcium carbonate concentrations are lowest in regions of high MS and become
much greater in the Ashmore Trough where clastic input is least and pelagic deposition
dominates. The origin of part of the calcium carbonate in core-top sediments is from
neritic settling of coccolithophores, planktonic foraminifera, and pteropods. However,
some fraction may also come from coralgal debris reworked by wave erosion off the
89 Portlock and Ashmore pinnacle reefs as it did during the early Holocene [e.g. Carson et al., 2007; Francis et al., 2007].
TOC concentrations demonstrate the clearest trend from high to low concentrations, beginning on the upper slope of the Kerema Canyon and decreasing to the
Ashmore Trough. The distribution of TOC is consistent with previous studies of geochemical distributions, transport models, and accumulation rates [Bird et al., 1995;
Brunskill et al., 1995; Brunskill et al., 2003; Walsh and Nittrouer, 2003; Keen et al.,
2006; Muhammad et al., 2007]. However, part of the distribution may also be a consequence of enhanced preservation of total organic carbon contents in regions with higher accumulation rates. For example, Muhammad et al., [2007] also observed a similar pattern in which highest accumulation rates were found on the upper slope of the
Kerema Canyon and then decreasing seaward. Burial of organic carbon has been demonstrated to be strongly correlated with the flux of organic carbon to the sediments
[Müller and Seuss, 1979] and sedimentation rates [Heath et al., 1977; Emerson and
Hedges, 1988].
Regardless of total preservation, degradation of labile organic compounds is likely the result of oxidation within the zone of extensive bioturbation [Stein, 1990;
Hedges and Keil, 1995; Hartnett et al., 1998; Hedges et al., 1999]. Figure 3.9 shows
TOC profiles in the upper 12 cm of three representative multi-cores ranging in water depth of 91 m to 2023 m. Except for the core-top of 42MC, data within each profile do vary significantly and are within the range of analytical error, which is indicative of effective oxidation and sediment mixing.
90 Rock Eval (RE) pyrolysis data and kerogen palynofacies data seem at odds with one another. For example, the RE data indicate that organic residues in the sediments are primarily Type 3, or derived from a botanical provenance. However, kerogen residues are dominated by AOM (algal or bacterial in origin) and much fewer amounts of highly degraded phytoclasts. Two things can explain the apparent differences between the geochemical and optical data. First, almost all of the kerogen constituents lack autofluorescence, a sign of severe oxic degradation that can lead to hydrogen index data
<100 mg of HC per gram of TOC [Tuweni and Tyson, 1994]. Second, HI data can be
lowered by oxidation of organic matter, principally because hydrogen atoms are lost from
the hydrocarbons, leaving behind hydrogen-poor residual carbon (see also chapter 2).
TOC, wt. % 0.45 0.55 0.65 0.75 0.85 0.95 1.05 1.15 0 1 2 3 4 5 6 7 Depth, cm 8 9 30 MC 10 42 MC 11 47 MC 12
Figure 3.9. Distribution of TOC in the upper 12 cm of three multi-cores from the GoP (Table 3.1). Selected multi-cores (Figure 2.1) range in water depth from 91 m (42MC, Kerema Canyon), 399 m (47MC, Kerema Canyon), and 2023 m (30MC, Pandora Trough). Only 42MC demonstrates a decrease in TOC from the surface layer and deeper. All mult-core profiles exhibit values within the range error, which is indicative of efficient sediment mixing via bioturbation.
91 A marine dominance in offshore GoP kerogen is supported by numerous other isotope studies. A carbon isotope study from Bird et al. [1995] suggests that TOC in sediments beyond the outer shelf are of >90% marine origin. Similarly, Goni et al.
[2006] found that terrestrial organic matter is greatest in the Fly delta and inner-shelf sediments, but that marine organic matter dominates farther into the GoP. Burns et al.
[2004] found that organic detritus collected in sediment traps from the Pandora Trough contained few terrestrial plant biomarkers from fluvial sources.
It is surprising that more terrestrial phytoclasts are not observed in the study area, given the proximity of large and very active tropical rivers. However, this likely attests to the retention of particles near shore in the mangrove-rimmed coastline and in the inner- shelf clinoform. It also reflects effective recycling of allochthonous phytoclasts in oxygenated water within the outer topset and upper foreset of the inner-shelf clinoform where tidally-driven physical reworking is strongest [Aller, 1998; Aller and Blair, 2004].
3.6. Conclusions
In this manuscript, the detrital and biogenic components of outer shelf-edge and basinal sediment in the Gulf of Papua (GoP) are examined using magnetic susceptibility
(MS), calcium carbonate content, organic geochemical measurements, and organic petrography. A detrital gradient is observed in the MS in which highest values are found on the outer shelf-edge and upper slopes riming the Pandora Trough. Moderate MS values extend into the central and southern Pandora Trough and then decrease to lowest values in the Ashmore Trough. Calcium carbonate displays an inverse trend relative to
MS where highest concentrations are observed in the Ashmore Trough adjacent to the
Great Barrier Reef. Moderate calcium carbonate concentrations interfinger with detrital
92 surface sediments in the southern Pandora Trough, and then reach lowest concentrations
on the upper slopes in the vicinity of the Kerema Canyon. Together, the distribution of
these two surface sediment parameters demonstrates a mixing of carbonate and clastic
depositional systems in the central to southern Pandora Trough.
Total organic carbon concentrations are highest (>0.8 wt. %) on the outer shelf-
edge and upper slopes in the Kerema Canyon and then decrease into the Pandora and
Ashmore Trough. Rock-Eval Pyrolysis data indicate that the sedimentary organic matter is terrestrial Type 3 or 4 but kerogen petrography indicates <~30% to be of botanical provenance. Instead, kerogen is dominated by plankton or bacterially-derived amorphous organic matter, nearly all of which is nonfluorescent. Enhanced preservation of organic particles in the vicinity of the Kerema Canyon is probably favored by higher sediment accumulation rates found in that region [Muhammad et al., 2007] whereas surface oxidation and intense bioturbation of organic matter probably explains the severe degradation or organic particles that are preserved.
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99 CHAPTER 4 DETRITAL CONTROLS ON MAGNETIC SUSCEPTIBILITY AND CYCLOSTRATIGRAPHY RECORDS
4.1. Introduction
All materials are susceptible to becoming magnetized in an inducing magnetic
field, and magnetic susceptibility (MS) is a rapid, simple, and nondestructive indicator of
that magnetization. Because many marine sedimentary records demonstrate Milankovitch
climatic cycles, MS is used for the astronomical tuning, in part to better define geological
time scales [deMenocal et al., 1991; Hellar and Evans, 1995; Shackleton et al., 1999;
Weedon et al., 1999], and for correlation [i.e. Ellwood et al., 2007]. In addition, core and coretop MS datasets from marine settings have been used to assess sediment transport pathways and provenance [Ellwood and Ledbetter, 1977b; Bulfinch et al., 1982; Sachs and Ellwood, 1988; Andrews and Stravers, 1993; Ellwood et al., 2006]. To correctly interpret long-term cyclostratigraphic MS records and geological processes from cored marine sediments, however, it is important to understand the factors influencing MS
values.
Of particular concern here is to understand the proximate cause(s) responsible for
variability within marine MS datasets. The simplest explanation is that MS variations are proportional to changes in terrigenous concentrations and/or mineralogical composition
[deMenocal et al., 1991; Ellwood et al., 1997; Ellwood et al., 2000]. However, because an inverse relationship between calcium carbonate and MS was demonstrated in surficial marine sediment samples [i.e., Ellwood and Ledbetter, 1977b], it is the perception of
many in the geological community that calcium carbonate variability, and therefore
productivity, is an important proximate control on MS variability [Mead and Tauxe,
100 1986; Andrews and Stravers, 1993]. Along these lines, others have inferred that in the absence of magnetite, siliceous biological productivity may overwhelm and then lower weaker MS values [Warner and Domack, 2002].
Therefore, the two models for MS variability in marine sediments are that MS is controlled 1) by changes in detrital fraction flux or 2) by variable marine productivity.
This means that, depending on the model adopted, the ultimate cause (climate) of MS variations could be interpreted very differently. With the first model in mind,
Milankovitch climate cyclicity will influence MS by changing the rates of detrital sediment supply. For example, weathering processes that erode and transport terrigenous particles to marine basins will operate faster and/or more broadly during times of wetter climates. In contrast, if the second model is correct, then Milankovitch climate cyclicity must cause variations in marine biological productivity, perhaps through pulses of nutrient availability, and thus cause changes in the flux and accumulation of organic carbon, coccolithophores, foraminifera, and/or biosiliceous microfossils at the seabed. If the second model is correct, then it should be possible to discern productivity variations in the past from MS cyclostratigraphic records and aid in interpreting the global carbon cycle proposed by Berner [2003].
Here we discuss some theoretical aspects of MS values and then experimentally examine the two models for proximate causes of MS variability. We further test the two models by examining MS from recent marine sediments collected in the Gulf of Papua
(GoP) and two marine sediment outcrop records, one from the Paleocene of Spain, and the other from Upper Cretaceous Niobrara Chalk exposures in Kansas.
101 4.1. Magnetic Susceptibility (MS) of Marine Sediments
Marine sediments are composed of a complex mixture of abiotic and biotic ingredients that influence MS records. Terrigenous components in marine sediments are typically composed of a range of mineralogies that have a range of magnetic susceptibilities. For example, ferrimagnetic iron-bearing minerals such as magnetite, titanomagnetite, or maghemite commonly occur in marine sediments, resulting in high
MS values. When ferrimagnetic mineral content is low or absent, MS values may still be high because of the presence of relatively large amounts of weakly magnetic, paramagnetic minerals, mainly ferromagnesian silicates and clay minerals [Ellwood et al., 2000]. Paramagnetic compounds, such as the clay mineral illite, can dominate the detrital MS signal in marine sedimentary rocks [Ellwood et al., 2007].
The most common minerals in marine sediments are calcite, quartz, feldspar, and chert, and they are diamagnetic, exhibiting a negative MS. The question we are testing here is whether variations in these abundant diamagnetic minerals can significantly lower overall sediment MS in the presence of ferrimagnetic and/or paramagnetic compounds.
4.2. Materials and Methods
4.2.1. MS Instrumentation
Bulk low-field MS was measured on experimental, core, and outcrop samples using a susceptibility bridge in a magnetically shielded room in the LSU Rock
Magnetism Laboratory. The MS Bridge is sensitive to <1 (±0.2) x 10-9 m3 kg-1 [Ellwood et al., 1996], calibrated using values for standard salts reported by Swartzendruber
[1992], and we report the mean of three replicate measurements of each sample. Because many MS data sets in the literature are measured using a MS2 Bartington instrument,
102 commonly the MS2C Loop Sensor configured on a GeoTek Multi Sensor Core Logger
(MSCL), we examine how measurements made from a Bartington MS2C compare with
the LSU MS bridge measurements. Bartington units are reported in the volume-corrected
units of 10-5 SI, while values from the LSU bridge are reported in m3 kg-1 (mass corrected
units).
4.2.2. Experimental Samples
Three mineral samples were prepared to simulate the range of MS values in
marine sediments (Table 4.1). The first sample is composed of ~1-2 g of powdered
Iceland basalt [from samples collected by Ellwood and Ledbetter, 1977b] and represents
a ferrimagnetic end-member sample. The second sample contains ~1 g of an illite
standard [Ellwood et al., 2007] and represents a paramagnetic clay mineral found in
marine sediments. The third sample is composed of a ~3 g 50/50 mixture of the illite
standard and diamagnetic quartz sand. This third sample simulates the value of a marine
sediment sample containing both a paramagnetic clay mineral and diamagnetic
siliciclastic quartz sand.
Prepared samples were placed in separate small sample bags, weighed, and then
measured for initial MS. To simulate the effect of marine productivity on MS, samples
were diluted with powdered reagent-grade calcium carbonate (calcite) in 0.1 g increments
up to 0.5 g and then in 0.25 g increments. After enough calcite was added to the samples
to represent 100% dilution, samples were then diluted with an additional 1 g of calcite in
succeeding increments until at least one order of magnitude shift (e.g. 1x10-7 → 1x10-8) was achieved. MS was measured after each dilution increment.
103 We chose calcite as the diluting mineral because it is common in the marine
environment, is diamagnetic (Table 4.1), and because its concentration is often reported
in the literature in conjunction with MS datasets [Ellwood and Ledbetter, 1977a; Mead
and Tauxe, 1986, Andrews and Stravers, 1993]. We chose not to use pure powdered
silica to simulate biosiliceous productivity because of its unusual electromagnetic
properties in an applied field, and its MS value is very similar to calcite (Table 4.1).
TOC is also diamagnetic but not considered significant here because TOC concentrations in sediments (except for coal) are generally very low, typically <2% [Tissot and Welte,
1984].
Table 4.1. Minerals, initial mass, and initial MS values for samples used in the experiments conducted in this research. Sample # Mineral(s) Initial mass, g Initial MS, m3 kg-1 1 Iceland Basalt 1.8 7.58x10-6 2 Illite Standard 1.0 1.20x10-7 3 Illite/Quartz sand 2.7 4.53x10-8 Diamagnetic calcite -3.00 x10-9 Quartz sand -3.25x10-9
4.2.3. Core-top, Piston-core, and Outcrop Samples
A variety of published and unpublished datasets are used in this manuscript to test
for controls on MS [e.g., Febo et al., 2007; Ellwood et al., In Review]. Piston core samples include two radiocarbon-dated jumbo piston cores collected from the GoP [Febo
et al., 2007, see also Chapter 2] and contain sediment archives deposited in >800 meters
water depth (mwd) for the past 38,000 Cal. yrs. B.P. A core-top dataset from the GoP
includes 42 surface samples (0-2 cm) from multi-cores, box-cores, and trigger-cores
104 [carbonate reported in Febo et al., 2006, see also Chapter 3]. The GoP is a modern
example of a mixed siliciclastic-carbonate depositional setting situated between the Great
Barrier Reef and Papua New Guinea.
Outcrop samples were collected from two localities within the Upper Cretaceous
and the Paleocene. The Upper Cretaceous dataset includes 1,840 samples collected at 10
cm increments from a ~200 m composite sequence of the Niobrara Chalk Formation
(Fort Hays and Smokey Hill Chalk members), Kansas. This section was selected because the entire sequence is composed of chalk and therefore provides an excellent test of carbonate controls on MS variations. The Paleocene section is a 20 sample subset we measured for MS and calcite out of a larger collection (n=175) across the proposed
Danian-Selandian Global Stratotype Section and Point (GSSP) at Zumaia, Spain
[Ellwood et al., In Review a]. The Zumaia sequence represents a change from carbonate- dominated sediments during the Danian to siliciclastic-dominated sediments during the
Selandian and contains a cyclostratigraphic record controlled by Milankovitch climate cycles [Ellwood et al., In Review].
4.2.4. Calcium Carbonate Measurements
Calcium-carbonate concentrations were analyzed using the vacuum-gasometric
technique of Jones and Kaiteris [1983] at the LSU Rock Magnetism Laboratory. A ~2 g specimen from each sample was ground to a fine powder, and then dried for four hours
on a hotplate at 40°C. A 0.25 g sub-sample was then taken and reacted with phosphoric
acid for one hour under a vacuum in a glass reaction vessel. After complete reaction, the
pressure difference was then read from a pressure gauge and used to calculate the sample
carbonate concentration from a calibration equation. Replicate measurements on several
105 samples resulted in a precision of +/- 0.1%. Tests of reagent grade (99.9%) carbonate and a mixed sample standard (95% feldspar 5% carbonate) gave accurate results to +/-
1% and precision to <0.5%.
4.2.5. Numerical Methods: Un-mixing MS Values
Because MS values are a function of many minerals, they can potentially be decomposed into mineral fractions if one or more fractions are known and if sediment samples represent relatively simple mixtures [Potter et al., 2004]. A mass balance equation (equation 1) illustrates that the sample MS is a function of numerous mineral fractions and their mass-dependent MS values.
MS χ t )( = χ + χ 2211 + ...FFF χ nn
where χ t = total sample MS (1)
F1 = component fraction
This equation makes it possible to examine only the terrigenous MS of a sample if the non-terrigenous fractions, such as calcium carbonate, are known (equation 2).
χ − F χ cct χ nc = Fnc
where χ nc = noncarbona MSte (2) Fc = fraction carbonate
If we separate just the calcite fraction, there is still an unresolved fraction of diamagnetic siliceous material in the new MS value, both biological and detrital in origin. Therefore,
106 the new MS is conservatively labeled the non-carbonate MS. Then it is possible to
consider the importance of the ferrimagnetic and paramagnetic components. In most
-9 cases (MS > 1x10 ), χc is so small that it can be ignored and therefore equation 2
becomes:
χ t χ nc = (3) Fnc
4.3. Results
4.3.1. Instrumental Comparisons
In general, there is a very good agreement between MS measured using the LSU
bridge and a Bartington MS2C loop sensor (Figure 4.1). However, the LSU bridge
detected significant variability in MS in core MV-51JPC from 3-9 m that was not
detected using the MS2C sensor (Figure 4.1). This may be attributed to low sensitivity
in the MS2C sensor and therefore limited instrumental detection in the very narrow range
below ~5x10-5 SI.
4.3.2. Experimental Results
All three mixing experiments reveal that very high amounts of calcite must be
added to offset the initial paramagnetic MS of a sample. The basalt sample required
~700% of additional calcite to be added before the value shifted by one order of
magnitude (Figure 4.2). The illite sample required slightly more dilution by calcite
(~860%) to change MS by one order of magnitude. The mixed illite-sand sample
required ~880% dilution by calcite to reduce the MS by one order of magnitude. A
power curve fits all experimental data ( y=axb, regression coefficient R=0.99).
107
Figure 4.1. Comparison between Bartington MS2C loop sensor and the LSU bridge measurements for two piston cores from the Gulf of Papua. The R value is from a logarithm regression of the combined core datasets presented in the central cross-plot. Individual core profiles on the left (MV-51JPC) and right (MV-54JPC) illustrate similar trends in data. Gaps in the Bartington MS are deleted erroneous values measured at core section breaks during logging. Bartington data below ~ 5x10-5 SI do not detect the wide range in MS values observed by the LSU bridge because of reduced instrumental sensitivity in that range.
The results shown in Figure 4.2 consistently demonstrate a non-linear, inverse
correlation between calcite concentrations and MS. In addition, MS is not easily reduced
by diamagnetic calcite. Figure 4.2 illustrates the difficulty in producing diamagnetic values by dilution of illite.
108
Figure 4.2. (A) Results of experimental dilution with calcite of ferrimagnetic (basalt), paramagnetic (illite), and a mixed paramagnetic/diamagnetic sample. The illite curve shows that >850% dilution by calcite is required to lower the MS by one order of magnitude, representative of typical variations observed in MS records. (B) A schematic representation illustrating the extremely large difference between diamagnetic calcite and paramagnetic illite.
4.3.3. Geological Samples
MS and detrital non-carbonate (100-% carbonate) data are shown in Figure 2 from two contrasting locations and geological ages. The Spanish Danian-Selandian samples and the GoP core-tops both show highly significant and positive correlations (Figure
4.3). The Danian-Selandian dataset is linearly correlated whereas the GoP dataset is best
109 fit with a non-linear regression line. However, piston-core data are not correlated
because the range in calcite is very low (<10%) [Febo et al., 2007].
Figure 4.3. MS and non-carbonate percent (100-% carbonate) data for geological samples. The Danian-Selandian data show a positive and highly significant linear regression (R=0.84) whereas core-tops from the Gulf of Papua show a strong nonlinear b correlation, also highly significant (power curve, y=y0+ax , R=0.80). The two piston cores show little variation because they contain extremely low amounts of calcite.
The Niobrara Chalk dataset is complex and is best broken down into subsets for
analysis. Overall, there is a slight positive correlation (Section A, R=0.13) between non-
carbonate and MS (Figure 4.4). However, there are portions of the sequence that, while appearing to be inversely correlated, actually show no sample to sample correlations
(Section B, R=0). In addition, some intervals show a statistically significant sample-to-
sample positive correlation, (e.g. Section C, R=0.50).
110
Figure 4.4. MS and non-carbonate percent data for the Niobrara Chalk Formation. Overall (A), there is a slight positive correlation (R=0.13), however data subsets demonstrate variations in the correlation (see text). Section B shows an apparent inverse correlation while section C shows a positive correlation. In contrast, the section between 40 and 80 m is characterized by an increase in detrital component while the MS is relatively non-varying. These data illustrate the complex relationships between terrigenous fractions and MS.
4.3.4. Numerical Un-mixing
MS values for all geological samples were recalculated using equation 3 to determine the MS for only the non-carbonate fraction. Data in Figure 4.5 demonstrate statistically significant positive correlations for all data sets with both piston-cores demonstrate the strongest linear correlation (R=0.99). The GoP core-tops are best fit
111 with a non-linear regression line and also exhibit a strong correlation (R=0.81). Most
important to note here is the highly significant, strong linear correlation found between
these parameters for the whole Niobrara Chalk dataset (R=0.61) that was not evident in
the parameters plotted in Figure 4 (R=0.13). Correlations would be even stronger if it
were possible to extract and remove the diamagnetic siliceous component in these
samples.
Figure 4.5. Non-carbonate MS versus the whole sample MS for the geological data sets examined. Non-carbonate MS values were recalculated from the initial measured sample MS using equation 3. Piston-core samples form an approximate 1:1 line whereas lower slopes formed by other data sets reveal variable contribution from other noncarbonate diamagnetic fractions.
4.4. Discussion
4.4.1. Non-Carbonate Controls
The experimental data demonstrate that calcite can change MS values, but only at
very high levels of dilution (Figure 4.2). For example, a marine sample with a small
amount of basalt or illite (as in our experimental samples) would require thousands to
112 tens of thousands of percent dilution to cause a significant reduction in MS (Figure 4.2).
It is clear from these data that detrital (non-carbonate) fluxes added to marine sediments
control the observed MS. The main reason for this is that diamagnetic, siliceous, and
carbonate minerals exhibit such a low MS that slight additions of detrital paramagnetic
and ferrimagnetic minerals to the sample overwhelm the diamagnetic components
(Figure 4.2B).
The rock record of productivity indicates that natural variations in calcite, opaline
silica, or organic carbon rarely operate at the scale of thousands of percent dilution,
except in extreme cases. Chalk and diatomite are two rock types that form thick
accumulations (up to a few kilometers) in ancient basins that were both detrital sediment-
starved and contained favorable oceanographic conditions for marine productivity
[Hattin, 1982; White et al., 1992]. The Niobrara Chalk from the Western Interior Basin
is one example of enhanced Upper Cretaceous calcite productivity and reduced clastic
input. Here, even though productivity is high the overall correlation between detrital
component and MS is positive.
In addition to the work reported here for carbonate-dominated samples, several
samples of the well-known Oamaru Diatomite (Bain’s Farm site) from the South Island of New Zealand were measured with the LSU MS bridge. Diatom and radiolarian
(siliceous, diamagnetic) microfossils in these samples record extensive productivity
during the Late Eocene in a sediment-starved warm tropical to subtropical basin in water
depths of ~75-150 m [O’Connor, 1999]. Samples from this diatomite, just like the chalk,
are in middle range of marine sediment MS (4.59x10-8 m3 kg-1) and not diamagnetic as
might be predicted. This is because the diatomite is only composed of ~60% biosiliceous
113 microfossils, leaving the other 40% as detrital components such as clay and fine silt
[O’Connor, 1999].
One of the best examples of detrital controls on MS is from the GoP. A good correlation occurs (Figure 4.3) because there is a range in calcite concentrations in the
core-tops that is the result of cores taken from the “carbonate-dominated” western region of the Ashmore Trough adjacent to the Great Barrier Reef (GBR) platform and the clastic dominated Pandora and Moresby Trough. Today, sediments draining from rivers on the
Queensland margin are deposited on the inner-shelf, allowing the mid to outer-shelf GBR communities to thrive [Scoffin and Tudhope, 1985; Belperio and Searle, 1988; Niel et al.,
2002; Larcombe and Carter, 2004]. However, sediments from the Fly and adjacent rivers inundate the GoP inner-shelf and are then transported into the Pandora Trough via hemipelagic settling and episodic nepheloid layer transport, thus diluting carbonate in the sediments [Walsh et al., 2004; Muhammad et al., 2007]. Sediments deposited during the late Pleistocene, recovered in piston-cores from the central and northern Pandora Trough, show a very limited range in calcite and therefore there is no correlation between MS and non-carbonate fraction.
The Danian-Selandian data set is also an excellent example of non-carbonate controls on MS. The Danian stage sequence contains inter-bedded limestone and marl that becomes dominantly marl in the overlying Selandian stage. Here, a range of calcite concentrations are present in sediments because sediment accumulation rates changed across the stage boundary, lower in the Danian and higher in the Selandian [Ellwood et al., In Review].
114 Perhaps the best evidence of detrital controls on MS is shown in Figure 4.5.
Non-carbonate MS is easy to recalculate once the whole sample MS and the calcite
concentrations are known and this can be used to tease apart relationships not observed
with either the calcite or the non-carbonate fractions (Figure 4.5). While all correlations
are significant, not all are perfectly correlated (e.g. Niobrara Chalk) because of
diamagnetic compounds (i.e. quartz and feldspar) comprising an unknown portion of the
non-carbonate (inferred detrital) sediments.
4.4.2. Interpreting MS and Cyclostratigraphic Records
Comparing MS and calcite can understandably lead to misinterpretations because
sometimes the datasets are correlated and sometimes they are not (Figures 4.3 and 4.4).
As was discussed earlier, it is common to interpret sections with high MS values and
those with low values differently. For example, intervals of high MS values are often
interpreted as being strongly controlled by terrigenous input or provenance, while
intervals of low MS are inferred to be controlled by biological productivity [e.g. Andrews
and Stravers, 1993; Warner and Domack, 2002]. This then leads some to conclude that
MS may be used as an indication of past biogenic productivity, particularly for
cyclostratigraphic MS interpretation.
The proximate causes of MS variations are changes in the detrital concentrations
and mineralogy. The ultimate cause is the process(s) that govern changes to detrital input and compositions, such as climate change, provenance, topography, tectonics, and volcanism. MS variations producing cyclostratigraphic records with orbital frequencies
are best attributed to climatic changes that mitigate the detrital flux to marine basins.
Specifically, variations in wet-dry and warm-cool climates cause changes in weathering
115 rates and river discharge that transport detrital particles to ocean basins [Milliman and
Meade, 1983; Milliman and Syvitski, 1992; Hovius, 1998]. It is much more difficult to argue that long-term climate cyclicity is due to variation in productivity than to argue that this cyclicity is responding to detrital fluxes. This is particularly true when very large productivity rates are required to produce the range in MS values.
4.5. Conclusions
We demonstrate experimentally that dilution of detrital minerals with diamagnetic calcite cannot easily reproduce MS values within the natural range of variation of calcite compositions in marine sediments. Three experimental sediment sample sets were prepared to cover the range of marine sediment MS intensities, including ferrimagnetic
Icelandic basalt, paramagnetic illite standard, and a 50/50 mixed diamagnetic quartz sand and illite. All samples require a calcite dilution of at least 700% to shift MS values by one order of magnitude. In addition, we tested non-carbonate (100%-% calcite) concentrations and MS for several natural marine data sets, covering the range from detrital-dominated (GoP) to calcite-dominated (Niobrara Chalk) samples. Strong correlations were found in data sets where a range of non-carbonate values occurred, but no correlations in data sets where non-carbonate values were >95% (GoP piston cores) were observed. In this context, MS cyclostratigraphic records cannot be interpreted to represent changes in marine productivity but rather changes in climate that act to increase detrital sediment flux.
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120 CHAPTER 5
CONCLUSIONS
Surface sediment and sediment core samples examined in this study span the time
interval from recent to ~15,000 - 33,000 yrs B.P. and offer a unique window into
depositional events in the northeastern GoP. At present, a detrital gradient is observed in
the MS in which highest values are found on the outer shelf-edge and upper slopes riming
the Pandora Trough and lowest values are observed in the Ashmore Trough. Calcium
carbonate displays an inverse trend relative to MS where highest concentrations are
observed in the Ashmore Trough adjacent to the Great Barrier Reef and lowest values are found on the upper slopes of the Pandora Trough in the vicinity of the Kerema Canyon.
The distribution of these two surface sediment parameters demonstrates a classic mixing between carbonate and clastic depositional systems in the central to southern Pandora
Trough.
Total organic carbon concentrations are generally low in surface sediments from
the GoP. A gradient of organic carbon is observed from highest values on the outer
shelf-edge and upper slopes in the Kerema Canyon to lower concentrations in the
Ashmore Trough. Organic geochemical data indicate a terrestrial Type 3 or 4 organic matter, but kerogen petrography indicates <~30% to be of botanical origin. Kerogen is dominated by marine derived amorphous organic matter, nearly all of which is highly degraded. Sea-bed oxidation and intense bioturbation of organic matter probably explains the severe degradation of organic particles that are preserved.
121 In core records, geophysical, geochemical, and micropaleontological evidence indicate different oceanographic conditions in the GoP >32,000 14C yrs B.P. The
following are significant new findings for the late Pleistocene to recent depositional
history in the GoP.
1) Higher total organic carbon mass accumulation rates (TOC MAR) and benthic
foraminiferal accumulation rates (BFAR) prior to 32,000 14C yrs B.P. suggest that
greater amounts of organic carbon were being delivered to the seabed compared
to today. Increases in the densities of productivity taxa such as
U.peregrina/hispida-C. pachyderma-S. bulloides suggest some fraction of this
organic carbon was in a labile form.
2) Low hydrogen index values suggest highly oxidized organic matter of terrestrial
origin. If any labile compounds were originally present in sediments older than
32,000 14C yrs B.P., then they were mostly oxidized and not buried. Carbon
isotope data indicate that during this time, terrestrial organic carbon was more
abundant in the organic carbon pool than marine organic carbon and also suggest
that a significant fraction of the input was from C3 vascular plant matter.
3) A major change in the clastic sediment source is suggested by a shift magnetic
susceptibility values (MS) at ~32,000 14C yrs B.P. In contrast, an increase in
detrital flux rather than a change in mineralogy is implicated by thermomagnetic
results. During this time, other sediment properties (e.g. BFAR and carbon
isotopes) in core MV-51 are very different from modern sediments suggesting
that they were derived from a different source. The change in MS may be a
response to a different sediment provenance, such as a river(s) source draining the
122 Papuan Peninsula, or more direct dispersal pathways of central Papua New
Guinea rivers at lower sea level.
4) The time from ~32,000 14C yrs B.P. to 20,400 yrs B.P. corresponds to the
transition into the Last Glacial Maximum (LGM) and is marked by relatively
lower sediment and TOC accumulation rates but higher percentages of marine
organic carbon. This contrasts with many global records indicating enhanced
organic carbon burial during the LGM.
5) During the time from 15,000 to 20,400 yrs B.P., a period of high sediment
accumulation rates returned. Greatest accumulation rates are experienced in MV-
54 during late transgression (~15,000-18,000 yrs B.P.) when rivers were situated
on the exposed shelf and delivered sediments directly to the slopes. Stable-carbon
isotopes from MV-51 indicate enhanced terrestrial organic carbon delivery, but
the G. translucens, Gc. subglobosa, and M. barleeanus benthic foraminiferal
assemblage indicates higher seasonality of organic carbon flux during this time.
6) The significantly lower accumulation rates at MV-54 and MV-51 during the late
transgression to Holocene time are evidence in support of inland and coastal
sediment storage at higher sea level.
123 APPENDIX A: PUBLISHED AND SUBMITTED CHAPTER ABSTRACTS
CHAPTER 2
Lawrence A. Febo1*, Samuel J. Bentley2, John H. Wrenn**, André W. Droxler3, Gerald R. Dickens3, Larry C. Peterson4, and Bradley N. Opdyke5
1Department of Geology and Geophysics, Louisiana State University, E-235 Howe- Russell Geosciences Complex, Baton Rouge, LA 70803
2Earth Sciences Department, Memorial University of Newfoundland, St. John’s, Newfoundland, Canada A1B 3X5
3Department of Earth Science, Rice University, 6100 Main Street, Houston, TX 77251- 1892
4Rosenstial School of Marine and Atmospheric Science, University of Miami, Miami, FL 33149
5 Department of Earth and Marine Sciences, The Australian National University, Canberra, ACT 020, Australia.
* Corresponding Author, Lawrence A. Febo, [email protected]
**Deceased
Status: In Press to JGR Earth Surface, print scheduled for March 2008
Abstract
We investigated sediment and organic-carbon accumulation rates in two jumbo
piston cores (MV-54, MV-51) retrieved from the mid slope of the northeastern Pandora
Trough in the Gulf of Papua, Papua New Guinea. Our data provide a first assessment of
mass fluxes over the past ~33,000 14C yrs B.P. and variations in organic-carbon sources.
Core sediments were analyzed using a suite of physical properties, organic geochemistry,
and micropaleontological measurements. MV-54 and MV-51 show two periods of rapid
sediment accumulation. The first interval is from ~15,000-20,400 Cal. yrs B.P. (MV-51:
124 ~1.09 m kyr-1 and ~81.2 g cm-2 kyr-1) and the second occurs >32,000 14C yrs B.P. (~2.70
m kyr-1 and ~244 g cm-2 kyr-1). Extremely high accumulation rates (3.96 m kyr-1; ~428 g
cm-2 kyr-1) characterize 15,800-17,700 Cal. yrs B.P. in MV-54 and likely correspond to
early transgression when rivers delivered sediments much closer to the shelf-edge. A
benthic foraminiferal assemblage in MV-51 from ~18,400-20,400 Cal. yrs B.P. indicates
a seasonally variable flux of organic carbon, possibly resulting from enhanced contrast
between monsoon seasons. The oldest sediments, >32,000 14C yrs B.P., contain TOC fluxes >200 g cm2 kyr-1, with >50% of it derived from C3 vascular plant matter.
Magnetic susceptibility values are two to three times higher and benthic foraminiferal
accumulation rates are six times higher during this interval than any younger time
indicating a greater influence of detrital minerals and labile organic carbon. The MS data
suggest more direct dispersal pathways from central and eastern PNG Rivers to the core
site.
125 CHAPTER 4
Lawrence A. Febo1, Brooks B. Ellwood1, and David K. Watkins2
1Department of Geology and Geophysics, Louisiana State University, E-235 Howe- Russell Geosciences Complex, Baton Rouge, LA 70803
2Department of Geosciences, University of Nebraska, 214 Bessey Hall, Lincoln, NE 68588
Status: Will be submitted to Geology
Abstract
When marine sediments are analyzed, it is often argued that magnetic susceptibility (MS) values exhibit an inverse relationship with calcium carbonate concentrations. As a result, MS records are often interpreted as being controlled by calcium carbonate, and therefore, that MS may be a reliable indication of carbonate
productivity. However, because carbonate MS values are extremely small, other studies
have argued (and shown) that varying terrigenous detrital fluxes control MS values. To
test these two models, we present new experimental results and several natural large data
sets in support of the second model, that changes in mineralogy and/or detrital input are the main controls on MS variability, not carbonate productivity.
To model the effect of variations in biological productivity on marine sediments,
we prepared three separate samples of iron-rich minerals found as detrital components in
marine sediments. These were then incrementally diluted with powdered reagent grade
calcium carbonate and the MS measured. Prepared samples were designed to cover the
126 common range of MS observed for lithified marine sediments. In order of decreasing
initial MS, the samples are: 1) ~2 g of powdered Iceland basalt; 2) ~2 g of a paramagnetic
illite standard; and 3) ~3 g 50/50 mixture of illite and quartz sand. The calcium
carbonate dilution required to reduce initial sample susceptibilities by one order of
magnitude is >700% for all samples, far exceeding natural variations in carbonate
productivity. In addition, we observe that MS and calcium carbonate data for a 200 m
section of the Upper Cretaceous Niobrara Chalk, Kansas, show no significant correlation
(R=0.14, n=1840) when the whole data set is considered.
127 APPENDIX B: SUPPLEMENTAL CHAPTER 2 DATA TABLES NOT IN TEXT
Table 2.3. Sediment and organic carbon accumulation rate date for MV-54. TOC mass accumulation rate data were calculated using equation 1 in the text. MAR and TOC MAR are not reported where either dry bulk density or TOC data are absent.
Depth Age DBD LSR MAR, TOC TOC MAR m kyrs g cm-1 cm kyr-1 g cm-2 kyr-1 g cm-2 kyr-1
0 0.694 11.99 8.32 0.64 5.33 0.2 0.698 11.99 8.37 0.65 5.44 0.4 0.753 11.99 9.03 0.71 6.41 0.6 0.680 11.99 8.15 0.68 5.54 0.8 0.761 11.99 9.13 0.71 6.48 1 0.787 11.99 9.43 0.63 5.94 1.2 0.794 11.99 9.52 0.67 6.38 1.4 0.839 11.99 10.06 0.76 7.64 1.5 0.755 11.99 9.05 1.6 0.906 11.99 10.86 0.76 8.25 1.8 1.147 11.99 13.75 2.30 31.56 1.9 15.85 1.239 11.99 14.86 2 0.781 396.83 309.81 0.69 213.77 2.2 0.745 396.83 295.53 0.68 200.96 2.4 0.741 396.83 294.22 0.71 208.89 2.6 0.753 396.83 298.93 0.76 227.19 2.8 0.786 396.83 311.74 0.72 224.46 3 0.809 396.83 320.86 0.73 234.23 3.2 1.103 396.83 437.58 1.59 693.57 3.4 0.927 396.83 368.02 0.87 320.18 3.6 0.862 396.83 341.94 0.75 256.45 3.8 1.088 396.83 431.57 0.58 250.31 4 1.061 396.83 421.17 0.75 315.88 4.2 0.969 396.83 384.52 0.92 353.76 4.4 1.134 396.83 449.95 0.90 404.95 4.6 1.137 396.83 451.17 0.92 415.07 4.75 17 396.83 4.8 1.154 396.83 458.11 0.90 412.30 5 1.091 396.83 432.85 0.95 411.21
128
Table 2.3b. Continued
Depth Age DBD LSR MAR, TOC TOC MAR m kyrs g cm-1 cm kyr-1 g cm-2 kyr-1 g cm-2 kyr-1
5.4 1.152 396.83 457.25 0.81 370.38 5.2 1.105 396.83 438.57 1.00 438.57 5.6 1.112 396.83 441.27 0.98 432.44 5.8 1.098 396.83 435.90 0.94 409.75 6 1.143 396.83 453.43 0.93 421.69 6.2 1.215 396.83 481.98 2.25 1084.46 6.4 1.155 396.83 458.33 0.98 449.16 6.6 1.161 396.83 460.89 1.03 474.71 6.8 1.166 396.83 462.83 0.95 439.69 7 1.095 396.83 434.66 0.79 341.21 7.2 1.195 396.83 474.29 0.95 450.57 7.4 1.220 396.83 484.13 0.91 440.56 7.6 1.188 396.83 471.25 0.91 428.83 7.8 1.203 396.83 477.42 0.94 448.77 8 1.201 396.83 476.53 0.97 462.23 8.2 1.198 396.83 475.55 0.93 442.26 8.3 17 396.83 8.4 1.173 396.83 465.33 1.01 469.98 8.6 1.192 396.83 473.02 0.90 425.72 8.8 1.207 396.83 479.07 0.94 450.33 9 1.221 396.83 484.38 0.92 445.63 9.2 1.176 396.83 466.69 0.99 462.02 9.4 17.74 1.228 396.83 487.29 0.87 423.94
129
Table 2.4. Sediment and organic carbon accumulation rate date for MV-51. TOC mass accumulation rate data were calculated using equation 1 in the text. MAR and TOC MAR are not reported where either dry bulk density or TOC data are absent. Depth Age DBD TOC LSR MAR, TOC MAR m kyrs g cm-1 cm kyr-1 g cm-2 kyr-1 g cm-2 kyr-1 0.01 0.669 0.73 7.28 4.87 3.56 0.21 0.694 0.76 7.28 5.05 3.84 0.41 0.691 0.70 7.28 5.03 3.52 0.61 0.718 0.74 7.28 5.23 3.87 0.81 0.698 0.71 7.28 5.08 3.61 1.01 0.712 0.64 7.28 5.19 3.32 1.21 18.6 0.723 0.79 7.28 5.27 4.16 1.34 18.405 7.28 1.41 0.703 0.90 108.77 76.42 68.78 1.61 0.701 0.84 108.77 76.26 63.68 1.81 0.745 0.91 108.77 81.07 73.77 2.01 0.730 0.79 108.77 79.43 62.75 2.21 0.790 0.69 108.77 85.94 59.30 2.41 0.801 0.41 108.77 87.18 35.74 2.61 0.741 0.63 108.77 80.61 50.78 2.81 0.754 0.70 108.77 82.07 57.45 3.01 0.744 0.68 108.77 80.94 55.04 3.21 0.742 0.67 108.77 80.73 54.09 3.41 0.752 0.61 108.77 81.84 49.92 3.51 20.4 0.764 108.77 83.08 3.61 0.709 0.57 51.20 36.28 20.50 3.81 0.761 0.64 51.20 38.96 24.93 4.01 0.773 0.58 51.20 39.56 22.94 4.21 0.822 0.61 51.20 42.11 25.48 4.41 0.775 0.67 51.20 39.67 26.58 4.61 0.773 0.68 51.20 39.57 26.91 4.81 0.957 0.44 51.20 49.00 21.56 5.01 0.796 0.69 51.20 40.75 28.12 5.21 0.811 0.67 51.20 41.53 27.83
130
Table 2.4b. Continued Depth Age DBD TOC LSR MAR, TOC MAR m kyrs g cm-1 cm kyr-1 g cm-2 kyr-1 g cm-2 kyr-1 5.41 0.840 0.70 51.20 43.03 30.12 5.61 0.849 0.74 51.20 43.48 32.17 5.81 0.888 0.72 51.20 45.46 32.73 6.01 0.852 0.68 51.20 43.60 29.65 6.21 0.874 0.64 51.20 44.75 28.64 6.41 0.871 0.67 51.20 44.57 29.87 6.61 0.872 0.70 51.20 44.64 31.25 6.81 0.874 0.69 51.20 44.74 30.87 6.91 0.877 51.20 44.91 6.935 1.058 0.42 51.20 54.17 22.75 7.01 0.872 0.64 51.20 44.65 28.58 7.135 0.873 0.75 51.20 44.70 33.53 7.21 0.883 0.68 51.20 45.20 30.74 7.41 28.017 0.904 0.82 51.20 46.28 37.95 7.61 0.868 0.76 27.56 23.91 18.17 7.81 0.884 0.76 27.56 24.37 18.52 7.94 29.94 27.56 8.01 0.879 0.75 56.52 49.69 37.26 8.085 0.854 0.96 56.52 48.29 46.36 8.21 0.904 0.80 56.52 51.07 40.86 8.285 0.967 0.87 56.52 54.67 47.56 8.41 0.962 0.76 56.52 54.38 41.33 8.485 0.982 0.84 56.52 55.53 46.65 8.61 0.895 0.73 56.52 50.58 36.92 8.685 0.896 0.71 56.52 50.62 35.94 8.81 0.903 0.85 56.52 51.04 43.38 8.895 0.891 0.78 56.52 50.35 39.27 9.01 0.882 0.82 56.52 49.86 40.88 9.085 0.905 0.84 56.52 51.14 42.96 9.21 0.893 0.86 56.52 50.46 43.39 9.285 0.882 0.78 56.52 49.86 38.89 9.61 0.782 0.81 56.52 44.19 35.79 9.81 0.823 0.83 56.52 46.52 38.38 10.01 32.617 0.880 0.84 56.52 49.72 41.76 10.21 0.904 0.82 269.18 243.25 199.47 10.41 0.902 0.87 269.18 242.86 211.29 10.61 0.917 0.85 269.18 246.86 209.83 10.81 0.897 0.97 269.18 241.39 234.15 11.01 0.918 0.87 269.18 246.98 214.87 11.21 0.911 0.91 269.18 245.10 223.04 11.41 0.914 0.91 269.18 246.13 223.98 11.61 0.926 0.91 269.18 249.24 226.81 11.81 0.918 0.86 269.18 247.19 212.58 12.01 33.36 0.908 0.86 269.18 244.49 210.26 12.21 0.869 0.94 269.18 233.90 219.87
131
Table 2.5. MS bridge data and other physical properties measured from MV-51 sediments. - - - 3 3 3 kg kg kg 3 3 3 1 1 1 depth, m m depth, MS, m %>63µm %CaCO m depth, MS, m %>63µm %CaCO m depth, MS, m %>63µm %CaCO
0.01 7.93E-08 1.60 5.44 5.21 9.85E-08 1.04 3.63 8.81 8.56E-08 1.34 3.69 0.21 1.05E-07 1.04 5.35 5.41 8.09E-08 0.85 2.86 8.895 8.93E-08 0.41 8.35E-08 5.52 3.40 5.61 8.76E-08 0.68 3.79 9.01 9.38E-08 0.79 3.18 0.61 7.12E-08 1.14 5.42 5.81 1.03E-07 1.13 3.07 9.085 6.31E-08 0.81 1.04E-07 2.15 4.36 6.01 9.36E-08 .83 3.39 9.21 8.74E-08 1.52 3.01 1.01 9.97E-08 1.31 4.16 6.21 9.40E-08 0.62 3.73 9.285 9.10E-08 3.528 1.21 1.25E-07 8.26 4.23 6.41 9.55E-08 3.47 9.61 1.19E-07 1.50 4.24 1.41 7.86E-08 0.77 3.30 6.61 9.81E-08 0.82 2.60 9.81 1.52E-07 3.64 1.61 2.19E-07 3.32 4.31 6.81 9.00E-08 0.87 3.33 10.01 2.52E-07 0.51 3.23 1.81 1.03E-07 0.84 3.57 6.86 6.40E-08 0.98 3.04 10.21 2.31E-07 2.95 2.01 9.92E-08 0.93 5.12 6.91 7.26E-08 1.86 2.65 10.41 2.53E-07 1.00 3.20 2.21 8.97E-08 1.87 8.81 6.935 1.46E-07 9.80 1.21 10.61 1.68E-07 3.14 2.41 7.13E-07 6.30 4.29 7.01 8.85E-08 0.69 4.49 10.81 2.17E-07 3.21 2.61 7.31E-08 1.44 7.44 7.135 7.74E-08 3.63 11.01 1.64E-07 0.88 3.57 2.81 6.87E-08 1.00 7.90 7.21 9.04E-08 3.96 11.21 2.37E-07 2.94 3.01 7.04E-08 1.60 6.94 7.41 9.19E-08 0.74 2.75 11.41 2.42E-07 0.70 3.50 3.21 8.09E-08 0.92 5.04 7.61 7.51E-08 3.61 11.61 2.23E-07 3.17 3.41 1.02E-07 2.28 2.96 7.81 9.95E-08 1.29 4.60 11.81 2.61E-07 3.40 3.61 7.09E-08 0.77 4.47 8.01 8.28E-08 3.26 12.01 2.91E-07 1.31 3.44 3.81 8.89E-08 0.60 4.29 8.085 8.30E-08 1.01 2.95 12.21 1.53E-07 3.15 4.01 9.97E-08 0.56 3.58 8.21 8.88E-08 3.27 4.21 8.29E-08 4.10 8.285 9.95E-08 3.3 4.41 9.22E-08 1.19 4.91 8.41 8.59E-08 4.01 4.61 8.82E-08 0.62 3.55 8.485 8.27E-08 1.28 3.70 4.81 2.10E-07 20.44 1.54 8.61 1.10E-07 3.57 5.01 8.54E-08 0.56 4.35 8.685 7.59E-08
132
Table 2.6. MS bridge data for MV-54, other physical properties such as dry bulk density and TOC for this core are reported in Table 3. depth, m MS, m3 kg-1 depth, m MS, m3 kg-1 depth, m MS, m3 kg-1 0.01 1.92E-07 4.61 1.80E-07 10.21 2.91E-07 0.11 1.88E-07 4.81 1.23E-07 10.41 2.84E-07 0.21 2.25E-07 5.01 1.28E-07 10.61 2.49E-07 0.31 1.49E-07 5.21 1.36E-07 10.81 2.18E-07 0.41 2.65E-07 5.41 1.30E-07 11.01 1.76E-07 0.51 1.82E-07 5.61 1.46E-07 11.21 3.43E-07 0.61 2.81E-07 5.81 1.56E-07 11.41 3.88E-07 0.81 3.45E-07 6.01 1.82E-07 11.61 3.17E-07 1.01 2.17E-07 6.21 2.67E-07 11.81 2.01E-07 1.21 1.85E-07 6.41 1.34E-07 12.01 2.23E-07 1.41 2.41E-07 6.61 1.81E-07 12.11 1.94E-07 1.51 3.76E-07 6.81 1.64E-07 1.61 3.78E-07 7.01 1.04E-07 1.81 3.00E-07 7.21 1.32E-07 1.91 7.92E-07 7.41 1.18E-07 2.01 1.22E-07 7.61 1.11E-07 2.21 1.04E-07 7.81 9.77E-08 2.41 1.12E-07 8.01 8.25E-08 2.61 9.07E-08 8.21 9.32E-08 2.81 6.07E-08 8.41 9.56E-08 3.01 1.33E-07 8.61 8.86E-08 3.21 1.13E-07 8.81 1.14E-07 3.41 2.56E-07 9.01 1.33E-07 3.61 1.09E-07 9.21 1.96E-07 3.81 1.41E-07 9.41 2.37E-07 4.01 1.60E-07 9.61 3.11E-07 4.21 1.48E-07 9.81 2.85E-07 4.41 1.54E-07 10.01 2.52E-07
133
Table 2.7. Rock-Eval pyrolysis data measured from MV-51 sediments. S1, S2, and S3 refer to “source” data generated during pyrolysis and are in mg/g units. Tmax is the thermal maximum in degrees Celsius. HI is the hydrogen index calculated using the S2 and TOC values (Table 4) with equation 2, in mgHC gTOC-1 units. depth, m S1 S2 S3 Tmax HI depth, m S1 S2 S3 Tmax HI 0.01 0.42 0.56 2.51 393 76.71 6.41 0.75 0.90 2.00 414 134.33 0.21 0.52 0.63 2.84 383 82.89 6.61 0.44 0.65 2.04 414 92.86 0.41 0.55 0.69 2.50 394 98.57 6.81 0.18 0.30 1.99 388 43.48 0.61 0.55 0.75 2.52 392 101.35 7.01 0.50 0.66 2.34 408 103.13 0.81 0.55 0.74 2.33 398 104.23 7.21 0.38 0.54 2.19 396 79.41 1.01 0.46 0.56 2.16 392 87.50 7.61 1.64 1.16 1.65 399 152.63 1.21 0.55 0.77 2.22 394 97.47 7.81 0.98 0.86 2.15 399 113.16 1.41 0.44 0.60 2.97 399 66.67 8.01 0.68 0.61 2.11 403 81.33 1.61 0.59 0.80 2.55 397 95.81 8.21 0.91 0.86 1.96 403 107.50 1.81 0.56 0.74 2.62 394 81.32 8.41 0.90 0.73 1.68 396 96.05 2.01 0.54 0.59 2.95 392 74.68 8.61 0.80 0.52 2.48 387 71.23 2.21 0.56 0.56 2.84 402 81.16 8.81 0.02 0.06 0.45 492 7.06 2.61 0.26 0.30 2.74 391 47.62 9.01 0.74 0.59 2.01 395 71.95 2.81 0.41 0.48 2.56 398 68.57 9.21 0.19 0.31 1.88 400 36.05 3.21 0.17 0.35 2.32 507 52.24 9.61 0.41 0.46 2.30 400 56.79 3.41 0.14 0.35 1.95 557 57.38 9.81 0.54 0.65 2.01 555 78.79 3.61 0.12 0.17 3.09 384 30.09 10.01 0.62 0.64 2.03 410 76.19 3.81 0.41 0.52 2.41 397 81.25 10.21 0.40 0.52 2.12 398 63.41 4.01 0.13 0.15 2.81 418 25.86 10.41 0.72 0.70 1.99 399 80.46 4.21 0.19 0.33 2.31 406 54.55 10.61 0.73 0.76 2.01 401 89.41 4.41 0.34 0.45 2.21 403 67.16 10.81 0.71 0.58 1.91 384 59.79 4.61 0.21 0.32 2.21 462 47.06 11.01 1.01 0.84 2.21 402 96.55 5.01 0.13 0.12 2.24 393 17.39 11.21 0.18 0.30 2.17 401 32.97 5.21 0.53 0.72 2.38 407 107.46 11.41 1.10 0.73 2.14 416 80.22 5.41 0.29 0.54 1.98 451 77.14 11.61 1.36 0.91 1.75 415 100.00 5.61 0.17 0.27 2.03 392 36.49 11.81 1.43 0.86 1.84 412 100.00 5.81 0.33 0.52 1.93 415 72.22 12.01 0.22 0.35 2.05 380 40.70 6.01 0.26 0.34 2.17 394 50.37 12.21 1.06 0.83 1.91 418 88.30 6.21 0.20 0.29 2.15 390 45.31
134
Table 2.8. Elemental and stable isotopic data measured from MV-51 sediments. Equations 3-5 were used to calculate the fterrigenous and % terrigenous TOC values. % terrigenous 13 depth, m δ C C/N fterrigenous TOC 0.01 -23.91 20.38 0.57 41.50 0.41 -22.90 10.49 0.40 28.04 0.61 -23.74 11.30 0.54 40.01 1.01 -24.24 10.27 0.62 39.89 1.21 -23.85 13.12 0.56 44.12 1.41 -23.43 11.88 0.49 43.93 1.61 -23.60 11.26 0.52 43.19 1.81 -24.59 11.27 0.68 62.10 2.01 -23.35 11.41 0.47 37.48 2.41 -23.63 9.74 0.52 21.36 3.01 -22.64 11.07 0.36 24.26 3.41 -22.84 11.81 0.39 23.78 4.01 -22.01 12.83 0.25 14.62 4.81 -23.35 11.95 0.48 20.93 5.01 -23.19 12.24 0.45 30.91 5.61 -23.57 13.29 0.51 37.89 6.21 -23.30 13.65 0.47 29.90 6.81 -23.20 9.66 0.45 31.00 6.93 -28.93 17.51 7.41 -23.58 18.51 0.51 42.14 8.09 -23.98 19.31 0.58 55.60 8.49 -23.66 18.49 0.53 44.28 9.01 -23.82 16.94 0.55 45.39 10.01 -24.58 19.24 0.68 57.09 11.01 -23.73 19.55 0.54 46.77
135 Table 2.9. Benthic foraminiferal densities (BF#, # g sediment-1) from MV-51 sediments for the 12 most abundant species.
depth, m Total BF# peregrina/ U. hispida Bolivina robusta Bulimina aculeata Cibicidoides pachyderma Sphaeroidina bulloides Gavelinopsis translucens Globocassidulin a subglobosa Oridorsalis. spp. O. tener Melonis barleeanus Bolivinita quadrilatera Pullenia spp.
0.01 142.07 21.38 34.25 20.00 5.06 4.83 13.33 9.89 1.84 1.61 1.15 1.15 3.45 0.21 137.11 16.94 32.15 14.40 10.25 10.14 4.26 14.52 1.96 0.23 0.69 4.15 0.23 0.41 69.61 8.43 17.33 6.44 4.45 8.13 1.84 8.59 0.77 1.38 0.15 1.23 0.31 0.61 56.64 4.30 16.55 7.16 0.64 3.34 3.66 10.02 0.95 0.80 0.32 0.32 0.16 1.01 150.14 10.10 64.90 11.71 0.40 6.60 12.93 13.87 2.29 5.25 1.48 1.75 0.81 1.41 83.24 3.75 44.61 7.40 0.71 2.74 3.14 7.50 1.32 2.03 0.20 0.81 0.41 1.61 172.34 17.48 81.14 15.73 1.31 15.15 5.24 12.82 3.64 4.08 0.87 1.60 0.29 1.81 98.85 7.84 59.64 6.85 0.20 1.39 2.18 8.14 0.89 1.79 0.60 0.00 0.30 2.01 80.64 6.57 36.50 4.73 1.37 1.83 3.97 9.77 1.53 1.53 0.15 0.46 0.46 2.21 130.63 15.75 63.84 12.23 1.69 1.69 6.33 9.84 1.69 1.69 0.70 0.00 0.00 2.41 84.28 2.80 34.73 5.93 0.99 1.48 9.05 8.72 1.48 1.81 0.66 0.99 0.16 3.01 140.46 2.23 26.67 4.89 0.56 3.35 2.93 3.77 3.35 3.07 0.84 0.84 0.70 3.21 81.47 3.12 35.32 9.69 1.15 4.60 2.30 12.16 1.64 0.99 0.66 2.30 0.00 3.41 41.35 0.56 16.96 4.35 0.98 1.12 1.68 4.35 2.38 0.84 1.12 1.12 0.14 4.01 91.41 1.28 38.83 9.75 2.55 1.18 7.84 13.49 2.01 2.46 0.27 2.55 0.73 4.41 121.13 2.48 69.18 12.40 0.00 1.74 6.08 7.44 1.86 3.97 0.50 0.99 0.12 4.61 49.84 1.12 26.90 6.82 0.00 1.24 2.11 1.74 1.49 1.36 0.12 0.12 0.25 4.81 34.50 0.44 23.23 1.32 0.09 0.79 2.38 2.55 0.44 0.53 0.00 0.09 0.26 5.01 67.55 3.08 34.37 6.64 0.12 1.78 6.04 3.91 2.73 1.07 0.00 0.36 0.47 5.61 70.71 0.23 39.37 2.94 0.45 5.32 3.39 3.85 2.15 2.04 0.45 3.96 0.79 6.21 91.35 6.16 53.56 4.65 0.20 1.11 4.75 4.14 0.81 2.43 0.40 0.91 0.30 6.61 151.95 16.11 79.82 10.70 24.55 0.00 2.64 5.54 0.38 1.26 0.13 0.38 0.88 6.81 282.70 61.20 145.62 11.60 39.82 0.11 6.60 6.60 1.93 1.25 0.00 2.16 0.23 6.91 128.19 33.26 53.81 8.93 15.93 0.43 1.61 3.77 0.32 1.83 0.00 0.00 1.18 6.935 67.37 7.30 33.22 2.64 6.83 2.33 3.42 3.57 0.78 0.93 0.00 0.00 0.93 7.01 134.78 26.90 64.27 13.45 5.84 3.40 4.08 4.62 2.31 0.27 0.00 0.54 0.82 7.41 144.70 34.21 62.85 10.68 17.77 11.81 3.12 2.65 3.21 0.47 0.00 1.13 0.57 8.085 188.93 30.62 97.92 9.17 16.45 10.96 5.86 5.77 0.95 0.85 0.00 1.70 2.65 8.485 241.12 19.80 135.18 10.13 18.09 21.16 9.79 10.13 1.25 0.57 0.68 0.23 0.23 9.01 218.79 31.00 101.67 7.83 16.08 26.31 10.75 6.89 2.51 1.36 0.00 1.36 0.42 9.61 160.10 29.44 51.59 9.23 10.97 23.90 10.64 5.32 3.15 1.96 0.00 1.52 0.43 10.01 199.75 36.22 46.34 27.43 10.40 31.31 9.93 10.31 3.50 1.61 0.00 2.27 1.13 10.41 154.32 34.14 27.95 20.96 7.09 27.45 8.38 4.99 1.50 0.60 0.00 4.29 1.50 11.01 167.91 36.34 35.55 20.96 12.20 34.48 4.51 7.69 1.59 2.03 0.00 1.59 0.53 11.41 137.77 23.32 40.07 16.95 8.32 23.53 4.11 4.01 1.75 1.23 0.31 3.39 1.34 12.01 209.15 40.50 67.09 25.83 8.18 24.45 9.25 10.16 1.83 2.06 0.00 3.21 0.92
136 APPENDIX C: CORE IMAGES OF MV-51JPC
137 APPENDIX D
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155 VITA
Lawrence Febo was born in Live Oak, Florida, in 1977, and grew up in
Washington Court House, Ohio. Lawrence graduated from The Ohio State University in
1999 with a Bachelor of Science degree in geological sciences and then with a master’s degree in 2003, also from Ohio State. His master’s thesis was a benthic and planktonic foraminiferal paleoceanographic reconstruction of the Chukchi Borderland, Arctic Ocean, over the past 50,000 years.
Lawrence has an unusual passion for rocks, fossils, and the sea. In 2003, he was awarded a biostratigraphy internship at BP America, which led him to a full-time offer, which he accepted. During Lawrence’s research at Louisiana State University, he traveled to the Gulf of Papua in 2004 and again in 2005 on a second, more extensive cruise beginning in Bitung, northern Sulawesi, and ending in Darwin, Australia.
Lawrence will graduate with a Doctor of Philosophy degree in geology, in
December 2007.
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