Scientific objectives

The interpretation of past climate records to inform our future is one of the major themes of the IODP Science Plan (2013–2023). Expedition 356 will focus on the 5 m.y. marine history of a rifted margin downstream of a major arm of the global thermohaline conveyor (transported via the ITF) to understand the onset and variability of key oceanographic and climate features. We are particularly interested in the relationship between ocean gateways and marine conditions in the transition from the warmer early Pliocene to the cooler Pleistocene and in factors controlling the onset and variability of the Australian monsoon. Results from this expedition will also reveal insights into gateway tectonic history, paleobathymetry, and subsidence.

We will drill, core, and log a latitudinal transect from 18°S to 28°S (Fig. F1) to achieve three key objectives:

  1. Determine the timing and variability of the ITF, Indo-Pacific Warm Pool, and onset of the Leeuwin Current to understand the controls on Quaternary extra-tropical carbonate and reef deposition.
  2. Obtain an ~5 m.y. orbital-scale tropical to subtropical climate and ocean archive, directly comparable to deep-ocean oxygen isotope and ice-core archives, to chart the variability of the Australian monsoon and the onset of aridity in northwestern Australia.
  3. Provide empirical input into the spatio-temporal patterns of subsidence along the NWS that can be used to place fundamental constraints on the interaction between Australian plate motion and mantle convection and to groundtruth geodynamic models.

1. Determine the timing and variability of the ITF, Indo-Pacific Warm Pool, and onset of the Leeuwin Current to understand the controls on extra-tropical carbonate and reef deposition.

The ITF cannot be considered in terms of an “open” or “closed” gateway but instead as the interaction between (1) the source of the water (south versus north Pacific), (2) the location of the main outlet, and (3) variable sill depths and locations through time (Kuhnt et al., 2004). The current eastern (main) outlet through the Timor Sea probably originated around 2–2.5 Ma. Until then, the Bali-Lombok Strait was more important (Kuhnt et al., 2004; Hall, 2009). We intend to document ITF influence on the NWS using fossil biogeography/geochemistry and sedimentary/seismic facies.

Fossil biogeography and geochemistry

The early history of the Indo-Pacific Warm Pool in the Indo-Pacific region has been interpreted from planktonic microfossils. Kennett et al. (1985) suggested that the Indo-Pacific Warm Pool formed as a result of Indonesian seaway closure at ~8 Ma. In contrast, Srinivasan and Sinha (1998) used planktonic foraminiferal biogeography to suggest its formation at ~5.2 Ma. Jian et al. (2006) suggested a late Miocene (~10 Ma) Indo-Pacific Warm Pool with the “modern” warm pool developing at ~4 Ma. Karas et al. (2011) used comparisons between the Mg/Ca and stable isotope data from ODP Holes 709C and 763A on the western and eastern sides of the Indian Ocean, respectively (Fig. F1), to suggest ITF restriction from 3.5-3 Ma that caused 3°C cooling in the region (Fig. F4). Preliminary biogeographic analyses of NWS well cuttings by Gallagher et al. (2009) showed that particular shelfal benthic foraminifers species (Asterorotalia spp., Pseudorotalia spp., and Heterolepa margaritiferus) migrate to the NWS from the southwest Pacific with the waxing and waning of the ITF. Elsewhere, these taxa have proven useful tracers of the warmer Tsushima Current offshoot of the Kuroshio Current in the Japan Sea (Hoiles et al., 2012). Further foraminiferal analyses of 60 sidewall cores from the Fisher-1 well (near primary Site NWS-3A) reveal at least three periods of ITF restriction during the Pliocene (Fig. F4). Gallagher et al. (2009) also used these species to interpret the onset of a relatively more intense Leeuwin Current and ITF ~1 Ma during the Middle Pleistocene transition.

Palynomorphs are also useful tracers of Leeuwin Current flow. For example, pteridophyta spores with Indonesian affinities have been found in surface (modern) and subsurface Late Pleistocene marine sediment on the NWS (van der Kaars and De Deckker, 2002, 2003). Spooner et al. (2011) documented 500 k.y. of Leeuwin Current variability from a deep-sea core (MD002361, 1805 m water depth) (Fig. F1) using stable isotopes and planktonic foraminiferal proxies and suggested a weaker Leeuwin Current during glacial periods, when the West Australian Current was dominant, and a stronger Leeuwin Current during interglacials, especially marine isotope stage (MIS) 11. The southerly migration of Indo-Pacific mollusks and corals to southwest Australia has also been used to document the Late Pleistocene history of the Leeuwin Current (Kendrick et al., 1991; Greenstein and Pandolfi, 2008). Therefore, a combination of benthic foraminiferal, palynological, and macrofossil analyses will be helpful to chart ITF onset and variability and to trace Leeuwin Current activity through time. Further insights into this variability can be achieved using paired C/O isotope and Mg/Ca analyses of foraminifers from NWS glacial–interglacial cycles to interpret middle to outer shelf temperature variations that might reflect the waxing and waning of the ITF. These data will strongly complement nearby Pliocene–Pleistocene deep-sea records (Karas et al. 2009, 2011; Spooner et al. 2011; Stuut et al., 2014).

Transition to warm-water carbonate development

Warm-water carbonates are dominated by photozoan organisms (i.e., organisms relying on photosymbionts, such as corals). Two key features that distinguish tropical from non-tropical carbonates are the presence of coral reefs and ooids (Clarke, 2009; Bosence, 2009). Coral reefs are confined to seawater >18°C (Kleypas et al., 1999). Ooids form at temperatures >20°C with salinities >37‰ (Lees and Buller, 1972). We hypothesize that these two key diagnostic tropical features appear late in the history of the NWS, first developing during the Pleistocene south of 18°S, and that they become progressively younger southward.

Determining the timing and history of late Neogene reef development

The distribution and timing of coral reef development in west Australia is intimately related to the Leeuwin Current (Kendrick et al., 1991). Collins (2002) summarized late Neogene reef distribution in northwest Australia by describing a series of late Tertiary reefs that have developed discontinuously over time. The late Quaternary stratigraphic evolution of Scott (14°S), Rowley Shoals (17.3°S), Ningaloo (22.7°S), and Houtman-Abrolhos (28°S) reefs (Fig. F1) is related to a combination of increased subsidence amplitude toward the north, variability in the Leeuwin Current, and sea level change (Collins and Testa, 2010). Earlier reef development is less well constrained. Ryan et al. (2009) described a series of structurally controlled Miocene reefs from 15°S to 18°S (Fig. F1) using seismic data. These interpreted reefs drowned at the end of the late Miocene (Messinian) when they failed to keep up with sea level change. Further south (22°S), Cathro et al. (2003) interpreted possible Miocene “barrier” reefs or mounds in seismic data. Liu et al. (2011) also interpreted possible Miocene reefs based on seismically imaged mounds. Ryan et al. (2009) acknowledged the dearth of post-Miocene and pre-late Quaternary reefs in the region. However, Jones (1973) and Ryan et al. (2009) described an unnamed post-Miocene drowned “fossil” reef imaged using shallow- and deep-penetration seismic data close to the Rowley Shoals (Fig. F1). Another series of drowned fossil reefs are shown in seismic data from 20°S to 22°S (Figs. F1, F7, AF8, AF9).

We hope to date late Neogene reef re-initiation by sampling the platform/slope facies downdip or along strike from fossil reefs imaged by seismic data. For example, Sites NWS-1A and NWS-7A are updip from the buried fossil reef described by Ryan et al. (2009), and parallel, laterally-persistent reflectors connect these locations (Fig. AF15). Therefore, improved ages for these reflectors should allow us to date this reef. In addition, Site NWS-4A is downdip from a drowned reef (Fig. F7), so new age data should similarly constrain its onset (Gallagher et al., in press). Furthermore, downslope transported reefal detritus will enable analyses of reef development in response to variable sea level. Likewise, Site NWS-3A is along strike from several drowned reefs, and data from this site may constrain the ages of their formation. Farther south, Site NWS-6A is directly seaward of and downdip from the Houtman-Abrolhos main reef complex. Dating, coupled with seismic correlation to this complex, will provide insight into its pre-late Quaternary history.

The significance, timing, and onset of ooid formation

Warm waters supersaturated with respect to carbonate are required for marine ooid formation. Ooids are spherical to oval coated grains that typically form in shallow (<5 m), agitated, tide-dominated tropical environments with elevated evaporation and salinity (Simone, 1981; James et al., 2004). Globally, marine subtropical to tropical ooids have been interpreted to be direct evidence of physiochemical precipitation from seawater during periods of elevated alkalinity and supersaturation (Simone, 1981; Rankey and Reeder, 2009). As such, their occurrence in the NWS subsurface may be used as a sea surface temperature, paleobathymetry, and aridity index. Rankey and Reeder (2009) acknowledge the rarity of ooids in modern and pre-Holocene deposits in the Pacific region and suggest that this is due to the relative lack of regions with sufficiently elevated carbonate supersaturation. There is a similar dearth of ooids in the Indian Ocean (Braithwaite, 1994) for enigmatic reasons. One factor that may account for their rarity in the Indo-Pacific is the absence of particularly favorable conditions required for their formation. For example, ooids accumulate during relatively slow transgressions on flat carbonate platforms (Hearty et al., 2010). If the sea level rise is too fast on a flat platform, they will not form (Hearty et al., 2010). The oldest ooids previously described from the Indian Ocean formed 15.4–12 ka (James et al., 2004) on a low-angle ramp on the NWS (17°S–21°S). James et al. (2004) attributed their formation to increased Leeuwin Current activity (~12 ka) as sea level rose after the Last Glacial Maximum (LGM). Ooids are present in the carbonate supersaturated shallow water of the arid environment in Shark Bay (25.5°S) (Davies, 1970). Other ooids on the NWS (18.5°S) formed 3.3 ka during a period of slow sea level rise (Hearty et al., 2006). In the Maldives, ooids formed during the early cooling and late warming phase of the last glacial cycle (Braithwaite, 1994). Recent work (Gallagher et al., in press) showed that ooids are present in a core dated at 200 ka (Fig. F3) and in cuttings younger than 600 ka from the Maitland North-1 well (Fig. F6) near primary Site NWS-3A. These represent the oldest recovered ooids in the Indian Ocean, but numerous questions remain. For example:

  • When did ooids first occur?
  • Does the advent of ooid formation signify the onset of stronger aridity during the last 600 k.y.?
  • Is there a paleoceanographic trigger (the ITF?) for ooid formation that resulted in increased alkalinity in the Indian Ocean?
  • Is there a relationship between these physiochemical conditions and the onset of reef deposition?
  • What is the relationship between glacio-eustatic cycles, aridity, and ooid formation?

We expect to recover shallow-water facies suitable for ooid preservation in the Quaternary intervals at primary Sites NWS-1A/1B, NWS-2A, and NWS-3A and possibly at Sites NWS-5A and NWS-6A.

2. Obtain an ~5 m.y. orbital-scale tropical to subtropical climate and ocean archive, directly comparable to deep-ocean oxygen isotope and ice-core archives, to chart the variability of the Australian monsoon and the onset of aridity in northwestern Australia.

The aridity of Australia is alleviated by the Australian summer monsoon, which delivers substantial precipitation to the northern part of the continent (north of 25°S) from December to March (Suppiah, 1992; Herold et al., 2011). The winds blow predominantly from the northwest in the rainy season (austral summer), whereas in the dry season (austral winter) the winds blow from the southeast. These changes are associated with the seasonal migration of the subtropical high-pressure belt from 40°S to 30°S. In the austral winter, the Intertropical Convergence Zone (ITCZ) is north of Indonesia and moves south in austral summer to a position immediately north of Australia (Fig. F2). The ITCZ moves even farther south over tropical Australia in February and is associated with the peak of the northern Australian wet season (Williams et al., 2009). The Australian monsoon is thought to be caused by land-ocean temperature contrasts and inter-hemispheric flow from the Asian monsoon. Australian monsoon strength and timing is influenced by changes in insolation resulting from obliquity and precessional forcing (Wyrwoll et al., 2007; 2012). The Australian summer monsoon lacks the topographic influence that controls the Indian–East Asian summer monsoon and is therefore weaker and more sensitive to variations in insolation (Wyrwoll et al., 2007). The region affected by the Asian and Australian monsoon systems (70°E to 150°E) is one of the most significant heat sources driving global climate.

The paleomonsoon

An (2000) speculated that the histories of the East Asian and Australian monsoons are linked and that they originated before 7 Ma. Bowman et al. (2010) suggested (in the absence of any definitive northern Australia pre-Quaternary records) that “the (Australian) monsoon is of great antiquity” because of the pronounced diversity and strong adaptations of biota to the wet-dry tropical climate as well as their strong general adaptability. Herold et al. (2011) noted the incomplete knowledge of the nature and intensity of the pre-Quaternary Australian monsoon and investigated its potential impact on rainfall levels in the Miocene using a general circulation model constrained with a vegetation model. Herold et al. (2011) and Greenwood et al. (2012) compiled available paleontological proxy data for the Miocene and reconstructed a seasonally wet northern and interior Australia, supporting a biome (i.e., seasonally dry deciduous vine forests and sclerophyllous woodlands) consistent with a monsoonal precipitation regime wetter than today. The only pollen record from the semi-arid northwest Australian continent is west of the Cape Range Peninsula (Core GC17) (Fig. F1) and spans the last 100 k.y. (van der Kaars and De Deckker, 2002; van der Kaars et al., 2006). This location is at the southern extremity of the Australian summer monsoon and receives 200–300 mm of rainfall per year, making it ideal to record changes in the latitudinal position of the monsoon (Fig. F2). van der Kaars et al. (2006) used transfer functions to interpret rainfall from the pollen record and hypothesized that a marked reduction in summer rainfall occurred in the absence of monsoonal activity during the LGM. Other deepwater Quaternary records of the Australian monsoon have been obtained from farther north in the Timor Sea (13°S) (Holbourn et al., 2005) and in the Banda Sea (5°S, Beaufort et al., 2010; 8.5°S, Spooner et al., 2005). Holbourn et al. (2005) used foraminiferal and geochemical proxies to chart Timor Sea paleoproductivity over the last 350 k.y. and noted that the Timor Sea productivity record matches the 25°S summer insolation curve, which they interpreted to have strong precessional and eccentricity control. This observation indicates that tropical and/or Southern Hemisphere insolation forcing is an important modulating factor for Australian monsoon intensity. Spooner et al. (2005) combined stable isotope analyses with planktonic foraminiferal assemblages to interpret the 80 k.y. variability of the monsoon and concluded that it was less intense during the first 60 k.y., then intensified at ~15 ka. Beaufort et al. (2010) used calcareous nannofossils to suggest precessional control on primary productivity and Australian monsoon intensity over the last 150 k.y. The Australian monsoon is also interpreted to be broadly controlled by global glacial–interglacial variations (Wyrwoll and Miller, 2001). Strong variations in Australian monsoonal strength between glacial and interglacial periods (paced by orbital eccentricity and precession) have been documented over the last 460 k.y. off northwest Australia (Kawamura et al., 2006) with stronger monsoonal (wet) conditions prevailing during interglacial periods and a weakened monsoon (dry) during glacials. This pattern of glacial–interglacial precipitation variance was further suggested in a 550 k.y. dust record from offshore North West Cape (Stuut et al., 2014). Significant fluvial runoff and megalake expansion across northern and central Australia (Hesse et al., 2004) occurred during interglacials over the last 300 k.y. because of Australian monsoon enhancement. Conversely, reduced precipitation on the NWS (at 23°S) (van der Kaars et al., 2006) and megalake contraction typified glacial conditions (Magee et al., 2004), associated with decreased monsoonal activity.

Charting the 5 m.y. record of the Australian monsoon

Analysis of cuttings from industry well West Tryal Rocks-1 (adjacent to Site NWS-4A) indicates that recovery of a >4 Ma upper slope record of climate variability and monsoon history is achievable (Fig. F7). Cores recovered during Expedition 356 provide the potential to generate a record comparable to other global climate proxy records. High core recovery will be necessary to achieve this and other objectives fully (see “Risks and contingency”); however, even partial core recovery will improve our understanding of the Australian monsoon prior to 1 Ma.

Most global climate proxy records are from the deep ocean basins, but one benefit of coring in a continental-margin setting is that the pollen is likely to be more abundant, delivered by fluvial outflow during the rainy season (van der Kaars and De Deckker, 2003). In comparison, pollen in deep-ocean cores relies on eolian delivery and is consequently less abundant. The Expedition 356 sites are also close to the southern edge of Australian monsoonal influence and can therefore be used to chart its latitudinal variability. Studies on pollen content in modern NWS marine sediments have concluded that the bioclimatic zones of the adjacent Australian continent are extremely well represented (van der Kaars and De Deckker, 2003). We also anticipate good pollen preservation in the Pliocene–Pleistocene interglacials at the shelfal sites and where mudstones/marls are most common in the Pliocene sections.

Little is known about the timing of the development of the characteristic synoptic-scale division of Australia into a winter-wet south and summer-wet north. At a first approximation, the intensity and timing of the monsoonal northern Australia dry seasons and frontal-dominated southern Australia are both controlled by the intensity and seasonal migration of the subtropical anticyclone. Hence, the history of the southern subtropical anticyclone may be critical to our understanding of the evolution of synoptic systems at both ends of the continent. However, it is unknown whether the late Neogene evolution of the Australian monsoon was synchronized with contemporaneous climatic evolution in the southern part of the continent. Nevertheless, the timing of patterns seen in the gamma maxima of the NWS (Fig. F5) shows similarities with data from southeastern Australia. For example, fossil insect and pollen analyses from a small upland paleolake in southeastern Australia indicate that high annual and summer rainfall persisted there until at least 1.5 Ma (Sniderman et al., 2009, 2013), which is inconsistent with the modern climates and vegetation patterns. Drying of a megalake in what is now the semi-arid interior of southeastern Australia did not occur until the early Pleistocene (1.5–1.4 Ma; McLaren and Wallace, 2010; McLaren et al., 2011, 2012, 2014), just when the NWS gamma profile becomes subdued (arrow on Fig. F5). In central Australia, there is evidence that the final phase of aridification, marked by the presence of active dune fields, did not initiate until ~1 Ma (Fujioka and Chappell, 2010). Hence, it is possible that the onset of modern patterns of rainfall seasonality across the continent, as well as the initiation of full aridity in inland Australia, were synchronized. Coring will provide new data to constrain our understanding of Southern Hemisphere Pleistocene climate evolution across a tropical to temperate gradient by including both the monsoonal sites (northern Sites NWS-1A to NWS-4A) and Site NWS-5A. Further insight into the relative input from precipitation can be obtained by analyzing the clay mineralogy of the marly facies (Gingele et al., 2001a, 2001b). The percent kaolinite/illite/chlorite content in each horizon will vary depending on the relative intensity of precipitation and runoff from the source coastal hinterland. For example, reduced chlorite associated with a decrease in kaolinite is interpreted to indicate arid conditions on the NWS during the Holocene (Gingele et al., 2001a). In addition, geochemical analyses of dust in NWS sediments can be used as a relative aridity index (Stuut et al., 2014). Geochemical and foraminiferal proxy analyses at all sites may show paleoproductivity maxima and their possible relationship to the Australian monsoon and the position of the ITCZ. Particularly at upper slope Site NWS-4A, microfossil paleoproductivity analyses might also reveal the dominance of the West Australian Current over the Leeuwin Current during glacial periods (cf. Spooner et al., 2011) when the Australian monsoon is thought to have been weak. Site NWS-5A (~28°C) is located at the northern edge of the modern, winter rainfall–dominated regime of southwestern Australia. It is therefore ideally located to chart the onset of the mid-latitude, westerly wind regime that drives this winter-dominated precipitation regime through analysis of characteristic pollen assemblages.

Turney et al. (2006) suggested that late Quaternary climatic variability across northern Australia probably reflected changes in the latitude of the ITCZ, the westerlies, and ocean masses. However, these authors stated that “few local records are available that enable the frequency, timing, and latitudinal span to be reconstructed with great confidence.” They note that biological or geomorphic proxy evidence might often show a time-transgressive response to climatic variability. They concluded with a plea to perform quantitative reconstructions of past climates in this region with a refined chronology. Recent biodiversity studies emphasize the uniqueness of the Australian monsoon biota (Oliver et al., 2014; Crisp and Cook, 2013) as it evolved in response to the changing monsoon and increased aridity in the Neogene. These authors acknowledge that not only are climate records of the Australian monsoon sparse, but there are also disagreements as to its timing. Thus, coring long well-constrained climate archives in the NWS will greatly enhance our understanding of the evolution and biogeography of the diverse biota of northern Australia.

3. Provide empirical input into the spatiotemporal patterns of subsidence along the NWS that can be used to place fundamental constraints on the interaction between Australian plate motion and mantle convection and to groundtruth geodynamic models.

Lateral displacements of continents relative to mantle convection patterns have significant impact on the flooding history of continental platforms and margins (Sleep, 1976; Gurnis, 1990, 1993; Russell and Gurnis, 1994). Eustatic curves constructed from data from a single margin are known to misrepresent actual global sea levels because of the influence of dynamic topography (Müller et al., 2008b; Spasojevi? et al., 2008; Moucha et al., 2008). Since the breakup and dispersal of eastern Gondwana during the Cretaceous, the Australian plate has moved several thousand kilometers northward (Fig. F8) and recorded anomalous flooding patterns that cannot be reconciled with known eustatic variations (Fig. F9) (Russell and Gurnis 1994; Gurnis et al., 1998; Veevers, 2000; DiCaprio et al., 2009; Heine et al., 2010). Apart from these continental-scale observations, measurements of the amplitude, wavelength, and rate of dynamic topography resulting from circulation within the convecting mantle are rare. Subsidence anomalies along the NWS are ideal targets for investigation of dynamic topography, as the region lies across the gradient of the degree two geoid anomaly and on the fastest moving continent since the Eocene (about >35 Ma). These subsidence anomalies have long been known (Müller et al., 2000; Kennard et al., 2003) and can be ascribed to dynamic topography because both thermal subsidence and flexural effects are minimal (Czarnota et al., 2013).

Cooling of the lithospheric mantle as it returns to its pre-rift thickness drives post-rift thermal subsidence. As the final rifting phase on the NWS occurred more than 130 Ma, the lithospheric mantle should have completed re-thickening by the Neogene, and therefore thermal subsidence should have been insignificant during the Neogene and Quaternary. Flexural effects are likely to be negligible for two reasons. First, our sites are positioned beyond the ~200 km flexural response wavelength of plate boundaries. Secondly, the present elastic thickness of the NWS is ~5 km and therefore loads approach Airy isostasy (Fowler and McKenzie, 1989).

In the last five years, advancements in computer modeling have attributed subsidence anomalies along the NWS to dynamic drawdown of the Earth’s surface driven by Australia’s rapid northward motion over a generally stationary accumulation of subducted slabs within the mantle beneath southeast Asia (Lithgow-Bertelloni and Gurnis, 1997; Heine et al., 2010). These models predict that the NWS should be affected by a southward-propagating wave of subsidence related to Australia’s northward motion over this stationary cold and dense mantle anomaly. Because Australia’s northward motion is ~70 km/m.y. and the proposed drill sites span 10° of latitude, this model predicts a resolvable subsidence diachroneity of >10 m.y. between the northernmost and southernmost proposed drill sites.

In contrast to the diachronous results suggested by geodynamic modeling, recent backstripping of NWS clinoform rollover positions indicates that margin-wide anomalous subsidence was instead broadly synchronous and commenced ~10 Ma, with a down-to-the-north gradient equal in amplitude to adjacent oceanic floor residual depth anomalies (Czarnota et al., 2013). These data suggest that the mantle anomaly responsible for this subsidence may be transient and coupled to the plate motion. However, there is a lack of temporal resolution between 0-5 Ma because there are no well-defined clinoform rollovers of this age. Accurate, high-temporal-resolution subsidence analyses can directly resolve the discrepancy between the two end-member scenarios presented here, thereby providing fundamental insight into the interplay between plate motion and mantle convection.

Tectonic subsidence

Neogene and Quaternary subsidence histories for the NWS can be constructed with unprecedented accuracy using the Expedition 356 primary sites. To compare subsidence histories between drilling locations, 1-D water-loaded basement subsidence histories will be calculated assuming Airy isostasy. Flexural backstripping can also be performed because of the abundance of industry seismic profiles near the sites, but the low flexural rigidity of the region renders this degree of complexity unnecessary. Accurate knowledge of paleobathymetry, sediment compaction parameters, and lithology are essential for 0–5 Ma. In addition, the compaction parameters, lithology, and thickness of underlying sedimentary units are also needed because a large component of the accommodation space on passive margins results from the compaction of underlying sediments. Fortunately, the petroleum industry, government, and academia have extensively studied the underlying sediment pile and therefore the necessary data exist.

Paleobathymetric estimates

Paleobathymetry from foraminiferal and facies analyses is an essential input to backstripping analyses and is necessary to estimate relative sea level variations. Benthic assemblages, planktonic percentage, and sedimentary facies data may all be used to interpret paleodepths. Modern shelfal foraminiferal assemblages on Australia’s continental margin are similar to Pliocene–Pleistocene assemblages and can therefore be used as modern analogs for paleodepth and paleonutrient interpretations (Smith and Gallagher, 2003; Smith et al., 2001). Modern benthic foraminiferal assemblage analog data from across the region may be used for this purpose, including Sunda Shelf (5°N to 10°N) (Biswas, 1976), Banda Sea and Timor Trough (5°S to 10°S) (van Marle, 1988), Sahul Shelf and Timor Sea (8°S to 14°S) (Loeblich and Tappan, 1994), Exmouth Gulf (22°S) (Haig, 1997; Orpin et al., 1999), Ningaloo Reef (24.48°S) (Parker, 2009), and the western Australian continental shelf (20°S to 34°S) (Li et al., 1999; Betjeman, 1969; Quilty, 1977). These benthic foraminiferal assemblage analog data may be enhanced with paleodepth estimates from larger foraminifer distributions (Renema, 2006; James et al., 1999; Hohenegger, 1995; Langer and Hottinger, 2000; Hohenegger, 2005). The percentage of planktonic foraminifers in total assemblage data can also be used to estimate paleodepths (van der Zwaan et al., 1990) and has been used to obtain paleobathymetric estimates prior to backstripping and generation of subsidence curves (van Hinsbergen et al., 2005; Gallagher et al., 2013). Furthermore, paleoenvironmental analyses of ostracod assemblages may complement these analyses (Reeves et al., 2007). This holistic approach can be enhanced by interpretations of facies distributions (e.g., in situ ooid distribution) to generate robust paleobathymetric inputs into subsidence models.