Scientific objectives

1. Provide a testable record of eustatic variations

Backstripping is a proven method for extracting amplitudes of global sea level from passive margin records (e.g., Watts and Steckler, 1979). One-dimensional backstripping is a technique that progressively removes the effects of sediment loading (including the effects of compaction) and paleowater depth from basin subsidence. By modeling thermal subsidence on a passive margin, the tectonic portion of subsidence can be assessed and a eustatic estimate obtained (Kominz et al., 1998, 2008; Van Sickel et al., 2004). Backstripping requires knowing relatively precise ages, paleodepths, and porosities of sediments, and each of these criteria are best obtained from borehole transects; such transects also allow application of two-dimensional backstripping techniques that account for lithospheric flexural effects, increasing the precision of the eustatic estimates (Steckler et al., 1999; Kominz and Pekar, 2001). The eustatic component obtained from backstripping needs to be verified by comparing sea level records with other margins and those derived from δ18O estimates.

Drilling in Holes M0027A–M0029A allows us to make precise Oligocene to early middle Miocene eustatic estimates using backstripping as described above. One- (Kominz et al., 1998; Van Sickel et al., 2004) and two-dimensional (Kominz and Pekar, 2001) backstripping of onshore New Jersey sites has provided preliminary amplitude estimates of 10–60 m for million year–scale variations, but the estimates are incomplete, particularly for the Miocene, because most lowstand deposits are generally not represented (Miller, Sugarman, Browning, et al., 1998; Miller et al., 2005a) (Fig. F4). Amplitude estimates derived from δ18O studies require assumptions about temperature and the sea level/δw calibration; although the uncertainties are large, initial eustatic estimates based on δ18O records are consistent with backstripping results (Fig. F4). Holes M0027A–M0029A are precisely located to recover as nearly a complete set of Oligocene–middle Miocene sequences as possible and, through backstripping, provide a much more direct measure of the full range of amplitudes for this time interval.

When we have obtained precise eustatic estimates from Oligocene to lower middle Miocene records in Holes M0027A–M0029A, we will be able to extend our results to older and younger records. Middle Miocene through Holocene sediments record similar clinoform geometries on the middle to outer New Jersey shelf; by applying calibrations of seismic profiles and facies developed as part of this work, we should be able to derive eustatic estimates for the interval 16–0 Ma. In particular, deriving a firm, independent eustatic estimate from margin sediments will

  • Allow us to test temperature assumptions needed to make glacio-eustatic estimates from δ18O records (Figs. F3, F4);
  • Provide an estimate of the Oligocene–Miocene sea level/δw calibration; and
  • Evaluate the Pekar (1999) and Pekar et al. (2002) calibration of 0.09‰/10 m (versus 0.11‰/10 m for the late Pleistocene) that was based on backstripping an incomplete coastal plain record.

Although both the backstripping and δ18O methods make inherently large assumptions with inherently large uncertainties, the convergence of the two methods (Fig. F4) suggests that we will be able to produce a testable eustatic model for the past 35 m.y.

2. Test models of sedimentation on siliciclastic shelves

Shallow-water records contain unconformities observed in outcrop or in the subsurface at all spatial scales, whether they divide beds or basins. Unconformably bounded sequences are the fundamental building blocks of the shallow-water record (Sloss, 1963; Van Wagoner et al., 1990; Christie-Blick, 1991). Researchers at EPR (Vail et al., 1977; Haq et al., 1987; Posamentier et al., 1988; Van Wagoner et al., 1988) claimed that similarities in the ages of stratal unconformities pointed to global sea level (eustasy) as the overriding control. The resulting "eustatic curve" has remained controversial (e.g., Christie-Blick et al., 1990; Miall, 1991), largely because of basic assumptions about the stratigraphic response to eustatic change and because the work relies in part on unpublished data. In response to this controversy, Christie-Blick and Driscoll (1995), among others, pointed out that the fundamental activity of interpreting the origin of layered rocks does not really require any assumptions about eustasy. They emphasized that sequence boundaries attest to changes in depositional base-level. The timing of many of the EPR sequence boundaries has been validated onshore New Jersey and correlated to the δ18O proxy of eustatic change (Miller et al., 1998, 2005a), although other sequence boundaries on this and other margins may be tectonically derived. Whether or not sequence boundaries are caused by changes in eustasy, local tectonism, or sediment supply (Reynolds et al., 1991), disconformable surfaces irrefutably divide the shallow-water record into sequences. Whatever their cause, these stratal breaks are real and they provide an objective means of analyzing the rock record.

Facies between sequence boundaries vary in a coherent fashion, and various models have been proposed to explain observed spatial and temporal patterns in shelf settings (e.g., Posamentier et al., 1988; Galloway, 1989a, 1989b). Much work has been done by the exploration and academic communities in testing and applying these models, and much has been learned (see Catuneanu et al., 2009). Nonetheless, the complex interaction of processes controlling sequence architecture is not well enough understood for a single model to successfully predict facies successions in all depositional settings.

A major reason that models are poorly constrained and difficult to apply to a variety of settings is that there has been no publicly available study of continuous cores across a prograding siliciclastic clinoform deposit, which constitutes the central element of many facies models. As a result, the water depths in which clinoforms form and the distribution of lithofacies they contain are poorly known. It is widely debated whether clinoform tops ever become subaerially exposed during sea level lowstands and whether the shoreline ever retreats to (or perhaps moves seaward of) the clinoform rollover (Fulthorpe and Austin, 1998; Austin, Christie-Blick, Malone, et al., 1998; Fulthorpe et al., 1999; Steckler, et al., 1999). Settling these controversies will have significant implications on our understanding of how sequence boundaries develop and how much of the facies distribution within clinoforms can be attributed to eustatic variation. Some researchers assume that the shoreline is always located at the clinoform rollover (e.g., Posamentier et al., 1988; Lawrence et al., 1990; Van Wagoner, 1990). Others have presented models that suggest the shoreline and the clinoform rollover move independently of each other (e.g., Steckler et al., 1993, 1999). The sea level estimates of Greenlee and Moore (1988) argue that sea level falls expose an entire continental shelf and that strata onlapping clinoform fronts are coastal plain sediments deposited during the beginning of the subsequent sea level rise. Many researchers (e.g., Steckler et al., 1993) stress that if strata onlapping clinoform fronts were deposited at or near sea level, then the clinoform heights dictate that sea level occasionally fell hundreds of meters in less than a million years; such magnitudes and rates are beyond the reasonable scales of any known mechanism for eustatic change (Pitman and Golovchenko, 1983). Extracting the amplitude of sea level fluctuations from sequence architecture is critically dependent on whether the lowest point of onlap onto sequence boundaries is truly coastal or deeper marine. Determining water depths at the clinoform edge is essential to sequence stratigraphic models and to understanding this basic element of the dynamic land/sea interface. It can only be established by sampling, as done during Expedition 313.